77 10 13233 https://digitalcollections.lakeheadu.ca/files/original/1c7782936aa262c1fe4aef17e30d432a.pdf 5b46d679a8d6d3a763bca217af240a40 PDF Text Text Institute on Lake Superior Geology 59th Annual Meeting Houghton, Michigan May 8 - 11, 2013 Proceedings Volume 59 Part 1 - Program and Abstracts Editors: Allan R. Blaske and Theodore J. Bornhorst www.lakesuperiorgeology.org ��Institute on Lake Superior Geology 59TH ANNUAL MEETING MAY 8-11, 2013 HOUGHTON, MICHIGAN SPONSORED BY: A. E. Seaman Mineral Museum Michigan Technological University THEODORE J. BORNHORST AND ALLAN R. BLASKE Co-Chairs Proceedings Volume 59 Part 1 – Program and Abstracts EDITED BY ALLAN R. BLASKE AND THEODORE J. BORNHORST Cover Photo: Native copper from the Central Mine, Keweenaw Peninsula, Michigan. Collection of the A.E. Seaman Mineral Museum. Photograph by George Robinson. ��59TH INSTITUTE ON LAKE SUPERIOR GEOLOGY PROCEEDINGS VOLUME 59 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD TRIP GUIDEBOOK TRIP 1: GEOLOGIC OVERVIEW OF THE KEWEENAW PENINSULA, MICHIGAN TRIP 2: CALEDONIA MINE, KEWEENAW PENINSULA NATIVE COPPER DISTRICT, ONTONAGON COUNTY, MICHIGAN TRIP 3: GEOLOGY OF SILVER MOUNTAIN, HOUGHTON COUNTY, MICHIGAN TRIP 5: GEOLOGY OF THE KEWEENAWAN SUPERGROUP, PORCUPINE MOUNTAINS, ONTONAGON AND GOGEBIC COUNTIES, MICHIGAN TRIP 6: GEOLOGY AND ENVIRONMENTAL SITE CONDITIONS OF THE COPPERWOOD DEPOSIT, GOGEBIC COUNTY, MICHIGAN Reference to material in Part 1 should follow the example below: Cannon, W. F. Woodruff, L. G., and Schulz, K.. J., 2013, The Hiawatha Graywacke of the Iron River-Crystal Falls district, Michigan: a megaturbidite triggered by seismicity related to the 1850 Ma Sudbury impact [abstract]: Institute on Lake Superior Geology Proceedings, 59th Annual Meeting, Houghton, MI, v. 59, part 1, p. 14-15. Published by the 59th Institute on Lake Superior Geology and distributed by the ILSG Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca ILSG website: http://www.lakesuperiorgeology.org ISSN 1042-9964 i ��Table of Contents Institutes on Lake Superior Geology, 1955-2013 iii Sam Goldich and the Goldich Medal vi Goldich Medal Guidelines viii Goldich Medalists x 2013 Goldich Medal Recipient x Goldich Medal Committee x Citation for Goldich Medal Recipient xi Memorial to Glenn Morey xii Memorial to Paul Sims xiii Eisenbrey Student Travel Awards xiv Joe Mancuso Student Research Awards xv Doug Duskin Student Paper Awards xvii Student Paper Awards Committee xvii Board of Directors xviii Local Committee xviii Session Chairs xviii Banquet Speaker xix Report of the Chair of the 58th Annual Meeting xx Sponsors xxii Program xxiii Poster Presentations xxviii Abstracts 1-83 ii �Institutes on Lake Superior Geology, 1955-2013 95 o o 85 o Wabigoon subprovince90 o 80 48 o Wawa-Abitibi subprovince 48o Wawa-Abitibi subprovince o 45 45o Minnesota River Valley subprovince MEETING LOCATIONS Phanerozoic Mesoproterozoic Map by Mark Jirsa 95o Paleoproterozoic o 90 85o Archean Superior Province # Date Place Chairs 1 1955 Minneapolis, Minnesota C.E. Dutton 2 1956 Houghton, Michigan A.K. Snelgrove 3 1957 East Lansing, Michigan B.T. Sandefur 4 1958 Duluth, Minnesota R.W. Marsden 5 1959 Minneapolis, Minnesota G.M. Schwartz & C. Craddock 6 1960 Madison, Wisconsin E.N. Cameron 7 1961 Port Arthur, Ontario E.G. Pye 8 1962 Houghton, Michigan A.K. Snelgrove 9 1963 Duluth, Minnesota H. Lepp 10 1964 Ishpeming, Michigan A.T. Broderick 11 1965 St. Paul, Minnesota P.K. Sims & R.K. Hogberg 12 1966 Sault Ste. Marie, Michigan R.W. White 13 1967 East Lansing, Michigan W.J. Hinze 14 1968 Superior, Wisconsin A.B. Dickas 15 1969 Oshkosh, Wisconsin G.L. LaBerge 16 1970 Thunder Bay, Ontario M.W. Bartley & E. Mercy iii �# Date Place Chairs 17 1971 Duluth, Minnesota D.M. Davidson 18 1972 Houghton, Michigan J. Kalliokoski 19 1973 Madison, Wisconsin M.E. Ostrom 20 1974 Sault Ste. Marie, Ontario P.E. Giblin 21 1975 Marquette, Michigan J.D. Hughes 22 1976 St. Paul, Minnesota M. Walton 23 1977 Thunder Bay, Ontario M.M. Kehlenbeck 24 1978 Milwaukee, Wisconsin G. Mursky 25 1979 Duluth, Minnesota D.M. Davidson 26 1980 Eau Claire, Wisconsin P.E. Myers 27 1981 East Lansing, Michigan W.C. Cambray 28 1982 International Falls, Minnesota D.L. Southwick 29 1983 Houghton, Michigan T.J. Bornhorst 30 1984 Wausau, Wisconsin G.L. LaBerge 31 1985 Kenora, Ontario C.E. Blackburn 32 1986 Wisconsin Rapids, Wisconsin J.K. Greenberg 33 1987 Wawa, Ontario E.D. Frey & R.P. Sage 34 1988 Marquette, Michigan J. S. Klasner 35 1989 Duluth, Minnesota J.C. Green 36 1990 Thunder Bay, Ontario M.M. Kehlenbeck 37 1991 Eau Claire, Wisconsin P.E. Myers 38 1992 Hurley, Wisconsin A.B. Dickas 39 1993 Eveleth, Minnesota D.L. Southwick 40 1994 Houghton, Michigan T.J. Bornhorst 41 1995 Marathon, Ontario M.C. Smyk 42 1996 Cable, Wisconsin L.G. Woodruff 43 1997 Sudbury, Ontario R.P. Sage & W. Meyer 44 1998 Minneapolis, Minnesota J.D. Miller & M.A. Jirsa 45 1999 Marquette, Michigan T.J. Bornhorst & R.S. Regis 46 2000 Thunder Bay, Ontario S.A. Kissin & P. Fralick 47 2001 Madison, Wisconsin M.G. Mudrey & Jr., B.A. Brown 48 2002 Kenora, Ontario P. Hinz & R.C. Beard 49 2003 Iron Mountain, Michigan L. Woodruff & W.F. Cannon iv �# Date Place Chairs 50 2004 Duluth, Minnesota S. Hauck & M. Severson 51 2005 Nipigon, Ontario M. Smyk & P. Hollings 52 2006 Sault Ste. Marie, Ontario A. Wilson & R.Sage 53 2007 Lutsen, Minnesota L. Woodruff & J. Miller 54 2008 Marquette, Michigan T. Bornhorst & J. Klasner 55 2009 Ely, Minnesota J. Miller, G. Hudak & D. Peterson 56 2010 International Falls, Minnesota M. Jirsa, P. Hollings, T. Boerboom, P. Hinz & M.Smyk 57 2011 Ashland, Wisconsin T. Fitz 58 2012 Thunder Bay, Ontario P. Hollings 59 2013 Houghton, Michigan T. J. Bornhorst and A. R. Blaske v �Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970’s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 vi �INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL vii �Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. viii �Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. ix �Goldich Medalists 1979 Samuel S. Goldich 1996 David L. Southwick 1980 not awarded 1997 Ronald P. Sage 1981 Carl E. Dutton, Jr. 1998 Zell Peterman 1982 Ralph W. Marsden 1999 Tsu-Ming Han 1983 Burton Boyum 2000 John C. Green 1984 Richard W. Ojakangas 2001 John S. Klasner 1985 Paul K. Sims 2002 Ernest K. Lehmann 1986 G.B. Morey 2003 Klaus J. Schulz 1987 Henry H. Halls 2004 Paul Weiblen 1988 Walter S. White 2005 Mark Smyk 1989 Jorma Kalliokoski 2006 Michael G. Mudrey 1990 Kenneth C. Card 2007 Joseph Mancuso 1991 William Hinze 2008 Theodore J. Bornhorst 1992 William F. Cannon 2009 L. Gordon Medaris, Jr. 1993 Donald W. Davis 2010 William D. Addison & Gregory R. Brumpton 1994 Cedric Iverson 2011 Dean M. Rossell 1995 Gene La Berge 2012 Jim Miller 2013 GOLDICH MEDAL RECIPIENT Tom Waggoner Goldich Medal Committee Serving through the meeting year shown in parentheses. Laurel Woodruff (2013) United States Geological Survey Graham Wilson (2014) Turnstone Consulting Bernhardt Saini-Eidukat (2015) North Dakota State University x �Citation for Goldich Medal Recipient Tom helped Cliffs make the transition from underground mining of direct shipping ore to pellets, which are now the life blood of iron mining in the Lake Superior region. The success of his contributions is best measured by 40 years of continual production of pellets from the Empire and Tilden mines that together produce 13 to 14 million tons per year. The transition to pellets is public knowledge but much of the details are lacking. Before pellet production could begin iron formations had to be sampled for their recoverable iron content and ease of recovery. After mining, there had to be day to day adjustment in grinding time, depending on the hardness of taconite, and blending ores in order to control phosphorous contents. Tom’s contribution to preserving iron mining in the Lake Superior region has been documented by his colleagues because the skills of company geologists are seldom part of the public domain. During his long career with Cliffs (1965-1997), Tom actively participated in the teaching function of the Institute. He led 45 field trips for various organization and made numerous poster and oral presentations at annual meetings of the Institute. At his retirement in 1997, Tom was chief geologist. After his retirement he has continued to serve the mining industry. After the state of Michigan ended funding a curator for the core library at Harvey, Tom became the unpaid overseer of the facility and went looking for a building to assemble and preserve the large amount of additional core that exists at scattered location in the upper peninsula. Like the core library at Harvey it will be a repository for future exploration geologists who are essential for sustaining mining in the Lake Superior region. He found a building on the grounds of the closed K. I. Sawyer Air Force base that was suitable as a core library and has been actively seeking funding to acquire and maintain it. Submitted by Ronald E. Seavoy xi �Memorial to Glenn B. Morey 1935-2012 Glenn B. Morey, geologist, known to family, friends, and colleagues as “G.B.” or “Morey,” died at age 76 on August 2, 2012 in St. Paul, Minnesota. He spent the greater part of his 40-year career at the Minnesota Geological Survey in positions that ranged from junior geologist to associate director and chief geologist. He was a senior fellow of the Geological Society of America, a member of the Society of Economic Geologists, the History of Earth Sciences Society, and a lifelong member of the Mesabi Range Geological Society and the Institute on Lake Superior Geology. He received the institute’s Goldich Medal in 1986 for his contributions to geologic understanding of the Lake Superior region. G.B. Morey was raised in Proctor, a town on the Duluth, Mesabi, and Iron Range Railroad through which the iron ore and taconite from the Mesabi range passed on its way to the port of Duluth and steel mills of the lower Great Lakes. He learned of iron mining from his father who worked on the railroad, and he devoted much of his professional career to studies of the stratigraphy, mineralogy, and genesis of the Biwabik Iron Formation—the geological source of prosperity in the mining towns of the Mesabi Iron Range and communities along the routes between mine and mill. Morey completed an M.S. degree in 1960 and a Ph.D. in 1965, both at the University of Minnesota-Twin Cities under professor F.M. Swain. Sedimentology and stratigraphy were the principal foci of his graduate program and continued to be his primary geological interests throughout his professional career. His M.S. thesis entitled “Geology of the Keweenawan sediments near Duluth, Minnesota” was the first of many papers he authored or coauthored on the Mesoproterozoic sedimentary sequences within the Midcontinent Rift. Likewise, his Ph.D. thesis entitled “The sedimentology of the Precambrian Rove Formation in northeastern Minnesota” was the precursor to his many publications on Paleoproterozoic clastic rock units associated with iron-formation on the Biwabik, Cuyuna, and Gunflint iron ranges. Morey’s bibliography contains 100 refereed papers, 25 geologic maps, and many published abstracts and field trip guides—including those for the Institute on Lake Superior Geology. G.B. Morey’s professional accomplishments and adherence to high scientific standards are widely recognized and appreciated throughout the Great Lakes region. He mentored and critiqued many of us at the Minnesota Geological Survey and elsewhere, commonly playing the skeptic to extract the best from colleagues. G.B. was a credit to the survey, the University of Minnesota, and the geological profession as a whole. We who knew him celebrate his memory while we mourn our loss. David L. Southwick, Mark Jirsa, and Paul Weiblen xii �Memorial to Paul K. Sims 1917-2011 Paul K. Sims joined the field trip from which no geologist returns on October 29, 2011 in Denver, Colorado. Born in Newton, Illinois on September 8, 1918, Paul excelled in basketball, and entered University of Illinois Business School. The geology bug bit him and by 1940 he was actively engaged in his Master’s Degree (1942) based on rotary drilling in the coal beds of Illinois. After his degree, he began work with the USGS on zinc-lead deposits in Arizona and Washington before serving with distinction in the navy during World War II in the Pacific theater. Contacts at that time led him to Princeton University for his PhD (1950) on the Dover Magnetic District, New Jersey. He pursued many professional avenues while with the USGS, including international work, uranium geochemistry, editing, and geologic mapping. Lake Superior called, on September 1, 1961, he took leadership of the Minnesota Geologic Survey as Director. Within a year, PK initiated programs in all parts of the stratigraphic column, and developed an annual report series highlighting programs and accomplishments that were used to advance MGS activities. He identified geologic mapping as the backbone of research and he and his colleagues produced numerous reports of local to regional interest. After departing Minnesota in 1973, he continued his mapping and tectonic analysis in Wisconsin and Michigan before resuming his pursuit of the Precambrian in the Rocky Mountains. His geological mapping in the Lake Superior District and the western US led him to suggest that deformation during Archean and Proterozoic was not analogous to Phanerozoic plate-tectonic model, but instead consisted of oblique shortening and progressed from ductile to brittle. His work on Archean mantle gneiss domes resulted in his visualization of the Great Lakes Tectonic Zone north of which Archean volcanogenic assemblages prevail, and south of which Proterozoic continental collision and island arc basins dominated Significant professional accomplishments included President of the Society of Economic Geologists (and many committees), Secretary of the Subcommission of Precambrian Stratigraphy (IUGS). Numerous awards included SEG Thayer Lindsley Lecturer (1984-1985) and Ralph Marsden Award (1989); USGS Meritorious Service Award; and ILSG Goldich Medal (1985). PK was a reserved individual who was enormously unselfish and was quick to give credit to those with whom he collaborated. He encouraged all with whom he came in contact with. In his own work he was demanding and expected clear documentation of which page was fact and which was interpretation. He was a joy to be with, and always stimulated discussion. He was a credit to our profession, and we mourn his passing. Michael Mudrey xiii �Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xiv �Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive onehalf of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. xv �In 2012 the ILSG Board of Governors awarded five $500 awards from the Student Research Fund. The winners were: Jonathan E. Dyess University of Minnesota-Duluth Dept. of Geological Sciences, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 Current degree program: PhD Candidate (Advisor: Vicki Hansen) Shagawa Lake shear zone Elisa Piispa Michigan Technological University 1400 Townsend Drive, 630 DOW, ESE Building, Houghton, MI 49931-1295 Current degree program: PhD in Geology Paleomagnetism of the ~1140 Ma lamprophyre dykes in Ontario, Canada: Implications for the mantle plume hypothesis for Mid-Continental Rift origin Evgeniy V. Kulakov Michigan Technological University Department of Geological and Mining Engineering and Sciences 617 Dow ESE Bldg, 1400 Townsend Drive, Houghton, MI 49931-1295 Current degree program: PhD Paleomagnetism and Geochemistry of the Porcupine Volcanics and Lake Shore Traps: Implications for the Midcontinent Rift evolution. Mark Leatherman 1001 E. 10th Street, Department of Geological Sciences, Bloomington, IN 47405 Current degree program: PhD The Eagle and Tamarack Deposits Craig Caton Department of Geological Sciences
 229 Heller Hall, 1114 Kirby Drive, University of Minnesota Duluth, Duluth, MN 55812 Current degree program: Masters Petrogenesis and Metallogenesis of the Southern Troctolite Zone of the Bald Eagle Intrusion, Duluth Complex, Northeastern MN xvi �Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and long-time friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. Student Paper Awards Committee Helene Lukey – Cliffs Natural Resources Tom Fitz – Northland College Milt Gere – Michigan DNR (retired) xvii �Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Allan R. Blaske (Co-Chair) (2013-2016) – AECOM Theodore J. Bornhorst (Co-Chair) (2013-2016) – Michigan Technological University Tom Fitz (2011-2014) – Northland College Peter Hinz (2010-2013) – Ontario Geological Survey Pete Hollings - Secretary (2010-2013) – Lakehead University Mark A. Jirsa - Treasurer (2011-2014) – Minnesota Geological Survey Local Committee Chair Theodore J. Bornhorst – A. E. Seaman Mineral Museum, Michigan Technological University Allan Blaske – AECOM Volume Editors Theodore J. Bornhorst – A. E. Seaman Mineral Museum, Michigan Technological University Allan Blaske – AECOM Robert Barron – Department of Geological and Mining Engineering and Sciences, Michigan Technological University Session Chairs Jack Berkley – SUNY Fredonia Marcia Bjørnerud – Lawrence University Terry Boerboom – Minnesota Geological Survey Paula Leier-Engelhardt – HydroGeo Solutions Joe Maki – Michigan Department of Environmental Quality Glenn Scott – Cliffs Natural Resources xviii �Banquet Speaker Dr. James W. Ashley Postdoctoral Research Associate Lunar Reconnaissance Orbiter Camera Science Operations Center School of Earth and Space Exploration Arizona State University Rusty Metal at the Martian Equator: The Search for Life on the Red Planet The pursuit of an answer to the ancient question "Are We Alone in the Universe?" leads scientists down many paths that cross a multitude of scientific disciplines. In the planetary sciences, the quest often results in the careful engineering of robotic spacecraft designed to answer specific questions about the planets they are sent to explore. Mars is a world that is both easily accessible at reasonable costs, and potentially habitable. We are interested in the roles that water may have played in Mars' geologic history because of its importance to life on Earth. The Mars Exploration Rover (MER) mission was designed to last for 90 days on Mars in 2004. One of the two rovers (Opportunity) continues exploring today more than nine years later. Among the many discoveries made during this mission are several large, iron meteorites that show dramatic signs of corrosion and other effects of water interaction. MER science team member Dr. James Ashley lead the meteorite discovery and assessment campaign on the mission, and will show how rusty meteorites on Mars are giving us new insight into climate conditions at the red planet's equator. Dr. James Ashley, a Grand Rapids native, earned his BS in geology at Grand Valley State University, his MS in geological science at Michigan State University, and his PhD under Philip Christensen at Arizona State University. He is currently a postdoctoral research fellow at the Lunar Reconnaissance Orbiter Camera (LROC) Science Operations Center at ASU, where he studies Earth's Moon using LROC instruments. His most recent work has focused on Si-rich volcanic complexes on the Moon, potentially cavernous systems beneath the lunar surface, and the understanding of impact melt on the lunar far side. He is the executive director of Minor Planet Research, Inc., a non-profit company dedicated to mitigating the asteroid impact hazard, and has made many appearances on the History and Discovery Channels discussing the threat from near-Earth asteroids. Prior to his work in the planetary sciences, Dr. Ashley worked full time as a consulting hydrogeologist, and is a member of the American Institute of Professional Geologists. xix �Report of the Chair of the 58th Annual Meeting Thunder Bay, Ontario The 58th ILSG was held in Thunder Bay, Ontario on May 16-20, 2012. The meeting was chaired by Pete Hollings (Lakehead University) with the considerable assistance of the local organizing committee (Bill Addison, Mark Smyk, Peter Hinz, Al MacTavish & Dorothy Campbell), the meeting was attended by a total of 240 delegates. Thanks to very generous support from our corporate sponsors (Goldcorp Inc. – Musselwhite Mine, Osisko Mining Corporation, Activation Laboratories Ltd. (Actlabs), Cliffs Natural Resources Inc., Magma Metals (Canada) Limited, MMG Limited, Rio Tinto, Stillwater Canada Inc., Fladgate Exploration, Thunder Bay CEDC, Mega Precious Metals Inc., Rainy River Resources Ltd., Metals Creek Resources Corp., Midwest Institute of Geosciences and Engineering (MIGE), Benton Resources Corp. & Rare Earth Metals Inc.) we were able to provide free registration to the 60 students who attended. The two-day technical session began on Thursday morning with oral presentations on regional geology and then proceeded in chronological order from Archean topics through Friday afternoon’s presentations on Quaternary geology. A total of 27 talks were given, 7 of which were presented by students. A total of 31 posters were displayed, 15 of which were presented by student authors. The 2011 Goldich Medal was awarded to Jim Miller from the University of Minnesota Duluth. Mark Smyk presented the award during the annual banquet and supplied numerous examples of Jim’s passion for the geology of the Lake Superior region. David Overstreet gave the banquet address, discussing Human Adaptation to Late Pleistocene Landscapes - A View from Southeastern Wisconsin. For the first time we ran a core shack during the meeting with core from the Ring of Fire chromite deposits provided by Cliffs, from the Current Lake Ni-Cu-PGE deposit provided by Magma Metals Ltd. and core from Musselwhite Mine provided by Goldcorp Inc. The meeting offered 13 field trips that highlighted the geology of Thunder Bay region. Four premeeting trips were run on Wednesday, including the Lac des Iles Pd mine led by John Corkery (North American Palladium Ltd.), the Sudbury Impactoclastic Debrisites at Thunder Bay led by Bill Addison and Greg Brumpton, the Geology of the Sibley Peninsula led by Dr. Philip Fralick (Lakehead University) Mark Smyk & Riku Metsaranta (Ontario Geological Survey) and the Shebandowan greenstone belt led by Alan Aubut (Sibley Basin Group Geological Consulting Services Ltd.) and Dorothy Campbell (Ontario Geological Survey). On the Friday afternoon three trips were offered, the Geology of the City of Thunder Bay led by Mark Smyk (Ontario Geological Survey), the Panorama Amethyst Mine led by Steve Kissin (Lakehead University) and the Port Arthur building stone walking tour led by Peter Hinz (Ontario Geological Survey). Following the meeting, Goldcorp Inc. flew a lucky group of individuals up to the Musselwhite Mine for a tour led by John Biczok (Goldcorp Inc., Musselwhite Mine). Mark Puumala (OGS) led a trip to examine the Rehabilitation of the Past-Producing Shebandowan and North Coldstream Mines. Scott Hamilton (Lakehead University) took a group to look at the Geoarchaeology of Thunder Bay. Rob Cundari & Pete Hollings (Lakehead University) and Mark Smyk (Ontario Geological Survey) led a trip to look at the Midcontinent Rift intrusions. Greg Brumpton led a reprise of the Ejecta trip and Mark Smyk (OGS) and Philip Fralick (Lakehead University) took a group on a two-day trip to examine the geology along the Highway 527 xx �Transect. Many of the trips sold out and all were well-attended. On Friday evening the organizers hosted a barbeque social event for ILSG participants at the Whitewater Golf Club. The Institute’s Board of Directors met on May 17 to discuss the business of the Institute. The meeting was attended by Al MacTavish, Peter Hinz, Mark Smyk, George Hudak, Bill Addison, Dorothy Campbell, Mark Jirsa and Pete Hollings. ILSG Secretary Hollings took the minutes of the meeting, that are as follows: 1. Accepted the report of the Chair for the 58th ILSG, Ashland, Wisconsin; as printed in the 2012 Proceedings Volume, and minutes of last Board meeting, May 19, 2011. 2. Accepted the 2011-2012 ILSG Financial Summary. 3. Accepted the 2011-2012 ILSG Secretary’s report. 4. Appointed Al MacTavish as the on-going ILSG Board Member. 5. Discussed and approved 2013 (59th annual) meeting location. 6. Replaced Mary Louise Hill as “academic member” on Goldich Committee with Bernhardt Saini-Eidukat. 7. Discussed and approved the naming of the student research award after Joe Mancuso and student paper awards after Doug Duskin The Chair would like to thank all those who assisted with the running of this year’s meeting either by chairing sessions, running field trips or helping with the Registration desk. He is particularly appreciative of the work of the local organizing committee who made his job much easier through their tireless efforts. Respectfully submitted, Pete Hollings Chair, 58th Institute on Lake Superior Geology xxi �Sponsors The following organizations made general contributions to the 59th Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. For the past 59 years this organization has thrived as a result of the interest of individuals, corporations, universities and government agencies. The dedication to an exchange of scientific ideas and a passion for field trips has enabled the Institute to provide one of its primary objectives – to promote better understanding of the geology of the Lake Superior region. EAGLE MINE xxii �PROGRAM WEDNESDAY MAY 8, 2013 8:00 a.m. FIELD TRIP 1: GEOLOGIC OVERVIEW OF THE KEWEENAW PENINSULA, MICHIGAN Ted Bornhorst, Michigan Tech 8:00 a.m. FIELD TRIP 2: CALEDONIA MINE, KEWEENAW PENINSULA NATIVE COPPER DISTRICT, ONTONAGON COUNTY, MICHIGAN Robert Barron, Michigan Tech Richard Whiteman, Red Metal Minerals Ted Bornhorst, Michigan Tech 8:00 a.m. FIELD TRIP 3: GEOLOGY OF SILVER MOUNTAIN, HOUGHTON COUNTY, MICHIGAN Evgeniy Kulakov, Michigan Tech 5:00 p.m. Return of Trips 1, 2, and 3 4:00 p.m. - 8:00 p.m. Registration at Franklin Square Inn, 7th Floor 7:00 p.m. - 9:00 p.m. Ice Breaker Social and Poster Session, Franklin Square Inn, 7th Floor THURSDAY MAY 9, 2013 Note: Asterisk * denotes a student eligible for Best Student Paper Award + denotes students which qualify for travel awards, but not Best Student Paper awards 8:00 a.m. - 2:40 noon REGISTRATION 8:20 a.m. INTRODUCTORY REMARKS Theodore J. Bornhorst and Allan R. Blaske, Co-Chairs, 2013 ILSG TECHNICAL SESSION I Session Chairs: Marcia Bjørnerud – Lawrence University Terry Boerboom – Minnesota Geological Survey 8:30 a.m. Peter Hollings and Mark Smyk Preliminary geochemical analysis of the Nipigon Bay granites, northern Lake Superior xxiii �8:50 a.m. Emily Smyk*, Pete Hollings, and John Biczok Geochemical and petrographic study of a Mesoarchean felsic metavolcanic unit near Musselwhite Mine, North Caribou greenstone belt, northwestern Ontario 9:10 a.m. Aubrey Lee* and Jim Miller The Igneous Stratigraphy of the Bad Vermilion Intrusion, Mine Centre, Ontario, Canada: Which Way is Up? 9:30 a.m. Skylar Schmidt* and Mary Louise Hill North American Palladium’s Lac des Iles mine: Evidence for high temperature deformation and possible control on Pd mineralization 9:50 a.m. Ben Kuzmich*, Pete Hollings, and Michel G. Houlé Preliminary Investigations of the Fe-Ti-V-P mineralization associated with the Thunderbird and Butler gabbroic intrusions within the McFaulds greenstone belt, Superior Province, Northern Ontario, Canada 10:10 a.m. to 10:30 a.m. COFFEE BREAK AND POSTER SESSION 10:30 a.m. Ian R. Dasti* and Stephen A. Kissin The Geochemistry and Mineralogy of the Sulfides within the Ni-Cu-PGE Shakespeare Deposit, Ontario 10:50 a.m. Erik Haroldson and Philip Brown Fluid Inclusion study of the Magino Archean Gold Deposit; Implications for Regional Mineralizing Systems 11:10 a.m. Sarah Canning, Zoran Madon, and Keith Wallace A Geological Model and Resource Update for the Hammond Reef Gold Deposit 11:30 a.m. LUNCH BREAK – 2013 ILSG Board Meeting (by invitation) TECHNICAL SESSION II Session Chairs: Paula Leier-Engelhardt – HydroGeo Solutions Glenn Scott – Cliffs Mining Service Company 1:00 p.m. Phillip Larson Chemostratigraphy of the Biwabik Iron Formation: Implications for Basin Longevity and Evolution 1:20 p.m. Christopher Yip*and Philip Fralick Sedimentology and Geochemistry of a Regressive Surface in the Chemical Sediments of the Paleoproterozoic Gunflint Formation xxiv �1:40 p.m. Gordon Medaris Jr., Terry Boerboom, Brian Jicha and Brad Singer The McGrath metasaprolite: viewing Paleoproterozoic weathering through a veil of metamorphism and metasomatism 2:00 p.m. Mark Puumala Natural Groundwater Geochemistry in Bedrock of the Thunder Bay Area 2:20 p.m. – 2:40 p.m. COFFEE BREAK AND POSTER SESSION 2:40 p.m. William F. Cannon, Laurel G. Woodruff, and Klaus J. Schulz The Hiawatha Graywacke of the Iron River-Crystal Falls district, Michigan: a megaturbidite triggered by seismicity related to the 1850 Ma Sudbury impact 3:00 p.m. Monica M. Karman* and Philip W. Fralick Sedimentology and Paleographic Reconstruction of the Strata Adjacent to the Sudbury Impact Layer in a Cored Drillhole 3:20 p.m. Daniel LaFontaine* and Philip Fralick Sedimentology and geochemistry of the Espanola Formation, Huronian Supergroup 3:40 p.m. Breanne Beh*and Philip Fralick Depositional Processes Operating on the Paleoproterozoic Gowganda Ice Margin 4:00 p.m. SESSION ENDS 6:00 p.m. ICE BREAKER – MIXER – CASH BAR 7:00 p.m. ANNUAL BANQUET AND AWARD PRESENTATION • Announcement of 60th Annual Meeting Location • 2013 Goldich Award Presentation to Tom Waggoner • 2013 Banquet Address by Dr. James W. Ashley, Postdoctoral Research Associate LROC, School of Earth and Space Exploration, Arizona State University All registered participants are welcome to the banquet address xxv �FRIDAY MAY 10, 2013 8:25 a.m. INTRODUCTORY REMARKS Theodore J. Bornhorst and Allan R. Blaske, Co-Chairs, 2013 ILSG TECHNICAL SESSION III Session Chairs: Jack Berkley – SUNY Fredonia Joe Maki – Michigan Department of Environmental Quality 8:30 a.m. Teresa Johnson*, Richard Wendlandt, and James Shannon Geochemistry of reversely-polarized intrusions along the SW limb of the Midcontinent rift system, Carlton County, Minnesota 8:50 a.m. Raymond Anderson and Ryan Clark The Northeast Iowa Intrusive Complex; a Duluth Complex analog? What we know as the investigation begins. 9:10 a.m. Benjamin Drenth, Raymond Anderson, Val Chandler, William Cannon, Klaus Schulz, Joshua M. Feinberg, Paul Bedrosian, and Andy M. Kass High-resolution, multi-method geophysical imaging of a portion of the Northeast Iowa Intrusive Complex 9:30 a.m. Robert Cundari, Peter Hollings, and Mark Smyk Geochemistry of the Logan Igneous Suite and implications for the magmatic evolution of the northern part of the Midcontinent Rift 9:50 a.m. Jack Berkley Lithospheric Delamination during Midcontinent Rifting 10:10 a.m. – 10:30 a.m. COFFEE BREAK AND END OF POSTER SESSION 10:30 a.m. Connor Mulcahy, Edward Hansen, Theodore Bornhorst, and Dieter Rhede Chemical Zoning in Calc-Silicate Minerals Associated with Native Copper from the Keweenaw Peninsula, Michigan 10:50 a.m. Alex C. Brown Brine viscosity vs. temperature: A key to copper deposition in the finest-grained basal Nonesuch Formation, White Pine-Presque Isle district, northern Michigan 11:10 a.m. Stanley L. Vitton Glacial Lake Ontonagon and the Development of Large Scale Landslides 11:30 a.m. Bruce A. Brown Hydrofrac Sand: A major mining boom in the upper Midwest xxvi �11:50 a.m. Val W. Chandler and Richard S. Lively Passive-aggressive geophysics: An update on using the horizontal-to-vertical spectral ratio (HVSR) passive seismic method for determining glacial deposit thickness in Minnesota 12:10 p.m. Presentation of Student Awards Student Travel Awards Best Student Paper Awards 12:30 p.m. LUNCH BREAK AND END OF TECHNICAL SESSIONS 2:00 p.m.–7:00 p.m. FIELD TRIP 4: OPEN HOUSE AT A. E. SEAMAN MINERAL MUSEUM, MICHIGAN TECH, 1404 E. SHARON AVENUE, HOUGHTON Ted Bornhorst, Museum Director FIELD TRIP 5 PRESENTATION AT 8:00 P.M., AMERICINN, SILVER CITY, MI SATURDAY MAY 11, 2013 8:00 a.m. FIELD TRIP 5: GEOLOGY OF THE KEWEENAWAN SUPERGROUP, PORCUPINE MOUNTAINS, ONTONAGON AND GOGEBIC COUNTIES, MICHIGAN William Cannon, Laurel Woodruff, Klaus Schulz, Suzanne Nicholson, USGS LEAVES FROM AMERICINN, SILVER CITY 7:00 a.m. FIELD TRIP 6: GEOLOGY AND ENVIRONMENTAL SITE CONDITIONS OF THE COPPERWOOD DEPOSIT, GOGEBIC COUNTY, MICHIGAN Ted Bornhorst, Michigan Tech Allan Blaske, AECOM Dave Anderson and Tom Repaal, Orvana Resources US Corp. MEETS AT ORVANA RESOURCES US CORP OFFICE IN IRONWOOD, MICHIGAN AT 8:30 AM CDT LEAVES HOUGHTON FRANKLIN SQUARE INN AT 7:00 EDT 5:00 p.m. Return of Trips 5 and 6 END OF 59TH ANNUAL INSTITUTE ON LAKE SUPERIOR GEOLOGY xxvii �POSTER PRESENTATIONS Steven D. J. Baumann, Alex B. Cory and David Wilson Precambrian Faulting in the Ripon Wisconsin Area and Its Impacts on Groundwater Contamination, Originating at Superfund Site Ripon NN/FF Landfill Craig Caton Crystallization of Chrome Spinel in the Southern Troctolite Zone of the Bald Eagle Intrusion, Duluth Complex, Northeastern MN Jonathan Dyess* and Vicki Hansen Application of LiDAR to resolving regional tectonic and glacial fabrics in glaciated terrane: An example from an Archean granite-greenstone belt in NE Minnesota Jonathan Dyess* and Vicki Hansen Structural and Kinematic Analysis of the Shagawa Lake Shear Zone and Snowbank Lake Stock, Superior Province, NE Minnesota Ellen Fehrs+, Edward Kenny+, John Kuchma+, Sarah Sauer+, William Sylvester+, and George Hudak Bedrock Geologic Map of the Putnam Lake Area, St. Louis County, NE Minnesota – Precambrian Research Center Capstone Project George Hudak, Stephen Monson Geerts, Larry Zanko, April Severson, Allison Severson, Stuart Kramer and Bryan Bandli The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter - 2013 Update Mark A. Jirsa Bedrock geologic map of the western Gunflint Trail area, northeastern Minnesota Mark A. Jirsa, Dale R. Setterholm, and V. W. Chandler Minnesota River Valley subprovince as depicted on a new bedrock geologic map of Renville County, southwestern Minnesota Katrina Korman+, Suzanne Craddock+, Michael Doyle+, Jessica Walter+, Aubrey Lee+, and Mark Jirsa Geologic mapping of Neoarchean and Paleoproterozoic rocks near Ester Lake by students of the Precambrian Research Center's 2012 field camp – Precambrian Research Center Capstone Project xxviii �Mark Leatherman*, Edward Ripley, Dean Rossell, Andrew Ware, and Chusi Li Geochemistry and origin of slate-hosted massive sulfides of the Eagle Ni-Cu-PGE Deposit, northern Michigan: A preliminary study. Aubrey Lee*, and Jim Miller Field, Petrographic, and Geochemical Study of the Bad Vermilion Intrusion, Mine Centre, Ontario, Canada Adam Leu+, Lionel Djon+, Emily LaPietra+, Zech Martin+, Ricardo Martinez+, and Jim Miller 2012 Precambrian Field Camp Mapping in the Wilder Lake Intrusion, Lake County, Northeastern Minnesota – Precambrian Research Center Capstone Project Steven Losh and Ryan Rague Silica Remobilization in the Biwabik Iron Formation, Minnesota USA Brynley Nadziejka* and Marcia Bjørnerud Contrasting pressure-temperature-time paths for high-grade metamorphic rocks in the interior of the Penokean-Yavapai orogenic belt, southern Lake Superior region Matthew Schmus*, Prajukit Bhattacharyya, and David Hart Effects of Preexisting Fractures on Groundwater Flow Today Klaus J. Schulz, William F. Cannon, and Laurel G. Woodruff The Parent Lake Volcanics: Product of a phreatomagmatic eruption of basalt during deposition of the Michigamme Formation? Brent Trevisan*, Pete Hollings and Doreen Ames Petrology, mineralization, and alteration of the Thunder mafic to ultramafic intrusion, Midcontinent Rift, Thunder Bay Peter Voice, William Harrison, and Joyashish Thakurta A Preliminary Survey of the Geology of the Pre- Michigan Basin Rocks of the Southern Peninsula xxix �The Northeast Iowa Intrusive Complex; a Duluth Complex analog? What we know as the investigation begins. ANDERSON, Raymond and CLARK, Ryan Iowa Geological and Water Survey, 109 Trowbridge Hall, Iowa City, Iowa 52242-1319 The Northeast Iowa Intrusive Complex (NEIIC) is defined by a suite of gravity and magnetic anomalies that stretch from east-central Iowa to southeast Minnesota and are currently interpreted as mafic intrusions. They are characterized by a series of intersecting circular positive gravity anomalies and corresponding positive and negative aeromagnetic anomalies (Figs 1, 2). The NEIIC was described by Pals and Anderson (2011) who proposed that the complex was of Keweenawan age and analogous to the Duluth Complex. These anomalies were sampled by drilling in four locations (Figs 1, 2). Cores in Minnesota penetrated a "gabbroic rock" (B-1) and a metagabbro (BO-1) that was dated at 1760 Ma (Van Schmus et al., 2007) and may be country rock to the NEIIC. In northeast Iowa a 90 year-old oil test (Pioneer #1) penetrated 480 m of "troctolite" or "olivine gabbro" which yielded a Rb/Sr age of 1,130 Ma (Lidiak et al., 1966); no samples are currently available. A mineral exploration core recovered from northeast Iowa (A12) sampled 220 m of serpentenite and troctolite from a dike-like anomaly. Drill data and other interpretations suggest that the Precambrian basement surface lies about 350 m (north) to about 900 m (south) below the land surface. Many of the geophysical anomalies associated with the NEIIC have been surveyed and modeled by geology students (Fig 3). These anomalies have been interpreted as mafic intrusives, including lopoliths, dikes, and plug-like intrusions (see Dixt, 1984; Heathcote, 1979; Kittleson, 1975; Stepanek, 1978). Two of the modeled intrusives (Dixt, 1984; Heathcote, 1979) were interpreted as mafic lopoliths with maximum diameters of 47 km and 37 km with density contrasts of +0.3 g/cm3 with the felsic country rocks. Heathcote’s (1979) model displayed a Koeningsberger ratio (remanent to induced magnetism) of 6.78 and model remanent vectors consistent with Keweenawan directions and angles. Her model also featured a magnetic field reversal (normal to reverse) captured during the cooling of the intrusive (Fig 4). Kittleson’s (1975) analysis of the Osborne A1-2 core revealed an ultramafic composition composed of cyclic layers of olivine cumulates and olivine-plagioclase cumulates that constituted the upper portion of a dike-like intrusive with a width of about 300 m and a depth extent of about 4.8 km. The age of these intrusives is currently interpreted as Keweenawan based on several lines of evidence. The rocks into which the NEIIC was intruded were interpreted as Yavapai (geon 17) by Van Schmus and others (2007). The only subsequent major magmatism in the area was the widespread felsic anorogenic events (ca. 1,470 and 1,370 Ma) which produced low density plutons with gravity anomalies lower than regional values. Additionally, the 1,130 Ma age (Pioneer #1), the use of Keweenawan remanence vectors in pluton modeling (see above), and the trend of the NEIIC, subparallel to the Keweenawan Midcontinent Rift System, argue for a Keweenawan age. The U.S. Geological Survey, working with the state geological surveys of Iowa and Minnesota, has recently began investigation the potential of the NEIIC for Duluth Complex-like platinum group, nickel, and copper mineralization. Initial stages include acquisition of additional geophysical data and comprehensive modeling and interpretation, with possible core drilling to follow. 1 �Figure 1. Shaded relief total magnetic intensity map of NEIIC and Midcontinent Rift System (MRS) (http://www.mngs.umn.edu/nicegeo/niceimgs.htm) Figure 2. Bouguer gravity anomaly map of NEIIC (http://www.mngs.umn.edu/nicegeo/niceimgs.htm) Figure 3. Map identifying NEIIC intrusives that were modeled by geology students on Aeromagnetic Map of Iowa Figure 4. Magnetic model for north-south profile across Manchester Anomaly (Heathcote, 1979) References Cited Dixit, S.R., 1984, A geologic interpretation of the Vinton geophysical anomaly, in east-central Iowa: unpub. M.S. thesis, University of Iowa, 136 p. Heathcote, S.K.H., 1979, Geological interpretation of the Manchester geophysical anomaly, Delaware County, Iowa: unpub. M.S.. thesis, University of Iowa, 110 p. Kittleson, K.L., 1975, A gravity study of the Osborne magnetic anomaly, Clayton County, Iowa: unpub. M.S. thesis, University of Iowa, 81 p. Lidiak, E.G., Marvin, R.F., Tomas, H.H., and Bass, S.S.,1966, Geochronology of the Midcontinent Region No. 3: Journal of Geophysical Research, v. 71, p. 5,427-5,438. Pals, D.W., and Anderson, R.R., 2011, Reassembling Iowa: spatial and temporal evaluation of the mineral potential of the Iowa segment of the Micontinent Rift and related plutons: Geological Society of America Abstracts with Programs, v. 43, no. 5 p. 396. Stepanek, J.G., 1978, Geological interpretation of an aeromagnetic anomaly near Randalia, northeastern Iowa: unpub. M.S. thesis, University of Iowa, 105 p. Van Schmus, W.R., Schneider, D.A., Holm, D.K., Dodson, S., and Nelson, B.K., 2007, New insights into the southern margin of the Archean-Proterozoic boundary in the north-central United States based on U-Pb, Sm-Nd, and Ar-Ar geochronology: Precambrian Research, v. 157, p. 80-105. 2 �Precambrian Faulting in the Ripon Wisconsin Area and Its Impacts on Groundwater Contamination Originating at Superfund Site Ripon NN/FF Landfill BAUMANN, Steven D.J.1, CORY, Alex B.1 and WILSON, David2 1 Geology Section, Midwest Institute of Geosciences and Engineering, 2328 W. Touhy Ave. Chicago, IL 60645 2 Superfund Division, U.S. EPA Region 5, 77 West Jackson Blvd. Chicago, IL 60604 Faults are known to serve as conduits for the migration of contaminants, especially in the subsurface. One case of particular interest is the Superfund site designated as the Ripon NN/FF Landfill (Ripon Superfund), center of which is located 1,000 feet south-southeast of the junction of County Road FF and South Koro Road (GPS: 43.866850o, -88.870945o). This superfund site has historically shown vinyl chloride contamination in several off-site monitoring wells. The nearest potable well for the city of Ripon has also shown contamination. The contamination has spread in a manner that is inconsistent with traditional modeling techniques, suggesting that something unknown is occurring at the Precambrian-Cambrian contact. Thanks to information provided by David Wilson of the United States Environmental Protection Agency (U.S. EPA) new insight has been gained about the subsurface geology. The reason for the unusual migration of contaminants is most likely due to unmapped parallel faults in the Precambrian basement rock. A detailed bedrock geologic map was compiled by Steven Baumann in 2011 of Ripon and the surrounding area, to include the Ripon Superfund location. Faults in the basement rock were expected at the time the bedrock map was compiled, but could not be confirmed. In 1993 a deep borehole (P-107D) into the bedrock was drilled through the glacial cover, through the underlying Cambrian sands, and down into the Precambrian basement rock. During the compilation of the Ripon Wisconsin bedrock map (Baumann 2011), the superfund borings were unknown to the author. Even without the Superfund logs the surficial field work conducted in the Ripon area yielded several distinct and prominent structural features. No faults were observed extending up through the youngest surface rocks. Although faulting was suspected at the time of mapping, the available data did not provide direct evidence of the presence of faults. Based on the detailed log for boring P-107D (the only detailed log to significantly penetrate the Precambrian) the basement consists of a thin layer of purple quartzite on top of a thick red granitic or syenite sequence, similar to the rocks exposed in the Baraboo Wisconsin area. The Precambrian geology of Ripon is expected to be similar to that of Baraboo. At Baraboo, there is highly fractured Precambrian quartzite on top of igneous rocks, which in turn are covered by thick sequences of Cambrian and Ordovician rocks. The only real difference between Baraboo and Ripon is the thick cover of glacial outwash deposits present at Ripon. The western half of the Baraboo Precambrian exposures are part of the “Driftless Area” and were not glaciated. The similarities between Baraboo and Ripon, give justification for modeling structural features in the Ripon area in a similar manner to the structures observed in the Baraboo area. The presence of quartzite on the bedrock high penetrated by P-107D is a strong indication of local faulting. Based on the orientation of known local structures in the Cambrian and Ordovician a fault just south of P-107D is expected to trend N55W to N75W based on the orientation of the nearby Ripon Arch. However, near the theoretical north limit of the Ripon Arch (near Arcade Acres) does turn more north then west-northwest. The understanding of any local faulting in the Precambrian is of key importance due to the likelihood of the faults serving as a route of contamination, possibly leading to the City of Ripon’s #9 Supply Well (Wilson 2012). 3 �Any faults in the area are probably roughly parallel to each other and the Ripon Arch. The faults likely show fracturing 60 feet or so within the Precambrian-Cambrian (PC-C) boundary and most likely fade out within the Cambrian Wonewoc Formation. Although fracturing is very likely near the PC-C boundary, the faults are probably tight deeper than 30 feet below the PC-C boundary. The total displacement of the faults is likely on the order of 40 to 80 feet. This will greatly affect groundwater flow direction in the deepest hydrogeologic unit designated as “Layer 4” (Wilson 2012). “Layer 4” is the basal known confined aquifer in the Cambrian and it likely connects to the groundwater in the faults and fractures in the Precambrian basement. Weathering patterns at the PC-C boundary were not noted in the P-107D log. However, using the PC-C boundary at Baraboo, it is likely that the contact is significantly lithified, yet extremely porous. Although the groundwater in P-107D at the PC-C boundary is likely connected to “Layer 4” a definite connection cannot be ascertained without additional data. At present fault dynamics and contaminant migration cannot be definitively determined at present. Additional deep basement down gradient sentinel wells will need to be drilled in order to develop a good model and plan to stop additional contaminants from reaching the Ripon #9 Supply Well. References: Baumann, S.D.J., 2011. Geologic Bedrock Map of the Ripon Area, Green Lake and Fond du Lac Counties, Wisconsin U.S.A. Midwest Institute of Geosciences and Engineering M-102011-2A Baumann, S.D.J., 2011. Surficial Geologic Map of the Upper Narrows near Rock Springs, Sauk County, Wisconsin U.S.A. Midwest Institute of Geosciences and Engineering M-092011-3A Dalziel, I.D.W., Dott Jr., R.H., 1970. Information Circular No. 14: Geology of the Baraboo District Wisconsin. Wisconsin Geological and Natural History Survey Fassbender, J.L., Noel, M.R., Ronk, J.J. 1994. Remedial Investigation Report Ripon FF/NN Landfill Volumes I and II. Hydro-Search Inc. Contract SF-92-01 Wilson, D., 2012. Review of the Monitored Natural Attenuation for Ripon Landfill Site WI. U.S. EPA, Region 5, Superfund Division, Memorandum 4 �DEPOSITIONAL PROCESSES OPERATING ON THE PALEOPROTEROZOIC GOWGANDA ICE MARGIN BEH, Breanne1, and FRALICK1, Philip 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, bbeh@lakeheadu.ca Glacial sedimentary rocks of the Huronian Supergroup crop out along the north shore of Lake Huron and were likely deposited on what is thought to have been a divergent continental margin (Fralick and Miall, 1981; 1989). The rocks of the Gowganda Formation record one of three glacial events preserved in the Supergroup and are therefore of interest for developing further glaciomarine models (Puffett, 1974) and furthering understanding of this early stage in Earth’s history. Data has been collected in five main study areas in an attempt to cover as much of the ancient continental margin as possible. The study areas include Espanola; Elliot Lake; Thessalon; and Cobalt, Ontario; and Marquette, Michigan. There are two glaciogenic formations in the Marquette area of Paleoproterozoic age, the Reany Creek Formation and the Enchantment Lake Formation. The Enchantment Lake Formation has been chronostratigraphically correlated to the Huronian Supergroup based on U-Pb age determination on detrital zircon, 2317±6 Ma, and diagenetic xenotime, 2133±11 Ma (Vallini et al., 2006). As these formations are present in such close proximity to each other, and there are no Archean glacial events recorded in the rest of the Canadian Shield, it is reasonable to correlate them with the Gowganda Formation, the thickest and most commonly preserved of the three Huronian glacial events. Stratigraphic sections were compiled in each of the study areas and the sedimentary rocks were grouped into seven lithofacies associations (LA): 1) Planar Cross-Stratified Sandstone LA, 2) Basal Breccia LA, 3) Diamictite LA, 4) Interlayered Siltstone and Fine-Grained Sandstone LA, 5) Slump LA, 6) Heterogeneous Sandstone LA and 7) Quartz-Rich Sandstone LA. These lithofacies associations likely represent a sequence of depositional environments on a shallow continental shelf. Initially, the shelf was dominated by what were likely large-scale, low-angle sandwaves (Figure 1A), interbedded with successions of wavy bedding and possible hummocky cross-stratification indicating an open-water setting with tidal and storm processes reworking the sediments. The shelf then gradually evolved into an environment dominated by diamicite layers. The diamictite layers have dropstones (Figure 1B) as well as evidence of current activity indicating outsized clasts were likely being introduced into the environment as ice-rafted debris. Resedimentation events in the form of debris flows are thought to account for conglomeratic layers that are common in the Diamictite LA. The Interlayered Siltstone and Fine-Grained Sandstone LA, along with the Slump LA (Figure 1C), seem to indicate deposition from suspension in a prodelta setting where large slump events are common. The gradual transition into a more sandstone dominated LA, with an abundance of current-related sedimentary structures, is indicative of the shallowing and coarsening upwards succession common to deltaic deposits. A final transition into the Quartz-Rich Sandstone LA indicates a return to a sandy, current-dominated open continental shelf environment with abundant tidally generated sedimentary structures such as herringbone cross-stratification (Figure 1D). The Cobalt study area differs from this overall model in that evidence of grounded ice is present. Less exposure in the Marquette study area makes it difficult to draw overall conclusions on the evolution of the continental shelf but deposition in a subaqueous glacial outwash fan is hypothesized. 5 �Figure 1. A) Large-scale, low-angle planar cross-stratified sandstone. B) A dropstone compressing underlying laminations. C) A mud-block conglomerate of slumped prodelta deposits. D) Herringbone cross-stratified sandstone. REFERENCES FRALICK, P.W., AND MIALL, A.D., 1981. Grant 84: Sedimentology of the Matinenda Formation, in Pye, E.G., ed., Geoscience Research Grant Program, Summary of Research 1980 – 1981: Ontario Geological Survey, Miscellaneous Paper 98: 80-89. FRALICK, P.W., AND MIALL, A.D., 1989. Sedimentology of the Lower Huronian Supergroup (Early Proterozoic), Elliot Lake area, Ontario, Canada: Sedimentary Geology, 63: 127-153. PUFFETT, W.P., 1974. Geology of the Negaunee quadrangle, Marquette County, Michigan: U.S. Geological Survey, Professional Paper 788: 53 p. VALLINI, D.A., CANNON, W.F., AND SHULZ, K.J., 2006, Age constraints for Paleoproterozoic glaciation in the Lake Superior Region: detrital zircon and hydrothermal xenotime ages for the Chocolay Group, Marquette Range Supergroup: Canadian Journal of Earth Science, v. 43, p. 571-591. 6 �Lithospheric Delamination during Midcontinent Rifting BERKLEY, Jack, Department of Geosciences, Houghton Hall, SUNY Fredonia, Fredonia, NY 14063 USA The process of delamination, briefly defined as the physical peeling off and descent of a cold lithospheric slab into ascending hot mantle asthenosphere, has historically been applied mostly to convergent orogenic systems (Fig. 1a) involving continent-continent collision and subduction (e.g. Bird, 1979; Hamilton et al., 2004). Delamination has been invoked to explain perplexing tectonic-topographic expressions such as the Colorado Plateau, as well as massive collisional fold and thrust belts like the Himalayas, Alps, and more ancient terrains, notably the Mesoproterozoic Grenville Orogen (e.g., Wallner and Schmeling, 2010; Hamilton et al., 2004). Continental rift terrains (Fig. 1b), although dominated by extensional rather than convergent stress fields, are formed in the same general geophysical milieus as convergent systems (i.e., 30 km+ low-density crust plus upper mantle lithosphere -- overlying hot, low-viscosity asthenosphere). Rift terrains (Fig. 1b) are also subjected to upwelling heat sources with liberated volatiles (esp. water) that function to exacerbate the weakening of previously hyper-strained lithosphere. Given a few extenuating circumstances (see below), upper mantle delamination is predictable, or at least possible. (a) (b) Figure 1: (a) Convergent delamination of a mantle lithospheric root (gray) (from Schott and Schmeling,1998). (b) Extensional delamination in the Basin-Range province, USA (from Meissner and Mooney, 1998). If delamination played a controlling role in the production of magmatic suites of the Midcontinent Rift (MCR) specifically in the well-exposed Lake Superior region, what criteria can be used to detect that influence? Much depends on the precise sub-surface structural configuration underlying the rift environment, conditions that can vary widely from region to region – and, more importantly, along strike in the MCR. Recent and past studies show that the MR consists of two linear branches, the Western branch that extends northeast from Kansas to Lake Superior (active interior continental rift), and the Eastern branch (leaky transform fault) that extends southeast from Lake Superior to SE Michigan, and possibly as far as Kentucky. The bulk of igneous activity occurs in the Lake Superior basin due to the likely influence of a deep mantle plume (e.g., Merino et al., 2012; Davis and Green, 1997). Deep earth tomography reveals that early rift-induced normal faults have later transitioned to low-angle thrust faults in response to west-directed Grenville orogenic compression. In that light, conditions required for rift-induced delamination (“RID”, Wallner and Schmeling, 2010) include: (1) a thermal anomaly (e.g., plume), (2) low-strength lower crust, and (3) lateral density variations within the lower crust. 7 �Excess temperature and low yield stress at depth are especially critical parameters. The Lake Superior region meets the prima facie criteria required for delamination, although that fact does not guarantee that delamination has, in fact, occurred. On the other hand, the presence of late Grenville thrusting could provide a facilitating factor missing in most rift environments. Most confirmed or suspected delamination events occur during compressive, crust-thickening stress events, as in the relative nearby Grenville fold belt (as in the Adirondacks, e.g., Hamilton et al., 2004). Although not exhaustive, the following observations could be used to evaluate the role of delamination in the Lake Superior region: (1) Magmas produced by delamination would be primarily mafic (melting of hot mantle asthenosphere), but a variety of voluminous felsic magmas would also be expected from melting of Archean crust or differentiation of ponded mafic plutons, (2) Igneous rocks directly produced by delamination should show relatively young ages compared to, say, purely plume-generated magmas, and should be restricted in age, (3) Assuming a progressively peeling off and sinking lithospheric slab (or slabs) ages of pluton series may vary older-to-younger laterally across surface exposures, (4) Seismic studies show one or possibly two highly disrupted deep crustal zones under Lake Superior (Hamilton and Mereu, 1993) in which mantle material displays up-welling high up into crustal layers. Obviously, such areas should be considered prime targets for possible delamination effects on geographically associated igneous exposures. References Bird, P., 1979. Continental delamination and the Colorado plateau. Journal of Geophysical Research, 84: 7561-7571. Davis, D.W. and Green, J.C., 1997. Geochronology of the North American Mid-continent rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences, 34(4): 476-488. Hamilton, D.A. and Mereu, R.F., 1993. 2-D tomographic imaging across the North American Midcontinent Rift system; Geophysics Journal International. 112: 344-358. Hamilton, M.A., McLelland, J., and Selleck, B., 2004. SHRIMP U-Pb zircon geochronology of the Anorthosite-mangerite-charnockite-granite suite, Adirondack Mountains, New York: Ages of emplacement and metamorphism. Geological Society of America, Memoir 197: 337-355. Meissner, R. and Mooney, W., 1998. Weakness of the lower continental crust: a condition for delamination, uplift, and escape. Tectonophysics 296: 47-60. Merino, M., Keller, G., Stein, S. and Stein, C., 2012. Variations in mid-continent rift magma volumes consistent with microplate evolution. Pre-publication manuscript, Department Earth and Planetary Sciences, Northwestern University, Evanston, Ill. Schmeling, H., 2010. Dynamic models of continental rifting with melt generation. Tectonophysics, 480: 3347. Schott, B. and Schmeling, H., 1998. Delamination and detachment of a lithospheric root. Tectonophysics, 296: 225-247. Wallner, H. and Schmeling, H., 2010. Rift induced delamination of mantle lithosphere and crustal uplift: a new mechanism for explaining Rwenzori Mountains’ extreme elevation? International Journal Earth Science, 99: 1511-1524. 8 �Brine viscosity vs. temperature: A key to copper deposition in the finest-grained basal Nonesuch Formation, White Pine-Presque Isle district, northern Michigan BROWN, Alex C., 13250 rue Acadie, Pierrefonds, QC H9A 1K9, acbrown@polymtl.ca The sulfide-dominant main-stage stratiform copper (SSC) mineralization in the fine-grained carbonaceous Nonesuch Formation at White Pine, northern Michigan, has long been attributed to a vertical influx of oxidized low-temperature cupriferous brine from the underlying Copper Harbor Conglomerate aquifer during early Nonesuch diagenesis. However, despite the difficult influx of brine from a good aquifer into a fine-grained aquitard, the best-mineralized Nonesuch of the White Pine district is hosted by the finest-grained basal Nonesuch. This dilemma was addressed by White (1971) and Swenson et al. (2004) who showed, in 2-dimensional profiles, that brines driven southward out of the Lake Superior rift basin by compaction could have been forced upward into the basal Nonesuch where the Copper Harbor thins abruptly over a large volcanic dome (Porcupine Volcanics) located directly below the White Pine mine district. However, in 3-dimensions, those brines should have escaped largely to the east and west around the volcanic dome where the Copper Harbor is thick. Furthermore, compaction brines should have equilibrated with ferrous constituents of the aquifer and thus been too reducing to carry significant amounts of copper. On the other hand, copper may be carried in very significant amounts in oxidized lowtemperature brines such as generated during meteoric recharge (Brown, 2005). Furthermore, latent heat from a dormant or extinct resurgent caldera, represented here by the Porcupine Volcanics, could have locally increased the temperature of brine hosted by the Copper Harbor Conglomerate to ~100oC, the estimated maximum temperature experienced by the basal Nonesuch Formation at White Pine (Grigorita and Brown, 2002; Fig. 1). By analogy to modern calderas (e.g., the Valles caldera, New Mexico), heat may have emanated from the Porcupine Volcanics for more than one Fig. 1. a) Location map. b) A schematic illustration of resurgent caldera heat, from the recently dormant or extinct Porcupine Volcanics dome, inducing an exceptional brine infiltration into fine-grained basal Nonesuch beds in the White Pine-Presque Isle area. Lateral grain-size gradations within the Nonesuch are represented by shading: dark gray = very fine-grained, lighter gray = coarser grained. 9 �million years at rates (e.g., up to 2500 mW/m2; Goff and Gardner, 1994) far above the normal heat flow of intracontinental rifts (e.g., 60 to 120 mW/m2). Thermal blanketing by the fine-grained Nonesuch would have aided in the accumulation of anomalous heat in the Copper Harbor brine over the resurgent caldera. Heating of the brine from 20 to 100oC would have decreased its density by about 5% and thus encouraged upward buoyant circulation. More significantly, warming of the brine would have lowered its viscosity to ~30% of its viscosity at 20oC (Fig. 2). On the basis of density and viscosity changes alone, the vertical flux of brine into the basal Nonesuch in the area underlain by the resurgent caldera should have been ~3.5 times the flux in cooler distal areas. Applying Darcy’s law (Ingebritsen and Appold, 2012), a larger infiltration of brine into the basal Nonesuch is found even if the permeability of the finest-grained Nonesuch was 1/2 order of magnitude less than the permeability of more distal, coarser-grained Nonesuch. Furthermore, if brine infiltration into the basal Nonesuch was more rapid where it was warmed locally by a recently dormant or extinct resurgent caldera, replacement brine in the Copper Harbor aquifer should have been drawn laterally toward this area from surrounding areas (Fig. 3). This convergent supply of ore-forming brine in the Copper Harbor aquifer would have been more efficient than the ore-forming process of a linear meteoric recharge-driven scenario. Fig. 3. Sketch illustrating the convergence of meteoric recharge-driven brine (curved arrows) toward the base of the fine-grained Nonesuch Formation (due to anomalous heat accumulated there as a result of thermal blanketing) with a consequently high rate and focus of brine infiltration. Location of highland area is schematic only. Fig. 2. Viscosity (µ) vs. temperature from 20 to 100oC, for H2O and various brines. References Brown, A.C., 2005. Refinements for footwall red-bed diagenesis in the sediment-hosted stratiform copper deposits model. Economic Geology, 100: 765-771. Goff, F. and Gardner, J.N., 1994. Evolution of a mineralized geothermal system, Valles Caldera, New Mexico. Economic Geology, 89: 1803-1832. Grigorita, A. and Brown, A.C., 2002. A resurgent caldera model, rather than a compaction or gravity-driven model, for stratiform copper mineralization at White Pine, Michigan. Soc. Econ. Geol., Newsletter 50 (July 2002), p. 32. Ingebritsen, S.E. and Appold, M.S., 2012. The physical hydrogeology of ore deposits. Econ. Geology, 107: 559584. Swenson, J.B., Person, M., Raffensperger, J.P., Cannon, W.F., Woodruff, L.G. and Berndt, M.E., 2004. A hydro-geological model of stratiform copper mineralization in the Midcontinent Rift System, northern Michigan, USA. Geofluids, 4: 1-22. White, W.S., 1971. A paleohydrologic model for mineralization of the White Pine copper deposit, northern Michigan. Economic Geology, 66: 1-13. 10 �Hydrofrac Sand: A major mining boom in the upper Midwest BROWN, Bruce A. Badger Mining Corp., 409 South Church St., Berlin, WI 54923 The phenomenal growth of the hydrofrac sand industry in Wisconsin, Illinois, and Minnesota in the last five years has resulted in a mining boom the likes of which has not been seen since the discovery of lead-zinc and iron in the 19th century. In 2008 there were less than ten industrialsand mines in Wisconsin, including foundry sand mines. In 2012 the count was more than one hundred, almost exclusively hydrofrac sand operations. The sand boom is the direct result of the successful application of improved horizontal drilling techniques and hydraulic fracturing to hydrocarbon-rich shales that were previously too impermeable to develop by conventional vertical drilling. The hydraulic fracturing process involves applying high pressure to a well sufficient to fracture a hydrocarbon-bearing formation of low permeability. The pressure opens fractures around the well bore, and a proppant, usually sand grains, is injected into the well to prop the fractures open after the pressure is released, allowing oil or natural gas to flow into the well. To be useful as a proppant, sand must be pure quartz, of specific grain size, have high roundness and sphericity, and a high compressive strength. Hydrofrac sand has been produced in the region for over fifty years. The Jordan, Wonewoc, and Mount Simon sandstones of Cambrian age and the Ordovician St. Peter sandstone have long been recognized as excellent sources of proppant sand. Most of the new mines produce from the Wonewoc Formation, which is a good source of the finer (40-70) sand used for gas drilling. The Jordan and Mount Simon Formations are coarser and produce more of the 20-40 mesh sand favored for oil drilling. The St. Peter in Wisconsin and Minnesota is too fine for hydrofrac sand but is mined as a source for foundry sand. The rapid growth of natural gas drilling, particularly in Pennsylvania, Texas, and North Dakota resulted in an acute shortage of sand and high prices for proppant. Energy companies, established sand producers, and entrepreneurs with little experience in the industry rushed to the region to get into the proppant business. In late 2012 a surplus of natural gas resulted in a drop in demand for the finer grades of sand, just at the time that several large mines were coming on line. The industry is now facing the prospect of excess production capacity. The future of hydrofrac sand mining in the upper Midwest looks favorable. Hydraulic fracturing is the key to energy independence for the United States, and the Cambrian sandstones of Wisconsin and Minnesota are the best source of proppant sand in the country. The next few years will be interesting as the industry deals with overcapacity and with an ever-growing list of regulatory and land use issues that have resulted from rapid growth. 11 �A Geological Model and Resource Update for the Hammond Reef Gold Deposit CANNING, Sarah, MADON, Zoran, and WALLACE, Keith. Osisko Hammond Reef Gold Ltd, 101 Goodwin Street, Atikokan, ON P0T 1C0 The Hammond Reef Gold Deposit is located about 210 km west of Thunder Bay and 25 km north-east of Atikokan. On January 28, 2013, Osisko Mining Corporation released a new resource estimate of 7.2 Million ounces after completion of an aggressive definition drill program. The company drilled over 2,100 holes for a total of 629,000 m in the last 2 years and was able to upgrade 75% of the gold inventory into the Measured and Indicated category. Using a 0.5 g/t cut-off, the deposit contains 5.43 Million oz Measured+Indicated @ 0.86 g/t and 1.75 Million oz Inferred @ 0.72 g/t. Figure 1: Artist’s rendition of the Hammond Reef open pit and infrastructure. Permitting for mine development continues. The company recently submitted a Draft Environmental Impact Study/Environmental Assessment Report to both the Canadian Environmental Assessment Agency and the Ontario Ministry of Environment. The Project Feasibility Study is on schedule and expected to be finished by the second quarter of this year. Osisko has received letters of support from all Métis and First Nations communities affected by the project and continues its dialogue with them. 12 �Hammond Reef is situated in the Marmion batholith of the Wabigoon Subprovince near the north-east trending contact with the Finlayson greenstone belt. The Mesoarchean Marmion is a diverse assemblage of felsic intrusive rocks, predominantly tonalitic in composition that was later invaded by several intrusive pulses. These units are all variably sheared and altered along a 1 to 6 km wide anastomosing deformation corridor (Marmion Deformation Corridor - MDC) that is sub-parallel to the contact with the Finlayson volcanics. The MDC is the locus of numerous gold occurrences. As with many structurally-controlled gold deposits, Hammond Reef occurs within a significant flexure of the MDC. The MDC displays numerous characteristics of brittle-ductile deformation, including moderately to strongly sheared rocks, brecciation, veining and stockwork. The MDC is also characterized by a sericite-chlorite-ankerite-hematite alteration overprint. Gold was most likely introduced during this late Archean hydrothermal or metamorphic episode, along with pyrite and accessory sulfides and tellurides. Gold mineralization occurs in all lithological phases of the Marmion batholith, associated with fracture-controlled stockwork and pyrite. Native gold blebs are usually found within pyrite crystals – along py-py grain boundaries as well as healed fractures and inclusions. Free gold grains are found rarely on sericitic foliation planes. Accessory minerals include chalcopyrite, galena and hessite. Three main ore types were identified in the geological model that was developed for the Hammond Reef gold deposit. These include – 1) gold in structurally confined and pervasively altered tonalites, 2) gold impinging into partially altered tonalites, and 3) gold impinging into “unaltered” tonalites. 13 �The Hiawatha Graywacke of the Iron River-Crystal Falls district, Michigan: a megaturbidite triggered by seismicity related to the 1850 Ma Sudbury impact CANNON, W.F.1, WOODRUFF, L. G.2, SCHULZ, K.J1. 1 U.S. Geological Survey, Reston VA, 20191, wcannon@usgs.gov, kschulz@usgs.gov 2 U.S. Geological Survey, Mounds View, MN 55112, woodruff@usgs.gov The Hiawatha Graywacke is a coarse clastic unit, which includes breccias with fragments as large as a meter. It is underlain and overlain by chemical sediments and very fine clastic rocks of the Riverton Iron-formation and Stambaugh Formation respectively. The Hiawatha has long been interpreted as a submarine slump deposit generated by a strong earthquake (James and others, 1968). A connection to the 1850 Ma Sudbury impact event was established with recognition of shock metamorphic features (quartz grains with relict planar deformation features), as well as small fragments of devitrified glass (Cannon and others, 2009). The Hiawatha Graywacke was deposited in a backarc basin along the southern margin of the Superior craton and is the southernmost and deepest water occurrence of the Sudbury Impact Layer so far identified. It is about 20 m thick in the east, where it is almost entirely breccia, and 150 m or more thick in the west, where breccia is mostly near the base and is overlain by massive graywacke with interlayers of siltstone and breccia. The breccia is derived largely from the underlying Riverton Iron-formation. Clasts are chert and siderite, and rarely other sedimentary rocks. The matrix is predominantly fine-grained siderite. Sand-sized clasts of quartz are widespread and typically compose a few percent of the breccia matrix, indicating that material other than iron-formation also was incorporated into the breccias. Quartz grains that contain well preserved planar deformation features, indicative of impact shock, are widespread but generally sparse. A greater abundance of shocked quartz is found in the upper parts of the formation than in basal breccias. Massive siltstone beds in the upper part of the Hiawatha have the greatest abundance of shocked quartz that we have found in the Lake Superior region. This suggests that the breccias are largely locally derived by disaggregation and submarine slumping of the Riverton Iron-formation with incorporation of only minor amounts of other material, whereas the upper parts of the Hiawatha have a greater component of ejecta particles that arrived in the area while submarine slumping was in progress. In the western part of the district, a layer, as much as a few meters thick near the middle of the Hiawatha, is composed mostly of vesicular glass fragments with abundant shocked quartz, indicating that ejecta deposition locally overwhelmed deposition of terrestrial material. The Sudbury impact, about 550 km east of the Iron River-Crystal Falls area, generated an earthquakeof probably roughly magnitude 10.5 based on computer model results. This is about 30 times more powerful than the largest recorded earthquake. The first seismic wave arrived here about 1.5 minutes after the impact and disrupted the seabed, composed of partly consolidated chert and siderite. This material flowed down the paleoslope toward deeper parts of the basin. About 4 minutes later the first airborne ejecta arrived and settled through the water column to be incorporated to variable degrees into the still active submarine slumps. Thicker parts of the Hiawatha in the Iron River area have some stratigraphic variation in mineralogy suggesting that multiple sediment sources were sampled by multiple lobes of turbidites. A third catastrophic event caused by the impact should have been massive tsunami waves, arriving about an hour after the ejecta. We have not identified evidence of tsunami activity, but perhaps upper parts of the Hiawatha were reworked by tsunamis, followed by settling of suspended sediment from the water column to produce the more prominent bedding that characterizes the upper part of the formation. Accumulation of a large volume of slump debris in the western part of the basin requires a comparable amount of denudation elsewhere. Some areas in the southeastern part of the district are devoid of the Riverton Iron-formation and, more locally, of the underlying Dunn Creek Slate, 14 �so that the Hiawatha lies directly on older volcanic rocks. Perhaps these areas were the source for slump debris that was transported and deposited as the basal breccias farther west in the basin. Submarine slumps, well studied in many other parts of the world, commonly begin by mass wasting of large slabs of sediments, some of which become successively disaggregated and water-saturated, grade into debris flows, and eventually into true turbidites. We suggest that this process happened here as well and that there are unrecognized megaclasts of the Riverton within the western part of the basin. A peculiarity of the western part of the basin is a chaotic folding style shown both on maps and cross sections. The western folds are much more complex than in the eastern part of the district and all surrounding areas, where folding is intense but has a regular pattern of deformation. We suggest two possibilities for this unique structural style that might be related to massive slumping. (1) Parts of the Riverton in the western part of the basin may be megablocks of slump material that were folded during slumping. Extremely large blocks of this nature, some many kilometers in extent and hundreds of meters thick, are known elsewhere in the world (ref?). Perhaps part of the Riverton is such a block that was emplaced already folded, so that subsequent tectonic folding formed the uniquely complex pattern seen there. 2) The Hiawatha in the western part of the basin is a sedimentary mélange containing numerous megablocks of Riverton Iron-formation. Previous maps and cross sections of the area were interpreted on the assumption, perhaps unwarranted, of a continuous unit of Riverton and a coherent stratigraphy. By that assumption, all occurrences of iron-formation had to be connected to all others and stratigraphic relationships had to be maintained. To satisfy these requirements many folds had to be inferred. In contrast, if the Hiawatha is a sedimentary mélange containing megaclasts of the Riverton, many of these inferred folds are not required and the structural style, although still complex, may be simpler than previously interpreted. References Cannon, W.F., Schulz, K.J., Horton, J. Wright, Jr., and Kring, David A., 2009, The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA: Geological Society of America Bulletin, v. 122, p. 50-75. Dutton, C.E., 1971, Geology of the Florence area, Wisconsin and Michigan: U.S. Geological Survey Professional Paper 633, 54 p. James, H.L., Dutton, C.E., Pettijohn, F.J., and Wier, K.L., 1968, Geology and ore deposits of the Iron River–Crystal Falls district, Iron County, Michigan: U.S. Geological Survey Professional Paper 570, 134 p. 15 �Crystallization of Chrome Spinel in the Southern Troctolite Zone of the Bald Eagle Intrusion, Duluth Complex, Northeastern MN Caton, Craig, Department of Geological Sciences, University of Minnesota Duluth, 114 Kirby Drive, Duluth MN 55812 Duluth Metals had acquired mineral rights to portions the Bald Eagle Intrusion (BEI), located in the Duluth Complex about 35 km southeast of Ely, MN as a target for Ni-Cu-PGE potential. Exploration drilling by Duluth Metals in the BEI, a funnel-shaped (Weiblen, 1965; Green et al., 1966), differentiated mafic layered intrusion within the 1.1 Ga Duluth Complex in northeastern Minnesota, has identified multiple stratiform intervals enriched in chromium spinel (Cr-spinel) mineralization. Centimeter-scale layers enriched in up to 50 vol. % Cr-spinel occur within troctolite host rocks, which compose most of the southern BEI (Weiblen, 1965). The principal objective of this study is to characterize the mineralogy and textural occurrence of these chromite layers in order to understand their genesis and the magmatic history of the BEI. This information will be used as an exploration tool for further exploration of the BEI and other mafic intrusions within the Duluth Complex. This poster presents the observations made of textures and stratigraphic occurrences of Cr-spinel enriched intervals found in the troctolitic cumulates of the Bald Eagle Intrusion as well as the chemical variations throughout the study area. These observations are focused on three Duluth Metals drill holes that have been logged and assayed for whole rock geochemistry. Using these data along with petrographic descriptions and mineral chemistry from SEM of Cr-spinel layers and adjacent lithologies, possible models for emplacement and crystallization can be theorized. Core logging and follow-up petrographic observations show a sharp basal contact to all semi-massive oxide intervals that denote the start of a cyclic unit. In general, a cycle consists of a lower semi-massive oxide interval followed by an ultramafic section that grades to troctolitic followed by anorthositic compositions. Variations of a cycle sequence may occur. Inclusions are often present in oxide units, however the origin of these inclusions under interpretation. These descriptions suggest that individual cycles could be the product of additional magma inputs with fractional crystallization, or a result of changes to the stability field at the contact between cycles. Research is ongoing and further interpretations will be made using data collected via SEM and microprobe to constrain petrogenesis of cyclical units and metallogenesis of Cr-spinel in the southern BEI. References Weiblen, P.W., 1965. A Funnel-Shaped, Gabbro Troctolite Intrusion in the Duluth Complex, Lake County, Minnesota: PH.D. Thesis University of Minnesota, Minneapolis, MN. Green, J.C., Phinney, W.C., and Weiblen, P.W., 1966, Geologic map of Gabbro Lake quadrangle, Lake County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series, map M-2, scale 1: 31,680. 16 �Passive-aggressive geophysics: An update on using the horizontal-to-vertical spectral ratio (HVSR) passive seismic method for determining glacial deposit thickness in Minnesota Val W. Chandler and Richard S. Lively Minnesota Geological Survey, 2642 University Ave., St. Paul, MN 55114 chand004@umn.edu Considerable progress has been made on evaluating the horizontal-to-vertical-spectral ratio method (HVSR or sometimes H/V) for determining the thickness of glacial deposits in Minnesota. The HVSR method is used to estimate the primary resonant frequency (shear wave) of unconsolidated overburden. At this frequency, the horizontal components of oscillation are amplified relative to the vertical component. By dividing the averaged horizontal spectra by the vertical spectrum, a HVSR spectrum is produced, ideally with a pronounced peak at the primary resonant frequency. If the acoustic impedance (density*seismic velocity) at the overburdenbedrock contact differs by a factor of at least 2, the thickness (Z) of the unconsolidated materials can be estimated by the relationship: Z=af0b where f0 is the estimated primary resonant frequency, and a and b are parameters that are determined empirically for a given region from control points that have a range of known bedrock depths. The HVSR method has been used in several Minnesota Geological Survey (MGS) projects, with data acquired at over 675 sites state-wide. Roughly 40% of these were located at control points whose distribution allows the HVSR method to be evaluated in three different geologic environments; the seven-county Twin Cities Metropolitan area, south-central Minnesota, and the floodplains and terraces along major streams, including the Mississippi, Minnesota, and St. Croix Rivers. The 7-county Twin Cities Metropolitan area is particularly favorable for the HVSR method. Paleozoic strata comprising the bedrock surface are generally rigid, with little soft material, such as saprolith or Cretaceous sediment present and the glacial sequence is fairly simple, consisting chiefly of late Wisconsinan deposits. Analysis of 41 selected control points exhibiting single, high-amplitude (>3.5) HVSR peaks produced a fitted curve in the form of the above equation with an R value (Correlation Coefficient) = 0.968. Depth estimates at the 41control points have an average error of 13% and 90% are within +/-25% of the known bedrock depths. These figures serve as a proxy for expected error using HVSR peaks of similar quality in areas that lack well control. For low-quality peaks the error can be significantly greater; applying the same fitted curve to 20 control stations that exhibited flat-topped peaks, interfering multiple peaks, or low-amplitude (<3.5) peaks produced an average error of 26% and only 60% of the estimates were within +/-25% of known bedrock depth. Regardless of error for individual depth estimates however, the HVSR method is still very useful as a simple mapping tool, for example, it was highly effective for mapping the trace of a buried bedrock ravine in central Washington County which had been largely missed by drillhole and seismic data. The terrain of south-central Minnesota presents some sobering challenges to the HVSR method. Glacial deposits, which include both Wisconsinan and Pre-Wisconsinan materials, tend to be thick, complex, and have dense, over-consolidated tills in the lower parts of the section. These deposits may also overlie a bedrock surface composed of soft materials, such as saprolith or Cretaceous sediment. Hence, impedance contrasts that can produce a HVSR peak and depth information may actually occur within the glacial sequence and may miss the bedrock surface 17 �altogether. In some instances, the HVSR data may also produce a trough in addition to a peak (at 2 time the frequency of the peak) and occasionally the trough provides a better indication of bedrock depth than the peak, especially for depths of >100 meters. In other instances no useful HVSR signal can be confidently extracted from any of the data. The effect of soft bedrock was evaluated by comparing control points reporting either saprolith or Cretaceous strata against an idealized power-curve, based on 27 control points where these soft materials appeared to be either thin or absent. Correlation for this curve has an R value = 0.953. Use of this curve tends to overestimate the depth to saprolitic bedrock by as much as 30-50%, implying that the HVSR signal is picking up an interface below the bedrock surface, likely somewhere in the transition to fresh bedrock. For areas with a Cretaceous bedrock surface, the overestimate can be much worse, and it appears that the HVSR signal is actually responding to the Pre-Cretaceous surface rather than the Cretaceous bedrock surface. Overall the HVSR method must be applied cautiously in south-central Minnesota, and in other areas that are likely to have similar conditions. Nonetheless, useful information can still be derived at many locations. In contrast to south-central Minnesota the floodplains and terraces along major streams in eastern Minnesota provide an almost ideal environment for the HVSR method. The channels in this part of the state served as sluiceways for melt-water during the closing stages of the Wisconsinan glaciation, and the valley bottoms have been largely swept clean of soft materials such as Cretaceous strata, saprolith, or earlier glacial deposits, and replaced by poorly consolidated outwash, fluvial and lacustrine deposits. A strong acoustic impedance contrast is expected at the bedrock surface and the observed HVSR peaks, which usually have single, very high-amplitude (>5) signatures, are consistent with this. Using 37 control stations a curve was established with an R correlation value = 0.951. Curve-based depth estimates at the control points have an average percentage error of 20% and 82% are within +/-25% of observed bedrock depths. The greatest errors generally occur at depths less than 20 meters, and may reflect sloping or uneven bedrock surfaces near valley side-walls. In any case, the relationship derived here should be useful for rough estimates of bedrock depth along the bottoms and terraces of the major river valleys in eastern Minnesota and adjacent areas Although, the HVSR method does not quite match conventional seismic profiling for accuracy and derivative information, the advantages of passive seismic for determining depth to bedrock include rapid data collection (usually 16 minutes recording time), much lower equipment and field costs, relative ease of data analysis and large number of samples that can be collected within a given area. In addition, rough estimates of depth with a denser array of points are quite adequate for many kinds of geological applications. Finally, the HVSR method can be readily applied in areas of significant cultural noise, where conventional seismic profiling is difficult. ACKNOWLEDGEMENTS This study was supported by the Minnesota Legislature through the County Geologic Atlas Program of the Legislative and Citizens Commission on Minnesota Resources. Additional support was through The STATEMAP Program of the U. S. Geological Survey and the State Special Appropriation of the Minnesota Geological Survey. 18 �Geochemistry of the Logan Igneous Suite and implications for the magmatic evolution of the northern part of the Midcontinent Rift CUNDARI, Robert1,2, HOLLINGS, Peter2, and SMYK, Mark1 1 Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, 435 James St. S., Suite B002, Thunder Bay, ON, P7E 6S7 Canada 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Recent field work and interpretation of geochemistry has focussed on defining the emplacement sequence of Midcontinent Rift (MCR)-related units and evaluating their geochemical sources and contamination characteristics so as to better understand the magmatic evolution of the Logan Igneous Suite (Cundari, 2012). New geochemical discriminators applied to the Nipigon sills (e.g. Cundari et al., 2012) have been applied to a larger data set in order to identify previously unrecognized relationships between units. The plot of Nb/Ybpm versus Th/Ybpm (Fig. 1) illustrates a mixing model between mantle-derived melts and crustal sources. Separate trends with distinct ranges of Nb/Ybpm scattering away from the field of mantlederived melt compositions towards higher Th/Ybpm suggests the presence of separate, geochemically distinct magma source types within the Logan Igneous Suite which have undergone separate crustal contamination histories. The plot of Th/Ybpm versus Nb/Thpm (Fig. 2) also delineates distinct primary source compositions, as separate trends curve away from the field of mantle-derived melt compositions towards lower Nb/Thpm and higher Th/Ybpm. The position of the ultramafic units (i.e. Hele, Disraeli, Seagull, Kitto), which plot within or towards the field of mantle-derived melts in Figures 1 and 2, support previous work suggesting the ultramafic units are the best indicators for the primary magma formed from a deep-seated mantle plume (Nicholson et al., 1997; Hollings et al., 2007). Based on these criteria, four distinct distributions are noted (Figs. 1 and 2): the Nipigon sill trend; the Jackfish, McIntyre, Inspiration and Logan trend (JMIL); the Dyke trend (including the Pigeon River dykes, Cloud River dykes, Mount Mollie dyke and the Crystal Lake gabbro); and the Devon volcanics and the Riverdale sill trend (DR). Field work in the Logan Basin has delineated the following timing sequence between units, from oldest to youngest: Riverdale sill; Devon volcanics; Logan sills; Pigeon River dykes; Cloud River dykes; Mount Mollie dyke and Crystal Lake gabbro. In light of these relative ages, the mantle source characteristics are shown to become more depleted in incompatible elements as rift development progresses, each trend having been derived from a geochemically distinct mantle source. The mantle source of the Devon volcanic and the Riverdale sill magmas is the most enriched. The mantle source of the Logan sills is relatively more depleted and the mantle source of all dyke sets is the most depleted of all units in the Logan Basin. Furthermore, although all three dyke sets display similar source characteristics, younger dykes display progressively stronger crustal contamination signatures (i.e., higher Th/Ybpm and more negative ƐNd(t=1100Ma)) consistent with the magma having spent progressively more time in the magma chamber. A similar incompatible element depletion trend in mantle sources, akin to that in the Logan Basin units, is also apparent in the Nipigon Embayment. It is widely accepted through geochronological data and field relationships that the Nipigon sills post-date the ultramafic units of the Nipigon Embayment (e.g. Heaman et al., 2007). Geochemical source characteristics for the Nipigon sills and the ultramafic units show that the ultramafic units were derived from a more enriched source when compared to a more depleted source for the Nipigon sills (Figs. 1 and 2). Logan sills of the Logan Basin and ultramafic units of the Nipigon Embayment all show similar source characteristics (Figs. 1 and 2), suggesting an overlapping magmatic history. From the geochemical data presented here, in conjunction with known emplacement sequences, it is proposed that magma which produced units of the Logan Igneous Suite were derived from 19 �geochemically distinct mantle sources showing a progressive depletion in incompatible elements as MCR development progressed. Figure 1: Diagram showing variations in Nb/Ybpm and Th/Ybpm ratios for Midcontinent Rift-related mafic rocks. Normalizing values from Sun and McDonough (1989). Figure 2: Diagram showing variations in Th/Ybpm and Nb/Thpm ratios for Midcontinent Rift-related mafic rocks. Normalizing values from Sun and McDonough (1989). References Cundari, R.M., 2012. Geology and geochemistry of Midcontinent rift-related igneous rocks. Unpublished M.Sc. thesis, Lakehead University, Thunder Bay, ON, 142 p. Cundari, R.M., Hollings, P.N. and Smyk, M.C., 2012. Petrogenesis and crustal contamination of the Nipigon sills: a geochemical and spatial re-evaluation; 58th Institute on Lake Superior Geology, Annual Meeting, Thunder Bay, ON, May 16-20, 2012, Proceedings Volume 58, Part 1, p.22-23. Heaman, L.M., Easton, M., Hart, T.R., Hollings, P., Macdonald, C.A. and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario; Canadian Journal of Earth Sciences 44: 1055-1086. Hollings, P., Hart, T., Richardson, A. and MacDonald, C.A., 2007. Geochemistry of the mid-Proterozoic intrusive rocks of the Nipigon Embayment, northwestern Ontario; Canadian Journal of Earth Sciences 44: 1087-1110. Nicholson, S.W., Shirey, S., Schulz, K., and Green, J., 1997. Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development; Canadian Journal of Earth Sciences 34: 504-520. Sun, S. and McDonough, W. F. 1989. Chemical and isotopic systematics of oceanic basalts: Implications for mantle composition and processes. in A.D. Saunders and M.J. Norry (eds.) Magmatism in the Ocean Basins ; Spec. Publ. Vol. Geol. Soc. Lond. , No. 42, pp. 313-345. 20 �The Geochemistry and Mineralogy of the Sulfides within the Ni-Cu-PGE Shakespeare Deposit, Ontario DASTI, Ian R. and KISSIN, Stephen A. Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The 2217 Ma Shakespeare deposit, located 70 km west of Sudbury, Ontario, is part of the 2.2 Ga Nipissing Gabbro suite and is hosted within the metasediments of the 2.45-2.2 Ga Huronian Supergroup. The past producing Ni-Cu -PGE Shakespeare Mine, which has probable reserves of 11.8 Mt grading 0.87 g/t PGE+Au, 0.33% Ni, and 0.35% Cu, is hosted in the Shakespeare deposit (Prophecy Platinum, 2013). The majority of the mineral reserves are in the form of disseminated sulfides located within the Shakespeare melagabbro. Other significant sulfide mineralization is in the form of 2-5 cm heavily disseminated blebs to net-textured and locally semi-massive sulfides immediately above the disseminated sulfides. The aforementioned sulfides are collectively known as “interconnected sulfides” and straddle the contact between the Shakespeare quartz gabbro and melagabbro. The Shakespeare deposit is situated below the Mississagi Quartzite and above the unmineralized Nipissing Gabbro. Sulfide mineralogy within the disseminated and heavily disseminated portions of the deposit is largely dominated by pyrrhotite, pentlandite, and chalcopyrite. A detailed SEM study has also led to the discovery of rarer minerals, such as molybdenite, several compounds containing tellurium and bismuth, a rhenium sulfide, argentopentlandite, and gersdorfite . Agentopentlandite seemed to show an affinity for chalcopyrite, forming discreet, euhedral grains along the edges of chalcopyrite crystals. Interestingly, the bismuth tellurides also rarely contain abundant silver, up to 20 percent, and antimony. The SEM study also showed that gersdorfite was rarely encountered in its pure form but rather was usually in solid solution with cobaltite and arsenopyrite . Platinum and palladium minerals were not observed, suggesting they could be present in concentrations lower than 2-3% or as nuggets that were not encountered. However, a rhenium sulfide nugget was encountered on a few occasions, and because Re, Pt, and Pd are expected to act similarly under the geochemical conditions present, it is likely that Pt and Pd occur as minerals in the form of nuggets. Previous workers (Sproule et al., 2005) have suggested a contamination event lead to sulfur saturation in the magmas of the Shakespeare deposit, but those works do not comment as to whether the contamination event contributed significant sulfur to the magmas . Nine samples were taken from various stratigraphic depths in the deposit for sulfur isotope analysis and returned δ34S values from 0.01 ‰ to 2.38 ‰, averaging 1.14 ‰ (table 1). Sulfur data obtained by LECO analysis and preliminary selenium data obtained by ICP-MS provide S:Se ratios between 1245 and 3271, averaging 1810 (table 2). The data strongly suggest that the sulfur source for the Shakespeare deposit is dominantly magmatic, with little to no input from crustal sources. Additionally, the rocks of the Huronian Supergroup ((Mississagi through Matinenda Formations) are devoid of a sulfur source with which to contribute sulfur to the magmas that led to the formation of the Shakespeare deposit. 21 �Table 1. Sulfur isotope data Sample* δ34S ‰ 119-304.5 2.38 119-324 1.83 119-346 1.05 119-360 1 119-378 0.5 122.396.2 1.43 122-417 0.01 122-437.7 1.38 122-464.7 0.72 mean 1.14 First three digits = drill-hole, Next digits = depth in metres Table 2. Sulfur and selenium data and S/Se ratios ratios for mineralized samples Sample Se (ppm) S % S/Se ratio (ICP-MS) LECO 119-340 3.3 0.53 1607 119-350 8.4 1.44 1714 119-366 10.2 1.27 1245 119-378.6 9.4 1.35 1436 122-435.7 5.2 0.88 1692 122-440.7 9.2 3.01 3272 122-442.7 11.1 2.37 2135 122-444.7 7.5 1.32 1760 122.451.7 14.9 2.92 1960 122-454.7 6.9 1.05 1522 122-455.7 11.3 1.77 1566 mean 1810 References Lightfoot, P.C., Conrod, D., Naldrett, A.J. and Evensen, N.M., 1987. Petrologic, chemical, isotopic, and economic-potential studies of the Nipissing Diabase, Grant 230 in Milne, V.G. (ed.) Geoscience Research Grant Program, Summary of Research 1986-1987, Ontario Geological Survey, p. 4-26. Prophecy Platinum (2013) Prophecyplat.com, Accessed March 22, 2013. Sproule, R.A., Sutcliffe, R., Tracanelli, H., Lesher, C.M., 2008. Palaeoproterozoic Ni-Cu-PGE mineralization in the Shakespeare Intrusion: A new style of Nipissing gabbro-hosted mineralization, Transactions of the Institution of Mining and Metallurgy B. Applied Earth Science, 116: 188-200. 22 �High-resolution, multi-method geophysical imaging of a portion of the Northeast Iowa Intrusive Complex DRENTH, Benjamin1, ANDERSON, Raymond2, CHANDLER, Val3, CANNON, William4, SCHULZ, Klaus4, FEINBERG, Joshua M.5, BEDROSIAN, Paul 1, and KASS, M. Andy1 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 2 Iowa Geological and Water Survey, 109 Trowbridge Hall, Iowa City, IA, 52242 3 Minnesota Geological Survey, 2642 University Avenue W., St. Paul, MN, 55114-1032 4 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA, 20192-6320 5 Dept. Earth Sciences, Univ. Minnesota, 310 Pillsbury Dr. SE, Minneapolis, MN, 55455-0219 Numerous large amplitude regional aeromagnetic anomalies and ground gravity highs over northeast Iowa and southeast Minnesota (Fig. 1) suggest the presence of a buried intrusive complex made up of mafic/ultramafic rocks. This complex is known as the northeast Iowa Intrusive Complex (NE IIC) (e.g., Pals and Anderson, 2011). The NE IIC lies along the eastern margin of the Midcontinent Rift System (MRS) and occupies a minimum estimated area of 17,000 square kilometers, making it comparable in size to the Duluth Complex. Country rocks are thought to be accreted island arc terranes of the Paleoproterozoic Yavapai Province (1.7-1.8 Ga), implying at least a somewhat younger age for the NE IIC. While not yet directly dated, these considerations suggest that a Keweenawan (MRS) age for some or all of the NE IIC is possible and imply significant potential for undiscovered Ni-Cu-PGE deposits. Alternatively, the NE IIC could include Mesoproterozoic (~1450 Ma) gabbro-anorthosite-rapakivi granite intrusions like the Wolf River Batholith in Wisconsin. Only four boreholes are known to reach the complex, which is covered by 200-500 meters of Phanerozoic sedimentary rocks and sediments. Geophysical methods are thus critical to developing a better understanding of the fundamental nature and resource potential of the NE IIC. A high-resolution, multi-method geophysical mapping program was initiated in 2012 as a collaborative effort between the U.S. Geological Survey Mineral Resources Program, the Iowa Geological and Water Survey, and the Minnesota Geological Survey. An initial 3,333 line kilometer airborne data collection campaign in the region of Decorah, Iowa, included magnetic, gravity gradient (AGG), and time-domain electromagnetic (TDEM) data along flight lines spaced 400 m apart. At the time of this writing, only preliminary versions of these data are available for inspection. The preliminary data show numerous magnetic anomalies that are paired with AGG highs, indicating widespread strongly magnetized and dense rocks of likely mafic/ultramafic composition. In the Decorah region, a prominent horseshoe-shaped, 15 kilometer diameter magnetic- and gravity-field high is correlated with the occurrence of basement rocks that have been described as unaltered gabbro and troctolite, suggesting a ring-shaped anomaly source. A Yavapai age layered(?) metagabbro pluton (Van Schmus et al., 2007) is suspected to produce complex magnetic and gravity anomalies with different forms than the other basement rocks nearby. The TDEM data appear to image crystalline basement rocks in select locations, and it is expected that the data will ultimately provide important constraints on depths to the NE IIC and other Precambrian basement rocks. References Pals, D.W., and Anderson, R.R., 2011, Reassembling Iowa: spatial and temporal evaluation of the mineral potential of the Iowa segment of the Micontinent Rift and related plutons: Geological Society of America Abstracts with Programs, v. 43, no. 5 p. 396. Van Schmus, W.R., Schneider, D.A., Holm, D.K., Dodson, S., and Nelson, B.K., 2007, New insights into the southern margin of the Archean-Proterozoic boundary in the north-central United States based on U-Pb, Sm-Nd, and Ar-Ar geochronology: Precambrian Research, v. 157, p. 80-105. 23 �Figure 1: Regional aeromagnetic total-field (grayscale) and complete Bouguer gravity anomalies for the NE IIC region. Gravity contour (black lines) interval 10 milligals. White dots indicate wells that reach Precambrian rocks. Area of recent survey in region of Decorah, IA, shown as gray polygon. 24 �Application of LiDAR to resolving regional tectonic and glacial fabrics in glaciated terrane: An example from an Archean granite-greenstone belt in NE Minnesota DYESS, Jonathan and HANSEN, Vicki, Department of Geological Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth MN 55812 Regional tectonic fabrics define the broad structural architecture of an area and commonly have an associated topographic expression that may be identified via remote sensing (Chardon et al., 2002, 2008, 2009; Bedard et al., 2003). Glaciation commonly forms a geomorphologic fabric in the form of drumlins, flutes, valleys, eskers, and crag and tail features (Smith et al., 2006; Sharp, 1953). Overprinting of glacial fabrics, vegetation, and sediment cover commonly obscure topographic lineaments and hinder identification of lineament fabrics via high-resolution aerial photography and low to moderate resolution satellite imagery. Airborne LiDAR (light detection and ranging) systems provide high-resolution altimetry in vegetated areas (e.g., Haugerud et al., 2003). Although LiDAR altimetry is a useful tool for mapping small–scale geologic structures (e.g., Pavlis and Bruhn, 2011) and glacial geomorphologic features (e.g., Smith et al., 2006), previous studies do not constrain whether LiDAR altimetry may be used to differentiate between tectonic fabrics and overprinted glacial fabrics within the same area. In this study, we examine an Archean granite-greenstone terrane in NE Minnesota to illustrate the application of high-resolution LiDAR altimetry to mapping regional tectonic fabrics in glacially striated, forested areas. More specifically, we describe how to distinguish between tectonic and glacial fabrics and the effect of glaciation on the overall topographic expression of the tectonic fabric. A 1-m posted LiDAR derived bare-earth DEM (digital elevation model) collected as part of the Minnesota Elevation Mapping Project and shaded relief images constructed from the bare-earth DEM comprise the raw data for this study. Data processing and lineament mapping were done using ESRI ArcGIS software. Evaluation of the bare-earth DEM and shaded relief images revealed that shaded relied images provide the most potential for lineament mapping. In order to maximize the chance of mapping lineaments at all orientations, we constructed shaded relief images with a sun elevation of 45˚ and varying sun azimuth at 45˚ intervals. Using ESRI ArcScene, we draped shaded relief images over the bare-earth DEM to create a 3D perspective view of the field area and to visualize the topographic surface. Mapping revealed two suites of lineaments. Suite A consists of relatively short (1-2 km), discrete lineaments with a unimodal orientation distribution and a mean trend of 045. We recognize multiple striated deposits of sediment across the study area. Sediment deposits contain suite A lineaments only. Suite B consists of lineaments ranging in length from 1-30 km. Suite B lineaments are more continuous than suite A and have a quasi-bimodal orientation distribution. Suite B lineaments have a mean trend of 065 across the study area with local areas trending 090. In areas where suite A parallels suite B only one pervasive lineament set is visible. Where suite A and suite B are at high angles to one another, suite A lineaments are shorter and pervasive while suite B lineaments longer and spaced. We interpret suite A as a geomorphological fabric related to glaciation and suite B as the regional tectonic fabric. Field measurements of foliation trajectory (Goodman, 2008; Erickson, 2008; Dyess, unpublished field data) are largely consistent with suite B lineaments across the study area. Although not all suite B lineaments correlate to mapped structures, our analysis demonstrates that high-resolution LiDAR altimetry may be used to map regional tectonic fabrics in glaciated terrane. 25 �References Bedard, J.H., Brouillette, P., Madore, L., Berclaz, A., 2003. Archean cratonization and deformation in the northern Superior Province, Canada: an evaluation of plate tectonic versus vertical tectonic models. Precambrian Research, 127, 61-87. Chardon, D., Peucat, J., Jayananda, M., Choukroune, P., Fanning, C. M., 2002. Archean granitegreenstone tectonics at Kolar (South India): Interplay of diapirism and bulk inhomogeneous contraction during juvenile magmatic accretion. Tectonics, 21, no. 3, 1016. Chardon, D., Jayananda, M., Chetty, T.R.K., 2008. Precambrian continental strain and shear zone patterns: South Indian case. Journal of Geophysical Research, 113. Chardon, D., Gapais, D., Cagnard, F., 2009. Flow of ultra-hot orogens: A view from the Precambrian, clues for the Phanerozoic. Tectonophysics, 477, 105-118. Erikson, E., 2008. Structural and kinematic analysis of the Shagawa Lake shear zone, Superior Province, northeastern Minnesota. M.S. Thesis, University of Minnesota Duluth, MN. Goodman, S., 2008. Structural and Kinematic Analysis of the Kawishiwi Shear Zone, Superior Province. M.S. Thesis, University of Minnesota Duluth, MN. Haugerud, R.A., Harding, D.J., Johnson, S.Y., Harless, J.L., Weaver, C.S., Sherrod, B.L., 2003. High-resolution topography of Puget Lowland, Washington-A Bonanza for Earth Science: GSA Today, 13, no. 6, 4-10. Pavlis, T.L. and Bruhn, R.L., 2011. Application of LIDAR to resolving bedrock structure in areas of poor exposure: An example from the STEEP study area, soutern Alaska. GSA Bulletin, 123, 206-217. Sharp, R.P., 1953. Glacial Features of Cook County, Minnesota, American Journal of Science, 251, 855-883. Smith, M.J., Rose, J., Booth, S., 2006. Geomorphological mapping of glacial landforms from remotely sensed data: An evaluation of the principal data sources and an assessment of their quality. Geomorphology, 76, 148-165. 26 �Structural and Kinematic Analysis of the Shagawa Lake Shear Zone and Snowbank Lake Stock, Superior Province, NE Minnesota DYESS, Jonathan and HANSEN, Vicki, Department of Geological Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth MN 55812 The Archean (3.85-2.5 Ga) Superior Province, to a first approximation, consists of a series of east-west trending subprovinces of supracrustal rocks (greenstone belts) and granitoid rocks interpreted as a series of microcontinents, remnant arcs, oceanic terranes, and accretionary prisms that accreted to a growing continental block during dextral transpression driven by NW-directed oblique subduction (e.g., Percival et al., 2007, and references therein). Transpression platetectonics would predict the formation of faults/shear zones that record significant unidirectional strike-slip displacement at or near terrane boundaries (Sleep, 1992). The Vermillion District, southern Superior Province, is a Neoarchean (2.8-2.5 Ga) granite-greenstone terrane dominated by a series of NE-striking subvertical shear zones with ovoid to circular granitic bodies scattered throughout. Vermillion District shear zones have been interpreted as primarily dextral strike-slip shear zones formed during terrane assembly driven by NW oblique subduction (Hudleston et al., 1988; Bauer and Bidwell, 1990; Schultz-Ela and Hudleston, 1991). Others interpret Vermillion District shear zones as zones of dominantly oblique to dip-slip shear possibly formed during greenstone sagduction between rising granitoid diapirs (Erickson, 2008, 2010; Wolf, 2006; Goodman, 2008; Karberg, 2009). Differing interpretations of Vermillion District shear zones invoke different assumptions about displacement direction during non-coaxial shear. Displacement direction (flow direction) is genetically tied to foliation and elongation lineation orientation. Within the Shagawa Lake shear zone of NE Minnesota, displacement direction remains undetermined. Adjacent to the Shagawa Lake shear zone is the Snowbank Lake stock, a 30 km2 composite stock dominated by syenite and granodiorite (Sanders, 1929). Due to the proximity of the Shagawa Lake shear zone to Quetico and/or Wawa subprovince boundary, tectonic fabrics within the Shagawa Lake shear zone have implications for crustal assembly of the southern Superior Province. If the Shagawa Lake shear zone records significant unidirectional strike-slip displacement, then supported plate-tectonic models for Vermillion District formation will be further constrained. If the Shagawa Lake shear zone does not record significant unidirectional strike-slip displacement, then existing plate-tectonic and structural models of terrane amalgamation along the Vermillion District require reevaluation. We conducted a structural and kinematic analysis of the Shagawa Lake shear zone in three phases: 1) analysis of regional tectonic fabrics through Light Detection and Ranging altimetry data; 2) structural analysis of outcrop-scale structures through detailed field mapping; and 3) analysis of shear-sense indicators through kinematic analysis of thinsections. The Shagawa Lake shear zone contains a regional subvertical metamorphic foliation with an average strike of 065 but varies locally from 065 to 100. Near the Snowbank Lake stock foliation deviates from the regional trend and turns roughly parallel to the stock boundary. We recognize two types of elongation lineation within the Shagawa Lake shear zone. These include ridge-ingroove striations on C-foliation surfaces (Lc) as well as stretching lineations on S-surfaces (Ls) (Lin and Williams, 1992; Lin et al., 2007). The Shagawa Lake shear zone hosts unimodal Lc and Ls ranging from dominantly steep plunge to moderate plunge. Asymmetric fabrics occur in foliation-perpendicular, lineation-parallel planes and symmetric fabrics occur in foliationperpendicular, lineation-perpendicular planes, which is consistent with non-coaxial shear with lineation forming parallel to shearing. Therefore, in high pitch domains, displacement was vertical to oblique, and microstructures record both north-side-up and south-side-up displacement in different samples. Samples with oblique lineation record an apparent dextral strike-slip shearsense despite varying lineation orientation. Our data indicate the Shagawa Lake shear zone 27 �experienced both N-side-up and S-side-up dip- to oblique-slip with relatively minor apparent dextral strike-slip and does not record significant unidirectional strike-slip displacement. References Bauer, R. L and Bidwell, M. E., 1990. Contrasts in the response to dextral transpression across the Quetico-Wawa subprovince boundary in northeastern Minnesota. Canadian Journal of Earth Sciences, 27, 1521-1535. Erikson, E., 2008. Structural and kinematic analysis of the Shagawa Lake shear zone, Superior Province, northeastern Minnesota. M.S. Thesis, University of Minnesota Duluth, MN. Erickson, E., 2010. Structural and kinematic analysis of the Shagawa Lake shear zone, Superior Province, northern Minnesota: implications for the role of vertical versus horizontal tectonics in the Archean. Canadian Journal of Earth Sciences, 47, 1463-1479. Goodman, S., 2008. Structural and Kinematic Analysis of the Kawishiwi Shear Zone, Superior Province. M.S. Thesis, University of Minnesota Duluth, MN. Hudleston, P.J., Schultz-Ela, D., Southwick, D. L., 1988. Transpression in an Archean greenstone belt, northern Minnesota. Canadian Journal of Earth Sciences, vol 25, 1060-1068. Karberg, S M., 2009. Structural and Kinematic Analysis of the Mud Creek Shear Zone, Northeastern Minnesota. M.S. Thesis, University of Minnesota Duluth, MN. Lin, S., Williams, P.F., 1992. The origin of ridge-in-groove slickenside striae and associated steps in an S-C mylonite. Journal of Structural Geology 14, 315e321. Lin, S., Jiang, D., Williams, P., 2007. Importance of differentiating ductile slickenside striations from stretching lineations and variation of shear direction across a high-strain zone. Journal of Structural Geology, 29, 850-862. Percival, J.A., 2007, Geology and metallogeny of the Superior Province, Canada, in Goodfellow,W.D., ed.,Mineral Deposits of Canada:ASynthesis ofMajor Deposit-Types, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods: Geological Association of Canada, Mineral Deposits Division, Special Publication No. 5, p. 903-928. Sanders C.W., 1929, A composite stock at Snowbank Lake in northeastern Minnesota. Journal of Geology, 37, 135-149. Schultz-Ela, D.D., Hudelston, P.J., 1991. Strain in an Archean greenstone belt of Minnesota. Tectonophysics, 190, 223-268. Sleep, N., 1992. Archean plate tectonics: what can be learned from continental geology?. Canadian Journal of Earth Sciences, 29, 2066-2071. Wolf, D. E., 2006. The Burntside Lake and Shagawa/Knife Lake shear zones: Deformation kinematics, geochemistry and geochronology; Wawa Subprovince, Ontario, Canada. Masters Thesis, Washington State University. 28 �Bedrock Geologic Map of the Putnam Lake Area, St. Louis County, NE Minnesota FEHRS, Ellen1, KENNY, Edward1, KUCHMA, John1, SAUER, Sarah1, SYLVESTER, William1, and HUDAK, George1,2 1 Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Duluth, MN, 55811 2 Minerals Division, Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Duluth, MN 55811 Each year, students from the Precambrian Research Center (PRC) geology field camp complete “capstone” projects that encompass approximately one week of detailed field mapping followed by one week of mapmaking and map publishing. During the fifth and sixth weeks of the 2012 field camp, five PRC field camp students, under the direction of PRC Associate Director George Hudak, mapped Neoarchean rocks on the southwest side of the Tower-Soudan Anticline in the vicinity of Putnam Lake (Fehrs et al., 2012). This capstone mapping project sought to: 1) identify the lithologies and determine the detailed stratigraphy within the Neoarchean supracrustal strata in this area; 2) define and characterize the nature of the contacts between various units of the Neoarchean supracrustal strata and intrusive rocks; 3) obtain a better understanding of geological structures and their orientations within the area; 4) produce a detailed geological map in an area on the south side of the Tower-Soudan anticline, which has previously only been mapped at a regional scale; and 5) test the utility of LiDAR for mapping in heavily forested greenstone belt terranes. Prior to mapping, detailed 1:5000 scale laminated field mapping sheets were produced. One side of each field mapping sheet consisted of a georeferenced air photo, and the other side of the mapping sheet consisted of a corresponding georeferenced topographic map using recently released LiDAR- based elevation data available from the Minnesota Geospatial Information Office. Mapping was completed by means of numerous traverses through the bush, as well as traverses along the lakeshore of Putnam Lake. Following each day of field mapping, students and faculty transferred their field data to a master map, enabling the detailed geology of the region to be established by the middle of the fifth week of field camp. During the sixth week of field camp, students produced a digital version of the field map utilizing a variety of software (ArcView, AutoCad, Surfer, Adobe Illustrator). Previous mapping (Peterson and Jirsa, 1999; Peterson, 2005) established that the Neoarchean supracrustal rocks in this area comprise a NW/SE striking, SW facing homoclinal sequence that makes up the southern limb of the Tower-Soudan Anticline. This regional mapping indicated that the area consisted of rocks comprising the Lower Member of the Ely Greenstone Formation (Armstrong Lake / Central Basalt mafic-intermediate lava flows, dacite lava flows, and local Algoma-type oxide facies iron formations), the Soudan Member of the Ely Greenstone Formation (Algoma-type banded oxide-facies iron formation), and the Upper Member of the Ely Greenstone Formation (South Limb Basalts, comprising massive- to pillowed-facies basalt lava flows), as well as synvolcanic hypabyssal diabase/gabbro sills and dikes, and post-volcanic quartz-feldspar porphyry, hornblende-feldspar porphyry, and lamprophyre intrusions. The contact between the Soudan Member and Upper Member of the Ely Greenstone Formation was interpreted as a minor shear zone associated with regional D2 deformation (Peterson, 2005) In light of lithological characteristics obtained from mapping in the vicinity of Soudan Underground Mine State Park and Lake Vermilion State Park (Peterson and Patelke, 2003; Radakovich et al., 2010; Heim et al., 2011), we now believe that supracrustal rocks in this part of the Tower-Soudan anticline comprise part of the Soudan Member of the Ely Greenstone Formation. In this area, the Soudan Member is composed of: a) medium to light gray laminated to very thinly bedded chert; b) light gray to red to dark gray laminated to very thinly bedded interlayered horizons of chert, jasper, and magnetite-rich Algoma-type oxide facies iron formation; c) interbedded light green to medium green, sparsely amygdaloidal, sparsely plagioclase-phyric massive sheet flow- and pillowed-facies basalt lava flows locally interbedded with 0.10-1.0 meter thick laminated oxide facies iron formation horizons; d) light 29 �green-gray to medium green, sparsely amygdaloidal, sparsely plagioclase-phyric massive sheet flowfacies basalt lava flows; e) light green-gray to medium green, sparsely amygdaloidal, sparsely plagioclase-phyric pillow-facies basalt lava flows; and f) green-gray, locally amygdaloidal, plagioclasephyric dacite lava flows that are locally interbedded with chert and argillite horizons. Unfortunately, outcrop areal distribution (about 1% of the field area) was not sufficient to identify, or map out, individual lava flows with confidence. Locally, synvolcanic and post-volcanic sills and dikes intrude the supracrustal assemblage in the area. Synvolcanic diabase/gabbro sills and dikes are medium green-gray to dark green, fine- to medium grained, subophitic and locally plagioclase-phyric. Post-volcanic intrusions include: a) light gray to pinkish-gray hornblende- and plagioclase-phyric dacite sills and dikes; b) light gray to light pinkish-gray quartz- and plagioclase-phyric rhyodacite sills and dikes; and c) light to medium-gray, fine- to mediumgrained biotite-phyric lamprophyre dikes. Based on our mapping, the following sequence of geological events is believed to have formed the stratigraphic succession in this part of the Vermilion District: 1) early mafic volcanism dominated by massive sheet flow-facies lava flows; 2) deposition of several horizons (up to 50 meters thick) of Algomatype oxide facies iron formations during breaks in mafic volcanism; 3) resumption of extrusive mafic volcanism forming both massive sheet flow-facies and pillowed-facies basalt lava flows with local extrusion of dacite lava flows and minor, intermittent hydrothermal activity to form thin (less than 10 meters thick) Algoma-type banded iron formations; 4) extrusive mafic volcanism forming both massive sheet flow-facies and pillowed-facies basalt lava flows followed by local intrusion of diabase/gabbro sills and dikes; 5) intrusion of sills and dikes of both hornblende- and plagioclase-phyric dacite and quartzand plagioclase-phyric rhyodacite; 6) development of the Tower-Soudan anticline; 7) D2 deformation forming a penetrative east-west-trending foliation; and 8) intrusion of biotite-phyric lamprophyre intrusions. Based on the similarity of the basalt lava flows mapped in this area to the Soudan Basalts mapped in other parts of the Vermilion District, it appears that the Upper Member of the Ely Greenstone Formation may be absent on the southern limb of the Tower-Soudan anticline. More mapping, petrographic studies, and lithogeochemical studies will be necessary to further evaluate this preliminary interpretation. As well, we found that LiDAR-based topographic maps provided excellent base maps, and were extremely useful for identifying small topographic features in heavily forested areas that were later found to be outcrops that had never been previously mapped. References Fehrs, E., Kenny, E., Kuchma, J., Sauer, S., Sylvester, W., and Hudak, G., 2012, Bedrock Geologic Map of the Putnam Lake Area, St. Louis County, Northeastern Minnesota: Precambrian Research Center Map Series, PRC/MAP-2012/02, 1:5000 scale. Heim, N., Scott, H., Kilduff, R., Rahtz, C., Vial, A., Young, S., Mahr, C, and Hudak, G., 2011, Preliminary Bedrock Geology Map of the Eastern Part of Lake Vermilion State Park, St. Louis County, NE Minnesota: Precambrian Research Center Map Series, PRC/MAP-2011/01, 1:5000 scale. Peterson, D. M., 2005, Bedrock geologic and volcanogenic massive sulfide deposit mineral potential map of the Lower Ely Greenstone and Adjacent Areas: Soudan, Eagles Nest, and Bear Island 7.5” Quadrangles, St. Louis County, NE Minnesota: unpublished geologic map, 2005 North-Central Geological Society of America Field Trip 9, Minneapolis, MN May 2005, 1:10000 scale. Peterson, D. M., and Jirsa, M. A., 1999, Bedrock Geological Map and Mineral Exploration Data, Western Vermilion District, St. Louis and Lake Counties, Northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map Series M-98, scale 1:48,000. Peterson, D. M., and Patelke, R. L., 2003, National Underground Science and Engineering Laboratory (NUSEL): Geological Site Investigation for the Soudan Mine, NE Minnesota: Natural Resources Research Institute, Technical Report NRRI/TR-2003/29, 88 p. Radakovich, A. L., Parent, C. T., Partridge, M. E., Ritts, A. D., Pierce, R., and Hudak, G. J., 2010, Reconnaissance Bedrock Geological Map of the Northern Part of Soudan Underground Mine State Park and the Northwestern Part of Lake Vermilion State Park, St. Louis County, Minnesota: Precambrian Research Center Map Series, PRC/MAP-2010/04, 1:5000 scale. 30 �Fluid Inclusion study of the Magino Archean Gold Deposit; Implications for Regional Mineralizing Systems HAROLDSON, Erik1, BROWN, Philip1 1Department of Geoscience, University of Wisconsin-Madison, 1215 W. Dayton, Madison, WI, 53706 USA The Magino gold deposit is located approximately 45 km's Northeast of Wawa, Ontario, Canada in the Goudreau-Localsh area of the Michipicoten Greenstone Belt. The Archean Lode Gold deposit is a newly enlarged gold resource in a mining camp historically known to host deposits of ≤1 Moz. in total endowment. As of October 2012; the Magino resource is reported (NI 43-101) as having 6.25 Moz of gold, and there remains potential for expansion. Gold was first discovered on the Magino project in 1917. Production on the property has been from two generations of historical mining. The first generation was in the 1930’s and approximately 10,000 oz of gold were produced. The second more recent underground venture operated from 1987 to 1993 and an additional 105,000 oz of gold was produced. Recently Argonaut Gold Inc. has taken ownership of the property through the acquisition of Prodigy Gold Inc. Argonaut Gold is planning a Pre-Feasibility study for 2013 which will look to exploit the Magino deposit as an open pit mine. The objective of the study is to better understand the mineralizing system(s) Figure 1 shows the location of the Magino mine project in Northern responsible for forming this Ontario, Canada. Geology data is from OGS: Bedrock Geology of Ontario gold deposit and to put that 1:250,000 (Revised MRD 126) information in the context of the regional geology to better understand the economic potential of the deposit along with the surrounding greenstone belt. Fluid inclusion research will constrain the geologic conditions at the time of mineralization, as well as provide clues towards the water-rock interactions prior to and during ore deposition. Initial work involves microscopy of polished thick sections and microthermometry to establish fluid inclusion assemblages. To better aid the assignment to and interpretation of fluid inclusion assemblages, the fluid inclusion thick sections are being imaged in a Scanning Electron Microscope using variations in cathodoluminescence signatures to differentiate various hydrothermal quartz generations from earlier deformed primary quartz phenocrysts associated with the host trondhjemite intrusion. Raman spectroscopy will be 31 �utilized to interpret carbonic-rich fluid inclusions for CH4 content and to aid in pressure interpretations. The recent expansion of the Magino resource raises the exploration potential of the surrounding Michipicoten greenstone belt significantly. The opportunity exists to aid in discovery of new and expansion of known similar type gold deposits in the region. Fluid inclusion studies in the region have been somewhat inconclusive as to understanding the nature of the gold mineralizing system(s). In the Michipicoten Greenstone belt, fluid inclusion studies including my preliminary work on the Magino deposit have Figure 2 (A) shows a methane rich fluid inclusion at room temperature; notice small bright unknown solid and dark cluster of unknown solids near top. (B) shows the same inclusion after cooling to -100°C; notice a frozen CO2 solid (darker) and a Ch4 vapor bubble (lighter) which formed while cooling at ~ 98°C. hinted strongly at a regional genetic link between the various widespread study areas. Fluid inclusion study will form the foundation to more rigorous study of the Michipicoten Greenstone Belt evolution and most importantly the gold mineralization. REFERENCES Borthwick, R.W. 1987. The distribution and association of gold within quartz veins. Magino mine prospect, Wawa, Ontario; A thesis presented to the Department of Geological Sciences Brock University in partial fulfillment of the requirements for the degree Bachelor of Science with Honours in Geology, 49p Brown, P.E., and Hagemann, S.G., 1995, The program MacFlinCor and its application to geobarometry in Archean lode-gold deposits. Geochim Cosmochim Acta 59, 3943-3952. Brown, P.E., and Hagemann, S.G., 1994, MacFlinCor: A computer program for fluid inclusion data reduction and manipulation. In De Vivo and Frezzotti (eds) Fluid Inclusions in Minerals: Methods and Applications, VPI Press, 231-250. Götze, Jens. Application of Cathodoluminescence Microscopy and Spectroscopy in Geosciences. Microscopy and Microanalysis 18, no. 06 (2012): 1270–1284. Heather, K.B. and Arias, Z. 1992. Geological and structural setting of gold mineralization in the GoudreauLochalsh area, Wawa gold camp; Ontario Geological Survey, Open File Report 5832, 159p. Lu, Wanjun, I-Ming Chou, R.C. Burruss, and Yucai Song. A Unified Equation for Calculating Methane Vapor Pressures in the CH4–H2O System with Measured Raman Shifts. Geochimica et Cosmochimica Acta 71, no. 16 (August 15, 2007): 3969–3978. Sage, R.P. 1993. Geology of Augonie, Bird, Finan and Jacobson townships, District of Algoma; Ontario Geological Survey, Open File Report 5588, 286p Samson, I.M. Bulent, B., and Holm, P.E., 1997, Hydrothermal evolution of auriferous shear zones, Wawa, Ontario: Economic Geology, v. 92, p. 325-342. Studemeister, P.A., and Kilias, S., 1987, Alteration pattern and fluid inclusions of gold-bearing quartz veins in Archean trondhjemite near Wawa, Ontario, Canada: Economic Geology, v. 82, p. 429-439 Sutherland, K.S., 1987, A preliminary report on the Magino gold deposit, Wawa, Ontario, A report submitted to the Department of Geological Sciences in partial fulfillment of the requirements for the non-research Master’s of Science in Mineral Exploration, Queen’s University, Kingston, Ontario. 32 �Preliminary geochemical analysis of the Nipigon Bay granites, northern Lake Superior HOLLINGS, Peter1, and SMYK, Mark2 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada 2 Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, 435 James St. S., Suite B002, Thunder Bay, ON, P7E 6S7 Canada The granitic basement to exposed, overlying Mesoproterozoic Sibley Group sedimentary rocks in Nipigon Bay, northern Lake Superior, was first identified by diamond drilling in 1997 (Wells, 1997). A series of three, prominent magnetic anomalies delineates these magnetite-bearing granitoids in the subsurface (Fig. 1). These granitic intrusions occur at a confluence of three regional-scale structures: the northeast-trending Gravel River fault; the north-northeast-trending Jackpine River fault; and the west-southwest-trending North Shore fault, which forms the base of the Osler Group volcanic rocks of the Midcontinent Rift. The boundary between the Neoarchean Wawa and Quetico subprovinces has also been extrapolated beneath Proterozoic cover under Nipigon Bay (Williams, 1989). Sampling of drill core for geochemical and geochronologic analysis was recently undertaken in order to describe and characterize these intrusive rocks. Figure 1. A) Map of upper Great Lakes showing the location of the study area. B) Regional geology map showing the location of the Nipigon granites. Granitic rocks from beneath Nipigon Bay are characterised by enriched LREE (La/Smn = 2.9 to 7.7) and flat to weakly fractionated HREE (Gd/Ybn = 1.4 to 3.1) with pronounced negative Nb anomalies (Nb/Nb* = 0.1 to 0.2). The granites are metaluminous and plot within the field of volcanic arc granites on the granite discrimination diagrams of Pearce et al. (1984). When compared to other granite suites in the area, the Nipigon Bay granites most closely resemble the I-type granites of the Dog Lake chain (Kuzmich et al., 2012) rather than the S-type Neoarchean granites of the Pukaskwa batholith (Beakhouse et al., 2011) and nearby Georgia Lake area (Breaks et al., 2008), or the Mesoproterozoic anorogenic English Bay granites in the Nipigon Embayment (Hollings et al., 2004). 33 �The granites of the Dog Lake chain, 80 km west-southwest of Nipigon Bay, appear as a series of distinct aeromagnetic “highs” along the southern boundary of the Quetico subprovince. Kuzmich et al. (2011) interpreted the granites to have formed within a suprasubduction mantle and were subsequently emplaced along crustal-scale faults that form terrane boundaries. Similarities in geochemistry, magnetic signature and regional tectonic setting suggest that the Nipigon Bay and Dog Lake granites may have formed in a similar manner. Pending geochronologic and geochemical data for both the Dog Lake and Nipigon Bay intrusions will help to elucidate this possible relationship. References Beakhouse, G., Lin, S. and Kamo, S., 2011. Magmatic and tectonic emplacement of the Pukaskwa batholith, Superior Province, Ontario, Canada; Can. J. Earth Sci., 48, p.187–204. Breaks, F.W., Selway, J.B. and Tindle, A.G. 2008. The Georgia Lake rare-element pegmatite field and related S-type, peraluminous granites, Quetico Subprovince, north-central Ontario; Ontario Geological Survey, Open File Report 6199, 176p. Hollings, P., Fralick, P. and Kissin, S., 2004. Geochemistry and geodynamic implications of the Mesoproterozoic English Bay Granite-Rhyolite complex, northwestern Ontario; Canadian Journal of Earth Sciences, 41, p.1329-1338. Kuzmich, B., Hollings, P., Campbell, D. and Scott, J, 2012. Geochemistry and petrology of the Dog Lake granite chain, Quetico Basin, Northwestern Ontario; Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 1 - Proceedings and Abstracts, v. 58, part 1, 56-57. Pearce, J., Harris, N. and Tindle, A., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks; Journal of Petrology, 25, 956-983. Wells, K. 1997. Assessment report on the 1997 drilling program, Nipigon bay area; unpublished report 2.17382, Falconbridge Limited, assessment files, Thunder Bay South District, Ministry of Northern Development and Mines, Thunder Bay, 138p. Williams, H.R. 1989. Geological studies in the Wabigoon, Quetico and Abitibi-Wawa subprovinces, Superior Province of Ontario, with emphasis on the structural development of the Beardmore-Geraldton Belt; Ontario Geological Survey, Open File Report 5724, 189p. 34 �The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter - 2013 Update HUDAK, George1, MONSON GEERTS, Stephen1, ZANKO, Larry1, SEVERSON, April1, SEVERSON, Allison1, KRAMER, Stuart1, BANDLI, Bryan2 1 Natural Resources Research Institute, 5013 Miller Trunk Highway, Duluth, MN, 55811 2 Department of Geological Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 Since 2008, the Natural Resources Research Institute (NRRI) has been conducting a detailed characterization of mineral dust in northeastern Minnesota. The purpose of this research is to evaluate the effects of past and present emissions from taconite mining and processing on air quality throughout the Mesabi Iron Range (MIR) by characterizing airborne particulate matter within taconite operations, in communities surrounding taconite operations on the MIR, in population centers in other regions of northeastern Minnesota, and non-MIR locations, as well as particulate matter deposited in lake sediments (Figure 1). NRRI’s sampling and characterization work represents the community/environmental component of the Minnesota Taconite Workers Health Study, a broad University of Minnesota (UM) research effort investigating long-standing questions regarding the impact of dust derived from mining and processing of taconite (iron ore). The UM School of Public Health (SPH), with whom NRRI is collaborating, is responsible for the human health- and exposure-related components of that effort, which include: 1) an occupational exposure assessment; 2) a mortality study; 3) a cancer incidence study; and 4) a respiratory health survey of taconite workers and spouses. Figure 1. Locations of taconite processing plants on the Mesabi Iron Range being sampled during this study (after Oreskovich and Patelke, 2006) Air sampling is performed within taconite operations, MIR communities, and non-MIR communities by NRRI scientists during both winter and summer seasons. Sampling at taconite operations takes place at four locations: 1) secondary crushers; 2) magnetic separators/concentrators; 3) agglomerators/ ball drums; and 4) kiln/pellet discharge area. Sampling within MIR communities takes place on the rooftops of public buildings, whereas sampling in non-MIR communities occurs on rooftops or in remote locations so that background air quality can be evaluated. Airborne particles are collected using: 1) a micro orifice uniform 35 �deposit impactor (MOUDI) (Marple et al., 1991), which enables size-fractionated particulate matter collection; and 2) a total suspended particulate filter (TSP). Particulate matter is evaluated via gravimetric analysis and subsequently subjected to comprehensive particulate matter characterization that includes: 1) scanning electron microscopy (SEM) imaging; 2) energy dispersive x-ray spectroscopy (EDS); 3) electron backscattered diffraction (EBSD); 4) proton induced x-ray emission (PIXE); 5) the Minnesota Department of Health’s 852 Method Transmission Electron Microscopy (TEM) Analysis for Mineral Fibers in Air; and 6) the International Standard Organization’s Method 10312 Ambient air – Determination of Asbestos Fibers – Direct-Transfer TEM Method (ISO 10312, 1995). During the past year, the NRRI has been evaluating particulate matter physical data (including gravimetric data, and particulate matter morphology data), particulate matter mineralogical compositions, and particulate matter chemical compositions obtained from both MIR taconite operations, MIR communities, and non-MIR communities, including 14 sampling events at taconite operations and 79 sampling events at locations within communities and sites in northeastern Minnesota (73) and Minneapolis (6), as summarized in Table 1. Lake sediment analysis continues, and will provide important data regarding potential mineralogical inputs from iron mining and processing from ~1840 (which predates iron mining on the MIR) to the present, which includes the period where the transition from natural ore mining to taconite mining took place. Continued analysis, interpretation and reporting will take place in 2013. Taconite Facility United Taconite In-Plant Sampling Events Sampling Events Taconite Facility 2 events while active Keetac Hibbing Taconite Sampling Events 1 event while active, 1 event while inactive 1 event while inactive, 3 events while active 3 events while active 1event while active, Northshore 1 event while inactive Minntac 1 event while active Minorca Community Sampling Events Community Sampling Location Sampling Events and Number per Season Keewatin Elementary School 7 Events (3 Winter / 4 Summer) Hibbing High School 9 Events (4 Winter / 5 Summer) Virginia City Hall 10 Events (5 Winter / 5 Summer) Babbitt Municipal Building 16 Events (7 Winter / 9 Summer) Silver Bay High School 13 Events (4 Winter / 9 Summer) Ely Fernberg Site 6 Events (3 Winter / 3 Summer) Duluth NRRI Rooftop 12 Events (7 Winter / 5 Summer) UMN Mech. Eng. Rooftop 6 Events (3 Winter / 3 Summer) Table 1. Summary of in-plant and community sampling events. References ISO 10312, 1995, Ambient air – determination of asbestos fibers – direct transfer transmission electron microscopy method, 51p. Marple, V. A., Rubow, K. L., and Behm, S. M., 1991, A micro orifice uniform deposit impactor (MOUDI): description, calibration, and use: Aerosol Science and Technology, v. 14, p. 434-446. Oreskovich, J. A., and Patelke, M. M., 2006, Historical use of taconite byproducts as construction aggregate materials in Minnesota: A Progress Report: Natural Resources Research Institute Report of Investigation NRRI-RI-2006-02, 10 p. 36 �Bedrock geologic map of the western Gunflint Trail area, northeastern Minnesota JIRSA, Mark A. Minnesota Geological Survey (MGS), St. Paul, jirsa001@umn.edu Figure 1—Generalized map of northeast Minnesota showing geologic setting of map area along the western end of the Gunflint Trail (dashed line). The 2007 Ham Lake forest fire provided an opportunity for detailed mapping in a classic area of Precambrian geology along the western Gunflint Trail corridor into the Boundary Waters Canoe Area Wilderness (Fig. 1). Because the area is a favorite of campers and resorters, the map was designed somewhat differently from other products of the MGS to provide more “general interest” content, including highlights of the unique geologic features that lie along the many well maintained hiking trails (Fig. 2). The burn covers more than 120 mi2, but nearly 2/3 of it is in adjacent Canada—the Gunflint map encompasses most of the U.S. portion. Mapping was supported in part by U.S. Geological Survey STATEMAP element of the National Geologic Mapping Program, and the results were published in 2011 as MGS Miscellaneous Map M-191. The western part of the Gunflint trail provides a transect across well exposed rock units of Neoarchean, Paleoproterozoic, and Mesoproterozoic age. Archean granite-greenstone terrane of the Wawa subprovince of Superior Province is represented by a succession of metavolcanic rocks locally known as the Paulson Lake sequence (ca. 2720 Ma), and the Saganaga Tonalite (ca. 2690 Ma). Diabasic dikes of imprecisely known age cut the Archean bedrock. Both they and the Archean rocks are unconformably overlain by Paleoproterozoic metasedimentary strata of the Animikie Group (ca. 1870-1830 Ma), which includes the Rove Formation and Gunflint Iron Formation. The stratigraphic top of iron-formation is marked by a thin sequence of ejecta from a meteorite impact that occurred near Sudbury Ontario, ca. 1850 Ma. This sequence forms a discontinuous ejecta blanket that is present throughout the Lake Superior region in Ontario, Michigan, and elsewhere in Minnesota (Jirsa and others, 2011, and references therein). Mesoproterozoic rifting is manifest in hypabyssal dikes and sills of the Logan intrusions (ca. 1115 Ma), and several phases of the Duluth Complex (ca. 1100 Ma) emplaced into both Archean and Proterozoic rocks. Stratigraphic facing directions in Neoarchean supracrustal rocks, based on pillowed metabasalt flows, indicate that the sequence forms a northwest-trending, steeply south-dipping and younging homocline. Much of the temporal distinction between various geological elements in the Archean rocks is based on regionally consistent fabrics that resulted from three major phases of deformation, denoted D1, D2, and D3. All three deformation events are the result of NS- to NW-SE-directed compression. Regionally, D1 deformation folded, tilted, and thrust faulted large sections of the supracrustal rocks, but did not produce significant metamorphic mineral assemblages or cleavage and schistosity. The timing of D1 deformation is bracketed between deposition of the volcanic and clastic rocks at about 2722 Ma (Peterson and others, 2001), and emplacement of the Saganaga Tonalite at 2690.83 Ma (Driese and others, 2011). The effects of D2 deformation in this area include moderate to mild flattening of minerals and inclusions near the margins of the Saganaga Tonalite; and strong schistosity, metamorphism, folding of tonalite 37 �dikes, and flattening of pillow structures in the immediately adjacent supracrustal rocks. U-Pb dates of intrusions that bracket D2 place the regional deformation and metamorphic event between about 2674 Ma and 2685 Ma (Boerboom and Zartman, 1993). D3 deformation produced faults and shear zones. Much of the map’s apparent complexity in areas of Paleoproterozoic strata and Logan intrusions is a product of the shallow dip and differential erosion of these formations, local faults and shallowly plunging folds, and moderate to high topographic relief. Deformation associated with the Lookout Fault grades from a distinct fault structure on the west—where Paleoproterozoic iron-formation is folded and faulted against Archean rocks—to a sympathetic drape structure in the eastern part of the area that is dominated by higher stratigraphic levels of iron-formation and Rove slate. The drape structure is manifest in a shallowly east-plunging, anticline-syncline pair that shows progressive flattening of limbs toward the southeast. The depiction of the Mesoproterozoic Logan intrusions differs significantly from previous mapping that implied the intrusions were folded. New field work indicates that the intrusions are semi-concordant with adjacent strata where dips are gently southeastward, but are locally discordant, particularly along the north-dipping limb of the anticline-syncline pair. Although the sills generally mimic fold structures, their local discordance implies that much of the deformation associated with the Lookout Fault and related structures predated emplacement of Logan intrusions. Figure 2. Geology of the southern part of Gunflint map area showing sites of geologic interest. References Boerboom, T.J., and Zartman, R.E., 1993, Geology, geochemistry, and geochronology of the central Giants Range batholith, northeastern Minnesota: Can. J. of Earth Sci., 30: 2510-2522. Driese, S.G., Jirsa, M.A., Ren, M., Sheldon, N.D., Brantley, S.L., Parker, D., and Schmitz, M., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, 189:1-17. Jirsa, M.A., Fralick, P.W., Weiblen, P.W., and Anderson, J.L.B., 2011, Sudbury impact layer in the western Lake Superior region, in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 147–169, doi:10.1130/2011.0024(08). Peterson, D.M., Gallop, Christina, Jirsa, M.A., and Davis, D.W., 2001, Correlation of Archean assemblages across the U.S.-Canadian border: Phase I geochronology: (abstract): Institute on Lake Superior Geology Proceedings, 47th annual meeting, Madison, WI, v. 47, Part 1, p. 77-78. 38 �Minnesota River Valley subprovince as depicted on a new bedrock geologic map of Renville County, southwestern Minnesota JIRSA, Mark A., SETTERHOLM, Dale R., and CHANDLER, V.W. Minnesota Geological Survey, 2642 University Ave., St. Paul, MN 55114-1032 (jirsa001@umn.edu) The bedrock geology of Renville County has a long history of geologic study related to the apparent antiquity of the bedrock, and commercial interest in quarrying, clay mining, and early speculation about coal resources. The bedrock ranges in age from Paleoarchean (~ 3500 Ma) to Late Cretaceous (~90 Ma), and records a complex history involving multiple igneous, sedimentary, metamorphic, tectonic, and weathering events. The interpretation presented here focuses on the Archean and Paleoproterozoic geology. The map derives from published and unpublished mapping, augmented with high-resolution geophysical data, scant drill core, and field work by the authors in a narrow strip of scattered exposure along the Minnesota River that forms the county’s southern border. Figure 1. Generalized geologic map of Minnesota showing Renville County (black) adjacent to the Minnesota River (bold line) and within block subdivisions of the Minnesota River Valley subprovince of the Archean Superior Province. 39 �The Minnesota River Valley subprovince is divided into largely fault-bounded blocks, each having distinctive attributes. Renville County straddles two of those blocks—the Montevideo block to the north, and the Morton block to the south (Fig.1). Both blocks are characterized by gneissic bedrock, but they differ from one another in composition and grade of metamorphism of the gneisses contained therein. The Montevideo block consists of variably layered granitic to gabbroic gneisses, metamorphosed under granulite facies conditions and collectively referred to as the Montevideo Gneiss. The Morton block contains tonalitic to granitic migmatitic gneiss with amphibole-rich rafts, metamorphosed under upper amphibolite facies conditions; and granitic neosome occurring as diffuse pods, lenses, and discrete intrusions. Collectively, these rock types comprise the Morton Gneiss. The two blocks are separated by the geophysically distinct Yellow Medicine shear zone. The complex intrusive, metamorphic, and deformation history of the gneisses and intrusions has been partially unraveled by recent high-precision U-Pb geochronologic studies. The apparently oldest components of the gneissic bedrock in both the Montevideo and Morton blocks range in age from 3535±4 Ma (Bickford and others, 2006), to 3422±1 Ma (Schmitz and others, 2006). Collectively, the ages indicate that gneisses formed in the Paleoarchean (3600-3200 Ma), were multiply intruded, deformed, and metamorphosed during Mesoarchean (3200-2800 Ma) and Neoarchean (2800-2500 Ma) time, and intruded by granitic magmas at ca. 2600Ma. Geologic mapping by the authors and structural analyses by previous workers indicate that the gneissic bedrock underwent two major periods of regional deformation (designated D1, D2), and two later events that can be distinguished locally. The D1 event produced high-grade metamorphism and gneissic foliation (S1). The D2 event folded this foliation into shallowly northeast-plunging antiforms and synforms, localized doubly-plunging dome and basin structures, and zones of semi-ductile shear. Much of the apparent complexity of map patterns is inferred to be the result of these open, shallowly plunging F2 folds of the early-formed, nearly horizontal D1 metamorphic fabric. D2 also produced a localized foliation (S2) subparallel to fold axes. Movement along the Yellow Medicine shear zone was likely complex and protracted, but at least some of the deformation is inferred to represent north-over-south thrusting associated with crustal shortening during the Neoarchean (~2680 Ma) Minnesotan Orogeny. Drill core from two holes in the northern part of the county indicate that primary foliation and localized shearing in and near the Yellow Medicine shear zone dips northward, which is consistent with geophysical modeling conducted for this study. It is likely that the Yellow Medicine shear zone was reactivated during the Paleoproterozoic (2500-1600 Ma) to juxtapose strata inferred to be part of the Little Falls Formation against older bedrock during the Penokean and/or Yavapai orogenies. The gneissic bedrock in both blocks is cut by a variety of intrusions, including granite, granodiorite, and quartz monzonite of likely Archean age; discretely bounded felsic to mafic intrusions that could be either Archean or Proterozoic; and diabasic dikes of Paleoproterozoic and perhaps Mesoproterozoic age. The Bedrock Geology of Renville County is Plate 2 of the Minnesota Geological Survey’s County Atlas C-28 (http://www.mngs.umn.edu), which was supported by the Renville County Board of Commissioners and the Department of Natural Resources Division of Waters. References Bickford, M.E., Wooden, J.L., and Bauer, R.L., 2006, SHRIMP study of zircons from Early Archean rocks in the Minnesota River Valley: Implications for the tectonic history of the Superior Province: Geological Society of America Bull. v. 118, p. 94-108. Schmitz, M.D., Bowring, S.A., Southwick, D.L., Boerboom, T.J., and Wirth, K.R., 2006, High-precision U-Pb geochronology in the Minnesota River Valley subprovince and its bearing on the Neoarchean to Paleoproterozoic of the southern Superior Province: Geological Society of America Bull., v. 118, p. 82-93. 40 �Geochemistry of reversely-polarized intrusions along the SW limb of the Midcontinent rift system, Carlton County, Minnesota JOHNSON, Teresa1, WENDLANDT, Richard1, SHANNON, James2, 1 Department of Geology and Geological Engineering, Colorado School of Mines, 1516 Illinois Street, Golden, CO 80401 2 MMG, 390 Union Boulevard, Suite 200, Lakewood, CO 80228 The early rifting phase (1115-1100 Ma) of the Midcontinent rift system (MCR) is characterized by the emplacement of reversely-polarized extrusive and intrusive basaltic rocks. Comprehensive studies of the basalts from this early rifting period have delineated three distinct compositions, type I, II and III, which can be correlated throughout the rift system (Nicholson et al., 1997). These three basalt types are recognized in the Ely’s Peak Basalts along the southwest limb of the rift near Duluth, Minnesota. The reversely-polarized intrusions in this location are the diabase dikes of the Carlton Dike Swarm (CDS), the evolved Fe-Ti rich gabbro of the Esko intrusion and the Ni-Cu-PGE mineralized ultramafic Tamarack Intrusion. This study evaluates the diabase dikes of the Carlton Dike Swarm and their petrogenetic relationship to the other MCR reversely-polarized rocks. Previously, the CDS was characterized as a Ti- and Fe-enriched quartz tholeiite (Green et al., 1987; Reichhoff, 1987). This study evaluates four geochemical subgroups of dikes, types A-D (Table 1). Within the main NE-trending Carlton Dike swarm, three compositional subgroups are observed. The majority of dikes (type A) are classified as high-TiO2 (2.9-4.6 wt%) and outcrop mainly along the St. Louis River and include the prominent columnar-jointed dike outcropping south of the Thomson Reservoir. The two other dike types include: a similar high-TiO2 dike (type B) with a greater percentage of hydrous minerals (5%), and the steepest HREE slopes; and a lowTiO2 dike (1.2-1.3 wt%) (type C) locally known as the Cloquet dike. The Cloquet dike is recognized as the longest dike (55 km) in the swarm based on its aeromagnetic signature. In addition, a smaller subset of NW-trending dikes (type D) is also distinguished by aeromagnetics, with the most prominent dike located approximately 50 km to the southwest of the main CDS. This dike has major and trace element compositions that are analogous to type A dikes. The Esko Intrusion is a circular-shaped aeromagnetic feature located on the NE corner of Carlton County. The intrusion is also geochemically similar to the high-TiO2 CDS with a greater abundance of hydrous minerals similar to the type B dikes. The CDS diabase dikes have geochemical similarities to type II and III basaltic compositions, which is consistent with the timing of CDS dike emplacement near the end of the early reversal period (Green et al., 1987). The higher Mg#’s associated with the basaltic composition type I are not found in the diabase dikes sampled during this study or in previous research (Table 1). The type C dike is comparable to type II basalts with a similar Mg# and TiO2 but has a nearly horizontal HREE pattern and more pronounced negative Nb and Ta anomalies (Figure 1). Type A, B & D dikes are most similar to type III basalts. Type A & D dikes have the most evolved Mg#’s like type III basalts, while type B is less evolved. The trace-element and REE patterns of type D dikes overlap with the type A field, and both are more enriched in HREE than type III basalts (Figure 1). A significant difference, arguably, is the depletion of Nb relative to Th in the type A dikes, which is interpreted as indicating late contamination by a similarly sourced magma. The type B dikes and type III basalts have similar slightly negative Nb anomalies although the dikes are characterized by a positive Eu anomaly and steeper HREE pattern (Table 1). 41 �Nicholson and Schulz (2011) propose that type I and III basaltic compositions are possibly related by fractionation and variable amounts of contamination. Other investigators (e.g. Hou et al., 2011) have modeled sources of high-Ti basalts by fractional crystallization from highMg, low-Ti basalts. To evaluate possible genetic relationships between different dike groups, fractional crystallization models involving the most primitive samples (e.g. type C) along with basalt types I, II and III are being used. Thermodynamic modeling of the origin of high-Ti basaltic compositions is enhanced by low fO2 and a high Mg#, but possible influences of assimilation and magma mixing remain a complexity. Figure 1: Chondrite normalized trace-element patterns. Normalization factors from McDonough & Sun (1995) except for P from Thompson et al. (1983). Data for basalt type I, II and III from Nicholson et al. (1997). Table 1: Chondrite normalized as indicated (McDonough & Sun, 1995); FeO=0.85FeOT; Mg# = Mg/Mg+Fe (mole percent); n, number of samples REFERENCES Green, J., Bornhorst, T., Chandler, V., Mudrey Jr., M., Meyers, P., Pesonen, L., and Wilband. J., 1987. Keweenawan Dykes of the Lake Superior Region: Evidence for Evolution of the Middle Proterozoic Midcontinent Rift of North America. In Mafic Dyke Swarms. Geological Association of Canada. Special Paper 34: 289–302. Hou, T., Zhang, A., Kusky, T., Du, Y., Liu, J., and Zhao. A., 2011. A Reappraisal of the high-Ti and low-Ti Classification of Basalts and Petrogenetic Linkage Between Basalts and Mafic–Ultramafic Intrusions in the Emeishan Large Igneous Province, SW China. Ore Geology Reviews, 41: 133–143. Nicholson, S., Schulz, K., Shirey, S., and Green, J., 1997. Rift-Wide Correlation of 1.1 Ga Midcontinent Rift System Basalts: Implications for Multiple Mantle Sources During Rift Development. Canadian Journal of Earth Sciences, 34: 504–520. Nicholson, S., and Schulz. K., 2011. Geochemical Evolution of 1.1 Ga Midcontinent Rift Magmatism. Geological Society of America, Abstracts with Programs, 43: 228. Reichhoff, J., 1987. Two Keweenawan Basaltic Dike Swarms in the Duluth Area, Minnesota. University of Minnesota, Duluth, 206 p. 42 �SEDIMENTOLOGY AND PALEOGRAPHIC RECONSTRUCTION OF THE STRATA ADACENT TO THE SUDBURY IMPACT LAYER IN A CORED DRILLHOLE KARMAN, Monica M.1, and FRALICK, Philip W.1 1Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, mmkarman@lakeheadu.ca, pfralick@lakeheadu.ca Drill core BDQ-2, extracted adjacent to Highway 588 near Thunder Bay, Ontario, Canada revealed stratified layers of the 1878.3 1.3 Ma (Fralick et al., 2002) Gunflint Formation, the 1850 Ma (Krogh et al., 1984) Sudbury Impact Layer, and the overlying 1832 3 Ma (Addison et al., 2005) Rove Formation. A limestone unit in the drill core, situated just above the Sudbury Impact Layer, and immediately below the fine-grained siliciclastic sediment of the Rove Formation, was analyzed to determine its mode of formation. The approximately three meters of chemical sediment overlying the ejecta layer has evidence of silicification below the uppermost meter. The top meter is composed of euhedral and subhedral calcite crystals with thin wisps of dark material separating the crystals and forming a chicken-wire texture (Figure 1A). The crystals consist of, on average, CaO 46.1%, MgO 0.68%, FeO 0.54% and MnO 0.19%. Five points in the carbonate unit were analyzed: point 1 being at the top of the analyzed unit, and going sequentially downward, through to point 5 at the bottom of the analyzed unit. Point 1 is situated at a sharp contact with the Rove Formation (Figure 1B). This point contains the largest carbonate crystals in the unit, consisting of ~3-5mm euhedral and subhedral zoned calcite crystals, in encasing fine-grained material (Figure 1A). The size of the calcite crystals decreases down the drill core, in the direction of the Sudbury Impact Layer, and at point 5 are ~200 -1.0mm in size. Results from SEM-EDX analysis using elemental mapping show zoned calcite crystals (Figure 2A) with Fe enrichment (Figure 2B) adjacent to the Mg enrichment (Figure 2C). The finegrained sediment present between the calcite crystals was also analyzed and is an assortment of clay minerals and calcite. Results from ICP-AES and MS analysis show low vanadium and chromium abundances in samples from points 1 to 4, but with much higher values, two orders of magnitude and one order of magnitude respectively, for the sample from point 5. PAAS normalized rare earth element curves for points 1 to 4 (Figure 3) have similarities with curves for meteoric water, whereas the curve for point 5 is similar to Paleoproterozoic seawater. Somewhat oxygenated meteoric water is needed for the transport of the vanadium and it would have precipitated under more reduced conditions in the sub-aerially exposed marine carbonate sediments. The fine-grained siliciclastic sediments above this were more oxygenated resulting in no vanadium enrichment though calcite crystals were precipitated from the meteoric waters creating a chicken-wire fabric. No modern analogs to this type of diagenetic alteration exist. 43 �1A 1B 3 2A 2B 2C Figure 1A) Calcite crystal chicken wire texture. Figure 1B) Contact between Rove Formation and carbonate unit. Figure 2A) SEM image of zoned calcite crystal. Figure 2B) Fe enrichment. Figure 2C) Mg enrichment. Figure 3) REE's for Points 1-5 (PAAS). REFERENCES Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005, Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology, v. 33, n 3, p. 193-196. Fralick, P.W., Davis, D.W. and Kissin, S.A., 2002. The age of the Gunflint Formation Ontario, Canada: single zircon U-Pb age determinations from reworked volcanic ash. Canadian Journal of Earth Sciences, v. 39, p. 1085-1091. Krogh, T.E., Davis, D.W., and Corfu, F., 1984. Precise U-Pb zircon and baddeleyite ages for the Sudbury area. In, E.G. Pye ed., The Geology and Ore Deposits of Sudbury Structure. Ontario Geological Survey, Special Volume 1, p. 431-446. 44 �Geologic mapping of Neoarchean and Paleoproterozoic rocks near Ester Lake by students of the Precambrian Research Center’s 2012 field camp KORMAN, Katrina1, CRADDOCK, Suzanne1, DOYLE, Michael1, WALTER, Jessica1, LEE, Aubrey2, and JIRSA, Mark3 1 2012 Field Camp Students, Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Duluth, Minnesota 55811 2 University of Minnesota Duluth, Department of Geological Sciences, 1114 Kirby Drive, Duluth, Minnesota 5581 3 Minnesota Geological Survey (MGS), University of Minnesota, 2642 University Avenue W., St. Paul, Minnesota 55114(jirsa001@umn.edu) The University of Minnesota-Duluth’s Precambrian Research Center conducted its sixth annual field camp in 2012, and this presentation is one of a series that show some of the results. During the fifth and sixth weeks of camp, teams of students participate in “capstone projects” that test student skills by creating new geologic maps in areas of poorly known geology. This capstone project involved mapping an area of the Boundary Waters Canoe Area Wilderness accessed by 12 lakes, with Ester Lake at its center (Fig. 1). The map provides details about the complex depositional and deformation history of a Neoarchean, largely metasedimentary terrane that is part of the Wawa subprovince of Superior Province. Figure 1. Generalized bedrock geologic map of northeastern Minnesota showing the Ester Lake capstone area. The unit labeled “Knife Lake Group” also encloses older volcanic sequences that are not delineated separately at this scale. Dashed line is the border of the Boundary Waters Canoe Area Wilderness. The Ester Lake map area lies along and west of the boundary between the Saganaga Tonalite (ca. 2690 Ma), and sedimentary strata of the Knife Lake Group that are inferred to have been derived in part from it. Both rock units were tilted, folded, faulted, and metamorphosed to low 45 �greenschist facies during a regional deformation event at about 2680 Ma, which provides an approximate minimum age for the Knife Lake Group. Our mapping demonstrated that strata of the Knife Lake Group in this area form a broad, northeast-trending synclinorium that is bounded by the Saganaga Tonalite on the east, and an apparently uplifted fault-block of metabasalt on the west (see cross section on poster). The limbs of this large structure are marked by smaller sympathetic folds, and are dissected by several faults and shear zones. Mapping to the east revealed an erosional unconformity at the contact between Saganaga Tonalite and basal Knife Lake strata. The tonalite appears to have been weathered and is overlain by sandstone containing both subrounded and angular quartz pebbles—approximately the same size as quartz phenocrysts in the tonalite—in a sandy matrix of altered plagioclase grains and fine grunge that resembles reworked granitoid saprolite. To the west, this basal unit grades stratigraphically upward into irregularly interbedded sequences of arkose; conglomerate containing amoeboid clasts of tonalite (“Fish Lake conglomerate”); hornblende-phyric, trachyandesite-bearing coarse pebble to cobble conglomerate; polymictic conglomerate with metabasalt clasts; gritstone composed of subangular sand grains of hornblende and plagioclase; and graywacke and slate containing rare, thin lenses of banded iron-formation. Collectively, this arrangement of stratigraphic facies indicates that Saganaga Tonalite and the rocks it intruded were uplifted and subaerially eroded to provide detritus to nearby basins, now manifest as Knife Lake strata. However, the diverse lithologic character of these strata indicates that a great variety of sediment source regions and depositional settings existed. The Saganaga Tonalite was likely weathered to form saprolite, which was eroded by slopewash, alluvial fan, and fluvial processes to produce the quartz pebble sandstone, and to contribute clasts to various conglomeratic layers. We attribute the several hundred-foot-thick arkosic unit to reworking of saprolite in a near-shore marine or lacustrine environment. The Fish Lake conglomerate, containing matrix-supported amoeboid (paleosaprolitic) tonalite clasts, is interpreted to be channelized alluvium deposited in a rapidly subsiding basin. Conglomerate containing hornblende-trachyandesite clasts likely formed from erosion off the informally named Jasper Lake volcanic sequence of Knife Lake Group, which lies southeast of the map area (Jirsa and Starns, 2008). Graywacke and slate are interpreted to represent deposition in a lacustrine or marine setting, and the interlayered coarser polymictic clastic strata may represent braided stream, alluvial fan, or subaqueous fan deposition of sediment shed off the uplifted flanks of the basin. The layered strata exhibit chaotic soft-sediment deformation features, local growth faults, and abrupt facies changes, suggesting that deposition was synchronous with episodic basin subsidence. Thin layers and lenses of jasper-bearing iron-formation that are associated with graywacke and slate are interpreted as chemical precipitates into what may have been a shallow marine environment. The lithologic diversity and structural complexity is consistent with a model of deposition in rhombochasms created during strike-slip or extensional faulting. The general fining-upward character of Knife Lake strata, albeit irregular, may represent increasing water depth over time within the depocenter. The western, fault-bounded metabasalt unit contains abundant tightly packed mattress-size pillows that lack amygdules, implying deposition in fairly deep water. These attributes, along with the vertical dip of metabasalt flows, are similar to parts of the ca. 2720 Ma Ely Greenstone and Newton Lake Formation, and represent a very different tectonic and depositional setting from that of the adjacent Knife Lake Group. A small unit of metadiabase identified during mapping may be part of a swarm of mafic intrusions seen elsewhere in the region that are inferred to be Paleoproterozoic in age. This and other maps produced by capstone projects can be viewed at www.d.umn.edu/prc. Reference Jirsa, M.A., and Starns, E.C., 2008, Preliminary bedrock geologic map of the 2006 Cavity Lake fire area, northeastern Minnesota: MGS Open-File Report OF08-05, scale 1:24,000. 46 �Preliminary Investigations of the Fe-Ti-V-P mineralization associated with the Thunderbird and Butler gabbroic intrusions within the McFaulds greenstone belt, Superior Province, Northern Ontario, Canada KUZMICH, Ben1, HOLLINGS, Pete1, HOULÉ, MICHEL G.2 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada 2 Geological Survey of Canada, GSC-Quebec, 490 rue de la Couronne, Québec, Quebec G1K 9A9 The McFaulds Lake area (i.e., Ring of Fire) located in northern Ontario (Canada) has been the site of recent mineral exploration leading to the discoveries of several mineralization types such as chromite and nickel sulfide deposits. Although the majority exploration has been focused on chromium, this area also contains significant Fe-Ti-V-P mineralization associated with gabbroic intrusions, in which the Thunderbird and Butler occurrences are the best defined.   These gabbroic intrusions are widely distributed throughout the McFaulds Lake area and maybe grouped into two main types of occurrences: (1) large mafic-dominated intrusions and (2) subconcordant to slightly discordant mafic-dominated sills/dikes characteristic of the Thunderbird and the Butler intrusions respectively. These intrusions are composed of an evolved mafic suite termed the ‘Ferrogabbro’ characterized by the presence of Fe-Ti oxides. Through detailed core logging, it has been recognized that both intrusions are largely composed by very similar lithologies including iron-rich gabbros, leucogabbros, and anorthosites. Two types of mineralization occur in these intrusions: (1) Fe-Ti-V and (2) Fe-Ti-P mineralization. Fe-Ti-V mineralization occurred within both intrusions whereas the Fe-Ti-P mineralization have been only identified within the Thunderbird intrusion. The mineralization occurs dominantly as disseminated magnetite and ilmenite (1-10%), but also present as semi-massive (50-80%), to massive layers (>80%). These layers typically contain distinct sharp, stratigraphically lower contacts and gradational upper contacts typical of primary igneous layering. The massive oxide layers are composed of magnetite and ilmenite typically at a ratio of 10:1. The ilmenite occurs as anhedral to subhedral crystals and to a lesser extent, as very fine-grained exsolutions within anhedral magnetite grains. Vanadium grades range from 0.30% to more than 0.60% (V2O5) and titanium grades range from 2.5% to more than 4.5% (TiO2) whereas Vanadium grades, at Butler, range from 0.42% to more than 1.17% (V2O5) and titanium grades range from 0.46% to more than 11.3% (TiO2).   47 �Sedimentology
and
geochemistry
of
the
Espanola
Formation,
Huronian
 Supergroup
 LAFONTAINE,
Daniel
and
FRALICK,
Philip
 Department
of
Geology,
Lakehead
University,
955
Oliver
Road

Thunder
Bay,
ON
 P7B
5E1
Canada
 
 The
Huronian
Supergroup
is
a
southerly
thickening
wedge
of
Paleoproterozoic
 sediments
with
a
maximum
thickness
of
12
km
and
an
age
range
of
2450
Ma
to
2219
Ma
 (Bennett
et
al.,
1991).

The
area
is
hypothesized
to
be
a
divergent
continental
margin
with
 the
paleo‐ocean
directly
to
the
south
(Fralick
and
Miall,
1989).

The
Huronian
Supergroup
 contains
evidence
of
three
separate
glacial
related
formations,
including
the
Ramsay
Lake
 Formation,
Bruce
Formation
and
the
Gowganda
Formation.

The
immense
size
of
similar
 glacial
formations
and
their
sedimentary
deposits
at
tropical
paleo‐latitudes
led
geologists
 (ie.,
Paul
Hoffman)
to
put
forth
the
Snowball
Earth
theory.

Geochemistry
from
the
Espanola
 Formation
may
very
well
help
identify
possible
reflections
of
atmospheric
composition,
 specifically
related
to
carbon
dioxide
and
oxygen
content.

The
Espanola
Formation
is
a
lone
 cap
carbonate
sandwiched
between
the
Bruce
and
the
Gowganda
Formations.

It
is
 classically
divided
into
three
individual
members;
the
lower
Limestone
member,
the
middle
 Siltstone
member
and;
the
upper
Dolomitic
Cap.

The
sediments
of
the
Espanola
were
 deposited
on
a
carbonate
shelf
post‐glaciation
that
begin
to
reflect
a
shallowing
upwards
 sequence
later
in
stratigraphy.

This
can
be
attributed
to
post‐glacial
isostatic
adjustment.

 This
abstract
will
discuss
the
lithofacies
associations
of
each
member,
their
depositional
 environments
and
possible
oceanic
geochemical
signatures.

Sediments
deposited
consist
of
 a
terrigenious
sediment
source
with
a
carbonate
cement
as
well
as
carbonate
minerals
 precipitated
directly
from
seawater.

Partial
dissolutions
in
acetic
acid,
of
selected
samples,
 were
analyzed
by
ICP
MS
to
better
understand
the
geochemistry
of
the
carbonates.

Yttrium
 is
not
removed
from
seawater
like
it’s
geochemical
twin
Holmium
(due
to
differing
surface
 stabilities)
and
seawater
generally
displays
high
Y/Ho
ratios
that
may
range
between
44‐74
 (Nagarajan
et
al.,
2011).

The
Espanola
geochemical
results,
represented
by
Figure
1
in
a
 graph
of
Yttrium
verses
Holmium,
displays
a
low
ratio
of
~28
which,
is
indicative
a
 dominant
chondritic
terrigeneous
sediment
source
(Nagarajan
et
al.,
2011),
possibly
 attributed
to
river
run‐off
systems
fed
by
the
glacial
melt
waters.
 Using the same samples, the spider diagram in Figure 2 displays a distinct “hat-like” pattern of MREE enrichment. This is not consistent with carbonates precipitating from Paleoproterozoic seawater but analogous with the MREE enriched patterns of major modern river systems (Matsaranta, 2006). This MREE enrichment supports a non-marine water components in the near-shore waters from which the carbonates precipitated .

The
possibility
that
the
carbonate
precipitation
in
the
middle
 siltstone
member
was
simply
carbonate
cement
is
likely
with
the
high
amount
of
siliciclastic
 material,
while
the
upper
and
lower
members
are
dominated
by
carbonate
precipitates.

 Geochemical
evidence
does
not
indicate
a
stratigraphic
control
that
classically
divides
the
 Espanola
members
but
rather
suggests
a
lateral
geographic
change
away
from
the
shore‐ line
in
limestone/dolostone
precipitation.

The
Espanola
represents
a
period
of
minor
 increases
in
oxygen
content
and
future
thesis
work
will
hopefully
aid
in
drawing
 conclusions
regarding
paleo‐atmospheric
composition.
 
 48 �
 References
 Bennett,
G.,
Dressler,
B.O.,
and
Robertson,
J.A.,
1991.
The
Huronian
Supergroup
and
Associated
 Intrusive
Rocks.
Ontarion
Geologic
Survey,
v.4,
no.1,
549–586.
 Fralick,
P.
and
Miall,
A.D.,
1989.
Sedimentology
of
the
Lower
Huronian
Supergroup
(Early
 Proterozoic),
Elliot
Lake
area,
Ontario,
Canada.
Sedimentary
Geology,
v.63,
127–153.
 Metsaranta,
R.,
2006.
Sedimentology
and
Geochemistry
of
the
Mesoproterozoic
Pass
Lake
and
 Rossport
Formations,
Sibley
Group.
Unpublished
Masters
Thesis,
Lakehead
University,
p.217
 Nagarajan,
R.
et
al.,
2011.
Geochemistry
of
Neoproterozoic
limestones
of
the
Shahabad
Formation,
 Bhima
Basin,
Karnataka,
southern
India.
Geosciences
Journal,
v.15:
9–25.
 49 �Chemostratigraphy of the Biwabik Iron Formation: Implications for Basin Longevity and Evolution LARSON, Phillip Duluth Metals Limited, 306 W Superior Street #610, Duluth MN 55802 United States The Paleoproterozoic Biwabik Iron Formation (BIF), an ~200m thick iron-formation in northeastern Minnesota, USA, is traditionally subdivided into four conformable members: Lower Cherty (LC), Lower Slaty (LS), Upper Cherty (UC), Upper Slaty (US), comprised predominantly of granular (cherty) and banded (slaty) iron-formation. Basin geometry and depositional environment of the iron-formation in the BIF and other 1.88 Ga circum-Superior Craton ironformation has been a topic of speculation for decades. Recent work suggests deposition on a clastic sediment-starved stable platform is a viable model for explaining the sedimentological and geochemical characteristics of circum-Superior iron-formations, as well as their widespread distribution. Geochemical evidence from the BIF in the Virginia Horn area provides further evidence in support of this model, and offers additional insight into both basin longevity and basin evolution. The BIF is composed predominantly of chemically precipitated sediment, reflected in Fe, Si, Mn, P, Mg, and Ca concentrations. A minor detrital clastic component is reflected in Al, Ti, and K concentrations. Al2O3 (the predominant detrital component) concentration ranges from 0.03% in granular iron-formation to 3.79% in slaty iron-formation, averaging 0.35% (sd=0.38%) (Fig. 1). Assuming a PAAS shale-like composition (Taylor and McLennan 1985), this corresponds to a detrital component averaging 2.3%, comparable to the detrital component found in modern platform carbonates. In contrast to the variability in detrital content, the Al2O3:TiO2 ratio is remarkably consistent through the thickness of the BIF, with the exception of a break in ratio corresponding to the top of the ‘Variably-Bedded and/or Mottled Unit’ of the LC member of Severson and others (2009) (Fig. 2). Below this horizon, Al2O3:TiO2 in the lower LC is 29:1, while above this horizon Al2O3:TiO2 in the LC-LS-UC-LS is 10:1. These constant ratios suggest that detrital material input to the BIF was sourced from long-lived, homogeneous reservoirs, likely as fine-grained, windborne dust. The constant ratios also suggest that variation in detrital concentration is a function of variation in chemical precipitate accumulation rates, rather than variation in detrital deposition rates. The two Al2O3:TiO2 ratios suggest two fundamentally different detrital source areas contributed to the BIF; the 29:1 ratio is comparable to that of modern airborne particles sourced from mature continental sediment in the Sahara, while the 10:1 ratio may reflect a volcanic source. The abrupt transition between the high Al2O3:TiO2 lower LC and the overlying low Al2O3:TiO2 LC-LS-UCUS sequence suggests that this contact represents a significant depositional hiatus, accompanied by a fundamental reorganization of atmospheric circulation and (or) emergence of a new detrital source area. Modern windborne dust accumulation rates in south Florida are on the order of 1.25 g·m-2·yr-1.; assuming similar deposition rates during BIF accumulation, it is possible to calculate rates and the duration of iron-formation accumulation from detrital concentration. Assuming s.g. 3.34, instantaneous accumulation rates range from 1.5 to 189 m/m.y., averaging 32 m/m.y. (Fig. 3). Granular iron-formation subunits are notably characterized by significantly higher accumulation rates than banded iron-formation subunits, both in the LC and LS-UC-US members. 50 �Overall, for the BIF in the Virginia Horn area, 1.6 myr of accumulation in the 63 m thick lower LC and 13.6 myr of accumulation in the upper 150 m thick LC-LS-UC-LS are indicated. The presence of a significant disconformity internal to the LC member challenges the assumption that accumulation of iron-formation in the BIF (and in other circum-Superior iron-formations) is the product of a single conformable depositional sequence, and particularly suggests re-evaluation of the traditional four member (LC-LS-UC-LS) subdivision of the BIF is warranted. 125.0 Depth (ft) Relative to top of "Variably-Bedded and/or Mottled Unit" (Severson et al 2009) Depth (m) Relative to top of "Variably-Bedded and/or Mottled Unit" (Severson et al 2009) 125.0 100.0 75.0 50.0 25.0 0.0 -25.0 -50.0 75.0 50.0 25.0 0.0 -25.0 -50.0 0.0 1.0 2.0 3.0 Al2O3 (wt%) 0 Figure 1. Al2O3 in the Biwabik Iron Formation, Virginia Horn area, Minnesota. n=1341 30 60 90 120 150 Accumulation Rate (m/myr) Figure 3. Iron-formation accumulation rates in the Biwabik Iron Formation, Virginia Horn area, Minnesota. LC 0.14 LS/UC/US 0.12 TiO2 (wt%) 100.0 0.10 0.08 0.06 0.04 0.02 0.00 0.0 0.5 1.0 Al2O3 (wt%) 1.5 Figure 2. Al2O3 and TiO2 in the Biwabik Iron Formation, Virginia Horn area, Minnesota. n=917 51 �Geochemistry and origin of slate-hosted massive sulfides of the Eagle Ni-Cu-PGE Deposit, northern Michigan: A preliminary study. LEATHERMAN, Mark1, RIPLEY, Edward1, ROSSELL, Dean2, WARE, Andrew2, and LI, Chusi1 1 Department of Geological Sciences, Indiana University, 1001 E. 10th St., Bloomington, Indiana 47405 2 Kennecott Eagle Minerals-Rio Tinto, Ishpeming, Michigan 49849 The Eagle Ni-Cu-PGE sulfide deposit is located within the Baraga Basin in northern Michigan. It is also located just south of the 1.1 Ga Midcontinent Rift System axis. Sulfide mineralogy consists of pyrrhotite-chalcopyrite-pentlandite. Primary ore types consist of disseminated, semimassive (net-textured), and massive that are associated with the main feldspathic peridotitemelagabbro intrusion. Recent drilling has revealed three additional classes of sulfides that are hosted in black slates of the Michigamme Formation. Class 1 consists of massive sulfide analogous in appearance to that of igneous-related massive mineralization (Figure 1). Class 2 consists of sulfides with quartz veins thought to be hydrothermal in origin. Some of these occurrences are associated with localized soft-sediment deformation. Also, neither of these sulfide classes shows evidence of lobate margins and dissolution. Class 2 sulfides can further be divided based on their appearance. Class 2a sulfides have fragments of sulfide entrained in predominant quartz (Figure 2a), whereas class 2b has approximately equal quartz and fine-grained sulfide intermixed (Figure 2b). Class 3 consists of discrete lenses or pods of massive pyrrhotite (class 3a) or chalcopyrite (class 3b). Furthermore, class 3 sulfides are enigmatic due to there being no obvious emplacement mechanism within slate (Figure 3a, b). Sulfur isotope and petrographic evidence suggests a dual igneous-replacement and thermal-sedimentary origin for the slate-hosted sulfides. Class 1 sulfides show 34S = 1.2 – 3.9 and minimal quartz along the sulfide-slate contact (origin). Class 2 sulfides show  34S = 7.5 – 11.8 and sharp sulfide-slate contacts with no noticeable alteration mineralogy. Class 3a sulfides display  34S = 5.4 – 10.3 along with acicular muscovite forming along the sulfide-slate boundary, protruding into the former, and are distal from the igneous intrusion; whereas class 3b displays  34S = 1.6 – 2.9, acicular Mg-rich chlorite, and are located proximal from the intrusion. Class 1 sulfides are thought to be magmatic in origin with some minor interaction with crustal sediments at lower temperatures. Class 2a sulfides are hypothesized to form as a result of preEagle tectonic activity and remobilization whereas class 2b sulfides are inferred to have originated via aqueous fluids that were initiated by heat flow supplied by intrusives. The class 3b chalcopyrite is suggested to be magmatic in origin with some crustal interaction, produced as a result of fractional crystallization of an immiscible sulfide liquid. Minor crustal interaction is indicated given: 1) petrographic evidence showing quartz and other silicates entrained in sulfide, 2) presence of Mg-bearing phases along slate-sulfide contact, and 3) the development of lobate sulfide-silicate contact zones. A hypothesis on the emplacement mechanism of slate-hosted chalcopyrite is that they are replacements of older sedimentary sulfides (i.e. trace pyrite). The class 3a pyrrhotite is thought to be of a thermal-sedimentary origin based on the following: 1) moderately high  34S values, 2) lack of mafic alteration phases along the slate-sulfide contact, and 3) significant spatial separation from the intrusion. Sedimentary pyrite may be converted to pyrrhotite in C-rich sedimentary rocks via reactions such as: FeS2 + 3H2O+ 5/2C = FeS + H2S + 3/2CO2 + CH4. It is clear that multiple periods of sulfide generation occurred in the area of the Eagle intrusion, all related in some manner to thermal effects associated with rift-related magmatic activity. 52 �Figure 1 (top): Class 1 sulfide showing evidence of hydrothermal-sedimentary involvement. Figure 2a (middle-left): Class 2a sulfide – fragments in majority quartz thought to be from pre-Eagle tectonism. Figure 2b (middle-right): Class 2b sulfide intermixed with quartz showing localized soft-sediment deformation. Figure 3a (bottom left): Pod of massive pyrrhotite (class 3a sulfide) encased in slate with no noticeable sedimentary deformation. Figure 3b (bottom right): Bleb of massive chalcopyrite (class 3b sulfide) encased in slate. 53 �Field, Petrographic, and Geochemical Study of the Bad Vermilion Intrusion, Mine Centre, Ontario, Canada LEE, Aubrey, and MILLER, Jim Department of Geological Sciences, University of Minnesota Duluth, 1114 Kirby Drive, HH 229, Duluth, MN 55812. During the summer of 2011, Aubrey Lee was contracted by Numax Resources, Inc. to map the bedrock geology of the Bad Vermilion Intrusion (BVI) which lies within Numax subsurface claims in Mine Centre, Ontario, Canada (Fig. 1b, Lee et al., 2012). The objective was to establish continuity of massive oxide units along the northern shore of Bad Vermilion Lake which had previously been documented by White and Albers (2010) along the northern shore of Seine Bay (Rainy Lake) to the west. In addition to facilitating accurate predictions for future exploration targets, the field work was to provide data for a field, petrographic, and geochemical study of the BVI as Aubrey’s master’s thesis at the University of Minnesota, Duluth. This study seeks to characterize the lithologic, petrographic, and geochemical attributes of the BVI in order to evaluate the potential for economic concentrations of Fe-Ti oxides, as well as PGE reef mineralization. The map (Fig. 1b) was presented last year at the 58th annual ILSG meeting in Thunder Bay (Lee et al., 2012). This poster presentation displays geochemical data and petrographic observations which provides a more comprehensive depition of BVI geology. The BVI is a 14 km long and 1-3 km wide, sub-vertically dipping, sill-like, layered plutonic sequence of gabbroic rocks exposed along the shores of Seine Bay (Rainy Lake) and Bad Vermilion Lake in Northwest Ontario, Canada (Fig. 1b; White and Albers, 2010). The intrusion is situated at the boundary between the metavolcanicplutonic Wabigoon subprovince to the north and the metasedimentary Quetico subprovince to the south in the Superior Province of the Canadian Shield Precambrian Figure 1: a) Magnetic Anomaly Map and Oxide Unit Correlation of the BVI; terrain (Poulsen, 1986). Ontario Geological Survey (2008). b) Geologic Map with subsets showing sampling profiles, drill collar locations, and plan-view extents of drill holes. It occurs within a dextral 54 �wrench fault zone bounded by the Quetico fault to the north and the Rainy Lake-Seine River fault to the south. The BVI lithologically consists of a mafic layered intrusive package of anorthositic, gabbroic, pyroxenitic rocks with layers/lenses of semi-massive to massive Fe-Ti oxides that have been significantly altered and possibly metamorphosed. The intrusion is exposed between felsic intrusive rocks at its northern contact and mafic volcanic rocks at its southern contact. The nature of these contacts and the magmatic relationship between each unit is enigmatic. Magnetic surveys of the Mine Centre area reveal three distinct magnetic highs extending the length of the intrusion and merging in the northeast where the exposure of the intrusion narrows along Bad Vermilion Lake (Fig. 1a; Ontario Geological Survey, 2008). Correlation of bedrock mapping with magnetic surveys indicates that the anomalies correspond to massive, semi-massive, and disseminated Fe-Ti oxide layers which are generally persistent along the length of the intrusion, though they may pinch, swell, and anastomose. The economic importance of the BVI lies in these oxide layers which contain significant concentrations of iron, titanium, vanadium and/or phosphorous (White and Albers, 2010). Four Numax drill holes (Fig. 1b) in the BVI highlight Fe-Ti mineralization, though only two on the west end of the intrusion have been fully geochemically assayed. Two near the central and east areas will eventually be assayed and incorporated into this study. During 2011 mapping, samples were collected along three profiles in the west, central, and eastern areas to supplement core data. Geochemical assays of field samples and drill core have helped establish petrologic and mineralogic continuity along the strike of the intrusion and document the igneous stratigraphy of the BVI. The western drill holes both intersected a thick package of melagabbroic rocks with 10 cm – 30 m intervals of massive oxide among oxide-bearing melagabbros. There are at least four massive oxide intervals with continuity between the two holes (600 m apart along strike). It is likely that these units pinch and swell, evidenced by variations in unit thickness and strong shearing at oxide layer contacts. In addition, 200 thin sections from field samples and drill core are being viewed in both reflected and transmitted light to delineate mineralogy, texture, alteration and, most importantly, to establish oxide mineralogy. Oxide layers were originally thought to be hosted by pyroxenite or gabbro, but petrographic analysis shows that the rocks are significantly altered and metamorphosed. Chlorite, talc, and amphibole alteration is dominant, though quartz-carbonate veining is also common. Photomicrographs of samples taken from the “top” of the intrusion indicate that primary apatite is strongly associated with massive oxides in the upper zones of the BVI. The most common oxides are expected to be secondary magnetite and ilmenite after primary titanomagnetite which will be verified through reflected light petrography. REFERENCES Lee, A., Albers, P., Miller, J., Severson, M., Deen, T., 2012, Bedrock Geologic Map of the Seine Bay/Bad Vermilion Lake Intrusion, Mine Centre, Ontario. In: Hollings, P. (Ed.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 1 – Proceedings and Abstracts, v. 58, part 1, 1-2. Ontario Geological Survey, 2009. Ontario airborne geophysical surveys, Magnetic and Electromagnetic Surveys, Grid and Profile Date (ASCII and Geosoft Formats) and Vector Data, Mine Centre Area;. Ontario Geological Survey, Geophysical Dataset 1061a ISBN 978-1-4435-0854-4 [DVD] ISBN 978-1-4435-0855-1[ZIP FILE]. Poulsen, K. H., 1984. The Geological Setting of Mineralization in the Mine Centre – Fort Frances Area, District of Rainy River; Ontario Geological Survey Open File Report 5512, 126 p.  White, C. R. & Albers, P. B., 2010. Report on the Geology and Mineral Potential of the Seine Bay/Bad Vermilion Lake Intrusion, Mine Centre Property, Mine Centre, Ontario, s.l.: Unpublished report prepared for Numax Resources, Inc.  55 �The Igneous Stratigraphy of the Bad Vermilion Intrusion, Mine Centre, Ontario, Canada: Which Way is Up? LEE, Aubrey, and MILLER, Jim Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812. The Bad Vermilion Intrusion (BVI) is a layered plutonic sequence of gabbroic rocks exposed along the shores of Seine Bay (Rainy Lake) and Bad Vermilion Lake in Northwest Ontario, Canada (White and Albers, 2010). The intrusion is the subject of a field, petrographic, and geochemical study as Aubrey Lee’s master’s thesis at the University of Minnesota-Duluth. The BVI is sill-like, approximately 14 km long and ~1 km wide, and stretches along the northern shores of Seine Bay in the southwest and Bad Vermilion Lake in the northeast with the town of Mine Centre marking its northeast extent. The BVI lithologically consists of a mafic layered intrusive package of altered and metamorphosed anorthositic, gabbroic, and pyroxenitic rocks with layers/lenses of semi-massive to massive oxides. Primary apatite occurs near the upper (northern) zones and is strongly associated with massive oxide units. Within the east-west elongate block bounded by the Quetico and Rainy Lake-Seine River fault zones (Poulsen, 1986), the BVI dips sub-vertically (85˚ N-NW shallowing to 60˚ N-NW in central region) between felsic intrusive and mafic volcanic rocks (Fig. 1). The Mud Lake Trondhjemite to the north and the Bad Vermilion Tonalite to the south (along Bad Vermilion Lake) are concordant with the BVI while the mafic volcanic rocks (chiefly basalt) to the south (along Seine Bay) are weakly discordant. The nature of these contacts and the magmatic relationship between the BVI and the country rocks is poorly known. It is also unclear as to whether the intrusion tops to the north or south. Bedrock mapping (Lee et al., 2011), magnetic surveys, and petrographic analysis helped to establish subdivisions within the intrusion (Fig. 1). This was also made possible by the fact that the sub-vertical dip of the BVI exposes the entire stratigraphy of the intrusion in near-true thickness. On a large scale, the region can be divided into a Mafic Layered Series, bounded on the northern side by the Mud Lake Trondhjemite, in the southeast by the Bad Vermilion Tonalite, and in the southwest by mafic volcanics. The Mafic Layered Series can be further subdivided into three zones; the Leucogabbroic Zone to the south, the central Gabbroic Zone, and the Melagabbroic Zone to the north. Although all zones contain oxide-bearing lithologies throughout, each zone contains an oxide “unit” that is composed of multiple layers of massive to semi-massive Fe-Ti oxides Figure 1: Idealized igneous stratigraphy of the BVI showing the main zones of the mafic layered concentrated in discrete intervals, typically toward the base of the zones. The oxide “units”, series and the approximate locations of the three which range from 20 – 200 m in thickness, all oxide units. Vertical scale is approximate. 56 �contain more than one layer/lens of massive oxide which are typically 10 cm – 1 m thick. However, one oxide unit within the Melagabbroic Zone is up to 30 m thick, as observed in core and field exposures. Though all three oxide units are evident in the western and central areas of the BVI (White and Albers, 2010), they thin and appear to coalesce toward the northeast (Lee et al., 2011). Geochemical and petrographic analysis will hopefully help to resolve how the eastern extent of the BVI correlates with the thicker sequences to the west. In White and Albers’ (2010) mapping of the central and western BVI, they interpreted the southern Mafic Volcanic Sequence and Bad Vermilion Tonalite to be the footwall and the Mud Lake Trondhjemite the hanging wall of the BVI. North topping was originally determined based on graded layering of mafic to feldspathic lithologies which occur throughout the intrusion. This is especially evident in the northernmost Melagabbroic Zone near the contact with the Mud Lake Trondhjemite, and can locally be traced for up to 400 m. Typically, the base of each layer is dominated by clinopyroxene with plagioclase abundance increasing upward and a sharp contact with the base of the next layer. Moreover, apatite abundance in the massive oxides of the Melagabbroic Zone indicates P-enrichment to the north. However, if the BVI does indeed top to the north, one would expect the abundance of plagioclase to increase from south to north on a large scale. Instead, the general south-to-north progression is from plagioclase-rich rocks (Leucogabbroic Zone) to mafic rocks (Melagabbroic Zone) and may indicate that the intrusion is slightly overturned and that the Mud Lake Trondhjemite is actually the footwall of the BVI. The nature of the contacts between the BVI and its bounding rocks are difficult to interpret. There is no evidence for faulted contacts. Moreover, the felsic intrusive rocks on either side of the BVI are clearly not upper differentiates given their tonalitic/trondhjemitic composition and the abrupt transition from mafic to felsic lithologies. Rather, the mafic volcanics at the southwestern contact of the BVI show effects of contact metamorphism and thus imply an intrusive contact (White and Albers, 2010). The contact between the BVI and the Bad Vermilion Tonalite to the southeast is concealed beneath Bad Vermilion Lake, but has also been considered intrusive by Poulsen (1984). The nature of the northern contact between the BVI and the Mud Lake Trondhjemite is more enigmatic. Although the trondhjemite is concordant with the BVI, grades into the melagabbro zone over a narrow interval of quartz gabbro, and has the same strike length as the BVI, it is petrologically unlikely that fractionation of the BVI would generate a trondhjemitic differentiate. Rather, given that the trondhjemite is bounded on the north by an extensive region of felsic volcanics (Poulsen, 1984), it is possible that the trondhjemite represents a partially melted zone generated by underplating of the hot BVI magma beneath the felsic volcanics. Narrow compositional grading from melagabbro to quartz gabbro to trondhjemite at the northern contact of the BVI is perhaps evidence for a gradational partially melted upper contact, but back-veining of the trondhjemite into the BVI supports the idea that the trondhjemite is actually the basal intrusive contact. This latter interpretation also fits better with the southward mafic to felsic progression of BVI zones. It is hoped that evaluation of geochemical cryptic layering through the BVI will help resolve which way is up. REFERENCES Lee, A., Albers, P., Miller, J., Severson, M., Deen, T., 2012, Bedrock Geologic Map of the Seine Bay/Bad Vermilion Lake Intrusion, Mine Centre, Ontario. In: Hollings, P. (Ed.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 1 – Proceedings and Abstracts, v. 58, part 1, 1-2. Poulsen, K. H., 1984. The Geological Setting of Mineralization in the Mine Centre – Fort Frances Area, District of Rainy River; Ontario Geological Survey Open File Report 5512, 126 p.  White, C. R. & Albers, P. B., 2010. Report on the Geology and Mineral Potential of the Seine Bay/Bad Vermilion Lake Intrusion, Mine Centre Property, Mine Centre, Ontario, s.l.: Unpublished report prepared for Numax Resources, Inc.  57 �2012 Precambrian Field Camp Mapping in the Wilder Lake Intrusion, Lake County, Northeastern Minnesota LEU, Adam, DJON, Lionel, LaPIETRA, Emily, MARTIN, Zach, MARTINEZ, Ricardo, and MILLER, Jim Precambrian Research Center, University of Minnesota Duluth, Duluth, MN 55812 In the summer of 2012, the Precambrian Research Center of the University of Minnesota-Duluth held its sixth annual Precambrian field camp in northeastern Minnesota. As in years past, the fifth and sixth weeks of the camp are dedicated to student’s “capstone” mapping projects during which detailed geologic mapping is conducted in areas of poorly understood geology and digital geologic maps are generated. Three of last summer’s capstone projects focused on areas of the Duluth Complex affected by the Pagami Creek fire, which burned a 160-squaremile area within the Boundary Water Canoe area wilderness in the fall of 2011. The intense burn created a timesensitive opportunity to map inland exposures that were previous difficult to access due to thick woods and blowdown areas created from a 1999 windstorm. Our capstone mapping project focused an area centered on the Wilder Lake Intrusion (WLI), which had been previously reconnaissance mapped only along shoreline exposures (Phinney, 1972; Miller, 1986; Turnbull and Miller, 2004). When it was recognized that the most intense burn area of the Pagami Creek fire was centered in the Wilder Lake area, Jim Miller successfully applied for a USGS EDMAP grant to have a UMD graduate student map the newly created exposures of the WLI. He recruited Adam Leu, an alum of the 2011 PRC field camp, to make this mapping project the centerpiece of his MS thesis at UMD and to be a teaching assistant for the 2012 Precambrian field camp. The objective of the WLI capstone mapping project was to map now easily accessible inland areas of the western WLI. This new mapping would be intergrated with the previous shoreline mapping of Miller (1986) and Turnbull and Miller (2004) to create more complete and detailed geologic map of the western part of the WLI. This mapping would also serve as the anchor point for continued mapping of the WLI by Adam to the east. This area has no lake access and therefore could only be mapped by overland traverses. Adam completed much of the eastward mapping in the fall of 2012, and will return next summer to finalize the mapping. What will be presented at this poster presentation is the 1:24,000-scaled geological map created from the capstone mapping of a two square-mile area of the western end of the WLI (LaPietra et al., 2012). A pdf version of this and other Precambrian field camp capstone geologic maps can be downloaded from the PRC website: www.d.umn.edu/prc/fieldcamp/capstone. The Wilder Lake intrusion (WLI) is a mafic layered intrusion emplaced within the anorthositic series, and is part of the layered series of the Duluth Complex (Miller et al., 2002). The few studies conducted on the WLI show it to be one of the most distinctive intrusions of the layered series by virtue of its northward dip, emplacement entirely within the anorthositic series, reversed cryptic variation, and unique cumulate stratigraphy (oxide before augite) (Miller and Ripley, 1997; Miller et al., 2002). The WLI was first recognized by reconnaissance mapping by Phinney (1972), who documented exposures of well-foliated and layered gabbros and troctolites that extend from North Wilder Lake to the west and Arrow Lake to the east; a strike-length of about 10 kilometers. Phinney noted that internal layering and foliation dips to the north-northeast between 15° and 35°, which contrasts with the southerly to easterly (riftward) dip of most layered intrusions of the Duluth Complex. Miller (1986) conducted reconnaissance mapping of the western extent of the WLI in the Wilder Lake area and term the body the Wilder Lake gabbro. Miller and Ripley (1997) reported olivine and augite data which define a reversed cryptic variation in mg# for both phases. Unpublished mapping and geochemical data from Joy Turnbull 58 �acquired in 2002-04 verified the basic cumulate stratgraphy identified by Miller (1986) in the Wilder Lake area and the reversed cryptic variation defined by olivine and augite. Turnbull’s mapping in the South Wilder Lake area (Turnbull and Miller, 2004) revealed a variation in the thickness of cumulate units and the occurrence of a lower olivine gabbro unit that is not evident in the western part of the intrusion mapped previously by Miller (1986). 2012 capstone mapping verified the basic cumulate stratigraphy of the approximately 1-2 km-thick intrusion previously describe by Miller (1986), but added considerably more detail and insight to the origin of that stratigraphy.. As noted previously by Turnbull and Miller, (2004), the basal contact of the WLI is a varitextured olivine gabbro to augite troctolite in sharp contact with coarse-grained anorthositic series rocks. This taxitic basal unit grades into a thick sequence of troctolitic (Pl+Ol) cumulates. The troctolite progress from a lower ophitic augite interval, which locally shows well developed modal layering, into a more homogeneous troctolite containing abundant anorthositic series inclusions. Overlying the troctolite unit is a thin (20-30m-thick) oxide troctolite unit defined by the abrupt cumulus arrival of Fe-Ti oxide. This unit is then overlain by a well foliated, four-phase olivine oxide gabbro cumulate (Pl+Cpx+Ox+Ol). Miller (1986) noted that the four-phase cumulate abruptly gives way back to a troctolite and puzzled as to the significance of this cumulus regression. He speculated that it could represent a recharge event, or perhaps crystallization of troctolite from the roof zone down to a sandwich horizon represented by the four-phase olivine oxide gabbro. Detailed mapping this past summer showed unequivocally that the upper troctolite is a latter intrusive pulse that actually cut down into the four phase gabbro and is locally in contact with the oxide troctolite unit. Another revelation from last summer’s mapping is that the western contact of the WLI is not a fault as speculated by Miller (1986), but rather is an intrusive contact between the WLI and anorthositic series rock, wherein a taxitic margin is developed along the WLI. Adam Leu’s MS thesis will integrate this capstone map along with previous mapping by Phinney, Miller and Turnbull and new mapping of now well exposed areas in the eastern WLI in order to piece together a complete geological picture of this unique intrusion. This research aims to verify and evaluate the origin of some of its enigmatic petrologic features (early cumulus oxide arrival, reversed cryptic variation) by conducting detailed field mapping and sampling, petrographic analyses, mineral chemical analyses, and lithogeochemical analyses from a suite of samples collected along several profiles across the intrusion. References Cited LaPeitra, E., Martin, Z., Martinez, R, Leu, A., Djon, L., and Miller, J., 2012, Bedrock geologic map of the Wilder Lake intrusion, Lake County, Northeastern Minnesota. UMD Precambrian Field Camp Map Series, PRC Map 2012-4, 1:24,000. Miller, J.D., Jr., 1986, The geology and petrology of anorthositic rocks in the Duluth Complex, Snowbank Lake quadrangle, northeastern Minnesota. unpublished Ph.D. dissertation, University of Minnesota, Minneapolis, 280 p. Miller, J.D., Jr., 1999, Geochemical evaluation of platinum group element (PGE) mineralization in the Sonju Lake intrusion, Finland, Minnesota. Minnesota Geological Survey Information Circular 44, 32 p. Miller, J.D., Jr., and Ripley, E.M., 1996, Layered intrusions of the Duluth Complex, Minnesota, USA. In Cawthorne, R.G., ed., Layered Intrusions: Amsterdam, Elsevier Science, p. 257-301. Miller, J.D. Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.E., and Wahl, T.E., 2002, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota. Minnesota Geological Survey Report of Investigations 58, 207p. w/ CD-ROM Phinney, W.C., 1972, Northwestern part of Duluth Complex. In: Sims, P.K. & Morey, G.B. (eds.) Geology of Minnesota -A Centennial Volume. Minnesota Geological Survey, p. 335-345 Turnbull, J.A. and Miller, J.D., Jr, 2004, Preliminary geological map of the Wilder Lake Intrusion of the Duluth Complex, Lake County, Minnesota. Unpublished map submitted to USGS EDMAP program, 1:24,000. 59 �Silica Remobilization in the Biwabik Iron Formation, Minnesota USA LOSH, Steven and RAGUE, Ryan Dept. of Chemistry and Geology, FH 241, Minnesota State University, Mankato MN 56001 Excess silica in magnetite separated from iron formation in the Mesabi Range of northern Minnesota can diminish the quality of taconite pellets made there. Normally, silica comprises less than 5 wt% of magnetic concentrate, a desirable level, but it can exceed 10% in some instances. The source and nature of this silica has been enigmatic; it has been noted in and near oxidized zones and in particular stratigraphic units in the mined iron formation, but it has proven difficult to characterize in terms of its nature and distribution. To better understand silica behavior in magnetite ore, we applied petrographic, SEM, fluid inclusion, and bulk geochemical methods to samples collected from three mines, the Hibbing Taconite Mine, the Thunderbird Mine, and the Fayal Reserve Mine, to document the occurrence and origin of quartz that is likely included in magnetite in the separation process. We found that silica-filled microfractures and pits typically form in coarse magnetite within faulted iron formation near both low-angle and high-angle faults, the latter being everywhere associated with oxidation. Silica reprecipitation accompanied early oxidation, in which magnetite was variably oxidized to hematite (martite) and iron silicates were oxidized to goethite and recrystallized quartz near high-angle faults. In unoxidized magnetite in and near low-angle (bedding-parallel) faults, diagenetic inclusions and microfractures host quartz. For both highand low-angle faults, this microfracturing/silica remobilization event took place at diagenetic temperatures (150° – 200° C) and involved oxidizing, relatively high-salinity aqueous (late diagenetic) fluids. In contrast to quartz retention and recrystallization that is the hallmark of the early oxidation event, later supergene (lateritic) oxidation, the stage that produced the high-grade hematite/goethite ‘natural ores,’ extensively dissolved quartz. In magnetite grains in iron formation near faults, silica-filled microfractures are typically less than 5 microns in width; silica-filled pits are commonly less than 40 microns across, many less than 20. Ore is typically ground to -325 mesh (-44 microns), significantly coarser than these quartz inclusions. Image analysis shows that these features can comprise up to 11% of the magnetite by volume, potentially accounting for much of the excess silica in concentrate.   60 �The McGrath metasaprolite: viewing Paleoproterozoic weathering through a veil of metamorphism and metasomatism MEDARIS, Gordon Jr1, BOERBOOM, Terry2, JICHA, Brian1 and SINGER, Brad1 1 Department of Geoscience, University of Wisconsin-Madison, Madison, WI 53706 2 Minnesota Geological Survey, St. Paul, MN 55114 medaris@geology.wisc.edu, boerb001@umn.edu, bjicha@geology.wisc.edu, bsinger@geology.wisc.edu Paleosols are important indicators of ancient weathering processes and climatic conditions. However, many Precambrian paleosols have been metamorphosed and metasomatized, thereby obscuring their pedogenic features and modifying their original chemical compositions. A metamorphosed saprolite occurs in eastern Minnesota in the Archean McGrath Gneiss beneath the Paleoproterozoic Denham Formation. A detailed investigation has been undertaken of this metasaprolite to distinguish the effects of weathering from those of subsequent metamorphism and metasomatism and to evaluate the depth, characteristics, and magnitude of weathering. The McGrath Gneiss is granitic in composition, containing quartz, microcline, plagioclase, biotite, and muscovite. The upper ~400 cm of gneiss is devoid of plagioclase and rich in muscovite, consisting of quartz, microcline, muscovite, and minor biotite, and is interpreted to be a metasaprolite, from which plagioclase has been removed by weathering. The base of the Denham Formation is composed of metamorphosed siltstone, arkosic sandstone, and lenses of pebble conglomerate that contain abundant microcline grains and gneiss fragments derived from the underlying McGrath Gneiss. The age of weathering is thought to be ~2100 Ma, as constrained by the age of the McGrath protolith (2557 ± 15 Ma; Holm et al., 2005), the youngest age of detrital zircon from the base of the Denham Formation (2072 ± 17 Ma; Wirth et al., 2006), and the age of volcanic rocks in the Denham Formation (2197 ± 39 Ma; Beck, 1988). Metamorphism in the area, which reaches staurolite grade, was largely a Yavapai event, based on geon 17 U-Pb and 207Pb-206Pb ages of monazite and xenotime in the Little Falls Formation (Schneider et al., 2004; Holm et al., 2007), which overlies the Denham Formation. In a plot of molar Al2O3 (CaO+Na2O)K2O, a.k.a. ACNK, the compositions of metasaprolite deviate markedly from the trend expected for weathering of the McGrath Gneiss (Fig. 1). Such deviation is ascribed to K metasomatism, in which kaolinite formed by weathering has been transformed to muscovite by the introduction of K. Potassium metasomatism is a common phenomenon in paleosols and is exhibited by many Precambrian saprolites in the Lake Superior region (Medaris et al., 2012). Step heating of muscovite from the uppermost sample of metasaprolite (depth = 80 cm) yields a 40 Ar/39Ar spectrum with a well-defined plateau at 1742 ± 3 Ma, which includes 88% of 39Ar released. Thus, K metasomatism was probably associated with Yavapaiage metamorphism, although the 40Ar/39Ar result only places an upper limit on the age of muscovite growth, because metamorphism occurred at temperatures above the blocking temperature for Ar diffusion in muscovite. 61 �Potassium metasomatism precludes use of the Chemical Index of Alteration as an indicator of the intensity of weathering in the McGrath metasaprolite. Instead, the Plagioclase Index of Alteration [molar (Al2O3-K2O)/(Al2O3+CaO+Na2O+K2O)], which is a measure of plagioclase removal, may be used and yields values increasing from 67.4 at a depth of 370 cm to 93.5 at 80 cm. The % changes in selected oxides and elements with depth, relative to Al 2O3 in the mean McGrath Gneiss, are illustrated in Fig. 2, in which the compositional trends reveal that weathering extended to a depth of ~400 cm. The removal of CaO, Na2O, and Sr reflects the weathering of plagioclase, and the addition of K2O, Ba, and Rb is attributed to post-weathering metasomatism. SiO2 has been substantially removed over much of the weathering profile (10 to 20%), as has P2O5 (10 to 75%). The three protolith samples appear to be heterogeneous in their Fe 2O3 contents, with little significant change in Fe2O3 in the weathering profile, except for the uppermost sample. The magnitude of weathering of the McGrath metasaprolite, i.e. the mass removal of SiO2, CaO, and Na2O integrated over the depth of weathering, is 2.2 moles/cm 2. Assuming a duration of weathering of 100,000 years and following Sheldon's (2006) method, atmospheric pCO2 was 5 times that of the pre-industrial level. Although the McGrath metasaprolite represents an important episode of Paleoproterozoic weathering in the Lake Superior region, its degree of weathering was less than that of some other Paleoproterozoic paleosols. For example, the ~1700 Ma Baraboo paleosol has a larger magnitude of weathering, 4.3 vs. 2.2 moles/cm2, greater depth of weathering, 7.9 vs. 4 meters, and complete removal of potassium feldspar in addition to plagioclase, perhaps reflecting a warmer and wetter climate for Baraboo. References Beck JW (1988) University of Minnesota Ph.D. dissertation, 273 pp; Holm DK et al. (2005) Geological Society of America Bulletin 117, 259-275; Holm DK et al. (2007) Precambrian Research 157, 106-126; Medaris LG et al. (2012) 58th ILSG, Proceedings with Abstracts, 60-61; Schneider DA et al. (2004) Geological Society of America Special Paper 380, 339-357; Sheldon ND (2006) Precambrian Research 147, 148-155 62 �Chemical Zoning in Calc-Silicate Minerals Associated with Native Copper from the Keweenaw Peninsula, Michigan MULCAHY, Connor1, HANSEN, Edward1, BORNHORST, Theodore2, and RHEDE, Dieter3 1 Geological and Environmental Sciences Department, Hope College, Holland, Michigan, 49423 2 A.E. Seaman Mineral Museum, Michigan Technological University, Houghton, MI 49931 3 Helmholtz-Zentrum Potsdam, Deutsches GeoForschungsZentrum (GFZ), Telegrafenberg, D14473, Potsdam, Germany Calc-silicate minerals occur in clusters in hydrothermally altered basalts and rhyolite-pebble conglomerates associated with native copper deposits in the Keweenaw Peninsula. We studied chemical zoning in these minerals hosted by basalts from the Kearsage, Isle Royale, and Caledonia Mines and by rhyolite-pebble conglomerates and sandstones. All basalt and some conglomerate samples came from the research rock collection of the A.E. Seaman Mineral Museum, Michigan Tech. Additional conglomerate samples came from personal collecting from rock piles at the base of Bumbletown Hill. Samples were examined with the Scanning Electron Microscope at Hope College. Quantitative mineral analyses were carried out using a JEOL JXA-8500F field emission electron microprobe (hyperprobe) at the Deutsches GeoForschungsZentrum, Potsdam (GFZ). Calc-silicate masses in the basalt samples consist of pumpellyite, titanite, and epidote +/- prehnite. Veins of epidote cut pumpellyite, and pumpellyite occurs as inclusions in epidote, indicating that the formation of pumpellyite preceded the growth of epidote. Variations in brightness can be seen within epidote, prehnite and pumpellyite grains in SEM backscatter images (Figure 1, 2, and 3). Electron microprobe analyses indicate that the brighter regions are characterized by higher Fe/(Fe + Al) (fe ratio) (Figures 1, 2 and 3). In epidote these bright regions with relatively high fe ratios frequently exhibit regular geometric forms (Figure 1). Within these forms the fe ratio is relatively constant but decreases rapidly at the boundaries. Zones with regular geometric forms correspond to crystallographic faces (Figure 1) and areinterpreted as sector zoning in which the variation in fe ratios is due to the propensity of different crystal faces to incorporate different amounts of iron during mineral growth. However, some epidotes contain irregularly shaped bright patches and electron microprobe traverses across these patches show gradational changes in the fe ratio. This zoning may represent changes in the compositions of hydrothermal fluids during the growth of epidote. Zoning of rare earth elements (REE) in epidote was observed in samples from the Isle Royale mine where approximately rectangular patches are enriched in REE within larger epidote crystals (Figure 4). Hour glass zoning has been observed in pumpellyite grains (Figure 3) and this is also interpreted as sector zoning. Zoning in prehnite (Figure 2) is perpendicular to the elongated axes, and increases in the fe ratio may be a result of an increase in temperature, pH, the activity of Fe+3, or the activity of Ca+2 in the fluid phase. Calc-silicate clusters in conglomerates consist of epidote + titanite +/- andraditic garnet. Zoning was not detected in the garnet. The fe ratio zoning patterns in epidote hosted by conglomerate are very similar to patterns in epidote hosted by basalts. REE-zoning was found in epidote grains. Narrow zones enriched in REE outline crystal forms of epidote and are interpreted as a relatively brief increase in activities of REE elements during growth of the epidote (Figure 5). REE–enriched zones in epidote grains either at their margins or along fractures are interpreted as the replacement (by dissolution recrystallization) of ordinary epidote by REE-enriched epidote after the main period of hydrothermal epidote growth (Figure 6). This replacement could be due to a decrease in temperature, but could also be due to a decrease in the activities of F or Cl both of which form complexes with REE. The REE flurocarbonate synchysite was 63 �found from the Allouez mine. REE minerals appear to be more abundant in the conglomerate host rocks than in basalt host rocks suggesting that the REE may have been remobilized from the rhyolite clasts. Figures: Back-scatter-SEM images of calc-silicate minerals. The graphs below each image give either Fe/(Fe+Al) (fe ratios) or total REE concentrations (in wt%) across the traverse (A-B) marked in the image. Figure 1: Fe-Al sector zoning in epidote grains from the Kearsage mine. Figure 2: Fe-Al zoning parallel to growth direction in prehnite from the Kearsage mine. Figure 3: Fe-Al zoning in an epidote grain containing a pumpellyite inclusion with hourglass sector zoning from the Kearsage mine. The fe ratios of the two analyzed spots on the pumpellyite grain are given in white lettering directly on the image. Figure 4: REE-zoning in epidote from the Isle Royale mine. Figure 5: Epidote from a conglomerate showing narrow growth zones enriched in REE (bright lines). The broader zones (lighter to darker gray) with sharp rectilinear boundaries reflect variations in the fe ratios. Figure 6: Epidote grain from a conglomerate showing REE-enriched, BSE-bright veins following fractures. 64 �Contrasting pressure-temperature-time paths for high-grade metamorphic rocks in the interior of the Penokean-Yavapai orogenic belt, southern Lake Superior region NADZIEJKA, Brynley and BJØRNERUD, Marcia Geology Department, Lawrence University, Appleton, Wisconsin, 54911 USA Amphibolite-facies metamorphic rocks occur in a narrow band in the internal part of the Paleoproterozoic orogenic belt south of Lake Superior in Michigan, Wisconsin and Minnesota. Peak metamorphic temperatures were once thought to have been reached during Penokean tectonism (ca. 1830-1850 Ma), but recent monazite and 40Ar/39Ar dating has shown that these relatively high-grade rocks have a more complex thermal history that spans both the Penokean and Yavapai (ca. 1760-1740 Ma) events (Holm et al., 2007). Furthermore, the timing and causes of the thermal maxima appear to be different in different parts of this tectonic zone. The ‘metamorphic nodes’ in the Upper Peninsula of Michigan, such as the one near Republic, are classic examples of Barrovian metamorphism, with peak temperatures of 550-615°C and pressures of 0.2-0.3 GPa (Attoh & Klasner, 1989). The current consensus is that the nodes record post-Penokean development of gneiss domes and juxtaposition, via normal-sense shear zones, of sedimentary rocks against hot, remobilized Archean basement (Tinkham & Marshak, 2004). Gravitational instability may have been enhanced by tectonically thickened piles of dense iron formation. Monazite ages from the highest-grade rocks at Republic suggest that the thermal maximum occurred in Yavapai time at 1758-1768 Ma (Holm et al., 2007), consistent with petrographic observations that porphyroblasts overprint Penokean foliations. The full P-T-t path for these rocks would be clockwise, with Penokean burial preceding Yavapai heating. In east central Minnesota, rocks of the Denham Formation experienced comparable peak temperatures (520-590°C) but higher peak pressures (0.5-0.6 GPa) (Holm & Selverstone, 1990). Petrographic analyses point to one main period of porphyroblast growth, but monazite and zircon dates indicate distinct thermal pulses at ca. 1830, 1800, and 1780 Ma, with the third pulse apparently related to a combination of Yavapai doming and emplacement of the East-Central Minnesota batholith (Holm et al., 2007; Boerboom, 2010). The P-T-t path for these rocks is thus more complex than for the Republic node, but still consistent with an overall clockwise trend. The relatively high-pressure kyanite-bearing rocks of the Watersmeet Terrane in northern Wisconsin seem to have a significantly different metamorphic history from those in Michigan and Minnesota. In samples of schist taken near Powell, Wisconsin, the primary minerals are biotite, quartz, albitic feldspar, garnet, kyanite and rare staurolite. The rather coarse (2-3 mm) biotite grains define a crude planar fabric, with garnet, feldspar, and staurolite occurring in biotite-poor domains. Muscovite is notably absent, and quartz is less abundant than feldspar. The few staurolite crystals in the samples tend to be fragmented, and the feldspars typically have serrate grain boundaries, suggesting they have undergone partial recrystallization following ductile deformation. The garnets are relatively small (ca. 1 mm) and contain abundant inclusions, though not well-defined inclusion trails, of quartz, biotite, and in at least one case, staurolite. The kyanite crystals are as long as 4 cm and show a weak alignment. Most significantly, kyanite grains contain inclusions of quartz, biotite, feldspar and garnet and clearly overprint the planar fabric. Based on these textural relationships, we infer that staurolite and feldspar grew early in the metamorphic history, followed by garnet and finally kyanite. Biotite was likely an early phase that continued to grow over time as the later minerals formed. Most of the deformation occurred prior to the formation of both the garnet and especially the kyanite. Garnet-biotite geothermometry on rocks from the Powell area by Geiger and Guidotti (1989) placed Tmax in the range 630-680° C. Using the garnet-plagioclase-Al2SiO5-quartz barometer and the coexistence of sillimanite and kyanite in some specimens, they estimated Pmax at 0.75 GPa. Monazite dates from the Watersmeet terrane samples record two major thermal events at 1830 Ma and 1765 Ma (Holm et al., 2007). The fact that kyanite clearly postdates all the other metamorphic phases seems to preclude a clockwise P-T-t path for the rocks of the Watersmeet terrane. Unless 65 �nucleation and growth of kyanite had for some reason been kinetically suppressed earlier in the rocks’ history, the textural relationships require that at least one of the thermal maxima preceded the eventual pressure maximum. In addition, the combined geochronologic and textural constraints indicate that the pressure maximum post-dates Penokean time. These observations must be incorporated into orogenic models for the Penokean and Yavapai events. Schneider et al. (2004) proposed that orogenic ‘channel flow’ in Yavapai time could account for the exhumation of the high-pressure rocks of the Watersmeet terrane. They depict the channel as bounded by a south-dipping normal fault at the southern boundary of the terrane and a southdipping reverse fault at the northern boundary. Crustal thickening in Yavapai time could account for the late high-pressure metamorphism recorded by the petrographic relationships, but the kyanite-bearing rocks would initially have had to be in the footwall beneath the channel to experience elevated pressures, then subsequently incorporated into the channel, perhaps by northward migration of the lower channel boundary. Ductile extrusion of the Watersmeet terrane is predicted by the channel flow model, but the rocks from the Powell area show little evidence of deformation after the formation of the kyanite. However, deformation in channel flow is more concentrated at the boundaries, and it could be that the available outcrops represent the central part of the channel. The channel flow model for northern Wisconsin is also an imperfect match with the geometry of the ductile channel in the modern Himalaya, where the phenomenon was first recognized. In the Himalaya, the channel boundaries (Main Central Thrust, South Tibetan Detachment; Beaumont, et al. 2001) dip in the same direction as the subduction zone, while in the Watersmeet terrane, the proposed boundaries dip south while Yavapai subduction is thought to have been north-directed. Another requirement of the channel flow model is rapid erosion of the extruded rock mass. One interesting implication of the model could be the interpretation of the Baraboo Quartzite as the distal part of a clastic wedge formed by intense erosion of the Yavapai orogen. This is consistent with the ages of detrital zircons from the Baraboo and correlative quartzites (1782-1712 Ma) and with evidence that these units were deposited in a warm, humid climate (Medaris et al., 2003). Whether or not the channel flow model is appropriate for the Watersmeet terrane, there remains the question of why the P-T-t path for metamorphic rocks in northern Wisconsin differs so markedly from those in Michigan and Minnesota. Differences in the plate boundary configuration, location within the orogen, and nature of the Archean basement are among the factors that should be explored. Attoh, K. and Klasner, J., 1989. Teconic implications of metamorphism and gravity field in the Penokean orogen of northern Michigan. Tectonics, 8: 911-933. Beaumont, C., Jamieson, Ngyuen, M., and Lee, B., 2001. Himalayan tectonics explained by extrusion of a low-viscosity crustal channel coupled to facued surface denudation. Nature, 414: 738-742. Boerboom, T., 2010. Transect from Archean basement to the Animikie basin, east-central Minnesota. Institute on Lake Superior Geology Field Trip Guidebook, 57: 129-161 Geiger, C. and Guidotti, C., 1989. Precambirn metamorphism in the southern Lake Superior region and it bearing on crustal evolution. Geoscience Wisconsin, 13: 1-33. Holm, D., Schneider, D., Rose, S., Mancuso, C., McKenzie, M., Foland, K., and Hodges, K., 2007. Proterozoic metamorphism and cooling in the southern Lake Superior region, North America and its bearing on crustal evolution. Precambrian Research, 157: 106-126. Holm, D. and Selverstone, J., 1990. Rapid growth and strain rates inferred from synkinemative garnets, Penokean orogen, Minnesota Geology, 26: 166-169. Medaris, L.G., Singer, B. Dott, R., Naymark, A., Johnson, C., and Schott, R., 2003. Late Paleoproterozoic climate, tecotnics and metamorphism in the southern Lake Superior region and Proto-North America: Evidence form Baraboo interval quartzites. Journal of Geology, 111: 243-257. Scheider , D., Holm, D., O’Boyle, C., Hamilton, M., and Jercinovic, M., 2004. Paleoproterozoic development of a gneiss dome corridor in the southern Lake Superior region. GSA Special Paper 380: 339-357. Tinkham, D. and Marshak, S., 2004. Precambrian dome-and –keel structure in the Penokean orogenic belt of northern Michigan. GSA Special Paper 380: 339-357. 66 �Natural Groundwater Geochemistry in Bedrock of the Thunder Bay Area PUUMALA, Mark, Ontario Geological Survey, 435 James Street South, Suite B002, Thunder Bay, Ontario P7E 6S7 During 1978 and 1979, the Ontario Ministry of the Environment (MOE) carried out a groundwater quality sampling program in the Thunder Bay area (McMullen, 1985). This program involved the collection of 354 groundwater samples from private drinking water wells completed in both overburden (153 samples) and bedrock (201 samples). The data were used in the production of two groundwater resource evaluation reports for the privately-serviced rural areas of the City (McMullen, 1985; Trow Hydrology Consultants, 1988). Some significant geochemical variations between hydrogeologic units were noted in each report. However, limited work was done to understand the reasons for these differences. The purpose of this study was to take a more detailed look at the MOE data set to gain a better understanding of how the geology of the major bedrock formations in the Thunder Bay area influences groundwater geochemistry. The study area is located on the Canadian Shield, at the boundary between the Archean-age rocks of the Superior Province, and the Proterozoic-age rocks of the Southern Province. The northern half of the study area is underlain by Neoarchean-age metavolcanic, metasedimentary, and intrusive rocks (Brown, 1995), while the southern half is underlain by the flat-lying and relatively un-deformed Paleoproterozoic-age Animike Group sedimentary rocks of the Gunflint and Rove formations (Sutcliffe, 1991). These sedimentary rocks are intruded by Mesoproterozoic Logan diabase sills and dikes associated with the Midcontinent Rift (Sutcliffe, 1991). For the purposes of this study, the Archean-age rocks, Gunflint Formation, Rove Formation and diabase are considered to be four distinct hydrogeologic units. Although the Archean-age rocks include a diverse range of lithologies, they are grouped together because they are all sparselyfractured crystalline rocks with similar water-transmitting characteristics (McMullen, 1985). This study focussed on data from the 201 bedrock wells that were sampled by MOE in 1978-79. Each well was classified according to the 1:250 000 scale bedrock lithology mapped on the ground surface at that location (Ontario Geological Survey, 2011). The numbers of wells assigned to each hydrogeologic unit were as follows: 65 Archean; 81 Gunflint; 41 Rove; 3 Diabase. Eleven wells are in an area of deep overburden where the Gunflint/ Rove contact has not been defined and were classified as “Gunflint/Rove contact area” wells. All of the groundwater samples were analyzed for the following list of geochemical parameters: hardness, pH, colour, turbidity, conductivity, bicarbonate, sulphate, chloride, Na, K, Ca, Mg, Fe and Mn. 17 samples were also tested for the following additional list of parameters: Cd, Co, Cu, Zn, Pb, Ni, Hg and As. Data analysis for this study included the plotting of major ion data (Ca, Mg, Na, K, Cl, sulphate, bicarbonate) on Piper tri-linear diagrams, and the calculation of mean values and ranges for conductivity, hardness, chloride, sulphate, Fe and Mn. Both of these methods were used to evaluate the major ion geochemistry of each hydrogeologic unit, including variability and possible groundwater geochemical evolution trends. The results of this analysis showed that the Archean, Gunflint and Rove hydrogeologic units each have distinctive groundwater geochemical signatures that are attributable to differences in lithology. The data set for the diabase unit was too small to allow for any meaningful assessment of its groundwater geochemistry. 67 �Groundwater sampled from drinking water wells drilled into the Archean crystalline rocks typically had the lowest levels of dissolved solids and the least geochemical variability. The water is also typically of the Ca + Mg bicarbonate-type. The combination of low dissolved solids content and Ca + Mg bicarbonate dominated geochemistry is indicative of recently-recharged groundwater that has had limited time to interact with aquifer solids. This is consistent with what would be expected for a data set collected from relatively shallow wells in low permeability, competent bedrock units such as these. Although the major ion geochemistry of the Archean crystalline rock groundwater was relatively consistent regardless of lithology, there were some differences noted for iron and manganese, with higher mean concentrations in wells sourced from metasedimentary rock (possibly related to the weathering of biotite). Groundwater in the Gunflint Formation displays much more geochemical variability and has a much higher mean concentration of dissolved solids than in the Archean hydrogeologic unit. Although most groundwater in the Gunflint has a geochemical signature indicative of recent recharge, two important apparent geochemical evolution trends were noted. The dominant trend is from a Ca + Mg bicarbonate-type toward a Ca + Mg chloride-type groundwater. The second trend is toward a mixed cation chloride-type groundwater (i.e., higher relative proportion of Na). The second trend has a close spatial association with the argillite tuff horizon mapped by Moorhouse (1960), and is interpreted to be related to groundwater interaction with this stratigraphic unit. Because the Gunflint Formation is not known to contain evaporite minerals, the most likely source of chloride in this formation is connate brine that was trapped during diagenesis (Hem, 1985). Rove Formation groundwater has the highest mean dissolved solids content, and also shows significant geochemical variability. Similar to the Gunflint, there are two apparent geochemical evolution trends in the Rove. However, these trends are distinct from those seen in the Gunflint. The dominant trend is from Ca + Mg bicarbonate-type toward Ca + Mg sulphate-type groundwater, with a second trend toward sodium bicarbonate-type. The dominant trend toward a sulphate-type geochemistry is likely to be due to the weathering of pyrite, which is locally present in Rove Formation shale (Sutcliffe, 1991). The second trend is accompanied by elevated pH and may be related to the weathering of albite (Kehew, 2001). This trend also shows a close spatial association with diabase sills and appears to be characteristic of groundwater in altered/ metamorphosed contact zones between the Rove Formation and diabase. References Brown, G.H., 1995. Precambrian Geology, Oliver and Ware Townships. Ontario Geological Survey, Geological Report 294, 48p. Hem, J.D., 1985. Study and interpretation of the chemical characteristics of natural water. United States Geological Survey, Water-Supply Paper 2254, 263p. Kehew, A.E., 2001. Applied Chemical Hydrogeology. Prentice Hall Inc., Upper Saddle River, New Jersey, 368p. McMullen, R.F,. 1985. Groundwater potential in minimum service rural residential areas, City of Thunder Bay; Ministry of the Environment, 23p. Moorhouse, W.W., 1960. Gunflint Iron Range in the Vicinity of Port Arthur; Ontario Department of Mines, Volume 69, Part 7, 40p. Ontario Geological Survey, 2011. 1:250 000 scale bedrock geology of Ontario-revised. Ontario Geological Survey, Miscellaneous Release-Data 126-revision 1. Sutcliffe, R.H,. 1991. Proterozoic geology of the Lake Superior region: Geology of Ontario. Ontario Geological Survey, Special Volume 4, Part 1, p. 627-658. Trow Hydrology Consultants Ltd., 1988. A study of groundwater resources in the minimum service rural residential area designations of the City of Thunder Bay Official Plan, Thunder Bay, Ontario. unpublished report, 56p. 68 �North American Palladium’s Lac des Iles mine: Evidence for high temperature deformation and possible control on Pd mineralization SCHMIDT, Skylar and HILL, Mary Louise, Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, On, P7B5E1 Canada North American Palladium’s Lac des Iles Mine is located in a mafic to ultramafic intrusive complex just north of the Wabigoon-Quetico subprovince boundary. The Mine Block intrusion, host to the economic palladium mineralization, preserves evidence of a complex history of high-temperature deformation suggesting that the intrusive complex is pre- or syntectonic, not post-tectonic as commonly presumed. If so, deformation may be significant to mineralization and provide insight on future targets for exploration. The potential link between heterogeneous deformation and palladium enrichment will be further investigated. The Mine Block intrusion is elongate in a northeast-southwest direction, parallel to regional deformation and is composed of mainly gabbro-noritic rocks. Locations of investigation to date include the Baker zone, the North VT rim and the Sheriff zone. The Baker zone is the discovery outcrop of the property and is located near the centre of the intrusion as a topographic high. The North VT rim is located along the north edge of the intrusion and is characterized by variable grain size, and the Sheriff zone is an area of mineralization located to the south-east of the pit. In the Baker zone, northeast-striking mafic dikes are boudinaged and disaggregated, with narrow shear zones preserved along some of the boudin margins. In the Sherriff zone, felsic dikes and narrow ductile shear zones are mutually overprinting, indicating progressive brittle-ductile deformation. Economic Pd mineralization is commonly associated with chlorite-actinolite schist; some sulfide grains occur along cleavage planes in the metamorphic silicate minerals. The East Gabbro is mylonitized at the contact with the chlorite-actinolite schist and deformation decreases further from the contact. Microstructural analysis provides further evidence for deformation in the Mine Block intrusion. In the Baker zone, plagioclase has deformation twins and with some grains showing subsequent deformation of these twins. Subgrains in plagioclase, and small strain-free grains along margins of larger strained grains, provide evidence of dislocation creep in plagioclase in the Baker zone, Sherrif zone, and North VT rim. In the Baker zone, some sulfide grains have been fractured and healed, indicating subsequent deformation after initial formation. 69 �Effects of Preexisting Fractures on Groundwater Flow Today SCHMUS, Matthew1, BHATTACHARYYA1, Prajukti, and HART, David2 1 Department of Geography and Geology, University of Wisconsin-Whitewater, Whitewater, WI 53190 United States 2 Wisconsin Geologic and Natural History Survey, Madison, WI 53705 United States Crystalline rocks are not normally considered aquifers, but when joints and fractures are present, there is a very real chance for pathways to emerge as conduits for groundwater flow. Three bore holes were drilled in Pittsville, WI (Figure 1) within the Marshfield terrene. This area was the stage where orogeny and volcanism occurred when the Marshfield terrene collided with the Wausau terrene and Superior craton forming multiple episodes of deformation (Schulz and Cannon, 2007). Data from the boreholes was collected by Dr. David Hart, and was analyzed at University Wisconsin-Whitewater. The data includes orientations of five different surfaces, including (a), major and minor open joints or fractures, (b), partially open joints or fractures, (c), filled fractures or joints, (d), bedding, banding, or foliation planes, and (e), induced fractures. We investigated orientations of regional stress directions based on the fracture orientation data using T-Tecto software (Figure 2). We combined stereographic projection analyses and borehole gamma logs to investigate how the orientations of different types of fractures might have changed with depth, mainly focusing on the orientations of open and filled fractures as those might provide the best insights within past and present fluid flow patterns. We have determined the locations where the fractures might intersect with each other below surface by using apparent dip data of the fracture planes on vertical planes through any two of the three studied boreholes (Figures 3 and 4). Eventually we aim to create a three-dimensional model of the fracture network using three-dimensional visualization tools within the ArcScene® software package. Preliminary data shows that dominant fracture orientation patterns change with depth in each of the three bore holes, and the fracture orientation patterns in each of the three boreholes have little or no similarity with each other (Figure 5). Some of the fractures show evidence of past fluid flow in the form of filled veins. Since some of the fractures are not filled, this might indicate multiple episodes of fracture formation which would be consistent with the tectonic history of the area. This is also indicated by the T-Tecto plots, which show multiple stress directions (Figure 2). Also two dimensional models of fracture plane intersections have shown possible intersection points among different fracture types that may be conduits for groundwater flow. Here we will present our data, and discuss the potential implications of our analyses on understanding past and present groundwater flow in the studied region. References Hart, David J., (2011): Comparison of Groundwater Flows into Three Closely Spaced Crystalline Bedrock Wells. Geological Society of America Abstracts with Programs, Vol. 43, No. 5, p. 287 Schulz, Klaus, and Cannon, William (2007). The Penokean orogeny in the Lake Superior region. Precambrian Research, Volume 157, Issues 1-4, pages 4-25 70 �Figure 2: T Tecto Plot of all fractures in well 9 Figure 1: Map of Study Area Pittsville Wisconsin Figure 3: Calculations used to determine depth at which different fracture planes intersected the three boreholes Figure 4: 2D graph showing possible fluid flow connection. Figure 5: Stereographic Projections (left to right)Well 7 data filled fractures depth between 90-120 Well 8 data filled fractures depth between 80-110 Well 9 data filled fractures depth between 80-110 71 �The Parent Lake Volcanics: Product of a phreatomagmatic eruption of basalt during deposition of the Michigamme Formation? SCHULZ, K.J.1, CANNON, W.F.1, and WOODRUFF, L.G.2 1 U.S. Geological Survey, 954 National Center, Reston, VA 20192, kschulz@usgs.gov, wcannon@usgs.gov 2 U.S. Geological Survey, 2280 Woodale Ave., Mounds View, MN 55112, woodruff@usgs.gov. An unusual mafic volcanic layer occurs within an otherwise monotonous sequence of turbidites of the Michigamme Formation in a set of outcrops about 40 km south of L’Anse (NW ¼ Sec. 9 and NE ¼ Sec. 8, T. 48 N., R. 33 W), at the extreme west end of the Marquette trough in Michigan’s Upper Peninsula. The volcanic rocks are about 150 m thick and dip steeply (80-85 degrees) to the south. However, they are overturned and face north as indicated by both graded beds in adjacent graywacke and bedding-cleavage relations within the volcanic rocks. Typical Michigamme graywacke both underlies and overlies the mafic volcanic rocks, although the upper contact is intensely sheared and may not be a depositional contact. The mafic volcanic rocks are divided into two units. The lower unit is a fine-grained, mafic pyroclastic rock having fragments of highly vesicular basalt typically less than a few centimeters in long dimension. It is massive to crudely layered and shows greenschist metamorphic assemblages, including a colorless amphibole, chlorite, and clinozoisite as the most common minerals; secondary carbonate is abundant. The upper unit is a coarse breccia containing fragments of both vesicular basalt and exotic fragments of crystalline rocks, presumably derived from the underlying Archean basement and older Paleoproterozoic sedimentary rocks. These fragments are as much as 1 m diameter and commonly are well rounded. Lithologies include a variety of granitic rocks and quartzose sedimentary rocks. The breccia is crudely bedded in meter-scale layers; the matrix is intensely sheared and contains abundant secondary carbonate. A distinctive feature of some exotic fragments is a rind or shell of fine vesicular material that is more resistant to weathering than the matrix of the breccia (Fig. 1). These clasts look similar to cored bombs that formed during phreatomagmatic eruptions in some recent marr/diatreme complexes (Hanson and Elliot, 1996; Rosseel and others, 2006; Sottili and others, 2010). Cored bombs display a chilled shell of juvenile material, generally basalt, surrounding a lithic core, reflecting the thermal interaction of magma with fragments of wall rock in the vents of marr/diatreme complexes (Sottili and others, 2010). The highly vesicular, fragmental, and crudely bedded and graded nature of the Parent Lake mafic volcanic units, as well as the presence within them of angular to rounded clasts of basement rocks—some with the appearance of cored bombs—are all features compatible with formation by highly explosive basalt eruptions in a marr/diatreme complex (Hanson and Elliot, 1996). However, the location of the source vent for these deposits is unknown. Although not common, other largely fragmental mafic volcanic units are present in the Michigamme Formation, including the Clarksburg Volcanics to the east (Cannon, 1975) and unnamed mafic volcanics north of Iron River to the south (Cannon and Klasner, 1980). References Cannon, W.F., 1975, Bedrock geologic map of the Republic Quadrangle, Marquette County, Michigan: U.S. Geological Survey, Map I-862, scale 1:24,000. Cannon, W.F., and Klasner, J.S., 1980, Bedrock geologic map of the Kenton-Perch lake area, Northern Michigan: U.S. Geological Survey, Map I-1290, scale 1:62,500. Hanson, R.E., and Elliot, D.H., 1996, Rift-related Jurassic basaltic phreatomagmatic volcanism in the central Transantarctic Mountains: Precursory stage to flood-basalt effusion: Bulletin of Volcanology, v. 58, p. 327−347. 72 �Rosseel, J.-B., White, J.D.L., and Houghton, B.F., 2006, Complex bombs of phreatomagmatic eruptions: Role of agglomeration and welding in vents of the 1886 Rotomahana eruption, Tarawera, New Zealand: Journal of Geophysical Research, v. 111, B12205, doi:10.1029/2005JB004073. Sottili, G., Taddeucci, J., and Palladino, D.M., 2010, Constraints on magma-wall rock thermal interaction during explosive eruptions from textural analysis of cored bombs: Journal of Volcanology and Geothermal Research, v. 192, p. 27−34. Figure 1. Chert clast with a rind of fine vesicular basalt in the upper unit of the Parent Lake volcanics. Quarter for scale. 73 �Geochemical and petrographic study of a Mesoarchean felsic metavolcanic unit near Musselwhite Mine, North Caribou greenstone belt, northwestern Ontario SMYK, Emily1, HOLLINGS, Pete1, and BICZOK, John2 1. Geology Department, Lakehead University, 955 Oliver Rd, Thunder Bay, Ontario, P7B 5E1, Canada 2. Goldcorp Inc. Musselwhite Mine, P.O. Box 7500, Thunder Bay, Ontario, P7B 6S8, Canada The ~3 Ga North Caribou greenstone belt, North Caribou Terrane, Superior Province, comprises greenschist- to upper amphibolites-facies, ultramafic to felsic metavolcanic and metasedimentary rocks, intruded by 3000 to 2700 Ma felsic plutonic rocks. The study area is centered on a thin (~100m wide), north-trending, felsic metavolcanic unit on the western side of Opapimiskan Lake, 5 km northwest of Musselwhite Mine, in the South Rim Volcanic Unit (SRV). This rhyolitic unit has yielded a preliminary U-Pb age of 3053 Ma (McNicoll, unpublished data). The unit was mapped in detail over a distance of 2 km. Samples of the felsic rocks as well as the associated metabasalts and metapelites were collected. The metavolcanic and metasedimentary lithologies were classified by their petrography and whole rock geochemistry. The metarhyolitic units are characterised by a mineral assemblage of quartz + plagioclase + Kfeldspar with accessory muscovite + biotite + chlorite ± titanite ± clinozoisite ± zircon. Field observations identified some possible tuffaceous and pyroclastic felsic units as well as possible flows. The calculated anhydrous SiO2 values for the felsic metavolcanic rocks in the SRV range from 75 to 81 wt%. These highsilica, calc-alkaline rhyolites have been classified as FII-type rhyolites (based on a classification by Lesher et al. [1986]). The unaltered felsic metavolcanic rocks are characterised by LREE-enrichment and εNd values of -2.79 and -0.89 (Fig. 1). The tholeiitic and calc-alkaline basaltic and komatiitic-basaltic units have been metamorphosed to amphibolites consisting of plagioclase and hornblende with accessory titanite + chlorite + opaques ± rutile ± biotite ± zircon. The komatiitic-basalt has an MgO value of 14 wt% (compared to the range of 3 to 9 wt% MgO for the basalts) and has higher Ni and Cr contents. The basalts can be subdivided into two groups based their REE contents (Fig. 2). The εNd values are +0.78 and +0.58 for the flat REE basalts and -1.58, -1.61 and -3.35 for the LREE-enriched basalts. The three outcrops of metasedimentary rock found in this area have been metamorphosed into schists and a quartzite. The protolith of the schists was likely pelitic sedimentary rocks based on the mineral assemblage of strongly foliated quartz + biotite + muscovite + chlorite ± garnet ± zircon ± clinozoisite. These schists are associated with the felsic metavolcanic rocks. The quartzite is similar to the felsic metavolcanic units but does not contain feldspars. Its mineral assemblage contains quartz + muscovite + biotite + garnet. The rhyolites are calc-alkaline rocks with consistent LREE enrichment and negative Nb and Ti anomalies consistent with a suprasubduction zone environment. However, the negative εNd values indicate that these rocks have been contaminated by older continental crust suggesting emplacement in a continental arc. The tholeiitic mafic volcanic rocks with the flat REE is consistent with an oceanic island plateau environment generated by a mantle plume (Hollings and Kerrich, 1999). The positive εNd values support this conclusion as a positive value implies that the melt was derived from a depleted mantle source. 74 �1000.00 100.00 10.00 1.00 0.10 0.01 Th  Nb  La  Ce  Pr  Nd  Zr  Hf  Sm  Eu  Ti  Gd  Tb  Dy  Y  Ho  Er  Tm  Yb  Lu  Al V  Sc  Figure 1: Trace elements of the felsic metavolcanic rocks plotted on a primitive mantle‐normalized plot. Figure 2: Trace elements of the mafic metavolcanic rocks plotted on a primitive mantle‐normalized plot. References Biczok, J., Hollings, P., Klipfel, P., Heaman, L., Maas, R., Hamilton, M., Kamo, S., and Friedman, R. 2012. Geochronology of the North Caribou greenstone belt, Superior Province Canada: Implications for tectonic history and gold mineralization at the Musselwhite Mine; Precambrian Research, 192-195, 209-230. Hollings, P., Kerrich, R., 1999. Trace element systematics of ultramafic and mafic volcanic rocks from the 3 Ga North Caribou greenstone belt, northwestern Superior Province. Precambrian Research 93, 257–279. Lesher, C.M., Goodwin, A.M., Campbell, I.H., Gorton, M.P., 1986. Trace-element geochemistry of ore-associated and barren, felsic metavolcanic rocks in the Superior Province, Canada. Canadian Journal of Earth Sciences 23, 222–237. unpublished data, McNicoll, V., Geological Society of Canada, 2012. 75 �Petrology, mineralization, and alteration of the Thunder mafic to ultramafic intrusion, Midcontinent Rift, Thunder Bay TREVISAN, Brent1, HOLLINGS, Pete1, and AMES, Doreen2 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada 2 Geological Survey of Canada, Central Canada Division, 750-601 Booth St., Ottawa, ON K1A 0E8 Canada The northern Lake Superior region is host to a large segment of the North American Mesoproterozoic Midcontinent Rift (MCR). Since the discovery of high grade Ni-Cu-PGE mineralization hosted by mafic to ultramafic intrusions at the Current Lake deposit in Ontario, and the Eagle deposit in Minnesota, considerable exploration activity has been focused within the region. However, the small size of these buried deposits makes them difficult to locate both on the ground and on regional magnetic survey maps. The Thunder prospect is a mineralized mafic to ultramafic intrusion located on the outskirts of the City of Thunder Bay with mineral claims currently held by Rio Tinto (formerly Kennecott Canada Exploration Inc.). The intrusive body has been interpreted to be associated with the early magmatic stages of the MCR based on geochemical similarities with rocks of the Nipigon Embayment (Hollings et al., 2007). However, unlike other mineralized MCR intrusions, the Thunder intrusion has relatively low Ni grades (< 0.08%) and a large range in PGE tenors (combined platinum-palladium values > 0.5 g/t; Bidwell and Marino, 2007). The Thunder intrusion is also the only mineralised MCR intrusion hosted within the Archean Shebandowan greenstone belt as others, including Current Lake, intrude the Mesoproterozoic Sibley Group and/or the Archean Quetico metasedimentary subprovince (Williams et al., 1991; Hart and McDonald, 2007). This MSc study will characterise the petrology, mineralization, and alteration footprint of the Thunder intrusion and place it within the context of the MCR as a whole, in order to identify criteria for vectoring towards mineralization. The ~800m x ~800m Thunder intrusion consists of a mafic-ultramafic basal section that is overlain by a mafic sill-like body termed the “gabbroic cap”. The contact is xenolith-rich and lacks a well-developed chill margin. The mafic-ultramafic basal section of the intrusion consists of three transitional cumulate igneous phases: olivine websterite, olivine melagabbro, and olivine gabbro. Olivine websterite hosts up to 30% disseminated pyrrhotite, chalcopyrite, pyrite, rare bornite, and unknown platinum group minerals along its intrusive contact with significant drill intercepts including 20m at 0.22% Cu, 0.06% Ni, 0.25 g/t Pt, 0.29 g/t Pd, and 0.04 g/t Au (Bidwell and Marino, 2007). The Thunder gabbroic cap consists of three igneous phases: gabbro, pegmatitic gabbro, and leucogabbro all of which display a subophitic texture. The gabbro and leucograbbro are transitional and host up to 15% vein and 76 �disseminated chalcopyrite, pyrite and rare bornite however, no significant mineralisation was intercepted (Bidwell and Marino, 2007). The pegmatitc gabbro occurs along the contact between the two intrusive components of the Thunder and locally cross-cuts the olivine websterite. This unit has been interpreted to be a late magmatic phase of the Thunder. Marginal country rocks include metavolcanic and metasedimentary assemblages of the Shebandowan Greenstone Belt that have been structurally deformed and overprinted by regional greenschist facies metamorphism (Williams et al., 1991). The intrusive contact between the Thunder intrusion and the marginal country rock is sharp, xenolith-rich, lacks a well-developed chill margin, and contains blebs of granophyric material which suggests partial melting of the wall rock and/or wall rock assimilation. Within 100m of the Thunder intrusion a contact metamorphic aureole overprints the country rock consisting mostly of pervasive hornfels alteration. During Kennecott’s 2005 drill program a drill hole located ~400m southeast of the Thunder intrusion intercepted a mineralized unit 58m down depth. This peculiar unit is 4m thick, runs 1.7 g/t Au and 0.53% Cu, and was initially interpreted to be a massive sulphide magnetite garnet skarn associated with the emplacement of the Thunder intrusion (Marino and Bidwell, 2007). However, recent findings and discussions have challenged classifying this mineralized unit as skarn alteration and it is currently being investigated. Additional petrography, whole rock geochemistry, and SEM and S-isotopic analyses are currently being applied in order to further characterise the Thunder intrusion, mineralization and alteration assemblages both within the Thunder and the marginal country rock. In addition, this study will date the intrusion using U/Pb of zircon and/or baddeleyite. References Bidwell, G. E., and Marino, F., 2007., Geoinformatics Exploration Canada Limited, Thunder Project 2007 Field Program, Diamond Drilling on the 1245457 claim, Thunder Bay Mining Division, Ontario. Thunder Bay Resident Geologist’s Office, assessment files 20000002091-2.34638. Hart, T.R., and MacDonald, C.A., 2007. Proterozoic and Archean Geology of the Nipigon Embayment: implications for emplacement of the Mesoproterozoic Nipigon diabase sills and mafic to ultramafic intrusions. Canadian Journal of Earth Sciences, 44: 1021-1040. Hollings, P., Hart, T., Richardson, A., MacDonald, C.A., 2007. Geochemistry of the Mesoproterozoic intrusive rocks of the Nipigon Embayment, northwestern Ontario: evaluating the earliest phases of rift development. Canadian Journal of Earth Sciences 44: 1087–1110. Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L., and Sage, R.P., 1991. Wawa Subprovince in Geology of Ontario. Ontario Geological Survey, Special Volume 4 Part I: 485-539. 77 �Glacial Lake Ontonagon and the Development of Large Scale Landslides Vitton, Stanley J., Michigan Technological University, Houghton, MI, 49931 A massive landslide occurred in 2005 along the East Branch of the Ontonagon River in northern Michigan adjacent to US-45 (Figure 1A). The landslide initially blocked the river causing it to redevelop a new flow channel. While other massive landslides occur along this section of the river, they tend to be infrequent with respect to the general form of mass wasting such as slope regression due to river under cutting and surface erosion. An investigation of the landslide indicated two very distinct soil units that appear to correspond to the two phases of glacial Lake Ontonagon (Figure 1B). The two soil units have a relatively distinct boundary as seen in Figure 1C. The lower unit consists of a red till, which forms the floor of the valley, grading upward into alluvial sand (Figure 1D), while the upper unit is a distinct lacustrine soil deposit (Figure 1E). The massive landslide failure zone developed in the lower soil unit. It is unclear at this point as to whether the failure was due to softening of the lower red till or liquefaction induced failure caused by increased pore pressure development during the spring runoff in the alluvial sand. Due to the extensive development of soil liquefaction features (Figure 1F), however, it is believed that failure was induced via liquefaction in the transitional zone between the red till and the clean sand in the lower soil unit where the percent of fines in the sand prevent adequate drainage. Additional analysis of the soil’s strength and dynamic properties are needed, however, to make a more definitive determination (Smith, 2012). The origins of glacial Lake Ontonagon was first addressed by Leverette (1929) and later by Hack (1965), Farrand and Drexler (1985) and Attig, Clayton and Mickelson (1985). The formation Lake Ontonagon soils are believed to have developed in the post-Twocreekan time, around 11,800 Before Present (BP). The postTwocreekan glacier advance completely filled the Lake Superior basin with two ice lobes that were split by the Keweenaw Peninsula. The Superior lobe reached the position of the Nickerson moraine southwest of Duluth, while the Lake Michigan-Green Bay lobe moved southward across the northern peninsula of Michigan, ultimately reaching the Two Rivers moraine at Manitowoc, Wisconsin, about 11,800 BP. Following the Two Creek advance, de-glaciation formed lakes and drainage channels in front of the glacier lobes in which glacial lakes Duluth and Ontonagon formed. Lake Ontonagon drained westward into Lake Ashland and eventually to the St. Croix River, which drained southward to the Mississippi River at about 11,000 BP. Between 11,000 and 10,700 BP the glacier retreated into the Lake Superior Basin forming a much larger Lake Duluth and eventually as the ice retreated and the glacial rebound occurred lowering Lake Duluth to form Lake Algonquin. It is believed that the lower soil unit formed during this period of time. At about 10,000 BP, however, the last glacial re-advance, known as the Marquette Phase, advanced back into the Lake Superior Basin covering most of the northern portion of the Upper Peninsula. At about 9,900 BP the ice retreated again forming a series of lakes along the front of the ice sheet. Lake Ontonagon reformed at this time along with Lake’s Ashland and Nemadjic. Eventually the lakes became confluent and drained westward to the St. Croix outlet. At that time the lake levels for Ashland and Nemadjic dropped about 20 feet. Lake Ontonagon, on the other hand, dropped about 200 feet, (Leverett, 1929) leaving much of its lake bed dry land surface. It is believed that the upper lacustrine soil unit formed during this period of time. References Attig, W.J., Clayton, L. and D.M. Mickelson, 1985. Correlation of late Wisconsin glacial phases in the western Great Lakes area, Geological Society of America Bulletin vol. 96, no. 12; pp 1585-1593. 78 �Farrand, W.R. and Drexler, C.W. 1985. Late Wisconsin and Holocene History of the Lake Superior Basin, Quaternary Evolution of the Great Lakes, P.F Karrow and P.E. Calkin, editors, Geological Assoc. of Canada Special Paper 30. Hack , John, 1965. Postglacial drainage evolution and stream geometry in the Ontonagon area, Michigan, Geological Survey Professional Paper 504-B, Washington, D.C., 45 p. Leverett, Frank, 1929. Moraines and shorelines of the Lake Superior basin: U.S. Geological Survey Professional Paper 154-A, 72 p. Smith, J. 2012. Large Scale Landslide on the Ontonagon River, Michigan, Masters of Science Report, Michigan Technological University, Houghton, Michigan, 17 p. A B US-45 Upper Unit Lower Unit East Branch Ontonagon River C D Soil Unit Interface Lower LowerUnit Unit––Alluvial AlluvialSand Sand E F Upper Unit – Lacustrine Lateral spreading Figure 1. A) Large scale landslide adjacent US-45 on the East Branch of the Ontonagon River. B) Two soil units. C) Interface between soil units. D) Lower soil unit - alluvial sand. E) Upper soil unit – lacustrine soil unit. F) Liquefaction induces lateral spreading. 79 �A PRELIMINARY SURVEY OF THE GEOLOGY OF THE PRE-MICHIGAN BASIN ROCKS OF THE SOUTHERN PENINSULA VOICE, Peter, HARRISON, William, and THAKURTA, Joyashish, Department of Geosciences and the Michigan Geological Repository for Research and Education, Western Michigan University, 1903 W. Michigan Ave, Kalamazoo, MI 49008 The Michigan Geological Repository for Research and Education (MGRRE) is the premier core repository for Michigan Basin sedimentary rock materials. The collection holds roughly half a million linear feet of core. Archived with the MGRRE collection are multiple wells with cores (9 total) and a larger number of wells with drill cuttings (35 total) from preBasin rocks (Fig. 1). The majority of wells identified with pre-Basin rocks are from southern and southeastern Lower Peninsula. A handful of age dates are available from these wells, though most of the analyses for geochronological results were determined in the 1960s. Cuttings rich in biotite were age dated with both Rb-Sr and K-Ar from the McClure #2 State Beaver Island well and obtained ages of 1,040 and 1,090 Ma respectively (Lidiak et al. 1966). These dates suggest a relationship with the Mid Continent Rift System. The St. Blair 2-24 from Grand Traverse County was age dated with U-Pb from zircons from granite and yielded an age of 1,472 Ma (Hoppe et al. 1983). The St. Blair 2-24 well does have core though it is not archived with MGRRE. A series of wells from the southeastern Lower Peninsula were sampled for granite and granite-gneiss cuttings. These samples were dated with both the Rb-Sr and K-Ar systems and yield ages between 840 and 970 Ma (Lidiak et al. 1966, Summerson, 1962). Hinze et al. (1975) interpreted these rock units as being part of the Grenville province. Initial investigation of the collection of available cores has shown a complex set of lithologies preserved in the basement of the Michigan Basin in Branch and St. Joseph counties. Four closely spaced wells alternate between porphyritic granite and biotite-rich gneisses. A deep stratigraphic test drilled in Gratiot Co. cored through fine-grained turbiditic sediments interpreted to being part of the Mid Continent Rift System (Fowler and Kuenzi, 1978). The southeastern Lower Peninsula (Arenac, Huron and St. Clair counties) wells with core exhibit a diverse assemblage of metasediments, granites, and granite gneisses. Hinze et al. (1975) published a basement provinces map for the Lower Peninsula interpreted on the basis of limited well data, mostly from descriptions on driller’s reports and the geochronological results published by Lidiak et al (1966) and Summerson (1962). The addition of more than twice as many wells into the Lower Peninsula basement rock data set provides the opportunity to update Hinze et al.’s (1975) province map and produce the first basement geologic map of Michigan’s Lower Peninsula. To aid this effort, we are attempting to generate detailed lithological descriptions and a new geochronology of events. 80 �Figure 1. Distribution of wells in the Lower Peninsula of Michigan that drilled into the basement beneath the Michigan Basin. References Fowler, J. H., and Kuenzi, W. D. 1978. Keweenawan Turbidites in Michigan (Deep Borehole Red Beds): A Foundered Basin Sequence Developed During Evolution of a Proterozoic Rift System. Journal of Geophysical Research, 83: 5833-5843. Hinze, W. J., Kellog, R. L., and O’Hara, N. W. 1975. Geophysical Studies of Basement Geology of Southern Peninsula of Michigan. American Association of Petroleum Geologists Bulletin, 59: 1562-1584. Hoppe, W. J., Montgomery, C. W., and Van Schmus, W. R. 1983. Age and signficance of Precambrian Basement Samples from Northern Illinois and Adjacent States. Journal of Geophysical Research, 88:7276-7286. Lidiak, E. G., Marvin, R. F., Thomas, H. H.; and Bass, M. N.1966. Geochronology of the Miccontinent Region, United States. 4. Eastern Area. Journal of Geophysical Research, 71: 5427-5438. Summerson, C. H. 1962. Precambrian in Ohio and Adjoining Areas; State of Ohio, Department of Natural Resources, Division of Geological Survey. Report of Investigation. 44, 16 pp. 81 �Sedimentology and Geochemistry of a Regressive Surface in the Chemical Sediments of the Paleoproterozoic Gunflint Formation YIP, Christopher and FRALICK, Philip, Department of Geology, Lakehead University, Thunder Bay, ON, Canada, P7B 5E1, philip.fralick@lakeheadu.ca The 1878 ±1Ma Gunflint Formation is a chemical-sedimentary unit deposited in the Animike Basin; it shows a sequence of transgressive-regressive cycles. Wolff (1917) and Broderick (1923) divided the Gunflint into several individual members; lower cherty, lower slaty, upper cherty and upper slaty. These members were then grouped into two different sequences; the upper and lower sequences. The first and most extensive transgressive- regressive cycle is made up of the lower cherty member, while overlying transgressive-regressive cycles are made up of the lower slaty member, the lower cherty member and the upper slaty member. An outcrop present near Mink Mountain UTM: 329,520 E/5,338,163N, shows a complete section through the peak lower regressive-transgressive sequence. A detailed description of the section was logged through this sequence, which was divided into three main units; 1) a grainstone unit is the bottom unit and lies directly below 2) stromatolites which is capped off by a 3) oncolithic unit. The grainstone directly below the stromatolites is brecciated and shows injections of jasper and hematite throughout (Fig 1A). Microscopically the grainstone unit is composed of angular to rounded grains of chert. The cement is predominantly chert with some blocky quartz found forming at grain boundaries. The stromatolite unit is present above the pre-lithified grainstone unit and contains distinct stratiform and columnar stromatolites. The top unit is an oncolith-rich grainstone. The grains have a nucleus composed of either microquartzrich chert or blocky quartz. The cement of the unit is composed of a combination of a chalcedony-rich chert and blocky quartz. Several samples were taken up through this section and sent to the OGS lab in Sudbury for ICP-MS analysis for rare earth elements. The results were normalized to Taylor and McLennan (1985) Post Archean Australian Shale values and plotted (Fig 2A,B). All the samples taken from the grainstone layer and three samples taken from the stromatolite show a characteristic europium anomaly and a distinct positive cerium anomaly. The Ce anomaly is indicative of an oxidized environment where Ce (IV) was being precipitated and scavenged by the sediments (Peters, 2003). This requires oxygen production in the near-shore, and precipitation of Ce from sea-water that had not been previously exposed to significant oxygen. A B Figure 1. A) The lithified surface directly beneath the stromatolite section at Mink Mountain. B) A close up photograph of stromatolite lamiae and oncoliths in the depression between stromatolites from a polished hand sample. 82 �A B Figure 2. A) Rare earth element spider diagram for the grainstone unit at Mink Mountain. B) Rare earth element spider diagram for the stromatolites from Mink Mountain. References Broderick, T.M., 1920, Economic geology and stratigrsphy in the Gunflint iron district, Minnesota: Economic Geology 15: 422-452. Peters, J.M., 2003. Ancient iron formation: their genesis and use in the exploration of strataform base metal sulphide deposits, with examples from the Bathurst Mining Camp, in Lentz, D.R., ed., Geochemistry of Sediments and Sedimentary Rocks: Evolutionary Consideration to Mineral Deposits-Forming Environments: Geological Association of Canada. GeoText 4. P. 145-176 Taylor , S.R., and Mclennan, S.M., 1985. The continental crust; its composition and evolution; an examination of the geochemical record preserved in sedimentary rocks. Blockwell, Oxford, p. 312 Wolff, J.F., 1917, Recent geologic developments on the Mesabi Iron Range, Minnesota: Am. Institute of Mining and Metallurgical Engineers, Transactions 56: 229-257 83 �Sponsors The following organizations made general contributions to the 59th Annual Meeting. We thank the for their commitment to the Institute on Lake Superior Geology. For the past 59 years this organization has thrived as a result of the interest of individuals, corporations, universities and government agencies. The dedication to an exchange of scientific ideas and a passion for field trips has enabled the Institute to provide one of its primary objectives - to promote better understanding of the geology of the Lake Superior region. Eagle Mine � https://digitalcollections.lakeheadu.ca/files/original/74297d1a8a15b17000b67423e21aa2ba.pdf f54f6098f3af779b3b9135dc2d2d400e PDF Text Text Institute on Lake Superior Geology 59th Annual Meeting Houghton, Michigan May 8 - 11, 2013 Proceedings Volume 59 Part 2 - Field Trip Guidebook Editors: Theodore J. Bornhorst and Robert J. Barron www.lakesuperiorgeology.org ��Institute on Lake Superior Geology 59TH ANNUAL MEETING MAY 8-11, 2013 HOUGHTON, MICHIGAN SPONSORED BY: A. E. Seaman Mineral Museum Michigan Technological University THEODORE J. BORNHORST AND ALLAN R. BLASKE Co-Chairs Proceedings Volume 59 Part 2 – Field Trip Guidebook EDITED BY THEODORE J. BORNHORST AND ROBERT J. BARRON Cover Photo: Native copper from the Central Mine, Keweenaw Peninsula, Michigan. Collection of the A.E. Seaman Mineral Museum. Photograph by George Robinson. ��59TH INSTITUTE ON LAKE SUPERIOR GEOLOGY PROCEEDINGS VOLUME 59 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD TRIP GUIDEBOOK TRIP 1: GEOLOGIC OVERVIEW OF THE KEWEENAW PENINSULA, MICHIGAN TRIP 2: CALEDONIA MINE, KEWEENAW PENINSULA NATIVE COPPER DISTRICT, ONTONAGON COUNTY, MICHIGAN TRIP 3: GEOLOGY OF SILVER MOUNTAIN, HOUGHTON COUNTY, MICHIGAN TRIP 5: GEOLOGY OF THE KEWEENAWAN SUPERGROUP, PORCUPINE MOUNTAINS, ONTONAGON AND GOGEBIC COUNTIES, MICHIGAN TRIP 6: GEOLOGY AND ENVIRONMENTAL SITE CONDITIONS OF THE COPPERWOOD DEPOSIT, GOGEBIC COUNTY, MICHIGAN Reference to material should follow the example below: Cannon, W. F. Woodruff, L. G., and Schulz, K.. J., 2013, The Hiawatha Graywacke of the Iron River-Crystal Falls district, Michigan: a megaturbidite triggered by seismicity related to the 1850 Ma Sudbury impact [abstract]: Institute on Lake Superior Geology Proceedings, 59th Annual Meeting, Houghton, MI, v. 59, part 1, p. 14-15. Published by the 59th Institute on Lake Superior Geology and distributed by the ILSG Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca ILSG website: http://www.lakesuperiorgeology.org ISSN 1042-9964 i ��Table of Contents TRIP 1 – GEOLOGIC OVERVIEW OF THE KEWEENAW PENINSULA, MICHIGAN 1 TRIP 2 – CALEDONIA MINE, KEWEENAW PENINSULA NATIVE COPPER DISTRICT, ONTONAGON COUNTY, MICHIGAN 43 TRIP 3 – GEOLOGY OF SILVER MOUNTAIN, HOUGHTON COUNTY, MICHIGAN 59 TRIP 4 – A.E SEAMAN MINERAL MUSEUM – NO GUIDE TRIP 5 – GEOLOGY OF THE KEWEENAWAN SUPERGROUP, PORCUPINE MOUNTAINS, ONTONAGON AND GOGEBIC COUNTIES, MICHIGAN 69 TRIP 6 – GEOLOGY AND ENVIRONMENTAL SITE CONDITIONS OF THE COPPERWOOD DEPOSIT, GOGEBIC COUNTY, MICHIGAN 97 ii ��Field Trip 1 Geologic Overview of the Keweenaw Peninsula, Michigan Theodore J. Bornhorst A.E. Seaman Mineral Museum, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Robert J. Barron Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 Introduction The geology of the far western Upper Peninsula of Michigan consists of three temporally distinct episodes. During the Mesoproterozoic, between about 1.15 and 1.03 Ga, up to 30 km of Keweenaw Supergroup volcanics and clastic sediments filled an intracratonic rift, the Midcontinent Rift (MCR) (Figs. 1 and 2) (Heaman et al., 2007; Davis and Paces, 1990; Cannon et al., 1989). After a 500 million year period of erosion, the MCR rocks were buried by Phanerozoic sedimentary rocks from about 500 Ma to 175 Ma (Catacosinos et al., 2001). Pleistocene continental glaciations, beginning about 2 million years ago, removed the Phanerozoic rocks from the Keweenaw Peninsula leaving only a few outliers. About 10,000 years ago, as the last remaining glaciers retreated, and they left behind a variety of unconsolidated clastic sediments. The geologic evolution of the far western Upper Peninsula is illustrated in cartoon form in Figure 3. 75o 50o Midcontinent Rift 100o o 50 Precambrian bedrock below unconsolidated glacial sediments Canada Canada Minnesota Canada Wisconsin Grenville Tectonic Zone Iowa Nebraska Kansas 0 Shaded area = Phanerozoic bedrock belowunconsolidated glacial sediments o 35 100o North Slate Shoreline Islands Superior Shoal Bg/Og Bg/Og Ks 5 Plv Paleoproterozoic? Manitou Island Js Manitou structural zone 35o 75o South Shoreline Js Plv/Ks Ks A A 400 kilometers Fault blocks/ intrusives 10 A 15 Ks = Keweenawan Supergroup older than Portage Lake Volcanics Js = Jacobsville Sandstone Og = Oronto Group Bg = Bayfield Group Plv = Portage Lake Volanics A = Archean rock Stratigraphic column given in Fig. 4 Figure 1: Generalized bedrock geologic map of the Midcontinent Rift. Grenville tectonic zone after Cannon (1994) and interpretative cross-section across the Lake Superior segment of the Midcontinent rift after Cannon et al. (1989). 1 �Mesoproterozoic Midcontinent Rift Around Lake Superior Native Copper occurrences Sedimentary Rocks Igneous Rocks Several Major Faults Abundant Ontario Isle Royale Minnesota Lake Superior Keweenaw Peninsula native copper district Ontario Upper Peninsula of Michigan Wisconsin Lake Michigan 0 100 200 kilometers Archeanmetamorphosed sedimentaryandigneous rocks N Paleoproterozoic metamorphosed sedimentary and igneous rocks Phanerozoic sedimentary rocks Figure 2: Generalized bedrock map showing the exposed rocks of the Midcontinent Rift around Lake Superior and the bedrock of the Upper Peninsula of Michigan. Locations of concentrations of native copper are shown around Lake Superior. Modified from Bornhorst and Barron (2011). In the strictest sense, the geographic area of the Keweenaw Peninsula proper extends northeast of a NW-SE line drawn through L’Anse (Fig. 4), however, the term Keweenaw Peninsula has also been applied to the area containing MCR rocks farther to the south. Bornhorst and Barron (2011) used the name Keweenaw Peninsula native copper district to describe native copper deposits hosted by MCR rocks as far south as the White Pine Mine. The geologic description in this field trip guide is restricted to the Keweenaw Peninsula proper and does not include the southern bedrock such as at Silver Mountain described in Field Trip 3 (this volume) or Cambrian to Devonian rocks at Limestone Mountain described by Milstein (1987). The descriptions of the geology of the Keweenaw Peninsula provided here were modified from a combination of Bornhorst and Barron (2011), Bornhorst and Lankton (2009), and Bornhorst and Rose (1994), and Bornhorst et al. (1983). Specific citation or quotation is not given in all instances. 2 �D. Late compression and rift-flanking basin ~ 1.06 to 1.03 Ga H. Continental Glaciation Rift-flanking basin < 2 Ma Isle Royale Lake Superior Keweenaw Peninsula Reverse inversion of normal faults C. Sedimentary infillingduring sagging G. Burial by Phanerozoic sedimentary rocks ~ 1.092 to 1.05 Ga ~ 500 to 175 Ma Clastic sedimentary rocks B. End of Magmatic Phase ~ 1.098 to 1.092 Ga Sheeted dikes F. Extended period of erosion Basalt lava flows ~ 1.15 to 1.098 Ga Downward percolating groundwater Diabase dike Gabbro A. EarlyStage of Rifting ~ 1.03 Ga to 500 Ma (0.5 Ga) Basalt lava flows E. Formation of native copper deposits during late compression ~ 1.06 to 1.04 Ga Extension Crust Deep Magma Chamber Native copper precipitation Mantle Ore mineralizing fluids Figure 3: Cartoon NW to SE cross sections from Minnesota (left) to the Upper Peninsula (right) illustrating the progressive geologic evolution. Modified from Bornhorst and Lankton (2009). Midcontinent Rift Lithologic Units The Keweenaw Peninsula is located on the southern margin of the Lake Superior segment of the MCR (Figs. 1and 2). The rock units that are associated with the MCR have been termed the Keweenawan Supergroup (Fig. 5). These rocks were deposited from about 1.15 and 1.03 Ga (Heaman et al., 2007; Davis and Paces, 1990; Cannon et al., 1989). The MCR beneath Lake Superior is filled with up to about 30 km of volcanic rocks (Figs. 1 and 3) (Hinze et al., 1990; Cannon et al., 1989; Cannon, 1992). The MCR geology of the Keweenaw Peninsula can be divided into northwest-dipping, rift-filling volcanic and clastic sedimentary rocks located on the northwest side of the Keweenaw Peninsula and flat to low-dipping, rift-flanking clastic sedimentary rocks located on the southeast side (Fig. 4). These two contrasting lithologic settings are separated by the Keweenaw Fault which was originally a graben-bounding fault, but today is a high-angle reverse fault (Fig. 3). 3 �4 �Figure 5: Lithostratigraphic bedrock units of the Mesoproterozoic Midcontinent rift system of Michigan. Portage Lake Volcanics The Portage Lake Volcanics (Figs. 4 and 5) is a 2,500 to 5,200 m thick formation dominantly composed of subaerial basalt lava flows with less than 1 % by volume intermediate to felsic volcanic and subvolcanic rocks located stratigraphically near the base of the exposed formation. Less than 5 % by volume is stratigraphically scattered interflow reddish-colored conglomerate and sandstone units that are greater in abundance towards the top of the formation (Butler and Burbank, 1929; White, 1968). The base of the formation is truncated by the Keweenaw Fault. The lavas flowed from fissure vents that tended to be located nearer the axis of the rift zone which produced a layered succession of flood basalts comparable to the rift zones of East Africa and Iceland (e.g., Nicholson, 1992 and reference therein). The Portage Lake Volcanics erupted over 2 to 3 million years from 1,096.2+/-1.8 (Copper City flow, Fig. 6) to 1,094.0+/-1.5 (Greenstone flow, Fig. 6) (Paces and Miller, 1993; Davis and Paces, 1990). There are more than 200 individual basaltic lava flows in the exposed Portage Lake Volcanics which are typically aphyric, Mg-rich, high-Al olivine tholeiites (Paces, 1988). The most abundant type of basalt flows are olivine tholeiites, followed by primitive olivine tholeiites and quartz tholeiites. Iron-rich olivine tholeiites are generally lesser in abundance (Table 1). The thicker lava flows are compositionally stratified due to magmatic differentiation after eruption, especially the Greenstone flow, which is the thickest individual flow in the formation (Cornwall, 1951a and b; Broderick, 1935; Broderick and Hohl, 1935). The composition of the basalts is cyclical with minor and major cycles superimposed on an overall trend toward more primitive compositions towards the top of the formation. The basalt magmas were derived by partial melting of sub-continental upper 5 �mantle with an overall stratigraphically upwards trend towards younger, to more primitive basalt compositions as a result of less contamination by crustal rocks (Paces, 1988; Paces and Bell, 1989). The repeated magmatism at the rift axis and progressive crustal thinning provided pathways for transport of magma to the surface creating less extended contact with crustal rocks and hence, less contamination. The youngest rocks of the Portage Lake Volcanics in the Keweenaw Peninsula have compositions similar to MORB suggesting the MCR nearly formed an ocean basin. The major geochemical cycles are due to fractional crystallization and replenishment in large magma chambers near the crust/mantle interface whereas the minor cycles are due to closed system fractional crystallization in small magma chambers within the crust (Paces, 1988). The Portage Lake Volcanics were likely derived by partial melting of trace element enriched plume-related mantle (Nicholson et al., 1997; Nicholson and Shirey, 1990; Paces and Bell, 1989). Table1: Average and representative geochemical data for least altered lavas of the Portage Lake Volcanics (from Paces, 1988). Tholeiites were grouped by Ni content. Primitive Intermediate Olivine Olivine olivine olivine tholeiite tholeiite tholeiite tholeiite Ni (ppm) Wt.% SiO₂ Al₂O₃ FeOt MgO CaO Na₂O K₂O TiO₂ P₂O₅ MnO PPM Ni Cu Zr Iron-rich olivine and Andesite Dacite Rhyolite quartz tholeiites 400-300 n=5 300250 n=9 250200 n=14 200-100 n=8 100-15 n=6 n=1 n=1 n=1 47.82 15.89 9.77 12.44 10.58 2.04 0.19 0.98 0.16 0.14 47.34 15.27 11.82 11.69 10.24 2.10 0.22 1.13 0.19 0.16 48.03 15.32 12.32 9.85 10.16 2.25 0.33 1.35 0.22 0.16 48.55 15.12 12.86 9.06 9.65 2.31 0.42 1.60 0.25 0.18 49.94 13.28 14.91 7.78 6.64 2.91 1.43 2.34 0.36 0.24 56.39 13.78 9.87 5.52 5.10 3.94 2.27 1.83 1.00 0.30 68.44 15.17 4.46 1.14 1.40 4.74 3.86 0.51 0.19 0.08 77.89 12.77 1.11 0.17 0.04 3.67 4.28 0.08 0.01 0.01 231 73 101 172 86 126 54 126 212 10 5 430 7 13 573 5 61 145 326 279 37 51 78 85 FeOt=total Fe as FeO 6 �All observed basalt lava flows in the Portage Lake Volcanics were erupted subaerially and consist of a massive (vesicle-free) interior capped by a vesicular and/or brecciated flow top. The only evidence for involvement of water during eruption is a single thin hyaloclastic unit in the upper part of the formation (locally termed the ashbed). Subaerial eruption resulted in degassing of volatiles, notably SO2 (Cornwall, 1951c). The lava flows range in thickness from 1 to 450 m with most of them between 10 to 20 m thick (Paces, 1988; White, 1960). Most of the lava flows cannot be traced along strike with confidence although a few such as the Scales Creek, Kearsarge, and Greenstone flows have well documented lateral continuity flows (Fig. 6). The Greenstone flow has been correlated down dip across the Lake Superior syncline to Isle Royale (Longo, 1982; Huber, 1975). The uppermost 5 to 20% of the tops of most (89 %) individual lava flows are vesicular with between 5 and 50% vesicles (White, 1986). The tops of 21 % of the flows are brecciated with clasts of vesicular basalt. The vesicles in most lava flows within the Portage Lake Volcanics, except for the stratigraphically uppermost, are filled with secondary minerals and are amygdules. Thus, amygdaloids are lava flows with vesicle-only tops and fragmental amygdaloids are those with vesicular and brecciated tops. There are the minor amounts of andesite, dacite, and rhyolite lava flows and subvolcanic plutons that interfinger with and cross cut the basalts of the Portage Lake Volcanics (Table 1). Most of these occur in the stratigraphically lowermost portion of the Portage Lake Volcanics. A few dikes of intermediate composition and a diorite stock at Mt. Bohemia intrude the exposed Portage Lake Volcanics. The rhyolitic volcanic setting is analogous to the shield-type central volcanoes of Iceland (Nicholson, 1991). Interflow sedimentary units are important stratigraphic markers in an otherwise monotonous succession of basalt lava flows that can be traced up to 90 km along strike and so many of them are given informal names (see Fig. 6). They consists of red-colored, well-lithified, pebble-to-boulder conglomerates with lesser amounts of interbedded sandstone and occasional significant amounts of siltstone and shale ranging in thickness from a few cm up to about 40 m (Merk and Jirsa, 1982; White, 1968; Butler and Burbank, 1927). The typical conglomerate has an exposed interflow lithology characterized by sub-rounded to angular pebbles in a sandy matrix. Clast size varies from pebbles to boulders and clast lithologies are predominantly felsic, although there is considerable variation within and between specific beds reflecting diversity in source terrane. Within the interflow Calumet and Hecla Conglomerate, Kalliokoski and Welch (1985) interpreted a subunit as a caliche soil profile. The interflow clastic sedimentary beds were deposited during intervals of volcanic quiescence most likely in terrestrial alluvial fans in an arid to sub-arid climate. Deposition was on top of the shallow-dipping to flat-lying lava flows by streams flowing from the topographic high on the margin of the MCR toward the center of the rift basin (now under Lake Superior) (White, 1968). 7 �NE M eters 300 600 0 Top of the Portage Lake Volcanics Hancock Conglomerate Ashbed Flow Pewabic West Conglomerate Ashbed Flow Greenstone Flow Pewabic Flow Allouez Conglomerate 300 600 900 Evergreen Flows Kearsarge Flow Winona Flow 1200 Isle Royale Flow Wolverine Sandstone 1500 Scales Creek Flow 1800 Baltic Flow Gratiot Flow Copper City Flow Upper Limit of Epidote in Flows Upper Limit of Quartz in Flows Lower Limit of Prehnite in Flows 0 3 6 km Exceptionally Thick Lava Flows Location of Mine within Stratigraphic Strike Parallel Section Figure 6: Generalized stratigraphic position parallel to strike of informal units within the Portage Lake Volcanics. Modified from Stoiber and Davidson (1959). Figure 5 shows location of GreenlandMass subdistrict (Michigan, Caledonia, Mass, Adventure Mines) and Copper Harbor. Copper Harbor Formation The Copper Harbor Formation is the oldest formation in the Oronto Group and conformably overlies and interfingers with the Portage Lake Volcanics (Figs. 4 and 5). It consists of red-brown clastic sedimentary rocks with a maximum exposed thickness 2,000 m. The Copper Harbor Formation in the Keweenaw Peninsula includes a succession of subearially deposited lava flows informally named the Lake Shore Traps. The depositional environment of the Copper Harbor Formation was deposited in a prograding coalescing alluvial fan complex with proximal-to-distal braided stream and sheet flood facies on the alluvial fans to distal sand flats and flood plain facies (Elmore, 1984). The climate was probably arid with flashy seasonal streams. The highlands from which the Copper Harbor Formation were derived to the southeast, now buried under the Jacobsville Sandstone. 8 �Conglomerates and sandstones are the dominant lithologies in the Copper Harbor Formation. The formation fines distally and up section, reflecting a waning sediment supply due to progressive erosion of the source area (Elmore, 1984). The poorly-sorted clasts in the conglomerates range in size from granules to boulders that are subrounded to rounded and are mostly volcanic in origin and have a ratio of mafic-to-intermediate + silicic composition of about 2:1 (Daniels, 1982). The conglomerates include clast-supported and matrix- supported varieties; some of the latter are diamictites. The conglomerates are interpreted as high-energy channel deposits on a coalescing alluvial fan (Elmore, 1984). The diamictites are debris flow in origin. Sandstones are predominantly red-brown, subangular-to-angular lithic graywackes with volcanic lithic fragments. These exhibit current-ripples, trough-cross beds, current and parting lineations, and reduction spots. Sandstone interbeds are more common in the upper 2/3 of the formation. The abundant calcite cement in the conglomerate and coarse sandstone was probably deposited as vadose carbonate or caliche (Kalliokoski, 1986). Thin red-colored siltstone and shale interbeds have desiccation cracks and are interpreted as filling abandoned channels on the alluvial fan surface. In the Copper Harbor area, there are also laminated cryptoalgal carbonate beds and ooid lenses occurring within the same general stratigraphic position. These are laterally-linked contorted layers in shale-siltstone that are draped over cobbles and are found as poorly developed mats in coarse sandstone (Elmore, 1983). The laminated carbonate beds are algal stromatolite (genus Colleria). The stromatolites formed in shallow, medial fan lakes and possibly abandoned channels on the alluvial fan surface (Elmore, 1983). The Lake Shore Traps (Lane, 1911), an informal member of the Copper Harbor Formation (Fig. 5), are well exposed near the tip of the Keweenaw Peninsula where the unit is composed of 31 lava flows and one interflow conglomerate about 600 m thick (Paces and Bornhorst, 1985). The composition of the Lake Shore Traps is different than the underlying Portage Lake Volcanics reflecting the change from active rift-filling magmatism to active rift-filling clastic sedimentation with little to no magmatism. The rocks range from Fe-rich olivine tholeiitic basalt at the base to Ferich olivine-bearing tholeiitic basaltic andesites to tholeiitic andesites. Strato-geochemical relationships can be explained by a combination of fractional crystallization, parental magma replenishment, and wall rock assimilation (Paces and Bornhorst, 1985). Davis and Paces (1990) report a U-Pb age on zircon of 1087.2 +/- 1.6 Ma for the Lake Shore Traps. Nonesuch Formation The Nonesuch Formation conformably overlies and locally interfingers with the Copper Harbor Formation (Figs. 4 and 5). It consists of dominantly black-to-gray-to-green fine clastic sedimentary rocks with a maximum exposed thickness 240 m. Exposures in the Keweenaw Peninsula are limited with the best exposure at the Hancock Campground on M-203. This formation will not be visited for this field trip. The Nonesuch Formation was deposited in a generally anoxic lacustrine environment ranging from marginal lacustrine (sandflat-mudflat) to lacustrine to lacustrine-to-fluvial subenvironments (Elmore et al., 1989). Siltstone and shale are the dominant lithologies with lesser very-fine sandstone and minor carbonate laminates. While gray (reduced) color characterizes most of this formation, the stratigraphic upper beds have more red-brown colors (Bornhorst and Williams, in press). Well-laminated to massive 9 �black to dark-gray siltstone and shale were deposited in the lacustrine subenvironment. The lacustrine lithologies at the base of the Nonesuch Formation host economic quantities of chalcocite and native copper at the now closed White Pine Mine (Mauk et al., 1992) and chalcocite at the Copperwood project (Bornhorst and Williams, in press; and Field Trip 5 this guidebook). A thin carbonate laminate yielded a Pb-Pb isochron age of 1,081 ± 9 Ma (Ohr,1993) Freda Formation The Freda Formation is the youngest formation of the Oronto Groupand overlies the Nonesuch Formation with no explosed top (Figs. 4 and 5). The contact between the Freda and Nonesuch Formations is gradational. The exposed thickness is greater than 3,700 m, however, it is poorly exposed except along the Lake Superior shoreline. This formation will not be visited for this field trip. The Freda Formation is presumably overlain by the Jacobsville Formation. The Freda Formation was deposited in an environment characterized by shallow meandering streams (Daniels, 1982). Red-brown fine to very-fine sandstone, siltstone, and mudstone are the dominant lithologies in the Freda Formation. Fining-upward sequences occur on the scale of a few meters. Based on regional correlations the age of the Freda Formation is likely 1,060 to 1,040 Ma (Cannon, 1992). Jacobsville Sandstone The Jacobsville Sandstone is the youngest Mesoproterozoic bedrock formation in the Keweenaw Peninsula (Figs. 4 and 5). Its stratigraphic relationship with other units is uncertain. It occurs in a contiguous geographic region bound on the northwest by the Keweenaw Fault and to the southeast, it angularly unconformably overlies Paleoproterozoic and Archean rocks (Fig 3). The Jacobsville Sandstone is estimated to be more than 2,900 m thick and the top is not exposed (Kalliokoski, 1982). The Jacobsville Sandstone was deposited in an environment characterized by shallow meandering streams (Kalliokoski, 1988). The formation occurs in a rift-flanking basin and at least part of it was deposited during active reverse movement along the Keweenaw Fault. Red to red-brown sandstone is the dominant lithology with lesser amounts of red-brown conglomerate, siltstone, and shale. The sandstone varies from subarkose to quartz sublithic arenite although there are some beds of arkose and quartz arenite (Kalliokoski, 1982). Rounded-tosubrounded, very-fine to coarse sand grains of quartz, feldspar, and lithic fragments occur in massive to cross-bedded, fining-upward sequences. Quartz grains show evidence of volcanic and metamorphic origin. Ripple marked bedding surfaces and cross-bedding are common in some localities. The sandstone varies in color from red to a cream-white or purplish-red color; creamwhite color occurs as spherical reduction spots and layers that tend to follow bedding or fractures. Conglomerate is more common in localities near the Keweenaw Fault or near the unconformable contact. Near the Keweenaw Fault, pebble to boulder sized clasts in the conglomerates are composed of felsic and mafic volcanic rocks, similar to Keweenaw Supergroup lithologies. Near the unconformable contact, clast lithologies are of locally derived chemically resistant debris such as quartz and iron formation. There are no interbedded volcanic rocks or cross-cutting igneous dikes within the Jacobsville Formation and while the older age is constrained the upper age is not. The 10 �Jacobsville Formation is approximately 1.06-1.04 Ga to 1.03? Ga (Cannon, 1992). Midcontinent Rift Structure The last episode of the MCR was characterized by a compression of the continent. This compression transformed original graben-bounding normal faults into reverse faults, reactivated other extensional rift-related faults/fractures, and produced new compression-only faults/fractures and folds. Cannon et al. (1993) have determined that compression occurred at about 1,060+/-20 Ma. The probable cause of this event was continental collision along the Grenville front (Fig. 1) beginning as early as 1.08 Ga and ending by 1.04 Ga (Cannon, 1994; Cannon and Hinze, 1992; Hoffman, 1989). The rift-filling Keweenawan Supergroup strata dips moderately northwesterly toward the center of the rift (Lake Superior) (Figs. 3 and 7). Their dip angles increase toward the exposed stratigraphic base which is truncated by the Keweenaw Fault. The present day dip of the strata within the MCR is a combination of syn-depositional downwarpage and tilting in response to reverse movement along Eagle Harbor 28 3 area of fissure deposits Strike and dip of bedding Major copper deposits 24 Eagle River 30 1 Fault or fissure 28 U=up thrown side D=down thrown side U D 2 50 3 4 26 5 6 3 7 Calumet 83 8 19 9 5 Hancock Houghton 11 N 10 0 12 10 20 58 Kilometers 1 Name of Deposit Given in Table 2 Figure 7: Simplified geologic map showing the location of the major deposits within the Keweenaw Peninsula native copper district, Michigan. Table 2 provides the names and production for the numbered deposits. The areas shown on the map are the mined out down-dip portion projected to the surface. All of the native copper mines are hosted by the Portage Lake Volcanics. Modified from Bornhorst and Barron (2011). Table 2: Production from 1845 to 1968 of refined copper from native copper deposits (after Weege and Pollock, 1971). 11 �Million lbs Produced Refined Copper Location Number Shown on Figure 7 Calumet & Hecla Conglomerate 4,229 7 Kearsarge Flow Top 2,263 3 Baltic Flow Top 1,845 12 Pewabic Flow Top 1,077 9 Osceola Flow Top 578 8 Isle Royale Flow Top 341 10 Atlantic Ashbed 143 11 Allouez Conglomerate 73 6 Houghton Conglomerate 38 4 Kingston Conglomerate 20 5 Greenland-Mass Subdistrict 72 See Figure 3 Other Flow Top and Conglomerate Deposits 137 Cliff Fissure 38 1 Central Fissure 53 2 Other Fissure Deposits 123 Name of Deposit District Total 11,030 the Keweenaw Fault produced by continental compression. Bedding in the rift-flanking Jacobsville Sandstone dips less than 5O in most areas, except near the Keweenaw Fault, where dips steepen in response to drag along the fault. Compression-related deposition produced the Jacobsville Sandstone. There are many faults/fractures in the Keweenaw Peninsula. Some of these were exclusively formed during extension of the MCR when graben-bounding normal faulting was prominent along the margin (Fig. 3). However, most faults/fractures were likely reactivated by or related to the compressional event that inverted the major graben-bounding fault, the Keweenaw Fault, into an overall high-angle reverse fault (Cannon et al., 1989; White, 1968). The Keweenaw Fault strikes and dips more or less parallel to the bedding of the truncated Portage Lake Volcanics (Fig. 7) and is not necessarily one fault, as it is a zone with branches up to 0.8 km from the main fault (Butler and Burbank, 1929). Although the Keweenaw Fault would make an ideal conduit for movement of hydrothermal fluids, there are no native copper deposits along it similar to other ore bearing districts where the main faults are not well mineralized. The rocks within and adjacent are altered especially by paragenetically late hydrothermal fluids. Several reverse faults occur oblique to the strike of bedding. In the Eagle River area, high-angle faults with displacement from 0 to 200 m, faultcontrolled native copper veins are common (Butler and Burbank, 1929). The Allouez Gap fault bisects the largest lava flow top hosted native copper deposit in the district and was likely a significant conduit for native copper mineralizing hydrothermal fluids (Bornhorst, 1997). The Allouez Gap fault may have been a reactivated original rift fault. Faults were the principal pathway for the upward movement and focusing of ore fluids into the stratabound lava flow tops in the Baltic and Isle Royale deposits (Broderick, 1931) as well as those in the Greenland-Mass subdistrict (Field Trip 2, this guidebook). Faulting occurred before, during and after deposition of native copper along 12 �with associated alteration minerals based on fault brecciation and re-cementation of alteration minerals. There is a close relationship between faulting/fracturing produced by or reactivated by compression. These compressional structures acted as pathways for native copper mineralizing hydrothermal fluids. Broad open synclines and anticlines, with wavelengths of around 10 km and various orientations, are superimposed on the regional dip. Faults with displacement and mineralized tension breaks are common near the crests of anticlines (Butler and Burbank, 1929). These post-depositional folds are likely related to the Keweenaw Fault (White, 1968). Keweenaw Peninsula Native Copper District Active copper mining occurred from 1845 to 1968 in the Keweenaw Peninsula native copper district. The estimated pre-mining geologic resource for the district is 19.7 billion lbs of copper and small quantities of temporally and spatially associated native silver (Bornhorst and Barron, 2011). The major ore producing horizons are located in a 45 km-long belt in the Keweenaw Peninsula (Figs. 5 and 7) and in a subdistrict to the southwest (Field Trip 2, this guidebook). Accompanying native copper and silver, the only economic metallic minerals, were a suite of nonmetallic alteration minerals (Fig. 8). Sulfide minerals are uncommon in the native copper deposits; chalcocite only occurs in trace amounts. Pyrite, an acid-producer when exposed to oxygenated waters, is absent. Several chalcocite deposits of unknown connection to the native copper deposits are hosted by the stratigraphically older Portage Lake Volcanics; the largest of these contains roughly 230 million lbs of copper (Maki and Bornhorst, 1999). These will not be discussed here. Native Copper Ore Bodies Ore bodies in the Keweenaw Peninsula are tabular, stratabound concentrations of native copper in Portage Lake Volcanics host rocks with sufficient original porosity including brecciated and amygdaloidal flow tops (58.5% of production) and interflow conglomerate beds (39.5% of production). Secondary porosity occurs along fractures/faults host veins (about 2% of production). Since the deposits represent important stratigraphic horizons, the host rocks were given informal member names (Butler and Burbank, 1929). Several mines with different names often worked the same deposit or same lithostratigraphic unit. About 85% of the total district production came from four deposits: Calumet and Hecla Conglomerate, top of the Kearsarge lava flow, top of the Baltic lava flow, and the top of the Pewabic lava flow. The most common host rocks for native copper deposits are brecciated flow tops (fragmental amygdaloid) as their original porosity was typically much greater than vesicular (amygdaloidal) flow tops (White, 1968). The stratabound flow top deposits are “sandwiched” between a footwall consisting of the same flow as the mineralized flow top and hanging wall of barren massive basalt interior in the succeeding lava flow. Native copper is often more abundant near the top and bottom of the brecciated/fragmental amygdaloid interval of the flow top, however, in rich ore shoots, the entire brecciated/fragmental amygdaloid flow top contains significant amount of copper. As brecciated/fragmental amygdaloidal grades downward into massive basalt, it becomes deficient in native copper. In some cases, ore shoots are located in tongues of brecciated flow tops within massive basalt (Weege and Schillinger, 1962). The lateral and vertical distribution of 13 �brecciated/fragmental amygdaloid within the top of the lava flow is irregular and hence, so is the grade of copper. In general, mined stope heights are from 3 to 5 m. Ore shoots are elongate, but also occur in a wide variety of shapes, with widths of 30 to 150 m and down dip lengths from 50 m to 600 m (White, 1968). The strike length for major ore bodies ranges from 1.5 to 11 km with down dip mineralization extending from 1.5 to 2.6 km below the surface on the inclined deposit. (Butler and Burbank, 1929; White, 1968). Although interflow conglomerate beds make up only a small volume of the Portage Lake Volcanics, about 40 % of the district productions were hosted by them. These deposits were also tabular and stratabound, just like the flow top deposits. They are “sandwiched” between a footwall consisting of the top of the underlying lava flow and hanging wall of barren massive basalt interior in the underlying lava flow. The porosity of underlying brecciated/fragmental amygdaloid lava flow top is greatly decreased by silt and sand filling the origin open space between fragments. Native copper tends to be concentrated along specific stratigraphic bands that are 0.5 to 5 m thick (Weege et al., 1972). Interflow conglomerates overall host a significant fraction of the districts production. Specifically, the Calumet and Hecla Conglomerate was by far the largest single native copper deposit producing 4.2 billion lbs, as compared to the next largest deposit hosted by the Kearsarge flow top which produced 2.3 billion lbs (Fig. 7 and Table 2). The Calumet and Hecla Conglomerate was mined along a strike length of 4.9 km, and down-dip 2.8 km. The productive area corresponds to a thickening of the conglomerate from less than 1 m up to 6 m (Butler and Burbank, 1929; Weege et al., 1972). Ore grades decrease with depth where the width of the conglomerate is greater where essentially the same amount of copper is distributed throughout a greater volume. (Butler and Burbank, 1929). The highest grades correspond to beds where there is relatively little fine interstitial material or where interstitial spaces are filled with coarse sand or small pebbles (Weege et al., 1972). Thus, localization of native copper ore is dependent on sedimentary environmental factors. The first mines in the district were developed on tabular steeply dipping deposits that cross cut bedding at high angles. The veins have widths of up to 3 m or more (Butler and Burbank, 1929). Veins are not single tabular bodies, but rather a series of parallel of anastomosing filled open spaces. While brecciation within the deposit is common, gouge is not present (Butler and Burbank, 1929). These adjacent lava flow tops and conglomerates are mineralized. The distribution of native copper in veins is more erratic than in either lava flow top or conglomerate deposits. The richest ore veins tend to be spatially associated with the intersections of the vein and well-oxidized lava flow tops (Butler and Burbank, 1929). Native copper occur as both finely disseminated with associated alteration minerals and as masses weighing many tons. These vein deposits are of slight economic importance in the district. Several small vein deposits are localized just beneath the thickest basalt flow in the district. A good example is the Greenstone flow in which hydrothermal fluids moved up along the cross fractures until blocked by the very thick impermeable massive interior of the Greenstone Flow. 14 �There are veins spatially and genetically associated with the stratabound lava flow top or conglomerate deposits; these veins occur along faults that intersect major deposits such as the Baltic and Isle Royale (Broderick, 1931). This suggests that ore fluids moved upward along faults and outward into the permeable flow tops. The intersection of subsidiary faults with locally thick permeable horizons is a key factor in concentrating ore such as the Kearsage deposit (see Stop 4 and Fig. 12). White (1968) suggested that for the movement of ore fluids to occur, permeability due to fracturing was more important than primary permeability. Faults and small fractures cutting massive interior of lava flows were likely important for upward transport of ore fluids also. Overlapping of successive lava flows and minor unconformities suggests that simple up-dip movement of ore fluids was not likely without a network of fractures. Secondary Hydrothermal Minerals The rocks within the Keweenaw Peninsula native copper district were pervasively altered by lowtemperature, low-pressure hydrothermal/burial metamorphic fluids. Alteration was most intensely associated with the native copper deposits, although to some degree, secondary hydrothermal minerals occur in all rocks of the Portage Lake Volcanics. Areas in the Keweenaw Peninsula more distal to the native copper deposits were less altered. The intensity and degree of alteration also varies as a function of position within lava flows; the massive interiors of lava flows being much less altered whereas the lava flow tops are relatively more altered. Lava flows in close proximity to cross cutting features tend to be more altered. The minerals occur as amygdule and vein fillings, and as whole rock replacements. Within the Portage Lake Volcanics, some original igneous minerals are present in the massive interiors of flows, but secondary minerals exist in the massive interiors of all flows regardless of their thickness. While the massive interiors of lava flows contain secondary minerals, their original igneous geochemical composition is often only slightly modified by secondary hydrothermal processes. There are over 100 different secondary alteration minerals in the Keweenaw Peninsula; most of them are related to hydrothermal process and some are related to supergene processes. Only about 24 alteration minerals are common. Native copper with small quantities of native silver represents over 99% of the metallic minerals in the mined ore bodies of the district. Most of the native copper carries a small amount of arsenic in solid solution (typically less than 0.2 % arsenic in total copper + silver + arsenic; Broderick, 1929). Copper-nickel arsenides occur in veins that are paragenetically late (Moore, 1971; Stoiber and Davidson, 1959; Butler and Burbank, 1929). Within the native copper deposits, paragenetically late chalcocite occurs as small veins cutting lava flow top deposits, and as coatings on joints containing calcite in conglomerate deposits (White, 1968). 15 �Flow Top Deposits and Veins Microcline Chlorite Epidote Pumpellyite Prehnite Native Copper Datolite Silver Ankerite Quartz Sericite Calcite Arsenides Sulphides Albite Adularia Chlorite Laumontite Analcite Sulphates Conglomerate Deposits Little Little (barite, anhydrite, gypsum) Relative Age Abundant Relative Age Not abundant Figure 8: Paragenesis and relative abundance of secondary hydrothermal alteration minerals in the Keweenaw Peninsula native copper district. Modified from Butler and Burbank (1929). District-wide there is a well-defined mineral paragenesis (Fig. 8), although individual deposits may not exactly follow the district-wide timing of precipitation. There is a general spatial variation of hydrothermal minerals in the Calumet area of the district (Fig. 9). Epidote and the appearance of quartz are spatially associated with major native copper deposits (Stoiber and Davidson, 1959). A detailed study by Stoiber and Davidson (1959) of the Kearsarge deposit shows that native copper is much more irregularly distributed than secondary mineral zones, but there is a general correlation with the abundance of native copper associated with the variation of quartz and microcline (see Stop 4 and Fig. 14). On the tip of the Keweenaw Peninsula, the suite of hydrothermal alteration minerals consists of low temperature zeolite minerals except within about 750 m of the Keweenaw Fault where there is epidote (Cornwall, 1955; Cornwall and White, 1955). Amygdule-filling minerals are equivalent to zeolite and prehnite-pumpellyite metamorphic facies. The hydrothermal /metamorphic mineral zones dip more gently towards Lake Superior than the strata, implying that the strata were tilted prior to hydrothermal alteration (Livnat, 1983; Broderick, 1929). The paragenetic succession of alteration minerals at the Kearsarge deposit begins with low-temperature (80 to 100oC) agates followed by higher temperature native copper and temporally associated minerals (about 225OC) and the final stage is superimposition of lower temperature late-stage barren laumontite and calcite. This progression represents a waxing and waning of the hydrothermal system. Much later, the native copper has been altered by oxidized groundwaters generating copper oxide minerals. Native copper mineralization is younger than the Copper Harbor Conglomerate, which hosts rare veins of calcite and native copper (see Stop 6). White (1968) interpreted the age of native copper mineralization as after the deposition of parts or all of the Freda Sandstone. Minor amounts of 16 �native copper occur within the lower beds of the Jacobsville Sandstone near Rice Lake. Based on field relations, hydrothermal alteration is younger than deposition of rift-filling strata and at least some of the rift-flanking Jacobsville Sandstone. The absolute age of hydrothermal alteration is between 1060 and 1047 Ma (+/- ~ 20 Ma) (Bornhorst et al., 1988). This age is consistent with the approximate age of 1060 Ma for regional continental compression that caused reverse faulting along the Keweenaw Fault (Cannon et al., 1993). Thus, the age of hydrothermal alteration is about 1070 to 1040 Ma contemporaneous with regional continental compression and some 10 to 30 million years after eruption of the Portage Lake Volcanics. Calumet Cross Section Geographic Position Top of Portage Lake Volcanics ~ 200 oC Stratigraphic Position of Native Copper Mines %of District Production Epidote & quartz present 10%Quincy 38%C&H 500 ~200 to 250 oC 5%Osceola 21 %Kearsarge 0 250 to 300 oC Disappearance of ferrian prehnite ~ 3%Isle Royale 17%Baltic No actinolite oC < 300 to 325 ~ KeweenawFault Figure 9: Distribution of prominent secondary hydrothermal alteration minerals in the Portage Lake Volcanics in a cross-section in vicinity of Calumet at the center of the major deposits of the Keweenaw Peninsula native copper district. Data compiled from Livant (1983) and Stoiber and Davidson (1959) and modified from Bornhorst and Rose (1994). Genesis of the Native Copper Deposits Native copper occurs throughout the MCR in Wisconsin, Minnesota, and Ontario (Fig. 2) which suggests a regional distribution of mineralizing hydrothermal waters. The regional Cu-bearing hydrothermal fluids can be best explained by their generation during burial metamorphism of rift-filling basalts with temperatures reaching a thermal maximum 10 to 30 million years after the end of widespread rift magmatism (Fig. 3). The coincidence of regional continental compression with the thermal maximum provided an integrated paleohydrologic system through reactivated and new faults and fractures. This allowed the upward movement of hydrothermal fluids to 17 �focus in sites of future copper deposits at the very time of greatest fluid availability (Bornhorst 1997). During generation of the regional hydrothermal ore fluids, a few ppm of copper was leached from the rift-filling basalt strata (Jolly, 1974; White, 1968; Stoiber and Davidson, 1959). Simple calculations demonstrate that only a 3 km down dip volume of the 10 km thick rift-filling basalts along the 45 km strike length of the major copper deposits need be leached to generate sufficient copper for deposition, or in other words, the rift-filling basalts are a viable source rock for the copper. The hydrothermal fluids were low in sulfur since most of the sulfur in the buried rift-filling source rocks was degassed during eruption so subsequent deposition up dip in the same sulfurpoor rocks preferentially lead to deposition of native copper rather than copper sulfides. Precipitation of native copper was caused by mixing of ore fluids with cooler, more dilute shallower fluids, ore fluid-rock reactions, and cooling of ore fluids. The localization of large native copper deposits within the Keweenaw Peninsula rather than elsewhere in the MCR may be controlled by favorable geometric orientation within the regional continental compression stress field. Since much of the MCR strata is buried, perhaps another area of native copper deposits remains hidden. Phanerozoic The last events in the geologic development of the MCR in the Keweenaw Peninsula were the formation of the native copper deposits and deposition of the Jacobsville Sandstone during regional continental compression at 1.06 to 1.04 Ga; the Jacobsville deposition may have continued after compression until 1.03? Ga (Fig. 3). The Keweenaw Peninsula was subsequently subjected to a 500 million year period of erosion, from about 1.03 Ga to 0.5 Ga, 500 Ma and multiple kilometers of rock were eroded exposing the native copper deposits at the surface (Fig. 3) (Bornhorst and Robinson, 2004). Downward percolating groundwaters supergene altered native copper and produced a suite of including cuprite, tenorite, malachite, and chrysocolla. The rocks of the Keweenaw Peninsula were subsequently buried by Paleozoic sedimentary rocks associated with the Michigan basin beginning about 500 Ma (Fig. 3). Over the past two million years, the Keweenaw Peninsula was subjected to several continental glacial periods which removed all of the overlying Paleozoic sedimentary rocks with the exception of a Paleozoic outlier slightly south (Fig. 3). After the last glacial episode, the native copper deposits were exposed at roughly the same erosional level as at 500 Ma or the end of the Precambrian. The continental glaciers sculpted the bedrock of the Keweenaw Peninsula and when the last glacier retreated about 10,000 years ago, it left behind a variety of unconsolidated glacial-related sediments that included entrained boulders of native copper. The glaciers carved out the topographic low the Lake Superior basin corresponding to the less competent clastic sedimentary rocks under the center of the MCR. After the glaciers retreated, very large volumes of water filled this topographic low and initially all but the highest land elevations were underwater of a large glacial lake. The glacial lake levels successively dropped over time to the current level of Lake Superior (Farrand 1960). As the lake levels receded humans populated the area. Objectives of Field Trip 18 �This field trip is designed to provide a geologic overview of the Keweenaw Peninsula and the Keweenaw Peninsula native copper district. The Mesoproterozoic MCR bedrock and hosted native copper deposits are unconformably overlain by unconsolidated Pleistocene glacial sediments. The descriptions and stops of this field trip provide glimpses into both of these distinct geologic events. The MCR rocks and native copper deposits are the focus of this field guide. Figure 10 provides the regional geologic setting of the Stops. Figure 10: Geologic map of the far western part of the Upper Peninsula of Michigan showing field trip stops. 19 �Stop 1: Razorback Center Directions: Drive west through downtown Houghton on US-41to south M26. Drive 0.9 miles (1.4km) to Sharon Ave. at first stoplight. Turn left on Sharon Ave. for 0.1 miles(0.2km) to Razorback Dr. Turn right on Razorback Dr. for 0.1 miles (0.2km) to strip mall on left built on top of a small hill. Turn left and outcrop exposure is behind the strip mall. [UTM 5218745N 160379747E (NAD27 CONUS)] Figure 11: Photograph of the rock cut at Razorback Center looking northwest. The rock cut at the edge of the parking lot in the rear (south side) of Razorback Center, Houghton provides an excellent example of the characteristics of subaerial basalt lava flows that comprise the Portage Lake Volcanics, the host rock unit for native copper deposits of the Keweenaw Peninsula native copper district and the rock unit that holds up the spine of the Keweenaw Peninsula (Fig. 10). This of exposure of subaerial lava flows is located stratigraphically between the Calumet and Hecla and Kingston Conglomerates (Fig. 11). It is also located between the stratigraphic level of the Isle Royale and the Quincy Mines. The lava flows strike about N30oE and dip about 55o to the northwest (towards Lake Superior). The rock cut is at an oblique angle to the strike of the lava flows. When facing the exposure, the stratigraphic top is towards the northwest or toward the right. At the far eastern end, or far left, only the amygdaloidal top of the oldest basalt lava flow in this 20 �rock cut is exposed. Stratigraphically upwards, towards the west/right, this flow top is overlain by a thick section of dark-gray to black massive basalt representing the interior of a prominent lava flow. Progressively, the abundance of amygdules increases upwards and the color of the basalt changes to greenish tones reflecting an increased degree of alteration. This zone represents the amygdaloidal top of the prominent lava flow in this rock cut. The amygdules tend to be concentrated along layers and near the upper contact; they coalesce into a continuous now filled open space. The contact between the amygdaloidal top of this prominent lava flow and the massive basalt of the overlying flow is well exposed (Fig. 11). Along most of the exposed contact, amygdaloidal basalt lies directly below the massive basalt indicating the lava flow had a smooth top (pahoehoe lava flow), however, at the level of the parking lot, the planar contact bends and there is a small zone of brecciated flow top. The entire cross section of the prominent lava flow is exposed in the Razorback Center rock cut. A typical lava flow in the Portage Lake Volcanics is between 10 to 20 m thick, the prominent lava flow at Razorback Center. Stratigraphically further upwards, towards the west, the prominent lava flow is overlain by a thick section of dark-gray to black massive basalt representing the interior of the overlying flow (Fig. 11). On the far western end, there is amygdaloidal basalt representing the top of this overlying flow; an almost complete cross section of this flow is exposed here. Volcanic textures and structures at Razorback Center are typical of subaerial lava flows within the Portage Lake Volcanics. The basalts are mainly olivine tholeiites erupted as thick, ponded subaerial lava sheets. The very top and bottom of such lava flows typically consist of aphanitic chilled basalt. The contact between the underlying and overlying lava flows occurs where amygdules disappear abruptly and the overlying flow consists of massive basalt. The upper surface of the main flow was brecciated slightly by movement of lava after the formation of an upper crust, but rapidly grades downward to a non brecciated, highly vesicular flow top. The layered nature of amygdules in the prominent flow here at Razorback Center is likely a result of preferential accumulation of vesicles along laminar flow planes. The flow top breccia is laterally discontinuous for this flow. Slow cooling of the lava flow caused solidification toward the flow interior at a rate which allowed development of subophitic to ophitic textures (large oikocrysts of clinopyroxene enclosing a felted framework of An-rich plagioclase and intergranular olivine). The resulting massive, non-vesicular flow interior constitutes about two-thirds of the flow. The effects of regional hydrothermal alteration can be observed within the amygdaloidal flow tops. The massive interiors are much less altered except along fractures. The original plagioclase in the massive basalt has been replaced by albite and the mafic minerals by chlorite, pumpellyite, and iron oxides. The massive interior of the flow is much less altered than the flow top which represents a relatively impermeable horizon in the paleohydrologic system except in the vicinity of selected fractures. The pseudomorphic alteration minerals in the massive interior of the basalt are similar to those which fill the amygdules. The amygdules here are filled with a variety of secondary minerals including: calcite, chlorite, epidote, prehnite, pumpellyite, quartz (not in order of abundance), and traces of native copper. Late stage laumontite abundantly fills some amygdules. Stop 2: Float Copper US-41 Calumet 21 �Directions: Get back onto M26 at the traffic light and turn right. Stay on north M26/US41 and cross the bridge over to Hancock. Drive through downtown Hancock and continue 10 miles (16km) north to Calumet. Float copper is located on the left side of US41/M26 past the first traffic light just before Red Jacket Rd. [UTM 5233057N 160390423E (NAD27 CONUS)] A float copper boulder weighting 4,263 kg (9,392 lbs) is on display at Stop 2 (Fig. 10). This mass of glacially transported native copper was found in 1970 about 4.5 miles SW of Calumet in less than three feet of surficial sediments. Native copper deposits of the Keweenaw Peninsula were exposed at the bedrock surface at the time of the last period of Pleistocene glaciations. The glacial ice plucked masses of malleable native copper from the tabular lodes and fissures which were subsequently smoothed and flattened by abrasion from other rocks carried by the glacial ice. When the glaciers retreated about 10,000 years ago, unconsolidated rock debris (rounded boulders to clay sized material) were left behind by the melting ice including masses of native copper such as this one “floating” among the unconsolidated rock debris. While some of the rocks in the glacial deposits are from far north of the Keweenaw Peninsula, most of them are recognizable as from local MCR strata exposed in the Keweenaw Peninsula. The large float copper masses could not have moved far from their source, but smaller masses have been transported quite far and have been found in Lower Michigan and Wisconsin. The largest known float copper was discovered in the early 2000s and weighed about 25 tons (50,000 lbs) near the Houghton County airport; it was cut into smaller masses and sold to be smelted and refined. Most pieces of float copper are small, ranging from a few to 50 cm across. The famous example of float copper was the Ontonagon boulder, a 3,700 pound specimen visited by numerous explorers and finally removed from the Keweenaw to the nation’s capital in 1843. The Ontonagon boulder is part of the Smithsonian’s collection. This and other float copper masses have been surface altered by oxygenated groundwater and shallow precipitation since the glaciers retreated. This surface alteration consists of forms of copper including cuprite (copper oxide; Cu2O), tenorite (copper oxide; CuO), malachite (hydrated copper carbonate; (Cu2(CO3)(OH)2) and rarely azurite (hydrated copper carbonate, (Cu3(CO3)2(OH)2). Even when small cm sized masses of float copper are cut, the typical surface alteration is less than one mm thick; native copper is highly resistant to surface weathering. Float copper makes an attractive decorator specimen when a part of the surface is polished and buffed exposing shiny copper color. The basalt mine rock buildings are part of the Keweenaw National Historical Park. The park was established on October 27, 1992, by U. S. Congress Public Law 102-543. The enabling legislation ascertained that the Keweenaw was nationally significant because of: its unique geology; the prehistoric use of its copper by Native Americans; the importance of the region as a leading copper producer and developer of new technologies; its long history of corporate paternalism; and because it became home to so many European ethnic groups that migrated to the United States. Older mining districts typically had only single-industry economies and when the mines shut down, the communities suffered major contraction. In 1910, nearly 40,000 people resided within a few miles of Stop 2 whereas now, fewer people live in all of Houghton County. 22 �The idea that maybe the future of Calumet resided in its past was generated in the late 1980s; history could be “sold” to revitalize the community by increasing tourism. The national park itself only owns a few structures in Calumet, including these, and instead relies on public and private partners termed Keweenaw Heritage Sites. The heritage sites contain and interpret significant cultural and/or natural resources that together with park assets help tell the story of copper mining in the Keweenaw Peninsula. The Quincy Mine property on the edge of Hancock and the A.E. Seaman Mineral Museum on the campus of Michigan Tech are two among multiple Keweenaw Heritage Sites. Stop 3: Bumbletown Hill THE ROCK PILES DESCRIBED FOR THIS STOP ARE ON PRIVATE PROPERTY AND PERMISSION IS REQUIRED ACCESS THEM. Directions: Continue on US-41 3.6 miles (5.8 km) past headquarters of the Keweenaw National Historical Park denoted by park sign and large specimen of glacial float native copper to Bumbletown Rd. and turn left (west). Drive about 0.4 miles (0.6 km) to rock pile. [UTM 5237960N 160393345E (NAD27 CONUS)]. To access the overlook leave the rock pile and continue on Bumbletown Road west about 0.5 miles (0.8 km) to overlook at the top of the hill near towers. [UTM 5238215N 160392860E (NAD27 CONUS)] The description of this stop is reproduced from Bornhorst and Barron (2011). The Allouez conglomerate (informal member) is one of a small number of interflow clastic sedimentary horizons within the Portage Lake Volcanics visible in the rock pile at this stop (Fig. 10). This particular conglomerate bed can be traced along strike from the tip of the Keweenaw Peninsula, to at least the Mass area, a strike length of more than 120 km (Fig. 6). The Allouez conglomerate is stratigraphically just below the Greenstone flow, arguably the largest basalt flow in the world, within the Portage Lake Volcanics. The rock piles at the base of Bumbletown Hill are from the Allouez Mine. The Allouez conglomerate consists of mostly red-colored conglomerate with lesser amounts of sandstone and siltstone. The largest contained boulders at this locality are about 65 cm in diameter and the median size is about 8 cm. A pebble count of boulders more than 20 cm across gave the following results: mafic rock, mostly amygdaloidal basalt, 16%; quartz porphyritic rhyolite, 36%; feldspar porphyritic rhyolite, 11%; and granophyre, 37% (White, 1971). The mines on the Allouez conglomerate yielded only about 75 million pounds of refined copper (Table 2). Some evidence of native copper mineralization can be seen in rocks at this stop. Occasionally, one can find a specimen with native copper filling the void space between clasts and grains. Calcite and chlorite are the dominant pore-filling secondary minerals visible on this rock pile. Thin black veinlets cutting the Allouez conglomerate consist of calcite with chalcocite “dust.” While supergene alteration resulting from the downward percolation of groundwater is not common in most the native copper deposits, at this stop, supergene alteration minerals are common including chrysocolla, malachite, and cuprite. From the overlook on a clear day, Isle Royale may be seen 80 km to the northwest and the Huron 23 �Mountains may be seen beyond Keweenaw Bay, 60 km to the southeast. The land slopes very gradually to the northwest toward Lake Superior, as it does throughout most of the length of the Keweenaw Peninsula. The area is underlain mainly by conglomerates and sandstones of the Copper Harbor Formation dipping at about 20o. The southeastern flank of the Keweenaw Peninsula has a steeper slope at the skyline, following approximately the line of the Keweenaw fault. The low-lying plain between the fault and Keweenaw Bay is underlain by flat-lying Jacobsville Sandstone. Bumbletown Hill is located on the southwest side of the Allouez Gap, a NW- to SE-trending valley. The valley follows the Allouez Gap fault, a zone of faults and fractures, along which the Portage Lake Volcanics and Keweenaw fault, are offset. At this gap, the strike of the Portage Lake Volcanics swings from about N35oE to N50oE (Figs. 5 and 7). Almost every permeable horizon near the Allouez Gap fault contains above average amounts of native copper; nowhere else in the district are there so many mineralized beds (Fig. 7). About 60% of the district production can be linked to the fault as a primary pathway for ore fluids. The fault bisects the Kearsarge deposit (see Fig. 12), which was the second largest copper producer in the native copper district. The line of rock piles demarking the many mines along the Kearsarge deposit is a little more than 1,500 m southeast of Bumbletown Hill. The Kingston Mine, a small deposit that produced 20 million pounds of copper (1963 to 1968; one of the most recent native copper mines to open and last to close), is bisected by the Allouez Gap fault. About 1,200 m N65oE of the hilltop, the Houghton conglomerate and the Iroquois flow produced 33 million pounds of copper. Looking northeast along the strike of the Portage Lake Volcanics, one can see the cuesta form of the ridge upheld by the Greenstone flow. To the right of the ridge, the more distant hills are formed by lava flows lower in the Portage Lake Volcanics sequence. At Bumbletown Hill, the Greenstone flow is only 85 m thick, but the flow thickens abruptly to more than 400 m near end of the cuesta ridge. It dips northward at about 25o toward the center of the Lake Superior. The Greenstone flow can be traced along much of the Keweenaw Peninsula and has been stratigraphically and geochemically correlated with a similar unit on Isle Royale, 90 km away, on the opposite side of the rift. Thus, the areal extent of this great flow exceeds 5,000 km2, and its volume is on the order of 800 to1,500 km3 (Longo, 1983). The geochemical composition of the Greenstone flow magma is more evolved than typical olivine tholeiites of the Portage Lake Volcanics. 24 �Stop 4: Seneca Mine Rock Pile THE ROCK PILES DESCRIBED FOR THIS STOP ARE ON PRIVATE PROPERTY AND PERMISSION IS REQUIRED ACCESS THEM. Directions: Drive back to US41/M26 on Bumbletown Rd. and turn left. Continue northeast on US41/M26 0.6 miles(0.9km) to B St. Turn Left on B St. Drive about 0.3 miles(0.5km) to rock pile. [UTM 5238775N 160394119E (NAD27 CONUS)] The Kearsarge lode was worked by the Seneca Mine, one of multiple mines which produced native copper from the top of the Kearsarge basalt lava flow over a strike length of more than 12 km and down-dip as much as 2,500 m (Figs. 10 and 12). About 1,026 million kg of refined copper were produced at an average grade of 1.05% Cu, making the Kearsarge deposit the largest flow top hosted deposit and the second largest producer in the district behind the C&H Conglomerate mines (Table 2). Production of copper from the Kearsarge lode began in 1887 and stopped in 1967. The Kearsarge lava flow has been recognized for a distance of about 55 km along strike and dips between 35 and 40o NW (Fig. 12). It lies directly above the Wolverine Sandstone (Fig. 6). The amygdaloidal and/or brecciated top of the Kearsarge flow ranges from near zero up to 10 m in thickness. The productive top has an average thickness of around 2 m and consists of brecciated basalt (individual fragments of amygdaloidal basalt are generally less than 15 cm in greatest dimension). The brecciated basalt grades downward into amygdaloidal basalt with amygdules concentrated in layers. Further downward, the top grades into a zone of fewer and larger amygdules, and then into aphyric massive basalt in the interior of the flow. Just below the brecciated and/or amygdaloidal top of the flow, there is distinct plagioclase porphyritic basalt. The abundance and size of the plagioclase phenocrysts in this zone is variable, but they can make up a large percentage of the rock, with phenocrysts up to 2.5 cm in length. This zone is probably the result of plagioclase in situ floating during surface crystallization of the flow. Specimens with abundant plagioclase phenocrysts can be found on this rock pile. The basalt itself in the Kearsarge flow is well oxidized. Albitized and pumpellyitized basalt consists of pseudomorphically replaced plagioclase set in a fine-grained to cryptocrystalline groundmass. Original igneous minerals were replaced in areas where alteration was intense. Olivine is almost invariably completely replaced while other igneous mineral are replaced by alteration minerals to varying degree. The amygdule and interfragmental space-filling gangue minerals in the Kearsarge lode are generally (in order of most to least abundant): calcite, epidote, K-feldspar, quartz, and lesser amounts of chlorite, prehnite, pumpellyite, laumontite, and sericite. Native copper is closely associated in time and space with the secondary amygdule minerals (Stoiber and Davidson, 1959). Paragenetically, chlorite; epidote; microcline; and prehnite are early-formed minerals, and the latest-formed minerals are quartz; native copper; calcite; and chlorite (Fig. 14). A zonal stratabound arrangement of amygdule minerals in the Kearsarge deposit is seen in the Ahmeek Shaft No. 3 (Fig. 15). The zoning may be explained by deposition of secondary minerals from a hydrothermal solution moving along a permeable channel. 25 �SW NE Productive Area 90 Thickness 60 meters 30 Top of Wolverine Sandstone Location along strike South Centennial Kearsarge North 1 2 2 Wolverine Kearsarge 1 4 3 2 11 2 3 Allouez Gap Fault Seneca 2 3 Mohawk Ahmeek 2 1 5 3 4 2 1 4 6 Limit mining of Very high grade copper ore 0 Upper limit of quartz 600 1200 meters Lower limit of microcline Figure 12: Thickness of the Kearsarge lava flow from showing the location of the productive area where the top of the flow is thickest. Modified from Butler and Burbank (1929). The most productive area corresponds to the thickest part of the flow which is bisected by the Allouez Gap fault. Bottom diagram is a down-dip strike parallel section project to vertical showing distribution of higher grade native copper ore and occurrence of important alteration minerals. Modified from Stoiber and Davidson (1959). Abundance of quartz in amygdules is greater than 10 % on the down-dip side (lower) of the line shown and K-feldspar is absent on the down-dip side (lower) of the line shown. The Kearsage flow dips about 35 to 40o NW and all data are projected. 26 �Chlorite Microcline Prehnite Hematite Epidote Pumpellyite Quartz Sericite Native Copper Calcite Relative Age Less abundant More Abundant Figure 13: Paragenesis of secondary hydrothermal alteration minerals in the Kearsarge deposit at the Wolverine No. 2 Mine. North South Hanging Wall chlorite Hanging Wall microcline-calcite calcite-epidote calcite-epidote quartz-epidote chlorite Contact between Kearsarge flow top and overlying massive flow bottom 5 0 meters Abundant native copper Mineral Assemblage Band Volume Percent Amygdule Filling chlorite Chlorite Microcline Epidote Calcite Quartz Pumpellyite 100 0 0 trace 0 0 chloritemicroclinecalcite microclinecalcite quartzepidote calciteepidote 69-74 15-25 0-1 0-5 0-5 0-6 0-3 45-82 5-10 0-47 0-8 0-trace 0 0 90-96 0-1 4-9 0 0 0 12 87 1 trace Figure 14: Cross section of the top of the Kearsarge lava flow (amygdaloid) deposit showing the distribution of secondary hydrothermal amygdule-filling alteration minerals at the Ahmeek Mine, 35th level, 400 to 500 ft south of the shaft. Modified from Stoiber and Davidson (1959). Data from the back and walls are projected to a horizontal plane. There is a barren laumontitequartz-calcite zone not shown here. 27 �Chlorite and microcline would have been deposited first, along the outer limits of the solution channel; followed by quartz and epidote in the center of the channel; and finally, deposition of calcite in the remaining openings. This observation is consistent with the paragenetic relationships seen in individual samples. No strict correlation exists between the stratabound zoning and the grade of native-copper mineralization (Stoiber and Davidson, 1959). The amygdule minerals and grade of copper mineralization vary with depth. Within the upper limit of quartz (Fig. 12), the quartz content is typically about 15 % of open space fillings although it is considerably less than 10% at shallower depths. The lower limit of microcline may also mark the limit of significant copper mineralization. The amount of native copper present is much more irregular than variation of the mineralized zones. The Allouez Gap Fault bisects the thickest segment of the Kearsarge Flow along its 55 km strike length (Fig. 12). Higher grades and production occur northeast of the fault where fractures with orientations that parallel the fault are more abundant. Within the Allouez Gap Fault zone, early epidote and quartz were brecciated and recemented by calcite, quartz, and native copper. After another episode of brecciation, the fault zone was recemented again with calcite; quartz; and lesser laumontite (Butler and Burbank, 1929). Movement along the fault occurred before, during, and after deposition of native copper. The fault apparently was a conduit for transport of ore fluids to the permeable flow top. The coincidence of this fault with the relatively thick flow top resulted in the second largest deposit in the district. The Seneca Mine is an excellent locality to study the character of a representative basaltic flow top hosted native copper deposit. Specimens of massive basalt, massive basalt abundant plagioclase phenocrysts,and amygdaloidal basalt can be found on this rock pile. Masses of native copper are readily collectable especially when using a metal detector. Open-space filling minerals (amydgules and between breccia fragments) that occur in the lode can be found on the rock pile. Stoiber and Davidson (unpublished data) made a quantitative analysis of open-space filling minerals for the Seneca Mine rock pile and found open-space filling minerals consisted of: calcite, 57%; red feldspar 8%; pink feldspar 15%; epidote, 17%; prehnite, trace; pumpellyite, trace and quartz, trace. Many specimens contain multiple minerals and illustrate paragenetic relationships. Stop 5: Eagle River Falls Directions: Continue northeast on US-41/M26 9.5 miles (15km) to Phoenix and turn left on M26. Continue 2.3 miles (3.7km) to Eagle River and park by the bridge. [UTM 5251824N 160402193E (NAD27 CONUS)] The water falls of Eagle River is near the contact between the top of the Portage Lake Volcanics and the base of the Copper Harbor Formation (Fig. 10). The contact dips about 30o NNW. The beds strike roughly parallel to the shoreline of Lake Superior; the orientation of the Keweenaw Peninsula changes from NE in vicinity of Houghton to ENE at Eagle River to E-W near the tip. The tholeiitic basalt subaerial lava flows just below the contact are pahoehoe with ropy upper surface. The orientation of the ropes indicates that the flow was erupted from a vent to the north geographically under Lake Superior. That the ropy flow top is preserved suggests that little erosion occurred between deposition of the last of the lava flows of the Portage Lake Volcanics and the Copper 28 �Harbor Formation. The Copper Harbor Formation consists of red-brown rhyolite-pebble conglomerate, but includes many sandstone and even some shale beds. Under the bridge, one can get a good view of the lithology of the lower part of the Copper Harbor Formation. The Copper Harbor Formation was deposited in an alluvial fan shed off of a highland area to the SE (opposite of Lake Superior) likely buried under the rift-flanking Jacobsville Sandstone (Elmore, 1984). This contact marks an abrupt change in the geologic evolution of the Midcontinent rift. Below this contact there is a thick succession of basalt subaerial lava flows with more than 200 individual flows and a cumulative thickness of about 5,000 m, thus magmatic activity dominated the Midcontinent rift at that time. Abruptly above the contact lava flows are strikingly absent and clastic sedimentation dominated the Midcontinent rift. While generally absent, a last gasp of magmatic activity will be seen at Stop 10 where a thin package of mafic to intermediate volcanic rocks, the Lake Shore Traps, are interfingered within the Copper Harbor Formation. Stop 6: Great Sand Bay Directions: Continue driving northeast on M26 through Eagle River for 4.8 miles(7.7km) until the Great Sand Bay overlook. [UTM 5253974N 160406985E (NAD27 CONUS)] The description of this stop is reproduced with minor modifications from Bornhorst and Barron (2011). The Great Sand Bay overlook provides a beautiful view of Lake Superior (Fig. 10). Very large volumes of water filled the Lake Superior basin as a result of melting of the glaciers, turning it into a glacial lake. The levels of the glacial lakes depended on the position of the ice front, outlets, and crustal rebound (a result of removing the weight of the ice). There are 15 lake stages recognized in the Lake Superior basin (Farrand 1960). As the lake levels receded to the current level of Lake Superior, more and more of the Keweenaw Peninsula emerged. At the road level, the sand dunes are remains of the Lake Nipissing Stage (4,000 to 5,000 years ago) when the lake level was about 9 m (30 feet) higher than today. After lake stages at about 3,200, 2,000, and 1,000 years ago, the waters receded toward the present level termed Lake Superior. The underlying bedrock is the Copper Harbor Formation. In the Keweenaw Peninsula there is a succession of basalt lava flows interbedded near the middle of the formation (see Stop 8). The massive interiors of these lava flows are more resistant to erosion than the underlying and overlying conglomerates and sandstones of the Copper Harbor Formation. As a result, harbors such those at Eagle Harbor and Copper Harbor are maintained by lava flows visible at their mouths. While not visible, lava flows occur at the mouth of Great Sand Bay too. There are many extensive underwater fissure vein deposits which cross cut the Eagle River shoals located about 0.5 to 1 km offshore. Many of them are often quite rich in native copper and can contain long continuous stringers protruding up to 1.5 m in height and extending almost 6 meters in length. Most of veins are less than 50 cm in width and are primarily composed of quartz or calcite with minor amounts of laumontite , datolite, prehnite, and traces of silver. Veins will locally contain 29 �clay pockets which can produce well defined copper crystal specimens. The largest copper specimen ever recovered underwater was a massive 17 ton unattached copper boulder in July of 2001. It was recovered from one of these vein deposits north of Jacobs Creek in about 9 m of water. To date, there have been 36 underwater copper veins discovered from the eastern tip of Great Sand Bay to Eagle River, about 3.2 km west. Stop 7: Hebard Park Directions: Continue driving 10.5 miles (16.9km) east on M-26 until arriving at Hebard Park conglomerate exposure on left. [UTM 5258659N 160428890E (NAD27 CONUS)] The description of this stop is reproduced from Bornhorst and Barron (2011). The Copper Harbor Formation is exposed along the Lake Superior shoreline at Hebard Park (Fig. 10) and is stratigraphically above the Lake Shore Traps (Fig. 5). The lithologies at at this stop consist of interbedded conglomerates and sandstones that characterize the Copper Harbor Formation. Clast-supported conglomerate beds consist of rounded, cobble- to boulder-sized clasts with a matrix of coarse sand-sized subangular grains cemented with carbonate and iron oxide. Clasts are predominantly of silicic volcanic rocks, with subordinate basalt, pyroclastic, plutonic, and metamorphic rocks. Several finer grained interbeds higher in the exposed section exhibit crossbeds, current lineations, current ripples, parting lineation, and reduction spots. In particular, one should note the calcite-rich cemented zones that may represent vadose carbonate or paleocaliche (Kalliokoski, 1986). There is a thin continuous zone of laminated cryptoalgal carbonate, laterallylinked stromatolite, that is draped over cobbles and contorted layers in mudstone-siltstone. 30 �Stop 8: Hunter’s Point Park Directions: Continue driving 2.4 miles (3.8km) east on M-26 to North Coast Rd. and then turn left. Drive 0.3 miles (0.4km) to Harbor Coast Lane and turn right. Drive 0.3miles (0.4km), park at the end of the road and walk down to shoreline. [UTM 5258263N 160432253E (NAD27 CONUS)] Lake Superior Copper Harbor Formation Hunter’s Point Park Lake Shore Traps Copper Porters Island N 0 500 feet Copper Harbor Ft. Wilkins State Park Garden Copper Harbor Formation Figure 15: Geologic map of the Copper Harbor area taken directly from Cornwall (1955) showing the location of Hunter’s Point (Stop 8), Brockway Nose (part of Stop 9), and Fort Wilkins Historic State Park (Stop 10). Hunter’s Point Park was established in 2005 when funding provided by the Michigan Natural Resources Trust Fund and many generous private donors (www.hunters-point.org) allowed the land to be purchased. Prior to becoming an official park the point was a popular hiking destination for visitors (Fig. 10 and 15). The land owners subdivided the area for residential housing which would have restricted public access without its conversion into a park. The name of Hunter’s Point is uncertain but it could have been named after A.W. Hunter, an early resident in the town of Copper Harbor who purchased the point from the U.S. Government. The Copper Harbor Formation is overall composed of volcanogenic clastic sedimentary rocks, dominantly conglomerates with lesser sandstone, siltstone, and shale such as observed at Stop 7. These rocks were deposited in a fining upward prograding alluvial fan complex (Elmore, 1984). Typically conglomerates are composed of clasts with a ratio of mafic-to-intermediate+felsic composition of about 2:1 (Daniels, 1982). Towards the tip of the Keweenaw Peninsula, the Copper Harbor Formation is informally subdivided into an inner (land side) “member” and an outer (lake side) “member.” Between these two “members” there is a thin succession of interbedded lava flows collectively known as the Lake Shore Traps. The Lake Shore Traps consist of Fe-rich olivine tholeiite, basaltic andesite, and andesite lava that were erupted during the waning stage of volcanism within the MCR; the youngest flows tend to be more intermediate in 31 �composition. At 1087.2 +/- 1.6 Ma (Davis and Paces, 1990), the Lake Shore Traps are among the youngest magmatism within the MCR. The thickest section of the Lake Shore Traps is about 15 km to the east at the tip of the peninsula. Volcanologically, the lower lava flows are interpreted as erupted as ponded sheets while the upper lava flows erupted on a low positive slope such as a shield volcano. The Lake Shore Traps were subaerially erupted pahoehoe lava flows. At Hunter’s Point, the top of the andesitic lava flows of Lake Shore Traps are conformably overlain by contact conglomerates of the Copper Harbor Formation (Fig. 15) simple geo map of Copper Harbor and Hunter’s point). The strike of bedding is about E-W and dip is about 35o to the north (towards the lake). The orientation of the contact is roughly parallel to the orientation of Hunter’s Point. From the Hunter’s Point parking lot, follow the walkway to beach towards the west side of the point. As the walkway ends, you will be on outcrops of lava flows of the Lake Shore Traps (Fig. 15). Walking to the east, the beach gives way to a rocky shoreline. In erosional coves, you can see contacts between lava flows, represented by vesicular to amygdalodoidal andesitic lava (top of the lava flow) overlain by massive andesitic lava (massive interior of the overlying lava flow). The massive lava flow interiors within the Lake Shore Traps often retain relict olivine and interstitial glass due to the overall low degree of alteration (weathering and hydrothermal). Highly visible red hematitic bands form circular patterns within the massive interior; this banding is interpreted to be the result of alteration. Secondary minerals filling amygdules include agate, chalcedony, quartz, laumontite, analcite, calcite, and smectite in amygdules; this suite of minerals is equivalent to zeolite facies metamorphism. In contrast, in massive lava flow interiors within the Portage Lake Volcanics the olivine and interstitial glass are completely replaced by Mg-Fe phyllosilicates and amygdule filling minerals are equivalent to higher degree of metamorphism, greenschist facies. The Lake Shore Traps are geographically more distal to the thermal high and increased hydrothermal activity that resulted in the native copper deposits, hence, lower degree and grade of burial metamorphic/hydrothermal alteration. To the west from the walkway, you can see a rocky point extending towards Lake Superior, the rocks in this point are conglomerates of the Copper Harbor Formation. The sharp contact between the uppermost lava flow of the Lake Shore Traps and the conglomerates can be viewed on the eastern edge of this rocky point. The conglomerate above the contact is dominated by rounded to sub-rounded boulders that are matrix-supported. There are proportionately more basaltic and andesitic clasts in this conglomerate bed than stratigraphically higher elsewhere along the Lake Superior shoreline such as at Stop 7 as these clasts are derived from erosion of the Lake Shore Traps updip towards the highlands on the edge of the rift (the updip rocks are now missing having been removed by erosion). The very poor sorting and fine matrix-supporting the clasts suggest this conglomerate could have been deposited as a debris flow. Sedimentary debris flows are common in alluvial fan depositional environments. The Copper Harbor Formation was deposited in an alluvial fan derived from highlands to the south in the vicinity of Keweenaw Bay. Additional outcrops of the Copper Harbor Formation can be seen on the far western end of the cobble beach. These outcrops consist of interbedded conglomerates and sandstone that are typical of the formation as a whole. The conglomerates are described at Stop 7. There are several prominent 32 �white-colored calcite -filled fractures (calcite veins) within these outcrops. The calcite veins are northerly oriented consistent with the orientation of faults cutting the Portage Lake Volcanics about 5 km to the south. Calcite veins are a common occurrence in the Copper Harbor Formation and some of them contain native copper such as those described at Stop 6, Great Sand Bay, and at Stop 10, Fort Wilkins. Stop 9: Brockway Nose and Brockway Mountain Viewpoints Directions: Continue driving 2.7 miles (3.8km) east on M-26 to Brockway Mtn. Drive. Turn right and drive 0.6 miles (0.4km) to Brockway Nose turnoff. [UTM 5257463N 160432304E (NAD27 CONUS)] Brockway Mountain Drive intersects M-26 just west of Copper Harbor. After a steep climb upwards there is a pullover at the second hairpin curve which is Brockway Nose viewpoint (Figs. 10 and 15). Brockway Nose provides an excellent view of Copper Harbor and Lake Fanny Hooe (Figure for Hunter’s Point). The top of Brockway Mountain is accessed by continuing upwards from Brockway Nose. Brockway Mountain is a conglomerate ridge that reaches and elevation of over 400 m, with excellent views of the ridge and valley topography of the northern shore of the Keweenaw Peninsula. From Brockway Nose viewpoint, the town of Copper Harbor is the prominent visible feature (Fig. 15). The town of Copper Harbor began as a boom town in 1843, following the discovery of copper in the vicinity. Porter's Island, at the mouth of Copper Harbor on the west side (left) was the site of the first government land office. Hunter’s Point is west of Porter’s Island. On the east side of the mouth of Copper Harbor, the Copper Harbor Lighthouse, built in 1866, is visible. Lake Fanny Hooe is located south of Copper Harbor. Fort Wilkins is located on the north shore of Lake Fanny Hooe on the thin strip of land between the lake and harbor. It was built in 1844, with the intent to protect the miners from potentially hostile Indians. Fort Wilkins is now a Historic State Park and is discussed more at Stop 10. Nearby, the Estivant Pine is a 2.06 km2 nature sanctuary established in 1973, containing one the last stands of virgin white pines in the Midwest and the last stand in the Upper Peninsula. Some of the trees are up to 600 years old (www.michigannature.org). In 1955, the white pine was designated the state tree of Michigan. Copper Harbor and several other harbors between here and Eagle River are located within the Lake Shore Traps. Dipping massive interior of the basaltic to andesitic lava flows of the Lake Shore Traps occur at the head of the harbors. From the Brockway Mountain viewpoint there are an excellent 360o views. Underfoot, the Copper Harbor Conglomerate dips about 20o to the north. Near the base of the ridge on the south side, opposite Lake Superior, there is an exposure of a single basaltic lava flow erupted as part of the Lake Shore Traps. With care, southwest of the gift shop at the high point, one can view the dipping conglomerates of the Copper Harbor Formation and see the lava flow near the base of the ridge. To the west, the Lake Shore Traps form island chains on a prominent ridge in the vicinity of Agate Harbor and Esrey Park. The rocks of the Copper Harbor Formation are found in the drowned 33 �valleys and along the outer ridge jutting into Agate Harbor and associated island chain. The ridges of the Lake Shore Traps and Copper Harbor Formation along the Keweenaw Peninsula’s north shore are also the site of numerous shipwrecks. Lake Bailey (with the small island) and Lake Upsom occupy a topographically low valley on a finer-grained clastic horizon (sandstone and siltstone) within the Copper Harbor Formation which is overall composed of conglomerates. Just to the south of Lake Bailey, is the ridge of Mt. Lookout, marking the contact between the basal conglomerates of the Copper Harbor formation and the uppermost basalt lava flows of the Portage Lake Volcanics. The inland lake almost directly south, is Lake Medora, and just before the lake is a prominent ridge which marks the stratigraphic position of the Greenstone flow (see Stop 4). In the distance, farther to the south across Lake Medora, is Mount Bohemia, a dioritic stock-sized intrusion within the lower section of the Portage Lake Volcanics. To the southwest, a distant ridge is Gratiot Mountain, which is a small shallow rhyolite intrusive body that cuts the Portage Lake Volcanics. To the east are the communities of Copper Harbor and Lake Fanny Hooe (better viewed from Brockway Nose), both of which occupy the same stratigraphic horizon as Lake Bailey. Just south of Copper Harbor is a golf course that is part of Brockway Mountain lodge. Brockway Mountain lodge was built during the Great Depression in the 1930’s by the WPA. To the north, Lake Superior is the prominent feature. On the skyline 65 km away, is Isle Royale National Park, which can be visible on a clear day. The skyline of Isle Royale is formed by the Greenstone Flow, as it is on the Peninsula. The beds on Isle Royale dip towards the Keweenaw Peninsula forming the Lake Superior “syncline.” Viewed from here, the Midcontinent Rift proper extends from the Keweenaw Fault, originally a graben bounding fault on the edge of the rift, just south of Mt. Bohemia to the Isle Royale Fault, also originally a graben bounding fault on the edge of the rift, just northwest of Isle Royale. Glacial erosion exposed Keweenawan and pre-Keweenawan relatively hard and competent bedrock on the edges of the Midcontinent rift system. Dipping well-cemented conglomerates of the Copper Harbor Formation are exposed at Brockway Mountain and basaltic lava flows of the Portage Lake Volcanics are exposed when viewing south. Both are relatively resistant to glacial erosion. On Isle Royale, on the southeast (Keweenaw side) are exposed the same conglomerates of the Copper Harbor Formation and on the northwest side, there are exposed basaltic lava flows of the Portage Lake Volcanics. In the center of what is now Lake Superior, much less competent, nearly flat lying, very fine sandstone and siltstone of the Freda Formation was at the bedrock surface. The latest glacial advance(s) preferentially eroded out the less competent rocks in the center of the rift, resulting in present day Lake Superior following the horseshoe shape of the MCR. Very large volumes of water filled the basin as a result of melting of the glaciers, turning it into a glacial lake. The Duluth Glacial Lake was the largest of these glacial lakes and only elevations above roughly 400 m (1,300 ft) were emergent. Brockway Mt. and Mt. Bohemia. 34 �Stop 10: Fort Wilkins Historic State Park Directions: Continue driving east on M-26 4.6 miles (7.4 km) until arriving at entrance to Fort Wilkins State Park on right. [UTM 5257338N 160434763E (NAD27 CONUS)] The description of this stop is reproduced with modifications from Bornhorst and Barron (2011). Fort Wilkins was built in 1844 by the U.S. Army to provide order on the Keweenaw frontier and to protect the copper resources during the Civil War (Figs. 10 and 15). The army built 27 structures to house two full strength infantry divisions. After the soldiers were needed in the Mexican War in 1846, the fort was abandoned. Fort Wilkins became a State Park in 1923. During the 1930s under the Work Project Administration, the fort underwent extensive restoration. Many of these structures still survive today and have been either been restored or rebuilt after archeological excavations. Today, the restored buildings are a museum and contain exhibits on the mining history of the area. Fort Wilkins is a popular destination in the summer for recreation and camping. Considerable exploration activity took place in the immediate vicinity of the fort, and there are shafts and exploration pits between Lake Fanny Hooe and the harbor, mostly from exploration during the period from1843 to 1846. Just north of the park store, several pits provide evidence of early mining activity by European settlers. The Pittsburgh and Boston Mining Company operated here in the 1840's on a vein of native copper within the Copper Harbor Formation; the vein was reported to be up to 0.3 m wide. This venture was not profitable. In 1853 and for several decades thereafter mining activity took place about 4.4 km south of the fort in a series of workings called the Clark Mine. The mineralization at the Clark Mine is hosted in both fissures and basalt flow tops. It consists of prehnite, epidote, analcite, quartz, laumontite, adularia, microcline, chlorite, datolite, calcite and several copper minerals including native copper, chalcocite, cuprite and tenorite. Agates are conspicuous as vesicle fillings in the Copper Harbor area especially in the Lake Shore Traps. Opposite Fort Wilkins, on the harbor shoreline is a view of the Copper Harbor Lighthouse, one of the first on Lake Superior built in 1866. Near the lighthouse on the Lake Superior shoreline is the famous "green rock". The "green rock" is a vein that was described by Douglass Houghton. Houghton himself may have never really understood the uniqueness of the district. Conventional wisdom at the time led him to the interpretation that the “green rock” was the surficial alteration of a sulfide ore (Krause, 1992). Nevertheless, Houghton had a profound impact in promoting the district. His report to the Michigan legislature started the first major mining rush in North America to the Keweenaw Peninsula where the first economic discovery in 1845 at the Cliff Mine (Stop 11) was followed by many more until mining ceased in 1968. Douglas Houghton drowned in 1845 near Eagle River, MI while leading a geological expedition. 35 �Stop 11: Cliff Mine Rock Pile Directions: Get back onto M26 and continue driving east to stoplight in Copper Harbor. Turn right and drive 21.8 miles(35km) southwest on US41 past Phoenix until Cliff Dr. Turn right on Cliff Dr. and drive 0.4 miles(0.6km) to mine site.[UTM 52247173N 160400875E (NAD27 CONUS)] The description of this stop is reproduced with minor modifications from Bornhorst and Barron (2011). Fissure (vein) deposits were of little importance to the overall copper production from the Keweenaw Peninsula native copper district (Figs 10 and 7). Only a few fissure mines, including the Cliff Mine, were profitable. The Cliff Mine worked the Cliff fissure (vein) from 1845 to 1887 and produced a total of about 38 million lbs of refined copper (Table 2). The Cliff fissure is nearly at right angles to the attitude of bedding and dips steeply to the east. The productive portion of the fissure is under the Greenstone flow. While most of the mineralization was confined to the fissure, some lava flow tops (amygdaloids) cut by the fissure contained native copper. Multiple large masses of native copper, some up to 100 tons, were taken out of the Cliff Mine. Among the fissure deposits, the Cliff Mine produced the most native silver. Minerals other than native copper and native silver include adularia, apophyllite, calcite, chlorastrolite, chlorite, datolite, epidote, laumontite, and prehnite (alphabetical). Many specimens contain multiple minerals and illustrate paragenetic relationships. Fissures ranges in size from tight cracks to more than 3 m wide. In this part of the native copper district, fissures strike across the lava flows and dip steeply. Fissures formed as tension cracks related to bending of the lava beds, transverse to the axis of the MCR (Butler and Burbank, 1929). The steep ridge near the Cliff rock pile is the Greenstone flow (see also Stop 9). Here it makes up the entire high ridge from bottom to top and with a northward dip of about 25o. The very thick massive relatively impermeable interior of the Greenstone flow likely played an important role in the localization of native copper. The fissures acted as efficient pathways for fluid movement. On a local scale, fluids migrating upward through these open fractures and were impeded beneath the massive interior of the Greenstone flow and were forced to move laterally into adjacent permeable horizons. In general, flows beneath the thicker section of the Greenstone flow in this area contain more dispersed native copper than elsewhere, but economic deposits are not common. 36 �Stop 12: Jacobsville Formation M-26 Tamarack Directions: Drive 6 miles (9.6km) south west on Cliff Dr. until it intersects with US41/M26. Turn right and continue 5.2 miles (8.3km) to Calumet and turn left at the second stoplight onto Lake Linden Ave/M26 south. Drive 3.9 miles(6.3km) downhill to Lake Linden(blinking light) and turn right on M26. Drive 5.8 miles (9.3km) through Tamarack City to sandstone roadside outcrops. [UTM 5222120N 16038915E (NAD27 CONUS)] The description of this stop is reproduced from Bornhorst and Barron (2011). The Jacobsville Sandstone is a red-bed succession consisting of feldspathic and quartzose sandstones, conglomerates, siltstones, and shales up to 1,000 m thick that were deposited by fluvial processes in a rift-flanking basin (Fig. 10). Overall, there are neither interbedded lava flows nor cross-cutting dikes and, thus, the age of the Jacobsville Sandstone is inferred to be ca. 1,060 to 1,020 Ma. Jacobsville sedimentation was the last Precambrian event associated with the development of the MCR. The Jacobsville Sandstone at this stop displays features characteristic of the unit as a whole. At the northeastern end of the outcrop, reddish shale and red-brown siltstone are exposed at the highway level. They are overlain by two fining-upward sequences of conglomerate and red, red-brown, and white cross-bedded sandstone. The lower conglomeratic bed is planar and can be traced 30 m to the southwest, along with the directly underlying shale and siltstone. Farther to the southwest, the section is almost entirely cross-bedded red sandstone; some beds are contorted and mottled. The sandstone consists of almost equal parts of rounded-to-sub-rounded quartz, feldspars, and lithic fragments. Clasts in the lower conglomerate are predominately sub-angular, and rhyolitic in composition, with subordinate mafic volcanic rocks. Mining history will be viewed along M-26 from Lake Linden to Mason as an extension of the actual Stop described above, history summarized here from Molloy (2007). Just before leaving Lake Linden on the left is the C&H Mill. The C&H Mill was first built in 1867 with several expansions as milling practices changed and closed in 1956. Like Quincy, C&H also reclaimed copper from the sand tailings. The Houghton County Historical Museum is located on the edge of Lake Linden and exhibits mining and local history. Just outside of Lake Linden on the left are the remains of the Calumet and Hecla (C&H) smelter. C&H was the largest native copper producer in the district. The large building at the north end of the site next to the highway was the C&H mineral storage building where crushed and concentrated copper ore was smelted. It is now occupied by Peninsula Copper Industries which primarily recovers copper from scrap copper such as printed circuit boards to make copper sulfate as a fungicide for the wood preservative industry and to make other specialty copper compounds. About 2 miles (3.2 km) southwest of Lake Linden, is the only remaining steam stamp that was part of the Ahmeek Mill. The steam-driven stamp hammers could deliver about 104 blows per minute and process 7,000 tons of ore a day. About 1.5 miles southwest of the Ahmeek Mill, the Quincy Mining Company built a reclamation plant in 1942 to 1943 to reprocess the stamp sand tailings along Torch Lake, and from 1943 to 1967 recovered approximately 50,000 tons of copper. One of the mining dredges used in the recovery process sank in a storm in 1956 and is located just offshore. The foundations between the road and the dredge are part of the Quincy Mills for processing native copper ore. 37 �Acknowledgements We thank Allan Blaske for valuable comment that improved this field guide. References Cited Bornhorst, T.J., 1997, Tectonic context of native copper deposits of the North American Midcontinent Rift system: Geological Society of America Special Paper 312, p. 127-136. Bornhorst, T.J., and Barron, R.J., 2011, Copper deposits of the western Upper Peninsula of Michigan: Geological Society of America Field Guide, v. 24, p. 83-99. Bornhorst, T. J., and Lankton, L. D., 2009, Copper mining: A billion years of geologic and human history: in Schaetzl, R., Darden, J., and Brandt, D. (eds.), Michigan Geography and Geology, Pearson Custom Publishing, New York, p. 69-90. Bornhorst, T.J., Paces, J.B., Grant, N.K., Obradovich, J.D., and Huber, N.K., 1988, Age of native copper mineralization, Keweenaw Peninsula, Michigan: Economic Geology, v. 83, p. 619-625. Bornhorst, T. J., and Robinson, G.W., 2004, Precambrian aged supergene alteration of native copper deposits in the Keweenaw Peninsula: Michigan: Institute on Lake Superior Geology Proceedings, v. 50, part 1, p. 40-41. Bornhorst, T.J., and Rose, W.I., Jr., 1994, Self-guided geological field trip to the Keweenaw Peninsula, Michigan: Institute on Lake Superior Geology Proceedings, v. 40, part 2, 185 p. Bornhorst, T.J., Rose, W.I., Jr., and Paces, J.B., 1983, Field guide to the geology of the Keweenaw Peninsula, Michigan: Institute on Lake Superior Geology, v. 29, part 2, 116p. Bornhorst, T.J., and Williams, W.C., in press, The Mesoproterozoic Copperwood sedimentary rockhosted stratiform copper deposit, Upper Peninsula, Michigan: Economic Geology. Broderick, T.M., 1929, Zoning in Michigan copper deposits and its significance: Econ. Geol., v. 24, p. 149-162, 311-326. Broderick, T.M., 1931, Fissure vein and lode relations in Michigan copper deposits: Economic Geology, v. 26, p. 840-856. Broderick, T.M., 1935, Differentiation in lavas of the Michigan Keweenaw: Geological Society of America Bulletin, v. 46, p. 503-558. Broderick, T.M., and Hohl, C.D., 1935, Differentiation in traps and ore deposition: Economic Geology, v. 64, p. 342-346. Butler, B.S., and Burbank, W.S., 1929, The copper deposits of Michigan: U.S. Geological Survey Professional Paper 144, 238 p. 38 �Cannon, W.F., 1992, The Midcontinent Rift in the Lake Superior region with emphasis on its geodynamic evolution: Tectonophysics, v. 213. p. 41-48. Cannon, W. F., 1994, Closing of the Midcontinent Rift - A far field effect of Grenvillian contraction: Geology, v. 22, p. 155-158. Cannon, W.F., and Hinze, W.J., 1992, Speculations on the origin of the North American Midcontinent rift: Tectonophysics, v. 213, p. 49-55. Cannon, W. F., Green, A. G., Hutchinson, D. R., Lee, M.W., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American mid-continent rift beneath Lake Superior from Glimpse seismic reflection profiling: Tectonics, v. 8, p. 305-332. Cannon, W. F., Peterman, Z.E., and Sims, P.K. 1993, Crustal-scale thrusting and origin of the Montreal River monocline - A 35-km-thick cross section of the Midcontinent Rift in northern Michigan and Wisconsin: Tectonics, v. 12, p. 728-744. Catacossinos, P.A., Harrison, W.B., Reynolds, R.F., Westjohn, D.B., and Wollensak, M.S., 2001, Stratigraphic lexicon for Michigan: Michigan Department of Environmental Quality, Geologic Survey Division Bulletin 8. Cornwall, H.R., 1951a, Differentiation in lavas of the Keweenawan series and the origin of the copper deposits of Michigan: Geological Society of America Bulletin, v. 62, p. 159-201. Cornwall, H.R., 1951b, Differentiation in magmas of the Keweenaw series: Journal of Geology, v. 59, p. 151-172. Cornwall, H.R., 1951c, Ilmentite, magnetite, hematite, and copper in lavas of the Keweenawan series: Economic Geology, v. 46, p. 51-67. Cornwall, H.R., 1955, Geologic map of the Fort Wilkins quadrangle, Michigan: U.S. Geol. Survey Geologic Quadrangle Maps of the United States Map GQ-74. Cornwall, H.R. and White, W.S., 1955, Bedrock geology of the Manitou Island quadrangle, Michigan: U.S. Geol. Survey Geologic Quadrangle Maps of the United States Map GQ-73. Daniels, P. A., 1982, Upper Precambrian sedimentary rocks: Oronto Group: Geological Society of America Memoir 156, p. 107-134. Davis, D.W., and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system: Earth and Planetary Science Letters, v. 97, p. 54-64. Elmore, R.D., 1983, Precambrian non-marine stromatolites in alluvial fan deposits, the Copper Harbor Conglomerate, upper Michigan: Sedimentology, v. 30, p. 829-842. 39 �Elmore, R.D., 1984, The Copper Harbor Conglomerate: A late Precambrian fining-upward alluvial fan sequence in northern Michigan: Geological Society of America Bulletin v. 95, p. 610-617. Elmore, R.D., Milavec, G.J., Imbus, S.W., and Engel, M.H., 1989, The Precambrian Nonesuch Formation of the North American Mid-Continent Rift, sedimentology and organic geochemical aspects of lacustrine deposition: Precambrian Research, v. 43, p. 191-213. Farrand, W.R., 1960, Former shorelines in western and northern Lake Superior basin: unpublished Ph.D. dissertation No. 5366, University of Michigan, Ann Arbor, 226p. Heaman, L.M., Easton, R.M., Hart, T.M., MacDonald, C.A., Hollings, P., and Smyk, M., 2007, Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario: Canadian Journal of Earth Sciences, v. 44, p. 1055-1086. Hinze, W.J., Braile, L.W., and Chandler, V.W., 1990, A geophysical profile of the southern margin of the Midcontinent rift system in western Lake Superior: Tectonics, v. 9, p. 303-310. Hoffman, P. F., 1989, Precambrian geology and tectonic history of North America: in Bally, A.W., and Palmer, A.R., eds., The Geology of North America-An overview, Boulder, Colorado, Geol. Soc. America, The Geology of North America, v. A, p. 447-512. Huber, N.K., 1975, The geologic story of Isle Royale National Park: U. S. Geological Survey Bulletin 1309, 66p. Jolly, W.T., 1974, Behavior of Cu, Zn, and Ni during prehnite-pumpellyite rank metamorphism of the Keweenawan basalts, northern Michigan: Economic Geology, v. 69, p. 1118-1125. Kalliokoski, J., 1982, Jacobsville Sandstone: Geological Society of America Memoir 156, p. 147-155. Kalliokoski, J., 1986, Calcium carbonate cement (caliche) in Keweenawan sedimentary rocks (~1.1 Ga), Upper Peninsula of Michigan: Precambrian Research, v. 32, p. 243-259. Kalliokoski, J., 1988, Jacobsville Sandstone: An up-date: in Upper Keweenawan rift-fill sequence Midcontinent rift system, Michigan, Michigan Basin Geological Society 1988 Fall Guidebook, p. 127136. Kalliokoski, J., and Welch, E.J., 1985, Keweenawan-age caliche paleosol in the lower part of the Calumet and Hecla Conglomerate, Calumet, Michigan: Geological Society of America Bulletin, v. 96, p. 1188-1193. Krause, David., 1992, The making of a mining district: Keweenaw native copper 1500-1870: Wayne State University Press, Detroit, MI, 305 p. Lane, A.C., 1911, The Keweenawan series of Michigan: Michigan Geological and Biological Survey Publication 6 (Geology series 4), 297p. Livnat, A., 1983, Metamorphism and copper mineralization of the Portage Lake Lava Series, northern Michigan: Ph.D. Dissertation, University of Michigan, Ann Arbor, 292p. 40 �Longo, A.A., 1982, A geochemical correlation, with correlative inferences from petrographic and paleomagnetic data, of the Greenstone flow, Keweenaw Peninsula and Isle Royale, Michigan: Institute on Lake Superior Geology Proceedings, v. 28, part 1, p. 22-23. Maki, J.C., and Bornhorst, T.J., 1999, The Gratiot chalcocite deposit, Keweenaw Peninsula, Michigan: Institute on Lake Superior Geology Proceedings, v. 44, p. 33-34. Mauk, J.L., Brown, A.C., Seasor, R.W., and Eldridge, C.S., 1992, Geology and stable isotope and organic geochemistry of the White Pine sediment-hosted stratiform copper deposit: Society of Economic Geologists Guidebook Series, v. 13, p. 63-98. Merk, G.P., and Jirsa, M.A., 1982, Provenance and tectonic significance of the Keweenawan interflow sedimentary rocks: Geological Society of America Memoir 156, p. 97-105. Milstein, R.L., 1987, Anomalous Paleozoic outliers near Limestone Mountain, Michigan: Geological Society of America Centennial Field Guide, v. 3, p. 263-268. Molloy, L.J., 2007, A Visitor’s Guide to the Historic Quincy Mine: published by Great Lakes Geoscience LLC, 61 p. Moore, P.B., 1971, Copper-nickel arsenides of the Mohawk No. 2 mine, Mohawk, Keweenaw Co., Michigan: American Mineralogist, v. 56, 1319-1331. Nicholson, S.W., 1991, Geochemistry, petrography, and volcanology of rhyolites of the Portage Lake Volanics, Keweenaw Peninsula, Michigan, U. S. Geological Survey Bulletin, 1970B, p. B1-B57. Nicholson, S.W., and Shirey, S.B., 1990, Evidence for a Precambrian mantle plume: a Sr, Nd, and Pb isotopic study of the Midcontinent Rift System in the Lake Superior region: Journal of Geophysical Research, v. 95, p. 10851-10868. Nicholson, S.W., Shirey, S.B., Schulz, K.J., and Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504-520. Ohr, M., 1993, Geochronology of diagenesis and low-grade metamorphism in pelites: Ph.D. dissertation, The University of Michigan, Ann Arbor, MI.,161 p. Paces, J.B., 1988, Magmatic processes, evolution and mantle source characteristics contributing to the petrogenesis of Midcontinent rift basalts: Portage Lake Volcanics, Keweenaw Peninsula, Michigan: Ph.D. Dissertation, Michigan Technological University, Houghton, 413p. Paces, J.B., and Bell, K., 1989, Non-depleted sub-continental mantle beneath the Superior Province of the Canadian Shield: Nd-Sr isotopic and trace element evidence from Midcontinent rift basalts: Geochimica Cosmochima Acta, v. 53, p. 2023-2035. 41 �Paces, J.B., and Bornhorst, T.J., 1985, Geology and geochemistry of lava flows within the Copper Harbor Conglomerate, Keweenaw Peninsula, Michigan: 31st Annual Institute on Lake Superior Geology Proceedings (Kenora, Ontario), p. 71-72. Paces, J.B., and Miller, J.D., Jr., 1993, Precise U-Pb ages of the Duluth Complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagmatic, and tectnomagmatic processes associated with the 1.1 Ga Midcontinent Rift system: Journals of Geophysical Research, v. 98, p. 13,997-14,013. Stoiber, R.E., and Davidson, E.S., 1959, Amygdule mineral zoning in the Portage Lake Lava Series, Michigan copper district: Economic Geology, v. 54, p. 1250-1277, p. 1444-1460. White, W.S., 1960, The Keweenawan lavas of Lake Superior, an example of flood basalts: American Journal of Science, v. 258A, p. 367-374. White, W.S., 1968, The native-copper deposits of northern Michigan: in Ridge, J.D., ed., Ore Deposits of the United States, 1933-1967 (the Graton Sales volume), American Institute of Mining, Metallurgical, and Petroleum Engineering, New York, p. 303-325. White, W.S., 1971, Field Trip A-2 – Houghton to Calumet via South Range quarry and Eagle River: Society of Economic Geologists, Guidebook for field conference, Michigan copper district, Sept. 30-Oct. 2, 1971, p. 68-75. Weege, R.J., and Pollack, J.P., 1971, Recent developments in native-copper district of Michigan: Society of Economic Geologists Field Conference, Michigan Copper District, September 30 - October 2, 1971, p. 18-43. Weege, R.J., Pollock, J.P., and the Calumet Division Geological Staff, 1972, The geology of two new mines in the native copper district: Economic Geology, v. 67, p. 622-633. Weege, R.J., and Schillinger, A.W., 1962, Footwall mineralization in Osceola amygdaloid, Michigan native copper district: A.I.M.E. Transactions, v. 223, p. 344-350. 42 �Field Trip 2 Caledonia Mine, Keweenaw Peninsula Native Copper District, Ontonagon County, Michigan Theodore J. Bornhorst A.E. Seaman Mineral Museum, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Robert J. Barron Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 Richard C. Whiteman Red Metal Minerals, 202 Ontonagon Street, Ontonagon, MI 49953 Directions: Leave downtown Houghton and head south on M-26 towards South Range. Stay on M-26(highway turns into M-38 past the Mass City turnoff) for 39 miles past Greenland(Pat’s Auto & Sports Center) to Ridge Road. Turn left (south) on Ridge Road approximately 1 mile to Caledonia Rd. Drive southwest 1.7 miles (2.7 km) to the Caledonia Mine. The Caledonia mine is privately owned and permission is required to enter this property [UTM 5179684N 160338022E (NAD27 CONUS)] The Caledonia Mine, surface and underground, is strictly private property. Permission is required to enter the property. Introduction The Caledonia Mine is part of the Keweenaw Peninsula native copper district of the western Upper Peninsula of Michigan (Fig. 1). The Caledonia Mine is located in the Greenland-Mass subdistrict about 40 km southwest of the Baltic Mine, the southernmost major native copper mine (Fig. 2). In total the mines on the Evergreen succession produced about 73 million lbs of copper at grades ranging from 0.5 to 1.25 % (Weege and Pollock, 1971) and, as compared to total district production of 11,000 million lbs of copper, this subdistrict was a minor producer. The largest producer among the Greenland-Mass subdistrict mines was the Mass Mine which produced about 51 million lb of refined copper from 1851 to 1923 (Butler and Burbank, 1929). The Adventure Mine produced about 11 million lb of copper. Despite its low production of copper, the geologic characteristics of the Greenland-Mass subdistrict deposits are typical of the native copper deposits elsewhere in the Keweenaw Peninsula. Since native copper mining in the Keweenaw Peninsula ceased in 1968, underground access to observe or study the native copper deposits is limited today to four mines, two of which are in the Greenland-Mass subdistrict (Caledonia and Adventure Mines). The Caledonia Mine is one of the few remaining localities where a typical native copper ore body hosted by the top of a basalt lava flow can be observed up close underground. 43 �This field trip guide relies on existing publications by Bornhorst and Whiteman (1992 and 1995) and Bornhorst and Barron (2011). The geologic and human history overview is summarized from Bornhorst and Lankton (2009). Since professional and collector oriented underground field trips to the Caledonia Mine are common, the introductory explanations have been expanded to provide a more readily stand alone field trip guide. Underground, the Caledonia Mine is relatively dry and regular field boots are usually sufficient; hard hats and lights are required. The field trip involves an easy walk underground to observe the character of native copper mineralization in an adit and a drift that parallels the strike of the tabular native copper ore body (lode) that is hosted by the top of the Knowlton lava flow. The cross-cutting adit provides opportunity to observe the Knowlton lode in cross-section as well as to observe underlying lava flows of the Evergreen succession. An optional more difficult segment of the field trip involves climbing up into an underground stope on the Knowlton lode to observe the character of mineralization and to collect specimens. Specimens of native copper and associated minerals from hosted by rocks of the Caledonia Mine can be collected on the rock pile adjacent to the adit. The Caledonia Mine is owned and operated by Red Metal Minerals as an educational facility and to recover specimens for resale in the general public and mineral collector markets. Mesoproterozoic Midcontinent Rift SystemAround Lake Superior Native Copper occurrences Sedimentary Rocks Igneous Rocks Several Major Faults Abundant Ontario Isle Royale Minnesota Lake Superior Keweenaw Peninsula native copper district Ontario 46.00 Wisconsin Lake Michigan 0 100 200 kilometers Archean metamorphosed sedimentaryandigneous rocks N Paleoproterozoic metamorphosed sedimentary and igneous rocks Phanerozoic sedimentary rocks Figure 1: Generalized bedrock geologic map of showing the Keweenaw Peninsula Native Copper District. 44 �Regional Geologic and Human History Overview The Keweenaw Peninsula is home to the largest known accumulation of native copper on the planet termed the Keweenaw Peninsula native copper district. The district is unique in comparison to copper mining districts elsewhere in that native copper comprises nearly all of the metallic minerals in the mined ore bodies. Approximately 11 billion pounds of refined copper from 380 million tons of ore were produced from native copper mines from 1845 to 1968 (Weege and Pollock, 1971). Small quantities of native silver occur with the native copper. Copper sulfides are uncommon in the Keweenaw Peninsula native copper district although chalcocite occurs in veinlets cutting the deposits (White, 1968). Near the tip of the Keweenaw Peninsula, there are several small unmined chalcocite-dominated deposits but their connection with the native copper deposits is uncertain (see Field Trip 3 this volume, Maki and Bornhorst, 1999). The native copper deposits of the Keweenaw Peninsula are hosted by rocks of the Mesoproterozoic Midcontinent Rift (MCR) (Fig. 1and 2). The MCR is filled with more than 25 km of volcanic rocks and 8 km of clastic sedimentary rocks (Cannon et al., 1989 and 1993). This thick succession of rocks was emplaced between about 1.15 to 1.03 Ga (Cannon et al., 1989; Davis and Paces, 1990; Heaman et al., 2007). Volcanic rocks were erupted on land surface initially over a broad area above a mantle plume and later, were erupted from fissure volcanoes within the normal fault-bounded rift graben. The volcanic rocks erupted during this syn-rift phase of the MCR were predominantly subaerial tholeiitic flood basalt lava flows. The subaerial basalt lava flows have a top which is vesicular (amygdaloid) and/or brecciated (fragmental amygdaloid) underlain by a massive (vesicle-free) interior. The typical flow is 10 to 20 m thick. Minor gravels and sands were deposited on top of the lava flows during hiatuses in volcanic activity and today occur as red conglomerate and sandstone that are interbedded with the lava flows. After active rifting and volcanic activity ended, the rift basin continued to sag. Rivers carried gravels, sands, silts, and muds to fill this sagging rift basin, and with subsequent burial they were lithified into clastic sedimentary rocks which occupy the center portion of the rift today (Merk and Jirsa, 1982). Volcanic rocks crop out around the margin of the rift (Fig. 1). The last and final phase of the MCR resulted from a regional compressional event due to collision of continental land mass along the eastern edge of North America at that time (Grenville Orogeny, Cannon, 1994). Compression inverted the rift-bounding normal faults into reverse faults as well as folding, faulting, and fracturing rift-filling volcanic and clastic sedimentary rocks. Native copper and related minerals were emplaced during this regional compressional event about 1.06 to 1.04 Ga (Bornhorst, 1997). For roughly 500 million years, from about 1.0 Ga to 500 Ma, there were no geologic events recorded by rocks of the Upper Peninsula. During this time interval, erosion exposed the native copper deposits to the surface and downward percolating oxidizing groundwaters had access to alter the native copper (Bornhorst and Robinson, 2004). After being buried beneath Phanerozoic sedimentary rocks (500 Ma to 175 Ma) (Catacosinos and others 2001), Pleistocene continental glaciations removed all but a few outliers of these rocks from the Keweenaw Peninsula and exposed the native copper deposits at the surface. 45 �Native people began exploiting native copper by ca. 7,000 years ago as the land surface of the Keweenaw Peninsula emerged above the retreating glacial lake levels. At first these prehistoric ancient miners likely found boulders of native copper (locally termed float copper) with their distinctive green weathered crust of malachite among the brown, gray, red, and white rocks. As they discovered the usefulness of native copper, they moved on to mining of bedrock. Because of the scars these early exploits left on the landscape, most mines of the Keweenaw Peninsula were rediscovered later including the Caledonia and those nearby. The first major mining rush in North America was started by Douglass Houghton through his report to the Michigan legislature in 1841. The Cliff Mine became the first profitable mine in the district in 1849. The Minesota Mine, in the southwest cluster near Caledonia Mine (Fig. 3) became profitable soon after the Cliff. In 1880, copper production from native copper mines of the Keweenaw Peninsula accounted for up to 80 % of the nation’s copper production. The peak copper production occurred in 1916 at 267 million pounds with mining ending in 1968. The Keweenaw National Historical Park was created in 1992 to preserve and interpret the historical importance of native copper mining to the history of the U.S. Native Copper Deposits of the Keweenaw Peninsula The pre-mining geologic resource of the Mesoproterozoic Keweenaw Peninsula native copper district totaled about 20 billion lbs of Cu (Bornhorst and Barron, 2011). Most of the native copper in the district is hosted in the permeable and porous brecciated and amygdaloidal lava flow tops (~58.5% of production) and interflow conglomerate-sandstone horizons (~39.5% of production). The ore is "sandwiched" above and below between barren massive basalt that lacks permeability and porosity and is geometrically found in tabular bodies between 3 and 5 m thick that have the same orientation as surrounding host rocks, i.e., stratiform lode. The typical lode has a lateral extent of 1.5 to 11 km and extends down-dip 1.5 to 2.6 km (Butler and Burbank, 1929; White, 1968). Native copper fills open spaces from a few cm across (e.g., vesicle-fillings) to small-to-moderately sized openings (e.g., space between lava flow top breccia fragments or between clasts in conglomerate) that contain native copper masses weighing up to several pounds and rarely weighing tons. A minor amount of native copper was produced, ~ 2%, from high-angle tabular veins that cut across the volcanic-dominated strata. Native copper is closely associated with over 100 different minerals in the Keweenaw Peninsula although only about 25 of them are common. These minerals fill the same open spaces along with and instead of native copper (Butler and Burbank, 1929; Stoiber and Davidson, 1959; White, 1968). The suite of minerals is similar to those found where rocks have undergone very low to low grade burial metamorphism, < 300OC. Overall, higher temperature assemblages are spatially associated with the area of native copper deposits where the thermal anomaly was greatest because of focused hydrothermal fluid flow. In areas more distal to the deposits, the open spaces are filled with lower temperature assemblages. At any one location, there is a recognizable sequence in the precipitation of minerals from hydrothermal fluids due to changing hydrothermal fluid temperature and composition. The absolute age of the hydrothermal activity is coincident with the age of the regional compressional event at about 1.06 to 1.04 Ga (Bornhorst et al., 1988). The compressional event provided the plumbing system, faults/fractures, that facilitated movement of the hydrothermal fluids into the sites of future mineable deposits of native copper (Bornhorst, 1997). 46 �47 �The native copper mineralizing event was widespread throughout the exposed MCR (Fig. 1). Burial metamorphism at depth of rocks down-dip from the deposits was the likely source of the mineralizing hydrothermal fluids; small amounts of copper and other constituents were leached from the rift-filling basalt-dominated rocks. Subaerial eruption of the basalt lava flows likely resulted in the degassing of most of the contained sulfur leaving them sulfur poor. Thus, the hydrothermal fluids generated from them were low in sulfur and the movement of these fluids through the same sulfur poor rocks resulted in sites of ore deposition where host rocks were also sulfur poor. This low sulfur environment favored the deposition of native copper rather than copper sulfide. The heating of the volcanic rocks during burial probably reached a maximum millions of years after they were erupted and it was most likely the coincidence of increased fluids generated at this temperature maximum with the regional compressional event which played a critical role in providing the plumbing system necessary for producing the deposits (Bornhorst, 1997). Geology of the Evergreen Series The native copper mines in the Greenland-Mass subdistrict produced native copper from the tops of rift-filling lava flows that comprise the Evergreen succession within the Portage Lake Volcanics (Fig. 3 and 4). The Evergreen succession of basalt lava flows have a total thickness of about 210 m. The individual copper-rich lava flows within the succession were each informally named. From bottom to top the Evergreen succession consists of: the Evergreen flow: a 3 to 15 m thick plagioclase porphyritic otherwise aphanitic basalt lava flow; the Ogima flow, a 30 to 43 m thick slightly plagioclase glomerophyritic basalt lava flow; the Butler flow, a 15 to 27 m thick plagioclase glomerporphyritic basalt lava flow; and a horizon of thin plagioclase glomeroporphryitic flows 75 to 90 m thick. This latter stratigraphic horizon of multiple flows includes the Mass flow. The South Knowlton flow overlies this horizon and is a plagioclase glomeroporphyritic lava flow up to 15 m thick and at the top of the Evergreen succession is the Knowlton flow, a 9 to 21 m thick plagioclase glomeroporphyritic lava flow (Calumet and Hecla, 1958). The volcanic rocks nearby underlying the Evergreen succession are ophitic and aphanitic basalt lava flows. The nearby overlying volcanic rocks are ophitic basalt lava flows. The Evergreen succession is stratigraphically at the level of the Isle Royale flow near Houghton. The tops of these flows were productive over a strike length of about 5 km. Of the lava flows in the Evergreen succession, the Butler flow top yielded the most copper followed by the Evergreen and Knowlton flow tops which also yielded significant amounts of copper. Most of the Evergreen succession basalt lava flow tops are brecciated (fragmental amygdaloid) with considerable lateral (along strike) variation in the degree of brecciation and thickness. In some areas, thin lava flows lack brecciated tops and are simply vesicular (amygdaloid). In general, the best grades of copper occur where the brecciated flow top thickens. Secondary minerals in all of the flow tops are quite similar. Quartz, feldspar, pumpellyite, chlorite, calcite, and epidote are abundant minerals filling amygdules and spaces between breccia fragments. There is less abundant native silver, prehnite, datolite, and laumontite. 48 �Greenland-Mass Subdistrict Bedrock Geology City Nonesuch Formation Native Copper Mine 45 Strike and Dip of Bedding Freda Sandstone Fault 12 Anticline 20 28 38 Syncline 35 43 44 45 North Lake Toltec Greenland A Adventure Belt 45 South Lake Lake Mass 45 Mass Old Mass Knowlton Caledonia 48 45 Rockland 45 Nebraska Nassau Superior Rockland Flintsteel Bumblebee National Minesota Michigan 55 60 Algomah 46o45' N 0 55 1 mile 0 60 5 Jacobsville Sandstone T51N T50N 2 kilometers A’ 20 NW 6 SE NonesuchFormation A' A FASL 0 Freda Sandstone Jacobsville Sandstone -2000 -4000 1 0 0 mile 1 kilometer Figure 3: Generalized bedrock geologic map of the Greenland-Mass subdistrict of the Keweenaw Peninsula Native Copper District showing the location of native copper mines. Geologic cross section shows the Evergreen Succession within the Portage Lake Volcanics; a lithostratigraphic column is given in Figure 4. Subdistrict bedrock geology and cross section modified from Whitlow (1974). 49 �The Evergreen succession in the Greenland-Mass subdistrict dips about 45o NW towards Lake Superior and forms a local broad open anticlinal structural bend (Fig. 3). The largest mine, the Mass Mine, occurs near the maximum bend in this anticline. Faults with significant vertical displacement are uncommon as most have displacement of < 1 m. There are multiple veins in tension fractures that cut perpendicular across the lava flows in association with the anticline (Butler and Burbank, 1929). However, some veins are parallel to strike of the lava flows but dip in the opposite direction. In the stratigraphically equivalent Isle Royale lode, Broderick (1931) describes similar strike parallel veins which he interpreted to be feeders of ore fluid into the top of the lava flow. Figure 4: Stratigraphic position of the Evergreen succession that hosts native copper deposits of the Greenland-Mass subdistrict as compared to lithostratigraphic units of the Keweenaw Peninsula, Michigan. Units from the Powder Mill Group to the Jacobsville are all Mesoproterozoic in age and related to the Midcontinent rift. Geology of the Caledonia Mine In the context of mines in the Keweenaw Peninsula native copper district, the Caledonia Mine was very small, producing only about 6.8 million lb of refined copper from the top of the Knowlton basalt lava flow (Knowlton lode), the youngest basalt lava flow of the Evergreen Series (Fig. 4). The near horizontal adit of the mine follows approximately along strike of the Knowlton lava flow where it connects with the Knowlton Mine (Fig. 5) and to where it connects to the stopes of the Mass Mine “C” shaft (not shown). At the Mass Mine, the stratigraphically lower Bulter lava flow top was the principal focus of native copper mining. In the GreenlandMass subdistrict the Butler lava flow top was developed for about 2000 m along strike and to a maximum depth of 300 m along dip. The most abundant secondary minerals in the Butler are quartz and calcite with slightly lesser amounts of K-feldspar and epidote. Prehnite and pumpellyite are usually much less abundant and chlorite is present in amounts < 1 %. The Butler contains a high number of veins, usually they strike subparallel to the strike of the Butler lava flow top and have dips both similar to the dip of bedding and at a high angle to bedding (Butler and Burbank, 1929). 50 �51 �In the Greenland-Mass subdistrict, the Knowlton flow top was developed for about 3000 m along strike and to a maximum depth of about 375 m. The Knowlton was the focus of native copper mining at the Caledonia Mine. The Knowlton lava flow top is a brecciated flow top or fragmental amygdaloid. Stopes at Caledonia were raised on the Knowlton lode upwards from the drift toward the topographic high of the Caledonia bluff (Fig. 6). The floor (footwall) of the stopes reflects the original depositional irregularities of the top of the underlying lava flow such as gentle flexures. The average thickness of the Knowlton lava flow top is about 2.5 m but locally it can thicken to around 6 m (Calumet and Hecla, 1958). In general, a thicker flow top results in better ore. While most of the ore occurs in the top of the Knowlton lava flow top, there are pockets of ore that extend into the underlying Knowlton massive flow interior footwall and are closely associated with strike-parallel fractures and veins. The grade of the ore in the flow top may correlate with these footwall pockets of ore; there is an approximate strike parallel (drift parallel) orientation of the grade of the ore (Fig. 5). The workings of the Caledonia Mine provide excellent access to observe the 3-D geometry of the Knowlton lode. An elongated volume of highly epidotized basalt characterizes a footwall ore pocket that was mined out in the 1990s by Red Metal Minerals and will be visited on this field trip. Veins are of two types. Veins within the Knowlton flow top that extend into underlying Knowlton massive flow interior and contain the same basic minerals as those found in the lode itself (including native copper) are considered synchronous with the native copper deposit at the Caledonia Mine. Native copper occurs as small to large masses in the fractures and seems to be more abundant in the overlying adjacent tabular ore body within the lava flow top. The fractures typically have little or no displacement. Adjacent to the fractures even massive basalt can be highly altered and host native copper. One main-stage vein that has been studied in more detail by Bornhorst strikes subparallel with the strike of the flow top but dips more steeply, about 80oNW dip of the vein as compared to 45oNW of the lava flow top. This vein has been traced along strike for over 100 m. It extends into the footwall, but is hard to identify as it enters the lode. Within this vein the intensity of alteration varies from slight to very high. Original basalt can be completely converted to a green soft epidote and lesser chlorite rock, or a hard epidote and lesser quartz rock. Overall mineralogy and paragenesis in the vein is similar to the lode. The vein contains pockets of a soft blue-green mineral identified as corrensite by XRD (mixed layered clay mineral with 50/50 chlorite and smectite unit cells stacked in perfect alternation). NW SW 1300 1200 Caledonia Mine adit 1100 1000 Stope 900 Drift 800 0 500 Horizontal Scale feet 5X vertical exxageration Figure 6: Cross-section sketch of the topography of Caledonia bluff showing the positions of the Caledonia Mine adit and the top of the Knowlton lava flow. 52 �This vein has yielded outstanding museum quality specimens of crystalline native silver (now part of the A.E. Seaman Mineral Museum collection). These native silver specimens were encased in white calcite; the calcite was removed by acid cleaning. This vein also yielded clusters of colorless calcite crystals internally laced with native copper from open vugs. Several masses of native copper weighing over 100 kg and small groups of copper crystals originally encased in white calcite (removed by acid cleaning) have also been recovered during exploration. Some of the copper crystals were coated with very small cubic native silver crystals. There are multiple other veins at Caledonia synchronous with native copper precipitation. One of these veins along the drift that will be visited is notable for hosting datolite; the datolite commonly contains very-fine inclusions of native copper. The occurrence of veins at the Caledonia Mine is quite similar to veins described by Broderick (1931) occurring in the Baltic and Isle Royale Mines near Houghton. At Caledonia, they are interpreted as pathways for ascending hydrothermal fluids and thus, played an important role in the deposition of native copper and associated minerals. At Caledonia, there are veins that crosscut the native copper mineralized lode. These post-copper mineralization veins contain calcite and laumontite and are barren of copper. Several of these post-mineral veins are readily visible along the down-dip side of the Caledonia adit. At the Caledonia Mine, the most abundant mineral filling amygdules and spaces between fragments in the Knowlton lode is calcite which is closely followed by subequal amounts of quartz, epidote and red K-feldspar. There are lesser amounts of prehnite, pumpellyite and chlorite. Native copper is present in small amounts with an average grade of about 1.2 % Cu. Native silver and datolite are present in much lesser amounts. Least abundant are laumontite, adularia, and corrensite (clay mineral). No major differences exist in the abundance of secondary minerals averaged over the scale of 100's of meters. In contrast, over the scale of a few meters the distribution of secondary minerals is variable to highly variable. While a secondary mineral may be completely absent in one zone and extremely abundant in another, the meter scale variation does display a degree of regularity. For example, within the Knowlton flow top secondary minerals may occur in overlapping bands; the bands are consistent with the progressive filling of open spaces indicated by amygdule paragenesis. The intensity of alteration is highest near both the hanging wall and footwall of the brecciated flow top lode where apparently there was preferential flow of hydrothermal fluids. Distribution of secondary minerals also has a poorly defined correlation with the occurrence of synchronous veins. In general, native copper tends to be more commonly associated with epidote, calcite, and quartz. Rarely is native copper abundant in areas with abundant K-feldspar. Paragenetically, K-feldspar is an early formed mineral followed by epidote and then datolite, prehnite, pumpellyite, chlorite, calcite and quartz. Native copper is found as inclusions in epidote, calcite, quartz, and datolite. Much of the calcite that is overall synchronous with native copper precipitation does not contain obvious inclusions of native copper. Post-native copper mineralization hydrothermal minerals in veins and open space fillings as coatings on earlier formed minerals include calcite, laumontite, adularia, and corrensite. These likely formed from superposition of later lower temperature hydrothermal fluids on earlier higher temperature formed minerals associated with native copper during collapse of the hydrothermal system. This relationship is found elsewhere in the broader district. 53 �Post-emplacement alteration of hydrothermal minerals is most obvious for native copper. At Caledonia, tenorite and cuprite (Cu oxide) often but not always occurs as a thin coating on native copper that is found in open space fillings. The tenorite and cuprite could have its origin during Precambrian weathering and downward-percolating ancient groundwater leading to supergene alteration of the native copper prior to Phanerozoic marine submergence or during more recently since Pleistocene glaciations eroded and exposed the Caledonia deposits at the surface. Today, only a few fractures within the stopes of Knowlton lode are damp with meteoric groundwater despite the shallow depth, an argument in favor of Precambrian tenorite and cuprite. However, the presence of modern groundwater flow into the mine, such as near the intersection of the cross-cut and the Knowlton drift, suggests Pleistocene age for the tenorite and cuprite cannot be precluded. In addition to tenorite and cuprite, there are occasional copper carbonate minerals (such as malachite), brochantite (hydrated Cu sulfate), atacamite (hydrated Cu chloride) and unknown green to blue-green minerals on native copper surfaces. These may also be supergene in origin. However, there is at least one mineral, gerhardtite (hydrated Cu nitrate) that is the result of postmining chemical reactions. An extension of the Caledonia adit cuts across the lava flows towards the Nebraska Mine; the cross-cut was a component of Calumet and Hecla’s exploration program (Fig. 5). This cross-cut intersects the South Knowlton lava flow top directly below the base of the Knowlton lava flow. An NSF-sponsored Teachers Earth Science Institute (educating middle and high school teachers about mining) drilled, blasted, and mucked out the South Knowlton lava flow top to its present expanded opening in the early 2000s. Below the South Knowlton, there are several thin basalt lava flows that can be identified before the stratigraphic level of the Butler lava flow. While the Butler lava flow top is readily identified in the cross-cut by abundant amygdules filled with Kfeldspar and calcite, it lacks significant native copper here. A small fault can also be seen along the cross-cut. This cross-cut and adit connects the Caledonia Mine to the Nebraska Mine where the Butler lava flow top was a principal target of mining. Mining History of the Caledonia Mine The Caledonia Mining Company began its operations in 1863 after acquisition of the mining rights of the former Nebraska Company and acquisition of the adjacent Kansas Properties. The workings at that time consisted of a horizontal adit driven about 90 m to the Butler deposit on the west end of the Caledonia bluff and two shafts about 60 m deep (Nebraska Mine) (Fig. 5). A vein showing mineralization was explored on the north side of the bluff by four adits and while the vein proved to contain too little copper, the adits intersected the Knowlton, South Knowlton, Mass, and Butler lava flow tops. About 900,000 lbs of refined was produced from 1863 to 1870 by the mining of native copper at a grade of about 1.25 % Cu from the Knowlton and Butler lava flow tops. Production halted in 1870 when a fire destroyed the processing facility. Subsequently, the Caledonia Mining Company acquired the Flintsteel properties in 1870 and despite investing in a new processing facility, the Caledonia operations closed before significant ore was processed. Captain Martin leased the Caledonia properties and from 1873 to 1881 produced more than 330,000 lbs of copper, including a single mass weighing 80,000 lbs (40 short tons). After the lease expired, mining ceased. In 1901 there was a failed proposal to merge the Caledonia properties with other mineral rights in the Greenland-Mass subdistrict and to construct a processing facility on Lake Superior some distance away. There was no reported mining from the Caledonia properties for 56 years from 1881 to 1937. 54 �The Calumet and Hecla Consolidated Mining Company was the major producer of copper from the mines north of Houghton. Calumet and Hecla did exploration core drilling and reopened the Caledonia adit. They drifted some 600 m along the strike of the Knowlton flow top and estimated the grade of native copper ore to be 1.45 % Cu. The workings from the nearby Nebraska Mine were connected to the Caledonia with a cross-cut and adit. Exploration of the Caledonia Mine by Calumet and Hecla ended in c.a. 1941 as a result of World War II. After World War II, Calumet and Hecla resumed exploration of Caledonia and removed a 200 ton bulk sample from the Knowlton lode in 1950. The sample had a very promising grade of 1.84 % Cu. From 1951 to 1958, Calumet and Hecla produced 5.55 million lbs of refined copper with an average grade of 1.24 %. This program included the stoping on the Knowlton lode visible above the drift today. Subsequently, Copper Range Company acquired the mineral rights at the Caledonia Mine. The Caledonia Mine was a candidate for in-situ leaching of Cu, hence a limited evaluation of the mine was completed by Copper Range and the U.S. Bureau of Mines from 1971 to 1972. The in-situ leaching option was abandoned due to potential problems with groundwater pollution. In 1985, Red Metal Minerals acquired the mineral rights for the Caledonia Mine and other properties in the Greenland-Mass subdistrict from the Copper Range Company. The Caledonia adit was reopened and reconditioned to undertake a limited program of exploration. Red Metal Minerals mucked out broken rock as well as drilled and blasted new areas to recover native copper and other minerals. Over 28 years from 1985 to present, Red Metal has removed a single mass weighing 3000 lbs. Masses of recovered native copper are sold to the general public and mineral collectors. Red Metal distributes native copper on a wholesale basis to retail outlets that distribute Caledonia native copper around the world. You may find native copper from the Caledonia Mine for sale in unexpected places, even visitor gift shops at other copper mines! A large mass of native copper is on exhibit at the A.E. Seaman Mineral Museum, on the campus of Michigan Tech in Houghton. When the native copper is dispersed through the rock, Red Metal uses this material to prepare decorative bookends and cut slabs. The Caledonia Mine has yielded minerals sought after by mineral collectors such as native copper crystals, native silver crystals, datolite nodules, copper in calcite crystals as well as adularia and epidote. At this time, Red Metal does not undertake in underground drilling and blasting to recover specimens from the mine. Instead, the Caledonia Mine serves as an educational facility, used by Michigan Tech and others, and to recover specimens from already broken rock for resale in the general public and mineral collector markets. Since most of the native copper mines of the Keweenaw Peninsula are closed and flooded, the Caledonia Mine is significant today because it provides rare access to observe the character of native copper mineralization underground in three dimensions. Acknowledgements We thank Allan Blaske for helpful reviews which improved the quality of this field trip guide. References Cited Bornhorst, T.J., 1997, Tectonic context of native copper deposits of the North American Midcontinent Rift System: Geological Society of America Special Paper 312, p. 127-136. 55 �Bornhorst, T.J., and Barron, R.J., 2011, Copper deposits of the western Upper Peninsula of Michigan: Geological Society of America Field Guide, v. 24, p. 83-99. Bornhorst, T.J., and Lankton, L.D., 2009, Copper mining: A billion years of geologic and human history: in Schaetzl, R., Darden, J., and Brandt, D., eds, Michigan Geography and Geology, Pearson Custom Publishing, New York, p. 150-173. Bornhorst, T.J., Paces, J.B., Grant, N.K., Obradovich, J.D., and Huber, N.K. 1988. Age of native copper mineralization, Keweenaw Peninsula, Michigan: Economic Geology, v. 83, p. 619625. Bornhorst, T.J., and Robinson, G.W., 2004, Precambrian aged supergene alteration of native copper deposits in the Keweenaw Peninsula: Michigan; Institute on Lake Superior Geology Proceedings and Abstracts, v. 50, part 1, p. 40-41. Bornhorst, T.J., and Whiteman, R.C., 1995, Native copper and associated minerals in basalts at the Caledonia Mine, western Upper Michigan: Institute on Lake Superior Geology Proceedings, v. 41, part 1, p. 3-4. Bornhorst, T.J., and Whiteman, R.C., 1992, The Caledonia native copper mine, Michigan: Society of Economic Geologists Guidebook Series, v. 13, p. 139-144. Broderick, T.M., 1931, Fissure vein and lode relations in Michigan copper deposits: Economic Geology, v. 26, p. 840-856. Butler, B.S., and Burbank, W.S., 1929, The copper deposits of Michigan: U.S. Geological Survey Professional Paper 144, 238 p. Calumet and Hecla, 1958, Unpublished report for Defense Minerals Exploration Administration, 29p. Cannon, W. F., 1994, Closing of the Midcontinent Rift - A far field effect of Grenvillian contraction: Geology, v. 22, p. 155-158. Cannon, W. F., Green, A. G., Hutchinson, D. R., Lee, M.W., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American mid-continent rift beneath Lake Superior from Glimpse seismic reflection profiling: Tectonics, v. 8, p. 305-332. Cannon, W. F., Peterman, Z.E., and Sims, P.K. 1993, Crustal-scale thrusting and origin of the Montreal River monocline - A 35-km-thick cross section of the Midcontinent Rift in northern Michigan and Wisconsin: Tectonics, v. 12, p. 728-744. Catacossinos, P.A., Harrison, W.B., Reynolds, R.F., Westjohn, D.B., and Wollensak, M.S., 2001, Stratigraphic lexicon for Michigan: Michigan Department of Environmental Quality, Geologic Survey Division Bulletin 8. Lansing, MI. 56 �Davis, D.W., and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system: Earth and Planetary Science Letters, v. 97, p. 54-64. Heaman, L.M., Easton, R.M., Hart, T.M., MacDonald, C.A., Hollings, P., and Smyk, M., 2007, Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario: Canadian Journal of Earth Sciences, v. 44, p. 1055-1086. Maki, J.C., and Bornhorst, T.J., 1999, The Gratiot chalcocite deposit, Keweenaw Peninsula, Michigan: Institute on Lake Superior Geology Proceedings and Abstracts, v. 45, part 1, p. 33-34. Merk, G.P., and Jirsa, M.A., 1982, Provenance and tectonic significance of the Keweenawan interflow sedimentary rocks: Geological Society of America Memoir 156, p. 97-105. Stoiber, R.E., and Davidson, E.S., 1959, Amygdule mineral zoning in the Portage Lake Lava Series, Michigan copper district: Economic Geology, v. 54, p. 1250-1277, p. 1444-1460. Weege, R.J., and Pollack, J.P., 1971, Recent developments in native-copper district of Michigan: Society of Economic Geologists Field Conference, Michigan Copper District, September 30 October 2, 1971, p. 18-43. White, W.S. 1968, The native-copper deposits of northern Michigan: in Ridge, J.D., ed., Ore Deposits of the United States, 1933-1967 (the Graton Sales volume). American Institute of Mining, Metallurgical, and Petroleum Engineering, New York: p. 303-325. Whitlow, 1974, Geologic map of the Greenland and Rockland quadrangles, Ontonagon County, Michigan: U.S. Geological Survey Miscellaneous Field Studies Map MF-596. 57 ��Field Trip 3 Geology of Silver Mountain, Houghton County, Michigan Evgeniy Kulakov Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931. Directions: Leave downtown Houghton and head east on Montezuma avenue and College Avenue. Continue on US 41south for 27 miles. In Baraga MI turn right and continue to follow M38 west for 13.5 miles. Make left turn on S Laird road and follow it for 5.7 miles. Take a left turn onto Forest Rd 2276 and continue 1.8 miles. Turn right onto Forest Rd 193/NF-2270. In 0.9 miles make a right turn onto Forest Rd 922 road. Arrive in 0.4 miles. [UTM 5171550N 367079 E (NAD27 CONUS)]. Introduction Figure 1: View from the top of Silver Mountain looking east. 59 �Silver Mountain is located at the south-eastern part of Keweenaw Peninsula, near Sturgeon River Falls, Houghton County, Michigan (sections 1 and 12 of T 49 N, R 36 W, and section 6 of T 49N, R 35 W). A former US Forest Service lookout tower, Silver Mountain is now designated as a scenic view. Silver Mountain is a 100 meter high glacially polished dome-shaped hill with a 360° view (Fig.1) of the surroundings. The low lying areas around Silver Mountain consist of unconsolidated glacial deposits unconformably overlying Mesoproterozoic Jacobsville Sandstone. Fourteen shallow-dipping tholeiitic lava flows are exposed on the top and sides of the mountain. Silver Mountain was first geologically mentioned by Burt (1849) who reported its location and small amounts of copper sulfides associated with calcite. Foster and Whitney (1851) described Silver Mountain as basaltic flows made up of labradorite and hornblende with scattered nodules of quartz and chalcedony. They interpreted the mountain as igneous rocks that intruded the surrounding sedimentary strata. In the mid 1850s, the National Company drove an approximately 30 m long adit. However, metallic mineralization was insufficient to justify additional mining. Lane (1909) concluded that the rocks of Silver Mountain were typical Keweenawan lava flows. Regional Geology The Midcontinent Rift (MCR) extends from NE Kansas northward to Lake Superior and through Michigan (Cannon et al., 1989). Rifting began ~ 1.1 Ga during an interval of reversed polarity of geomagnetic field with the oldest erupted material being reversely magnetized. These oldest lava flows include the Siemens Creek Formation of the Powder Mill Group, the lowermost part of North Shore Volcanics, Osler Volcanics, and the lower part of Mamainse Point Formation, located around Lake Superior (Paces and Miller, 1993; Davis and Green, 1997) (Fig.2 and 3). This early stage of magmatism occurred from 1109 to approximately 1105 Ma (Heaman et al., 2007; Paces and Miller, 1993) and was followed by a quiescence period when the geomagnetic field reversed to normal polarity (Davis and Green, 1997). Magmatism resumed by 1102 Ma (Paces and Miller, 1993) during the normal polarity interval. During this interval, the main stage of the rift-related magmatism was represented by a sequence of approximately 200 lava flows of the Portage Lake Volcanics that erupted within a short interval around 1095 Ma (Davis and Paces, 1990). Magmatism of the Portage Lake Volcanics ended and the rift basin was filled with clastic sedimentary rocks during continued sagging (Bornhorst and Lankton, 2009). The final phase of the MCR was continental compression at about 1060 Ma related to the Grenville Orogeny which inverted original rift-bounding graben normal faults into high angle reverse faults (Cannon, 1994). During late rift compression, rift-wide burial metamorphic/hydrothermal fluids altered rift-filling rocks and formed the native copper deposits of the Keweenaw Peninsula native copper district (Bornhorst, 1997; Bornhorst and Barron, 2011). 60 �Figure 2: a. Generalized geologic map of Keweenaw Peninsula. The inset shows the geology of the Lake Superior segment of the Midcontinent Rift. The inset from Ojakangas et al, (2001); b, Section of Keweenaw Peninsula along the Line A-A’ (B- B’part of the section is present on the map). Section is taken and modified from Cannon and Nicholson (2001). 61 �Fourteen tholeiitic lava flows have been recognized at Silver Mountain and that dip NE at about 15o. The lava flows of Silver Mountain are characterized by moderate to high magnetic anomalies and overall high positive gravity anomaly (Campbell, 1952). The magnetic and gravity anomalies are similar to those attributable to the Siemens Creek Formation. Based on this geophysical data, the lava flows at Silver Mountains were interpreted as being Keweenawan in age and part of the South Range Traps (termed Siemens Creek Formation today) (Campbell, 1952). However, Silver Mountain is an isolated knob and there is no direct geological contact with the rest of the Powder Mill Group. The mountain is located in the immediate vicinity of the Marenisco fault (Fig.2). Reverse and thrust faulting occurred during the latest compressional stage of the MCR evolution and resulted in uplifting the deeply buried strata to the surface. At Silver Mountain, the Marenisco fault has uplifted the lava flows stratigraphically older than the Portage Lake Volcanics which are exposed in the Keweenaw Peninsula because of uplifting along the Keweenaw fault. The Jacobsville Sandstone surrounds the uplifted basalt flows at Silver Mountain and vicinity. The Jacobsville Sandstone was deposited in a rift-flanking basin that while initiated by, and contemporaneous with compression, its deposition continued after compression. About 10 km west-southwest of Silver Mountain, there are gravity and magnetic anomalies resulting from the Echo Lake Gabbro which was also uplifted during compression along another fault. This geophysical anomaly was drilled in 1994 and confirmed as a layered intrusion (Waggoner, 1994). Additional drilling by Bitterroot Resources identified a 5.5 m thick interval containing 0.5 to1 ppm total Pt + Pd + Au (Cannon and Nicholson, 2001). There is at least two faults cross-cutting Silver Mountain (Roberts, 1940). One fault can be observed at the adit at the base of Silver Mountain. It strikes N85°E and dips approximately 60° N (Roberts, 1940). The other fault is exposed near the NE side of Silver Mountain. This fault strikes N45°W and dips at about 80° N. Figure 3: Stratigraphic column of Midcontinent Rift rocks in the western Upper Peninsula of Michigan. The available radiometric ages shown are from Davis and Paces, (1990). ”R” and “N” indicate reversed and normal polarity of remnant magnetization, respectively. 62 �The massive interiors of the lava flows are fine-grained with intergranular texture. The predominant rock forming mineral is plagioclase with laths up to about 2 mm in length with an aspect ratio of around 10:1 . Typically, more equant altered mafic minerals, less than 1 mm across, and opaques, less than 0.2 mm across, fit between the plagioclase laths. In most massive interiors, the mafic minerals are completely pseudomorphically replaced by chlorite, however, in some interiors; patches of original pyroxene have survived a combination of burial metamorphic and hydrothermal alteration. Some plagioclase has altered to sericite. The plagioclase laths and space between them contain irregular patches of calcite. The abundance of calcite is near zero in some flow interiors and much greater in others, but always less than a few percent. Amygdules are filled with quartz, calcite, chlorite, adularia, sericite, hematite, bornite, and chalcopyrite. Small amounts of copper sulfides are particularly noted in flow tops cropping out at the top of the mountain. The nonmetallic minerals are similar to those found throughout the Keweenaw Peninsula (Butler and Burbank, 1929). However, the occurrence of the copper sulfides chalcopyrite and bornite in amygdules is very uncommon in the Keweenaw Peninsula. Copper sulfides are reported by Robertson (1975) in the tops of lava flows at Mount Bohemia near the tip of the peninsula. New Paleomagnetism Data Rocks of the MCR are probably among the worlds most extensively studied by paleomagnetic methods (Halls and Pesonen, 1982). Reversed polarity of natural remanence was recently reported in the flows of Silver Mountain (Kulakov et al, 2012). Well-defined characteristic remanent magnetization in samples from 13 flows revealed a paleomagnetic mean direction typical for reversely magnetized Keweenawan rocks (Fig. 4) (Kulakov et al., 2012). The paleomagnetic direction was similar to that found in the Lower North Shore Volcanics that have been dated at 1107.9±0.8 Ma (Davis and Green, 1997) and the Powder Mill Group (Halls and Pesonen, 1982) dated at 1107.3±1.6 Ma (Davis and Green, 1997). Thus, the likely age of the Silver Mountain basalts is 1107 to 1108 Ma and the same as the Siemens Creek Formation, Powder Mill Group. Figure 4: Equal area projection showing the mean paleomagnetic directions for selected reversely magnetized rocks from the MCR. Open square - Silver Mountain (N=13) (Kulakov et al, 2012.); open triangle – Powder Mill Group (N=9) (Palmer and Halls, 1986); open circle – North Shore volcanics (N=21) (Halls and Pesonen, 1982). 63 �Geochemistry Major and trace element geochemical analysis were conducted on samples from seven flows (Fig 5.) These samples were characterized by very uniform composition for both major and trace elements. The major and trace element composition of the Silver Mountain flows are very similar to that reported for the Upper Siemens Creek formation and equivalent rocks of the lowermost part of the North Shore Volcanics and Osler Volcanics. These youngest rift-related flows belong to basalt type II of Nicholson et al, (1997). The compositional similarity of the flows of Silver Mountain to this group of basalts, and in particular to the Upper Siemens Creek rocks further confirms the close relationship of the Silver Mountain lava flows to the Powder Mill Group. Nicholson et al. (1997) concluded that these basalts were derived from a mantle plume, but were contaminated by continental lithospheric crust. Figure 5: Primitive mantle normalized plot comparing the average trace element composition of rocks of Silver Mountain (N=7) (open squares) and Basalt type II from the Upper Siemens Creek Formation (N=18) (open circle). Siemens Creek Formation data from Nicholson et al. (1997) and Silver Mountain data from Kulakov et al. (2012) Field Trip Stop This field trip provides the opportunity to observe the lava flows at Silver Mountain and to enjoy a scenic view from the top. The trip begins from the parking area at the bottom of Silver Mountain with a short hike up a trail consisting of a combination of rocky terrain and built in stairs to the top. The Silver Mountain adit is located near the parking area with the 1850s poor rock pile scattered near the entrance. It mainly consists of amygdaloidal basalt from the oldest exposed flow at Silver Mountain. The adit exists because the top flow here was more mineralized adjacent to a fault striking parallel to the adit. Fragments of fault breccia can be observed. 64 �After viewing the mineralized rocks from the adit, the trip will proceed laterally to view the relatively thick lava flows at the base of Silver Mountain. The trip will then continue with a climb to the top of the mountain and involves strenuous physical exertion. Along the way there will be an opportunity to observe the character of the massive interiors and amygdaloidal flow tops of several flows with different thicknesses. The thickest ( ~ 6 m) lava flow at Silver Mountain crops out approximately half way to the top. Towards the top, the thickness of the flows tends to decrease. At the top, there is an excellent scenic view of the surrounding area. The shallow-dipping lava flows roughly parallel the gentle sloping topography of the top of Silver Mountain making it more difficult to observe contacts between the cross sectional views of the lava flows. The highest point is on the updip side (southwestern side). Proceeding northeast from the top and slightly down the steep side one can observe the amygdaloidal top and underlying massive interior of a flow. The flow top is particularly notable because the amygdules are filled with copper sulfides (chalcopyrite) and calcite. The trip to the top is not recommended in stormy or wet weather. The glacially polished rocks surfaces at the top can be quite slippery when wet. REFERENCES CITED Bornhorst, T.J., 1997, Tectonic context of native copper deposits of the North American Midcontinent Rift System: Geological Society of America Special Paper 312, p. 127-136. Bornhorst, T. J., and Lankton, L. D., 2009, Copper mining: A billion years of geologic and human history: in Schaetzl, R., Darden, J., and Brandt, D. (eds.), Michigan Geography and Geology, Pearson Custom Publishing, New York, p. 69-90. Burt, A.W., 1849, Message from the President of the United States to the two Houses of Congress at the Commencement of the First Session of the 31st Congress, December 24, 1849: Pt.111, Geological Report of W.A. Burt, On Survey of Township Lines in 1846, p. 849. Butler, B.S., and Burbank, W.S., 1929, The copper deposits of Michigan: U.S. Geological Survey Professional Paper 144, 238 p. Campbell, R.E., 1952, Geophysical investigation of the Silver Mountain Area – Houghton County, Michigan: In Partial fulfillment of the Requirement for the Degree of Master of Science. Michigan College of Mining and Technology, 64 p. Cannon, W. F., Green, A. G., Hutchinson, D. R., Lee, M.W., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American mid-continent rift beneath Lake Superior from Glimpse seismic reflection profiling: Tectonics, v. 8, p. 305-332. 65 �Cannon, W. F., 1994, Closing of the Midcontinent Rift - A far field effect of Grenvillian contraction: Geology, v. 22, p. 155-158. Cannon, W.F., and Nicholson, S.W., 2001. Geologic map of the Keweenaw Peninsula and adjacent area, Michigan. United States Geological Survey, Geologic Investigations Series, Map I-2996. Davis, D.W., and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system: Earth and Planetary Science Letters, v. 97, p. 54-64. Davis, D.W., and Green, J.C., 1997. Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution: Canadian Journal of Earth Sciences, v.34, p. 476-488. Foster, K.W. and Whitney, J.D., 1851, Report on Geology of the Lake Superior Land District: Pt.11, 32nd Congress, Special Sessions Supt. Executive Documents, v. 41, p. 68-69. Halls, H.C. and Pesonen, L.J., 1982, Paleomagnetism of Keweenawan rocks: Geological Society of America, Memoir, 156, 173-201. Heaman, L.M., Easton, R.M., Hart, T.M., MacDonald, C.A., Hollings, P., and Smyk, M., 2007, Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario: Canadian Journal of Earth Sciences, v. 44, p. 1055-1086. Kulakov, E.V, Smirnov, A.V., Bornhorst, T.J., Cundari, R., and Hollings, P. N., 2012. Paleomagnetism and geochemistry of the Geordie Lake and Silver Mountain basalts: Implications for the Midcontinent Rift evolution. American Geophysical Union 2012 Fall meeting abstract. GP21A-1130 Lane, A.C., 1909, The Keweenawan Series of Michigan: Michigan Geological Survey, Publication, 6. Geological Series 4, p. 628-629. Nicholson, S.W., Shirey S.B., Schulz, K.J. and Green J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implication for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504-520. Paces, J.B., and Miller, J.D., Jr., 1993, Precise U-Pb ages of the Duluth Complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagmatic, and tectnomagmatic processes associated with the 1.1 Ga Midcontinent Rift system: Journals of Geophysical Research, v. 98, p. 13,997-14,013. Palmer, H.C. and Halls H.C., 1986, Paleomagnetism of the Powder Mill Group, Michigan and Wisconsin: A Reassessment of the Logan Loop. Journal of Geophysical Research, 91, 11, 11571-11580. 66 �Robertson, J.M., 1975, Geology and mineralogy of some copper sulfide deposits near Mount Bohemia,Keweenaw County, Michigan: Economic Geology, v. 70, p. 1202-1224. Roberts, E., 1940, Geology of the Alston District Houghton and Baraga counties, Michigan: In Partial fulfillment of the Requirement for the Degree of Master of Science, California Institute of Technology, Pasadena, California. 47p. Waggoner, T.D., 1994. Echo Lake Gabbro, Houghton County, Michigan. Abstract, Institute on Lake Superior Geology, 40th Annual Meeting, Houghton, Mich, 1994, Programs and Abstracts, pt. 1, p 70. 67 ��Field Trip 6 Geology and Environmental Site Conditions of the Copperwood Deposit, Gogebic County, Michigan Theodore J. Bornhorst A.E. Seaman Mineral Museum, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Allan Blaske AECOM, 401 S. Washington Square, Suite 100, Lansing, Michigan, 48933 Dave Anderson and Thomas D. Repaal Orvana Resources US Corp., 10199 Lake Road, Ironwood, Michigan 49938 97 �Introduction The Copperwood deposit is a stratiform copper deposit in Gogebic County, Upper Peninsula, Michigan that is hosted by gray to black shales and siltstones filling the Midontinent rift (MCR) (Fig. 1). Copperwood is a reduced-facies or Kuperschiefer-type sedimentary rock-hosted stratiform copper deposit (Bornhorst and Williams (in press)). Mineralization at Copperwood contains 33.2 million short tons of Canadian National Instrument 43-101 compliant Measured and Indicated resources with an average grade of 1.65 % Cu and 4.34 ppm Ag, There are 3.0 million short tons of inferred resources with an average grade of 1.07 % Cu and 2.01 ppm Ag (Ward 2011). The resources are based on a cutoff composite grade of 0.8 % Cu and thickness of 5 ft (1.5 m). The Porcupine Mountains sedimentary rock-hosted copper district (Bornhorst and Barron, 2011) encompasses the White Pine Mine and the Copperwood deposit (Fig. 1). The White Pine Mine produced approximately 4.5 billion lbs of Cu and 50 million ounces of Ag from 1953 to 1996 with a few interruptions in production. Copperwood was discovered in 1956 (a subsidiary of AMAX) after a USGS publication in Economic Geology (White and Wright, 1954) indicated potential for copper mineralization in the Western Syncline. In addition to Copperwood, the U. S. Metals and Refining Company exploration program discovered 3 lower grade mineralized areas within the Western Syncline with indicated and inferred resources of 1,348 million short tons with an average grade of 1.34 % Cu (Kulla and Thomas, 2011). During1957 to 1958, a 71 m vertical shaft, 635 m of drifting, and 3 small stopes were completed in the higher grade Copperwood deposit (Bornhorst and Williams, in press). A mine was not developed because of presumed ground stability issues that would force excess dilution during mining. Advances in mining technology combined with higher Cu prices make Copperwood an economic deposit today. Orvana Minerals Corp began exploration and environmental baseline studies at Copperwood in 2008. This was quickly followed by the first Canadian National Instrument 43101 compliant mineral resource reported in 2010 (Kulla and Parker, 2010), prefeasibility in 2011, and feasibility in 2012 (Keane et al., 2012). Copperwood was granted a mining permit by the State of Michigan in 2012 and in February 2013, the Michigan Department of Environmental Quality granted the wetlands permit which is the last major permit needed before construction and production can proceed. Production is projected to begin in the near future. The descriptions in this field guide are based on Orvana (2011) and Bornhorst and Williams (in press). This field guide is published with permission granted by Orvana Minerals Corporation; however, the content is the sole responsibility of the authors. Regional Geologic Setting The broad 300 km wide MCR in Michigan consists of more than 25 km thick succession of riftfilling tholeiitic flood basalts with minor interbedded red conglomerates and sandstones overlain by 8 km thick succession of rift-filling clastic sedimentary rocks (Cannon, 1993). A rift-flanking basin is filled with 3 km of sandstone. These rocks are make up the Keweenawan Supergroup in Michigan and were deposited from 1.15 to about 1.03 Ga (Fig. 2) (Heaman et al., 2007; Davis and Paces, 1990; Cannon et al., 1989). 98 �75 o 50o o 100 50 o Precambrian bedrock at surface Canada Minnesota Canada Lake Huron Midcontinent Rift Porcupine Mountains sediment-hosted copper district Wisconsin Canada Grenville Front Tectonic Zone Iowa Nebraska Lake Nipigon Calumet Copperwood project White Pine Mine Houghton Kansas 0 35o 100o Major native copper deposits 400 35o 75o kilometers Stipled area - Phanerozoic bedrock at surface Ontario Ontario Lake Superior Minnesota Wisconsin 46.00 Phanerozoic Sedimentary rocks Mesoproterozoic Midcontinent Rift Sedimentary rocks Igneous rocks Paleoproterozoic Metamorphosed sedimentary and igneous rocks Archean Metamorphosed sedimentary and igneous rocks Lake Michigan N 0 100 kilometers Major Faults Figure 1: Generalized bedrock geologic map of the Midcontinent Rift. Modified from Bornhorst and Barron (2011). The bedrock at Copperwood is part of the rift-filling clastic sedimentary rocks. Continental compression occurred at 1.06 Ga in response to the Grenvillian collision along the eastern edge of the North American continent inverting the original graben-bounding faults into reverse thrust faults (Fig. 2) (Cannon, 1994). A syncline at Copperwood is a result of this compressional event. Erosion followed continental compression from about 1.03 Ga to 0.5 Ga (500 Ma) during which time multiple km of bedrock was removed and likely exposing the Copperwood orebody at the bedrock. This would have allowed a period of downward percolating groundwater into the Copperwood deposit (Bornhorst and Robinson, 2004). Marine submergence beginning 500 Ma buried the Precambrian bedrock under multiple km of Phanerozoic sedimentary rocks; evidence of Phanerozoic rocks is missing at Copperwood. 99 �Pleistocene glaciation over the last 2 million years removed the Phanerozoic rocks overlying Copperwood and once again exposed the orebody at the surface. The retreat of the last glaciers about 10,000 years ago left behind unconsolidated glacial deposits that today overlie the Precambrian bedrock at Copperwood. Peterson (1985, 1986) determined that four distinct advances of glacial ice occurred in the western Upper Peninsula during the late Wisconsinan time (14,500 to 10,200 years B. P.). Two of these advances and retreats produced the current surficial features at Copperwood. The first advance moved out of the Lake Superior basin south to the Wisconsin-Michigan border and upon retreat, left behind the lower till. The youngest advance occurred approximately 10,200 years B.P., when the glacier overrode previous till deposits at Copperwood. The approximate southern limit of this last advance was approximately 3 km south of Copperwood. Following this final retreat, glacial Lake Duluth covered the Copperwood area. Several Lake Duluth shorelines are evident along the topography at Copperwood. o o 88o 89 90 Copper Harbor N Calumet Yj Yn Hancock Houghton 47o Copperwood Deposit 2- 6 Ontonagon L'Anse White Pine mine 1 Ironwood Wakefield 3 Stop Location 0 fault 10 20 30 kilometer Figure 2: Bedrock geologic map and lithostratigraphic column for the western part of the Upper Peninsula of Michigan. The location of the field trip stops are shown. 100 �Mesoproterozoic Bedrock Geology The Copperwood deposit is on the southwest limb of the open shallow-plunging Western Syncline (Fig. 3). The bedrock at Copperwood consists of clastic sedimentary rocks of the Mesoproterozoic Oronto Group (Fig. 2) that strike approximately east-west and dip approximately 10 degrees to the north. The MCR bedrock at Copperwood is overlain by unconsolidated Pleistocene glacial sediments. The lowermost lithostratigraphic layer at Copperwood is the Copper Harbor Formation that consists of primarily of reddish-sandstone. The Nonesuch Formation overlies and interfingers with the Copper Harbor Formation. It consists of gray-black shale and siltstone to gray-white siltstone to brownish-red siltstone that are subdivided into multiple informal subunits. The Freda Formation gradationally overlies the Nonesuch Formation and consists of brown siltstone at Copperwood. Lithostratigraphy Copper Harbor Formation. Overall, the Copper Harbor Formation is composed of red-brown conglomerates and sandstones with lesser siltstones in an upward- and basinward-fining sequence. In Michigan, there is a maximum exposed thickness of about 2,000 m (Elmore 1984). The Copper Harbor Formation sedimentary rocks are fluvial and deposited a coalescing alluvial fan environment. At Copperwood, the Copper Harbor Formation is the oldest lithostratigraphic bedrock formation. It is lithologically dominated by red-brown to white and gray fine- to coarse-grained, arkosic sandstone. The Copper Harbor Formation in one drill hole consisted of 140 m of sandstone, a 1 m thick red, matrix-supported conglomerate, and more sandstone. Outcrops of the Copper Harbor Formation along the southern portions of Namebinag Creek and an unnamed creek indicate an increase in conglomerate facies in the lower portions of the formation. Outcrops along these streams are predominantly conglomerates with lesser amounts of sandstones. The uppermost few feet of the Copper Harbor Formation intersects in all of the Orvana and legacy exploration drilling at Copperwood. The uppermost Copper Harbor Formation consists of interlaminated red-brown siltstones and shales with occasional beds of very fine-grained sandstones. Uncommonly, there are interbedded, thin beds of dark-gray shales and siltstones less than 1.5 cm below the upper contact indicative of interfingering overlying Nonesuch lithologies. Absent in some holes, is the red-brown siltstone at the top of the Copper Harbor Formation. It can be up to 1 m thick, but is typically less than 30 cm in thickness. Assay data has shown, the uppermost 1 m of the Copper Harbor Formation (red-brown siltstone and sandstone) does not carry significant amounts of copper. There is a dramatic and abrupt change from the reddish Copper Harbor Formation to dark-gray to black shales and siltstones of the overlying Nonesuch Formation. At Copperwood, this abrupt transition defines the change from the Copper Harbor Formation to the Nonesuch Formation. 101 �11 A’ Porcupine Mountains 22 21 20 11 6 29 30 28 27 33 34 7 32 31 36 4 6 8 2 1 5 6 4 3 20 A N Copper Harbor Formation 1 0 kilometers 11 12 25 8 7 10 9 Presque Isle River GeologyfromunpublishedmapsbyForbes(1959), Copper RangeCompany, andOrvanaexploration 0 M117 0 M57 Meter 500 1000 Feet 2000 N Orvana Drill Hole (C) M27 Western Sector of deposit M24 C68 M21 M13 C41 C42 C60 C62 C37 C111 1 Section Number M25 M63 C52 C39 C40 Section Line C100 36 M70 C16 M62 C44 M54 M23 C99 C58 C54 C101 P2 P5 C114 P6 P7 C31 P1 C73 C70 C55 C119 C116 C35A C36 C56 C66 M64 C32 C17 M48 M59 C79 M46 C78 C75 C71 C33 C50 P4 M69 C86 C81 P3 M28 C110 C29 P9 M22 C142 P8 C74 C49 C98 C13 C25 C30 C45 C28 C140 C83 C85 C80 C122 P10 C65 C136 C87 C104 C26 C106 C115 C82 C84 C138 C141 M53 C24 M31 C27 C47 C143 C95 C107 C97 C67 M72 C61 M11 C96 M52 M75 C108 M159 C88 M58 C102 C20 C63 C59 C77 P12 C92 C137 C139 C94 C121 C123 P11 M18 C21 C91 M9 P17 C133 C109 C93 C90 C43 C103 C69 C48 C105 P13 P18 M108 M12A C128 C23 C127 C126 M109 M80 M19 C38 M114 C132 M49 C130 M116 M4A P22 C89 C9 C22 C125 C129 C118 C11 C131 C113 C117 P16 P14 P19 P15 Nonesuch Formation at Bedrock Surface M103 C57 USMR Legacy Hole (M) Bear Creek Legacy Hole (P) M26 C34 C46 C112 M32 Freda Formation at Bedrock Surface C53 C72 C64 C51 C76 Eastern Sector of deposit Copper Harbor Formation at Bedrock Surface 2 Tabular CBSprojected to surface 11 Measured and Indicated Resources Inferred Resources Fault 1 1 6 12 12 7 T49N R46W P20 P21 T49N R45W Canadian National Instrument 43-101 Compliant Resource Classification Figure 3: Geologic map of the Western Syncline and Copperwood deposit. Modified from Bornhorst and Williams (in press). 102 �Nonesuch Formation. Overall, the Nonesuch Formation is composed of characteristically blackto-gray -green siltstones, shales, carbonate laminates, and minor sandstones with a maximum thickness of 215 m. Elmore et al. (1989) and Suszek (1997) interpreted the depositional environment of the Nonesuch Formation to be dominantly anoxic lacustrine ranging from marginal lacustrine (sandflat-mudflat) to lacustrine to lacustrine-to-fluvial subenvironments. At Copperwood, a completed stratigraphic section is exposed in the northeast part of the property, at a thickness of 200 to 215 m. The upper contact is missing due to erosion throughout most of Copperwood property. The formation has been subdivided into multiple informal members on the basis of lithologic variations (Fig. 4). All of these members have remarkable lateral continuity throughout the Copperwood area (Fig. 5). The Parting Shale member is at the base of the Nonesuch Formation and is further subdivided into units (Fig. 4). The three lower units of the Parting Shale are the host to copper mineralization at Copperwood and together are termed the Copper-Bearing Sequence (CBS). The lowermost unit of the Parting Shale and CBS is termed Domino after terminology used at the White Pine Mine. Domino averages 1.6 m thick in the western sector and thins to about 60 cm thick in the eastern sector. The overall average thickness is 90 cm with a range from 9 cm to 2.3 m. Domino is characterized by laminated dark-gray to black shales and siltstones. Domino hosts the highest-grade copper at Copperwood. The contact between Domino and the overlying Red Massive unit is sharp and easily recognized in drill core as an abrupt change from the dark gray/black (Domino) to red-brown (Red Massive). The Red Massive unit of the Parting Shale and medial unit of CBS averages 35 cm thick, ranges in thickness from near zero to 1.2 m, but is usually less than 50 cm thick. It is somewhat thicker in the eastern sector than in the western sector. Red Massive is characterized by massive dark red-brown siltstones with interbedded red-brown, fine-grained sandstones. The contact between Red Massive and the overlying Gray Laminated unit is gradational and is placed where the color changes from reddish gray to gray. The Gray Laminated unit of the Parting Shale and upper unit of CBS averages 1.1 m thick and ranges in thickness from 50 cm to 4 m. Gray Laminated is characterized by of light-to-medium, gray-to-reddish-gray laminated siltstones; some intervals are massive. The contact between Gray Laminated and the overlying Red Laminated is gradational and is placed where the color changes from gray-dominated to mixed maroon and gray. The Red Laminated unit of the Parting Shale and hanging wall of the CBS ranges in thickness from 10 cm to 3.4 m and is more typically 1.2 to 1.8 m thick. Red Laminated is characterized by laminated siltstones with bimodal color distribution of maroon/red-brown and gray. Typical Red Laminated has mottled or wavy maroon intervals interspersed with medium-gray to reddish-gray siltstones. The contact between Red Laminated and the overlying Gray Siltstone is gradational. 103 �0 Copperwood Terminology 0 10 Feet 100 Surface Meters 20 30 Glacial clay-rich till Alpha Code for Sections Freda Formation L M Brown-red and greyish-red siltstone K Reddish-black siltstone J Greyish-red Siltstone I H Cross-stratified grey siltstone Copperwood Terminology Upper Sandstone Grey and red-brown interbedded siltstone 9 8 Nonesuch Formation 7 G Black laminated siltstone 6 F Grey concretion siltstone E Dark grey laminated siltstone 5 4 Red Laminated 3 D C B A Gray Laminated Galaxy Stripey 2 Red Massive Upper Shale 1 Upper sandstone Domino Parting Shale 0 Copper Harbor Formation RedSiltstone Sandstone Red and white sandstone meters Figure 4: Lithostratigraphic column of Copperwood bedrock units with detail of Parting Shale member. Modified from Bornhorst and Williams (in press). 104 �Freda Formation. The top of the Nonesuch Formation gradually transitions into the Freda Formation over an interval of about 10 m where beds of coarse, light-brown siltstones and massive to cross-bedded, dark reddish-brown siltstones are intercalated with grayish-red siltstones. The contact is placed at the base of a brown to white, cross-bedded siltstone. At Copperwood, the Freda Formation is up to 120 m thick above and consists of brown siltstone. Only the base of the formation occurs at Copperwood as the rest has been removed by erosion. Structure The structure at Copperwood is simple and consists of bedrock units that dip gently to the north on the southwest limb of the Western Syncline (Fig. 5). Under the unconsolidated glacial sediments, dips for all bedrock units vary from 12° in the south near the subcrop to 8° in the north nearer the synclinal axis. The bottom surface of the CBS approximates a gently curved, dipping plane lacking significant undulations. One fault has been identified at Copperwood (Fig. 3). This fault is interpreted to be a shallow, north-dipping reverse fault with 3 to 7 m of vertical displacement. Minor fractures with less than 1 inch of displacement were observed in multiple holes and these fractures are typically healed by calcite. South A 229 Subcrop Base of Nonesuch Formation 168 1,300 meters or 4,263 feet C89 C87 C55 292ft East 265ft West C/D B North A’ C99 C41 211ft West 183ft East 162ft West 750 H K J G M J L I I E H C/D B 107 550 K H J G G I E H 350 E 46 -15 MASL 6 5 4 3 2 1 0 No Vertical Exaggeration C/D B C/D B 150 G E C/D B -50 FASL C89 C87 C55 Red Laminated Red Laminated Upper CBS Upper CBS Domino Domino C99 C41 RL Red Laminated Upper CBS D Domino 20 15 10 5 0 CBS- Datum- topof Copper Harbor Formation Vertically Exaggerated M117 M57 LakeSuperior M27 South-North Cross SectionA- A’ M24 M26 C34 M32 C57 C46 C41 C111 C112 B’ C37 M63 M25 C40 C42 C39 C52 C60 C62 C100 M62 C44 M70 M54 M23 C64 C16 C101 C99 C58 C54 C76 C114 P7 C31 C55 C70 C116 C35A C36 C56 C66 M48 M46 M59 C32 C17 C79 M64 C75 C78 C71 C33 C50 C29 C86 C81 M69 C83 M28 C110 P8 M22 C49 C74 C98 C28 C13 C25 C30 C45 C80 C65 C122 C140 C85 C87 C104 C26 C115 C106 C82 C84 C138 C24 M53 C95 M31 C107 C97 C27 C96 M11 C47 C67 M72 C61 C77 M52 M75 C102 M159 C88 M58 C108 C121 C20 C63 C59 C94 C 92 C123 C21 M18 C109 M9 C91 C93 C90 C43 C105 C103 C48 C69 P18 M108 C126 C128 C127 C23 M12A M109 M19 M80 M114 C38 M49 P22 M116 M4A C9 C22 C89 C118 C11 C113 C117 C68 36 C53 C119 M103 M21 M13 C72 C73 C51 B P9 P2 P5 P6 P1 P4 C142 P3 C143 P10 C136 C141 P17 C133 C137 C139 P11 P12 P13 C130 P19 C132 C125 C129 P16 C131 P14 P15 NoMineralizedHorizon 2 1 Copper 11 12 P20 P21 Harbor Formation1 6 12 7 HoleusedinCrossSection T49NR46W T49NR45W Figure 5: Geologic cross section of the Copperwood deposit. Modified from Bornhorst and Williams (in press). 105 �Copperwood Deposit At Copperwood, copper mineralization is hosted by gray to black shales and siltstones of the CBS (Fig. 4). The footwall consists of red-bed sandstones and minor siltstones of the Copper Harbor Formation and the hanging wall consists of maroon/red-brown to gray siltstones of the Red Laminated unit. Geologic cross sections using the base of the CBS as the horizontal datum demonstrate the high degree of lithologic continuity of the CBS. The copper orebody is a conformable and tabular with an average thickness of 2.5 m. In the western sector the CBS averages 2.9 m thick whereas in the eastern sector it averages 2.1 m thick. The Copperwood deposit contains a total (all categories) undiluted geologic resource of about 1.16 billion lbs of Cu and 4.4 million ounces of Ag. The deposit (CBS) is characterized by copper in the form of chalcocite with minor amounts of silver. Other copper minerals such as chalcopyrite and bornite occur above the CBS (Bornhorst and Williams, in press). Pyrite is virtually nonexistent in the CBS, but above the Parting Shale, pyrite does occur in low abundance. The CBS is dominantly siltstones that are composed of over 90% silicate minerals (quartz, clinochlore, muscovite, illite, K-feldspar, plagioclase), about 2% calcite, and 3% hematite. Overall, pyrite and other minerals with the potential to generate acids are lacking whereas calcite, which is an acid-neutralizing mineral species, is abundant. The White Pine Mine produced copper from almost the entire Parting Shale and from the overlying Upper Shale (Fig. 4) (Mauk, 1992; Ensign et al., 1968). Whereas the three layers that compose the CBS at Copperwood are lithologically similar to those at the White Pine Mine, their thicknesses and proportions are not. The total Parting Shale is much thicker at Copperwood and e.g., the Domino within the western sector is typically 2.5 times thicker than at White Pine. At the White Pine Mine, the ore minerals are chalcocite and native copper whereas Copperwood is devoid of native copper. At White Pine, the chalcocite and minor native copper is stratiform ore interpreted as being related to diagenesis (Brown, 1971; Ensign et al., 1968). The native copper represents a second stage of mineralization (Mauk et al., 1992) hosted in faults and fractures and is interpreted as being related to the native copper deposits of the Keweenaw Peninsula (Bornhorst, 1997; see Field Trip 1 this volume). Copperwood lacks the complexity of the White Pine deposit (Bornhorst and Williams, in press). The White Pine copper deposit straddles a right-lateral strike-slip fault and an anticline (Ensign et al., 1968; Johnson et al., 1995). Thrust faults, strike-slip faults, and normal faults are encountered throughout White Pine and these faults and folds are mostly related to late rift compression. Some clearly compression-related thrust faults host sheets of native copper. The second-stage mineralizing fluids were likely of the same origin as those related to native copper in the Keweenaw Peninsula (Bornhorst, 1997). The Copperwood deposit is an example of a reduced facies or Kupferschiefer-type sedimentary rock-hosted copper deposit (Bornhorst and Williams, in press). Bornhorst and Williams (in press) proposed that at Copperwood “chalcocite replacement of pyrite in unlithified sediments during diagenesis is a result of emanating upward-focused, compaction-driven, Cu-bearing saline basinal waters whose Cu was leached from the underlying red-bed paleoaquifer.” In comparison to most examples of reduced facies or Kupferschiefer-type (Hitzman et al., 2010, 2005; Cox et al., 2003), Copperwood is notable for its simplicity (Bornhorst and Williams, in press). 106 �Pleistocene Unconsolidated Glacial Deposits The unconsolidated deposits at Copperwood consist primarily of a reddish-brown glacial till. This glacial till unconformably overlies the bedrock and ranges in thickness from 0 (at outcrops along streambeds south of the subcrop of the Copperwood deposit) to 43 m; average thickness is approximately 25 m (Fig. 6). The top of the bedrock surface is generally smooth. Boulders or otherwise weathered bedrock were encountered in some borings at the bedrock surface, but at most locations the glacial till sits on a scoured bedrock surface. The surface of the bedrock is more or less parallel to the ground surface topography, and slopes toward the north-northwest. Figure 6: Schematic cross section through the Copperwood project with generalized geology. The glacial till at Copperwood is a mud-matrix supported diamictite. This diamictite is massive with no stratification, lamination or fining upward or downward and the matrix is a uniform mix of sand, silt, and clay. Grain-size distribution curves are typical of those associated with subglacial tills (Fig. 7). The characteristics lead to the interpretation of the glacial till at Copperwood as a subglacial diamictite (Kemmis, 2008), meaning that it was deposited beneath the glacial ice. The diamictite is dense and overconsolidated, characteristic of subglacial till, but unlike normally consolidated to slightly overconsolidated lacustrine deposits. The glacial till was described during the soil boring program as variations of a silty clay based on slight variations in the observed portions of silt, clay, and sand. Field classification ranged from silt, clay, silty clay, clayey silt, silty sand, and sandy silt. The till was found to contain trace (1 to 9%) to little (10 to 19%) to some (20 to 34%) amounts of sand and gravel. Soil samples tested for particle size distribution indicated that the average composition of the cohesive (silt/clay) was approximately 50% silt, 20% clay, and 30% sand (Fig. 8). Sample analysis of the till fine fraction (<200 sieve size) by X-ray diffraction indicate that quartz was the major component, 107 �Figure 7: Grain size distribution of glacial till samples from the Copperwood site. with very little clay minerals (kaolinite, illite, montmorillonite, etc.) present. Analyses also indicated the presence of small amounts of feldspar, micas, calcite, and hematite. Testing using dilute hydrochloric acid indicated the presence of very finely-ground calcite, volumetrically too low to be detected by X-ray diffraction. The gravel/cobble portion was difficult to estimate from the laboratory analyses, due to the mass of the samples subjected to sieve analysis. Some samples contained no gravel-sized fraction, while others contain up to 45% (by dry weight) of gravel (coarse and fine). Inspection of Rotosonic soil cores revealed as much as 10 to 20% coarse fraction (% gravel by volume) in the recovered soil core. Gravel, cobbles and boulders larger than 9 cm (diameter of drill sampling device) were encountered during the drilling as well as during site reconnaissance activities. Large boulders (up to 1 m in diameter) were present within the bluffs along the Lake Superior shoreline. The gravel portion of the till consisted of dark reddish-brown sandstone (between 60% and 70%), diabase (between 8% and 18%), granite/gneiss (between 7% and 13%), basalt/amygdaloid (between 1% and 5%), and other types of rocks (between 2% and 8%). The sandstone and basalt are likely to have been locally derived from the Keweenawan Supergroup bedrock. Beach stones on the Lake Superior shoreline represent a material washed from the entire thickness of the till outcrops along the lakeshore. The predominance of red sandstone (likely Freda Formation) in the till accounts for the overall red-brown color of the glacial material. 108 �Figure 8: Soil texture plot of glacial deposits from the Copperwood site. Figure 9: Stratigraphic section of the glacial overburden deposits at the Copperwood site. 109 �The glacial till can be divided into two units (upper till and lower till), based on matrix grain size, amount and size of gravel, and vertical distribution (Fig. 9). In addition, there are thin (< 3 m with an average thickness of < 1 m) and isolated layers of coarser (non-cohesive) sediments throughout Copperwood in various borings that are located predominantly between the upper and lower till units. These generally consisted of fine to medium sand with varying amounts of silt and clay. These granular deposits, when encountered, are not laterally extensive, and for the most part cannot be correlated between adjacent borings. They are interpreted as lacustrine, intra-till sands or subglacial melt water deposits. Additional granular deposits were found in three borings at the base of the overburden, on the top of the bedrock surface. Peterson (1985) described the glacial deposits of the area as thin (< 10 m) drift over bedrock. Hack (1965) described the Ontonagon Plain (located on the east side of the Porcupine Mountains) to be underlain by reddish-brown glacial lake sediments and till. He described three units within the glacial deposits – the lower, intermediate, and upper units. The lower unit is described as a stony till containing locally derived subangular boulders and fragments. The intermediate unit is described as till and laminated silt and clay in distinct layers that are believed to be lacustrine sediments. This unit is less stony than the lower unit. The upper unit is described as a clayey till that is much less stony than either of the lower units. Thin lacustrine deposits related to glacial Lake Duluth are described as a patchy thin (< 60 cm) overlying veneer of strongly-laminated clay, silt and sand of variable thickness. The three primary glacial units described by Hack (1965) are present at Copperwood: the lower till which is slightly coarser grained with more (and larger) clasts than the upper till, the intermediate unit, less well defined at Copperwood, but composed of the laterally inconsistent layers of silty and sandy sediments, and the upper till. A thin (< 1 m thick) of lacustrine deposits related to Lake Duluth exists at Copperwood. The glacial deposits at Copperwood are composed predominantly of subglacially deposited till that was streamlined by glacial movement into elongated drumlins and flutes. While the flutes or ridges are not obvious within at Copperwood, such features may have been the cause of the distinct modern Copperwood drainage pattern. 110 �Objectives of Field Trip This field trip is designed to provide a geologic overview of Copperwood deposit hosted by the Mesoproterozoic MCR bedrock through descriptions and observations of drill core. The environmental site conditions at Copperwood will be observed and discussed onsite, especially those associated with surface and groundwater. The unconsolidated Pleistocene glacial deposits that unconformably overly the deposit will be observed as they play an important role in the environmental site conditions. The environmental site conditions are a critical aspect of permitting a modern mine. The location of Field Trip stops depend on site activities and accessibility. Those described below are based on full access and moderately dry conditions. Participants will be provided a map at the time of the field trip. Access to the Copperwood site is strictly forbidden without express permission from Orvana Minerals US Corp. No samples of any kind are allowed to be removed from the Copperwood site during this field trip. STOP 1: Orvana Offices, Ironwood, Michigan The first stop of the field trip will be to the offices of Orvana Resources U. S. Corp. in Ironwood. An overview of the Copperwood geology and project will be provided. Core of the bedrock and unconsolidated glacial deposits will be available for inspection and discussion. Core drilling of the Copperwood deposit will be available for inspection. Since there is little lateral variation within the Copperwood orebody, core from only a few drill holes are necessary to observe the character of the orebody as a whole. The unconsolidated glacial deposits were sampled using Rotosonic drilling techniques. Core samples of the glacial till will be available for inspection. Slight differences in the matrix composition and gravel content can be observed in the various core samples. STOP 2: Copperwood Historic Test Mine Rock Pile and Weather Station The Copperwood historic mine rock pile is a result of testing mining in 1957 to 1958. This rock pile was once much larger, as prior to Orvana’s activities at Copperwood, this rock was used as fill in wet areas of roadways and along stream crossings. The rock pile was used by Orvana as a staging area for exploration and pre-development activities and for an environmentally focused rock pile study. The rock pile will be removed upon creation of the tailings disposal facility. The rocks in this 1957-58 rock pile are dominated by unprocessed ore (CBS), but include hanging wall and footwall rocks as well. None of the ore was processed during this 1957-58 test mining, except for bulk samples delivered to Michigan Tech (then known as Michigan College of Mining and Technology) for process testing. The rock pile has been subjected to slightly more than 50 years of weathering which continues today. There is no evidence of acid drainage from the rock pile. Blocks of black shale, likely Domino, which contains the most copper within the CBS, are present on the surface of the rock pile and can be identified by notable green surface coloration. The green mineral has been identified by X-ray diffraction as malachite (copper carbonate). These blocks readily separate into smaller fragments along bedding planes 111 �and, upon separation, the malachite is only visible for less than 2 cm from the edge of the block even when the bedding planes are visibly moist. An environmentally focused study was initiated by Orvana to validate bench scale laboratory determined rates of release for chemical constituents and evaluate the long-term environmental impact of weathering of Copperwood unprocessed ore-bearing rock. After 50 years of well aerated, well-drained, and high-infiltration leaching, the rock pile has been and continues to be acid-neutralizing since the precipitation today is acidic. The study of the rock pile will be discussed at this stop. Immediately to the south of the rock pile a weather station is located in a clearing on the east side of the entrance road. This station was installed in 2008 and used to collect site-specific temperature, wind, precipitation, ground temperature, and air quality information for the Environmental Impact Assessment. STOP 3: Monitoring Well Sites Monitoring well nests were installed across the site to determine groundwater and aquifer characteristics (flow direction, flow rate, vertical gradients, groundwater chemistry, etc.). As we travel towards the Lake Superior shoreline, several sets of monitoring wells can be observed, and we will stop briefly to discuss the data collected from them. At each monitoring well site, a well was installed into the glacial deposits, and a second well installed into the underlying bedrock. Groundwater elevation data collected from the well network indicates that the groundwater in the glacial deposits, upper portions of the Copper Harbor formation, and the Nonesuch formation flow to the north-northwest, toward Lake Superior, and generally follows the slope of ground surface topography. The indicated groundwater velocity is from 0.63 to nearly 2.8 feet per year (fpy) within the bedrock units and from 0.7 fpy to nearly 1.1 fpy in the glacial deposits. The highest concentrations of total dissolved solids (TDS) and chloride in the groundwater are located in the top of the Copper Harbor Formation and the bottom of the Nonesuch Formation; concentrations decrease upward through the Nonesuch (Fig. 10). The groundwater in the glacial deposits generally contains much lower TDS concentrations. There is very little connection between groundwater within the various units, except for the uppermost unit in the glacial deposits which illustrates characteristics of recharge from surface water. 112 �113 �Groundwater in the uppermost portions (upper 30 feet) of the glacial deposits has relatively low TDS (average of 477 mg/L). It is depleted in sodium, chloride, and sulfate, and is mainly calcium bicarbonate type water. This water type is typical of groundwater near a precipitationfed recharge zone that has had relatively short contact time with the geologic materials. Groundwater within the remainder of glacial deposits has an average TDS of 878 mg/L (ranging between 110 mg/L to 7,300 mg/L) and it is depleted in magnesium and sulfate. The ion composition varies from sodium chloride, calcium chloride, sodium bicarbonate/carbonate, to calcium bicarbonate/carbonate water types. These variations in water type and TDS concentrations indicate the groundwater is not in connection with a precipitation-fed recharge zone and lacks connection with or flow path between the overlying uppermost zone. Sodiumand/or chloride-type water typically indicates that groundwater is equilibrating with the surrounding geologic matrix. For wells screened in the Nonesuch Formation, TDS ranges between 115 mg/L to 34,000 mg/L, with an average of 6,800 mg/L, which is about eight times that for groundwater residing in the overlying glacial deposits. Calcium chloride is the dominant water type, particularly when TDS is elevated. Sodium is also a dominant cation in groundwater at some wells. This groundwater is not near a precipitation-fed recharge zone and is not connected with the overlying glacial deposits. Wells screened in the Copper Harbor Conglomerate have a range of TDS between 140 mg/L to 66,000 mg/L, with an average of 10,545 mg/L, which is about 1.5 times that for groundwater in the overlying Nonesuch Formation. Groundwater in one well has TDS at 66,000 mg/L, which is greater than the TDS of sea water (35,000 mg/L). The groundwater is calcium-chloride type. In addition to calcium and chloride, sodium and bicarbonate/carbonate ions are dominant in groundwater from wells with lower TDS. This groundwater lacks connection with other groundwater zones beneath the site. STOP 4: Surface Water Monitoring Points Monitoring of the flow characteristics of the surface water in streams was performed at several locations at Copperwood. As we travel towards the Lake Superior shoreline, streams can be observed, and we will stop briefly to discuss the data collected from them. The surficial drainage system at Copperwood is part of the Lake Superior watershed and is composed entirely of small streams, roughly parallel to one another, flowing to the northwest from higher ground towards the south directly into Lake Superior. There are no lakes at Copperwood. Water flow within the streams is flashy and significantly controlled by timing and duration of precipitation. No groundwater contribution has been observed in these streams. High flow in streams occurs during spring when the snow melts and after significant rain events. Flow increases and decreases quickly during rain events. All of the streams have periods of zero measurable flow either due to dry conditions or freezing in the winter. During the summer between rain events, a slight “trickle” of water can be observed flowing between cobbles in the stream bed which enters and exits multiple isolated pools. The many isolated pools found on all of the streams are also maintained by water flow just beneath the stream-bottom substrate. The 114 �upper reaches of the streams at Copperwood can be classified as ephemeral (flow only during or immediately after periods of precipitation) and the lower reaches as intermittent (flows only during certain times of the year). Perennial streams, which have continuous flow, are not present at the site. Ephemeral streams have no base flow and the stream beds are above the water table. Intermittent streams have base flow for at least some periods of the year. At the Copperwood site, this base flow is extremely small. No springs, seeps, or areas of wetland vegetation were observed that would indicate groundwater discharge to the surface water environment. Surface water at Copperwood has a neutral to slightly alkaline pH with most values are between 6.5 and 8.0. Lake Superior water is slightly alkaline (average and median pH of about 8). The average and median dissolved oxygen values for surface water (8.3 mg/L and 7.9 mg/L) and Lake Superior water (8.9 mg/L and 8.0 mg/L) indicate the waters are oxidized as is typical for surface water in contact with the atmosphere. The surface water is calcium-bicarbonate type, which is associated with precipitation and little or no contact with soil. The surface water contains TDS at levels of approximately 20% of that in the groundwater in the uppermost glacial till, which is consistent with very little groundwater contribution to surface water. STOP 5: Incised Stream Channels The streams at the site are deeply incised into the glacial overburden. Along the entrance road, the stream channels are only a few feet deep, yet at the north end near Lake Superior, the channels are as much as 12 m deep. As we travel towards the Lake Superior shoreline, streams can be observed, and we will stop briefly to discuss the characteristic of these channels. The incised stream channels are contained within steep-walled valleys. Active erosion is present within the steep-walled portions of the stream valleys. The floors of the valleys are generally flat and range in width between 15 and 60 m. The streams meander within the bottoms of the valleys with significant portions of the streams dammed by beavers creating tiered meadows within the valleys. The upper portions of the stream valleys are generally narrower and shallower than those further downstream. The overall gradient of the streams at Copperwood is approximately 20 m per km. STOP 6: Glacial Exposures along Lake Superior Shoreline At this stop we will examine the exposure of unconsolidated glacial material in the eroding bluff along the Lake Superior shoreline. The current shoreline of Lake Superior is dominated by a steep bluff which rises as much as 15 m above the lake surface. The entire face of the bluff is composed of slumped blocks of silt/clay-rich till. The majority of the bluff is not vegetated, but slumped blocks of soil often contain trees and surface vegetation that were brought down from the top of the bluff; the bluff is experiencing significant erosion. A beach of mixed sand and cobbles is present at the base of the bluff. The maximum width of this beach is 10 m and in many places is much narrower. The beach does little to protect the base of the bluff from wave action and in many places, wave action reaches across the narrow beach to the base of the bluff. 115 �The upper till unit of the glacial diamictite is exposed on the face of the bluff. The diamictite has a massive structure with no observed stratification or laminations. It is composed of a silty clay matrix with scattered gravel. The gravel is mostly cobbles and pebbles less than 8 cm in diameter, but there are boulders up to one meter in diameter. The cobbles and pebbles on the beach are predominantly red sandstone which is likely from the Freda Formation. The exposed diamictite is interpreted as a subglacial till. References Bornhorst, T.J., 1997, Tectonic context of native copper deposits of the North American Midcontinent Rift System: Geological Society of America Special Paper 312: pp. 127-136. Bornhorst, T.J., and Barron, R.J., 2011, Copper deposits of the western Upper Peninsula of Michigan: Geological Society of America Field Guide, v. 24, p. 83-99. Bornhorst, T. J., and Robinson, G.W., 2004, Precambrian Aged Supergene Alteration of Native Copper Deposits in the Keweenaw Peninsula, Michigan [abstract]: Institute on Lake Superior Geology Proceedings, v. 50, Part 1: pp. 40-41. Bornhorst, T.J., and Williams, W.C., in press, The Mesoproterozoic Copperwood sedimentary rock-hosted stratiform copper deposit, Upper Peninsula, Michigan: Economic Geology. Brown, A.C., 1971, Zoning in the White Pine Copper deposit, Ontonagon County, Michigan: Economic Geology , v. 66, p. 543-573. Cannon, W. F., 1994, Closing of the Midcontinent Rift - A far field effect of Grenvillian contraction: Geology 22, p. 155-158. Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American midcontinent rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332. Cannon, W. F., Peterman, Z.E., and Sims, P.K., 1993, Crustal-scale thrusting and origin of the Montreal River monocline - A 35-km-thick cross section of the Midcontinent Rift in northern Michigan and Wisconsin: Tectonics, v. 12, p. 728-744. Cox, D.P., Lindsey, D.A., Singer, D.A., and Diggles, M.F., 2003, Sediment-hosted copper deposits of the world: Deposit models and database: U.S. Geological Survey Open-File Report 03-107. Davis, D.W. and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent rift system: Earth and Planetary Science Letters 97, p. 54-64. Elmore, R.D., 1984, The Copper Harbor Conglomerate: A late Precambrian fining-upward alluvial fan sequence in northern Michigan: Geological Society of America Bulletin, v. 95, p. 610-617. 116 �Elmore, R.D., Milavec, G.J., Imbus, S.W., Engel, M.H., 1989, The Precambrian Nonesuch Formation of the North American Mid-Continent Rift: Sedimentology and Organic Geochemical Aspects of Lacustrine Deposition: Precambrian Research, v. 43, p. 191-213. Ensign, C.O., Jr., White, W.S., Wright, J.C., Partick, J.L., Leone, R.J., Hathaway, D.J., Trammell, J.W., Fritts, J.J., and Wright, T.L., 1968, Copper deposits in the Nonesuch shale, White Pine, Michigan: SME Graton Sales Ore Deposits of the United States 1933-1967, v. 1, p. 460-488. Hack, J. T., 1965, Postglacial Drainage Evolution and Stream Geometry in the Ontonagon Area, Michigan, U.S. Geological Survey Professional Paper 504-B. Heaman, L.M., Easton, R.M., Hart, T.M., MacDonald, C.A., Hollings, P., and Smyk, M., 2007, Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario: Canadian Journal of Earth Sciences, v. 44, p. 1055-1086. Hitzman, M., Selley, D., and Bull, S., 2010, Formation of sedimentary rock-hosted stratiform copper deposits through Earth history: Economic Geology, v. 105, p. 627-640. Hitzman, M., Kirkham, R., Broughton, D., Thorson, J., and Selley, D., 2005, The sedimenthosted stratiform copper ore system: Economic Geology 100th Anniversary Volume, p. 609642. Johnson, R.C., Andrews, R.A., Nelson, W.S., Suszek, T., and Sikkila, K., 1995, Geology and mineralization of the White Pine copper deposits: unpublished Copper Range Company Report. Keane, J. M., Milne, S., and List, D., 2012, Feasibility Study on the Copperwood Project, Michigan, USA NI 43-101 technical report: KD Engineering, SEDAR published report. Kemmis, T., 2008, Unraveling the Complexity of Glacial Successions: in Course Workbook for Improving the Description and Characterization of Glacial Successions for Environmental and Engineering Projects, Midwest Geosciences Group, Western Michigan University, June 2008. Kulla, G., and Parker, H., 2010, Copperwood project, Michigan, USA NI 43-101 technical report: AMEC, SEDAR published report. Kulla, G., and Thomas, D., 2011, Copperwood S6 and satellite project NI 43-101 technical report, Michigan, USA: AMEC, SEDAR published report. Mauk, J.L., 1992, Geology and stable isotope and organic geochemistry of the White Pine sediment-hosted stratiform copper deposit: Society of Economic Geologists Guidebook Series, v. 13, p. 63-98. 117 �Mauk, J.L., Kelly, W.C., van der Pluijm, B.A., Seasor, R.W., 1992, Relations between deformation and sediment-hosted stratiform copper mineralization: evidence from the White Pine part of the Midcontinent rift system: Geology, v. 20, p. 427–430. Orvana, 2011, Copperwood Mine Orvana Resources US Corp. Part 632 Mine permit application: submitted September 26, 2011, published online by Michigan Department of Environmental Quality. Peterson, W. L., 1985, Surficial Geologic Map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: USGS Miscellaneous Investigation Series, Map I-1390-C. Peterson, W. L., 1986, Late Wisconsinan Glacial History of Northeastern Wisconsin and Western Upper Michigan: U. S. Geological Survey Bulletin 1652. Suszek, T., 1997, Petrography and sedimentation of the Middle Proterozoic (Keweenawan) Nonesuch Formation, western Lake Superior region, Midcontinent Rift system: Geological Society of America Special Paper 312, p. 127-136. Ward, M.B., 2011, Resource estimate and NI 43-101 technical report for Copperwood project, Ironwood, Michigan for Orvana US Corp: Marston & Marston Inc., SEDAR published report. White, W.S. and Wright, J.C., 1954, The White Pine copper deposit, Ontonagon County, Michigan: Economic Geology, v. 49, p. 675-716. 118 �Geology of the Keweenawan Supergroup, Porcupine Mountains, Ontonagon and Gogebic Counties, Michigan Laurel G. Woodruff1, William F. Cannon2, Suzanne W. Nicholson2, Klaus J. Schulz2, and Robert Wild3 1 U.S. Geological Survey, St. Paul, Minnesota; 2U.S. Geological Survey, Reston, Virginia; 3 Porcupine Mountain Wilderness State Park, Ontonagon, Michigan Introduction This field trip examines the geology of the rocks of the Keweenawan Supergroup (1.1 Ga) and related intrusive rocks of the Midcontinent rift system (MRS) exposed in and around the Porcupine Mountains. Most stops on this trip were visited in a previous Institute on Lake Superior Geology field trip guidebook (Cannon and others, 1992). The stop descriptions here are taken largely from that field guide with minor updates, new location maps and photographs. Because of uncertainties of weather, road conditions, and remaining snow pack in early May, the specific stops that we will visit will not be known until the date of the trip. Latitudes and longitudes are from GPS readings using WGS 84 datum. General Geology The 1.1 Ga Midcontinent rift system (MRS) is a prominent 2500 km linear feature on gravity and magnetic maps that extends from Kansas north to Lake Superior and then southeast beneath the Michigan basin to where it is cut off by the Grenville Front near Detroit, Michigan (Fig. 1). The MRS cuts across Early Proterozoic and Archean terranes and is attributed to crustal extension resulting from upwelling and decompression melting of an anomalously hot mantle plume at the base of the continental lithosphere (Nicholson and others, 1997). In the Lake Superior region, nearly complete crustal separation accompanied emplacement of as much as 2 million km3 of extrusive basalt and possibly an equal volume of intrusive rocks from about 1108 Ma to about 1087 Ma (Hutchinson and others, 1990; Nicholson and Shirey, 1990; Allen and others, 1995). In the Lake Nipigon area, intrusive rocks that have been attributed to the Midcontinent Rift event suggest that magmatism may have started as early as 1115 Ma (Easton and others, 2007). The deepest part of the rift subsided along large normal growth faults to a depth of nearly 30 km, accommodating at least 20 km of rift-related volcanic rocks. Following the volcanic phase of rifting, thermal subsidence accommodated deposition of up to 10 km of overlying sedimentary rocks in central rift grabens and flanking basins. Rocks of the MRS in western Lake Superior are known as the Keweenawan Supergroup and contain a remarkably complete record of igneous intrusion, flood basalt volcanism, and clastic sedimentation (Fig. 1). Along the south shore of western Lake Superior, initial subsidence of the MRS is recorded by deposition of the Bessemer Quartzite (Fig. 2), a blanket of relatively pure, fluvial sandstone as much as 100 m thick (Ojakangas and Morey, 1982). The Bessemer Quartzite is overlain by a great thickness of subaerial flood basalt flows and lesser intermediate and rhyolitic rocks (Powder Mill Group and Portage Lake Volcanics). Conformably overlying the volcanic rocks are fluvial sedimentary rocks (Copper �Harbor Conglomerate and Freda Formation) and lesser lacustrine sedimentary rocks (Nonesuch Formation) which are as much as 8 km thick beneath Lake Superior (Cannon and others, 1993) and at least 5 km thick on shore in the field trip area. Along the south shore of western Lake Superior, post-rift movement along the rift-bounding Keweenaw fault has caused large block rotation so that much of the Keweenawan Supergroup section is steeply to vertically dipping. Figure 1. Regional bedrock geologic map of western Lake Superior showing the distribution of rocks related to the Midcontinent rift system (N is magnetically normal; R is magnetically reversed) and area of the geologic map in Fig. 3 (modified after Miller and Chandler, 1997). �Figure 2. Age in Ma, magnetic polarity (N is normal and R is reversed), and stratigraphic section for rocks of the Keweenawan Supergroup in western Michigan and northeastern Wisconsin. Field trip stop numbers are placed in their relative stratigraphic positions (modified from Zartman and others, 1996). �Volcanic rocks of the Keweenawan Supergroup (Fig. 2) range in composition from olivine tholeiite to rhyolite. By far the dominant rock type is high-Al olivine tholeiite (Al203 = 15 to 19 wt. %) followed by lesser high-Fe tholeiite and rocks of intermediate and felsic composition (Green, 1982; Paces, 1988; Nicholson and others, 1997). The basalts commonly are ophitic in texture and the dominant phenocryst is plagioclase. The most primitive Keweenawan basalts are geochemically similar to ocean island basalts and have incompatible trace elements ranging from slightly to strongly enriched compared to depleted or primitive mantle. Radiogenic isotope analyses (Sr, Nd, and Pb) of the main stage high-Al olivine tholeiites suggest that a likely source of the voluminous basalts was a trace element-enriched mantle plume (Paces and Bell, 1989; Nicholson and Shirey, 1990; Nicholson and others, 1997). Flows near the base of the exposed Portage Lake Volcanics north of the Keweenaw fault were erupted at about 1096 Ma and those near the top of the formation at about 1094 Ma (Davis and Paces, 1990). Thus, the great thickness of Portage Lake Volcanics, at least 8 km in this area, was erupted in only a few million years. Because synchronous volcanism occurred along the entire trend of the rift, the rift system as a whole was producing basalt at a rate unrivalled by any modern analog (Cannon, 1992). The Porcupine Volcanics (Fig. 2) create much of the topography of the Porcupine Mountains. The Porcupine Volcanics overlie the Portage Lake Volcanics, and represent a volcanic center that became active late in the volcanic history of the region, at about 1093 Ma. As much as 5 km of andesite, rhyolite, and basalt were erupted in a large shield volcano and deposited on top of a flat-lying lava plain composed of Portage Lake Volcanics. The unit has a lateral extent of about 35 km on the present erosion surface. The present arcuate shape of the Porcupine Mountains and the unusual hook-shaped map pattern of the Porcupine Volcanics are partly a reflection of the original shape of the volcanic shield (Fig. 3). Rhyolite near the top of the section has an age of 1093±1.4 Ma (Zartman and others, 1997). The Porcupine Volcanics consist of a sequence of subaerially deposited (in order of abundance) andesite, basalt, felsite, and quartz-porphyry lava flows, and minor interbedded volcaniclastic lithic sandstone, siltstone, and conglomerate (Hubbard, 1975). The abundance of felsic rocks, most common near the top of the formation, where they occur as both lava flows and domes (stops 3 and 8), and the predominance of andesite over basalt clearly distinguish the Porcupine Volcanics from underlying Portage Lake Volcanics. An abundance of intermediate and felsic volcanic rocks, such as the Porcupine Volcanics, is atypical of the MRS as a whole and is limited to only a few felsic volcanic centers associated with shield volcanoes. The major element compositions of the basalt, basaltic andesite, and andesite of the Porcupine Volcanics and the Portage Lake Volcanics are similar, but the Porcupine Volcanics are distinctly enriched in light rare earth elements (LREE) and Th compared to the Portage Lake Volcanics (Fig. 4). The two formations differ more significantly in their rhyolite chemistry and mineralogy. The rhyolite that occurs most commonly in the Portage Lake Volcanics is aphyric or may contain sparse quartz phenocrysts. In contrast, numerous rhyolite bodies in the Porcupine Volcanics range from rhyolites that are aphyric to those with abundant quartz and/or feldspar phenocrysts. Rhyolites of the Portage Lake Volcanics on the Keweenaw Peninsula typically have lower abundances of incompatible trace elements �Figure 3. Generalized geologic map of the Porcupine Mountains area, modified from Cannon and others (1995). Field trip stops are shown as red diamonds. �(such as LREE, Zr, Y, Hf, and Th) than rhyolite of the Porcupine Volcanics (Fig. 4A). Radiogenic isotope analyses suggest that most Portage Lake rhyolites were derived by partial melting of already erupted Keweenawan basalt with a minor contribution, if any, from older basement, whereas rhyolites of the Porcupine Volcanics have a much larger contribution from older basement. Figure 4. A. Spidergram illustrating the average compositions of Portage Lake Volcanics Type 1 rhyolites, Porcupine Volcanics average rhyolite, and a sample of rhyolite from the quarry at stop 8. B. Spidergram illustrating the average compositions of Portage Lake Volcanics average basalts, Porcupine Volcanics average basalts, and samples of the Lake Shore traps from the area of stop 7. �Abrupt changes in thickness of the Porcupine Volcanics are inferred along prominent structural breaks that are especially evident on the aeromagnetic map of the area (King, 1987). These breaks are believed to be synvolcanic normal faults, which outlined a central caldera. A major gravity low, centered just south of the Porcupine Mountains, was interpreted by Klasner (1989) as a large, shallow felsic intrusion. This intrusion may have been the subvolcanic felsic magma chamber that erupted much of the Porcupine Volcanics. Eruption of the Porcupine Volcanics marked an end of major volcanism in the Midcontinent Rift, with further events dominated by fluvial and lesser lacustrine sedimentation. The Copper Harbor Conglomerate is a sequence of red to brown arkosic conglomerates and sandstones, interpreted as a northward prograding alluvial fan complex (Daniels, 1982). The Copper Harbor crops out along a discontinuous belt from the east end of the Keweenaw Peninsula westward into Wisconsin. There is an inverse relationship between the thickness of the Copper Harbor and the Porcupine Volcanics such that the Copper Harbor thins as the Porcupines Volcanics become thicker (White, 1972; Cannon and Nicholson, 1992). This suggests that the broad shield volcano that formed the Porcupine Mountains was a persistent topographic high during the time of Copper Harbor deposition. In the vicinity of the Porcupine Mountains, true conglomerate is rare; the Copper Harbor typically is a platy, reddish, fine- to medium-grained sandstone/siltstone (stop 6). Within the Copper Harbor, there are up to 31 basaltic andesite to andesite lava flows interspersed with sediment, generally in the upper part of the formation (Fig. 2). These flows, informally known as the Lake Shore traps, have an age of about1087±1.6 Ma (Davis and Paces, 1990) and form the prominent north flank of the Porcupine Mountains (stop 7). A unit of dark gray shale and siltstone, the Nonesuch Formation (stops 1, 4 and 5), conformably overlies and interfingers with the upper Copper Harbor Conglomerate. The Nonesuch is generally interpreted to have been deposited in a perennial lake located at the toe of a transgressing-regressing alluvial fan complex (Daniels, 1982, Elmore and others, 1988; Suszek, 1997). Basal beds of the Nonesuch and locally the top of the Copper Harbor contain regional-scale low-grade stratiform copper mineralization (White and Wright, 1954). Economic–grade ore bodies are located on the western (Presque Isle) and eastern (White Pine) flanks of the Porcupine Mountains. Copper typically occurs as fine-grained disseminated chalcocite; at White Pine chalcocite is accompanied by minor native copper. The White Pine mine, a large underground mine just east of the Porcupine Mountains, produced more than 1.8 million metric tons of copper from 1955 until the mine closed in 1995. The Copperwood deposit, as the Presque Isle deposit is now known, has just completed the pre-mining permitting process. The historical Nonesuch Mine (stop 4) is an example of the early mining history in the region. Deposition of the Nonesuch Formation was succeeded by a return to fluvial redbed deposition during which at least several kilometers of red to brown sandstone and siltstone of the Freda Sandstone (stops 1 and 2) were deposited. The Freda, although still rich in volcanic detritus, is compositionally more mature than older sandstones, which probably indicates that Archean and Early Proterozoic basement in the source area was exposed by erosion of the overlying Keweenawan basalts. Structures along much of the MRS are generally simple, consisting of thick monoclinical sections of rock titled toward the rift axis by a combination of rift subsidence, and later compression andrift inversion. Near the Porcupine Mountains, the structure is somewhat more complicated. The �major structure along the south shore of western Lake Superior is the Keweenaw fault, a reverse fault which partly inverted the central graben of the rift. A seismic section just east of the area shows the Keweenaw fault to be a north-dipping listric thrust (Hinze and others, 1990). Thus, in the area of the field trip, all sedimentary and volcanic units are on the upper thrust plate. Several cross faults that probably developed during eruption of the Porcupine Volcanics complicate the geology of the upper plate as do the gentle folds in the Porcupine Mountains, which produce a repetition of stratigraphy. Thrust faulting and folding can be indirectly dated in the interval from approximately 1060 to 1040 Ma (Ruiz and others, 1984; Bornhorst and others, 1988; Cannon and others, 1993). The regional compression that reversed the sense of movement along rift-bounding faults is likely the result of the Grenville orogeny (Cannon, 1994). The bustling community of Nonesuch, drawn in 1884 by Agnes Hathaway, a miner’s daughter. The town grew up around the Nonesuch copper mine that operated sporadically from 1866 until 1912. Now a ghost town, Nonesuch in its heyday once had a post office, school house, lumber mill, boarding house, and baseball team. The site (stop 4) became part of Porcupine Wilderness State Park in 1988. Image provided by Robert Wild. �Stop 1: Nonesuch Formation and Freda Sandstone at the mouth of the Presque Isle River: 46.7087ºN -89.9734ºW The upper portion of the Nonesuch Formation and the base of the Freda Sandstone are well exposed in the gorge of the Presque Isle River near its mouth and along the shore of Lake Superior west of the river (Fig. 5). Continuous exposures along the picturesque gorge of the river extend from just upstream of Nawadaha Falls to the lakeshore. Exposures continue in bluffs along the lakeshore for about half a mile west of the river mouth. An examination of exposures near the river mouth and a short distance to the west along the shore require a round trip hike of nearly a mile, mostly on wellmaintained trails and stairways. Figure 5. Location and geologic setting for stops 1 and 2. Geology is generalized from Cannon and others (1995). Structures are dotted and contacts are thin lines. �The rocks exposed here are on the northeast limb of the Presque Isle syncline, a gentle northwest-plunging fold. Dips range from nearly flat to about 10º SW. The Nonesuch Formation is distinguished from other sedimentary units of the Keweenawan Supergroup by a predominance of gray, green, or black fine-grained sediments. Lower Keweenawan felsic, intermediate, and mafic volcanic rocks were the major contributor of detritus to the Nonesuch, with contributions from Early Proterozoic and Archean crystalline rocks increasing up section (Suszek, 1997). Many of the units here show trough cross-bedding, symmetrical and asymmetrical ripples, rib and furrow structures, and parting lineations (Fig. 6A-C). A good example of ball-and-pillow structure, probably indicative of seismically-generated slumping, is seen in a one meter thick bed best exposed on the west bank of the river just upstream from the lower gorge (Fig. 6B). Finer-grained rocks include well-laminated shales, which are most abundant lower in the section. The Nonesuch displays coarsening-upward sequences at scales ranging from a few meters to the entire thickness of the unit. On a smaller scale, fining upward sequences are common in units from a few centimeters to a few meters thick. The Nonesuch grades upward to the Freda Sandstone through a zone of dark gray laminated and small-scale cross-bedded siltstone and sandy mudstone are interbedded with medium to coarse-grained reddish brown sandstone. There is a gradual change in oxidation state, and grain-size and bedding thickness both increase, reflecting increased environmental energy (Daniels, 1982). Although not exposed in this area, the lower part of the Nonesuch section is strongly mineralized with very fine-grained chalcocite. Orvana Mineral Corporation is in the process of developing the Copperwood deposit, with an estimate of total proven and probable reserves of 27.42 at 1.41% Cu and 3.6 ppm Ag for contained metal of 852 million pounds of Cu and 3.2 million ounces of Ag (Bornhorst and Williams, 2011). While the lower, mineralized fine-scale stratigraphy at Copperwood is directly correlative with the stratigraphy at the White Pine mine, the upper stratigraphic sequence of the Nonesuch in the Presque Isle syncline does not correlate as well with the Nonesuch stratigraphy at White Pine. This indicates that at least during early Nonesuch deposition, sedimentary conditions were very uniform over the entire region surrounding the Porcupine Mountains. Copperwood also lacks the structural complexity and hydrothermal overprint characteristic of White Pine. The Copperwood deposit occurs along a single dipping plane and only one fault has been recognized; White Pine is cut by a major strike-slip fault and numerous smaller thrust faults. The widespread chalcocite mineralizing event is hydrothermally overprinted at White Pine by a second stage influx of Cu-bearing fluids that deposited native copper along faults and adjacent parting planes (Mauk and others, 1992). Stop 2: Freda Sandstone along the Presque Isle River: 46.6962ºW -89.9744ºN At this stop, near the axis of the Presque Isle syncline, reddish cross-bedded sandstone typical of the lower part of the Freda Sandstone is exposed (Fig. 5). The gently southwest dipping beds are probably 100 to 200 m above the base of the formation and slightly higher stratigraphically than those at stop 1. These are dominantly lithic somewhat micaceous sandstones. The Freda marks a return to fluvial redbed deposition following the lacustrine deposition of the underlying Nonesuch Formation. The Freda Sandstone is a very thick unit in much of the rift in the western Lake Superior region and volumetrically is the dominant unit of the post-rift sedimentary fill. Thirty kilometers to the west, at the mouth of the �A. Potholes along the lower gorge of the Presque Isle River eroded into nearly horizontal shale of the Nonesuch Formation. B. Ball-and-pillow structures. Approximately 1-meter-thick bed of disrupted sediments between laminated siltstone. C. Ripple marks in laminated siltstone. Figure 6. Sedimentary structures in the Nonesuch Formation exposed in the Presque Isle River during low flow. Photographs by Bill Cannon. �Montreal River, about 3500 m of the upper part of the Freda is exposed in lakeshore bluffs. Seismic sections indicate the Freda is even thicker under the lake. The Freda generally becomes finer-grained and more mature upwards. Daniels (1982) interprets the cyclic sandstone-mudstone sedimentation of the Freda as an alluvial channel-fill sequence. Stop 3: Porcupine Volcanics - Rhyolite at Summit Peak and Beaver Creek: 46.7433ºN -89.7711ºW Summit Peak is the highest point in Porcupine Mountains Park and one of the highest points in the state. The observation tower at the summit provides a panoramic view of the field trip area. To the south, the highlands are underlain by the Portage Lake Volcanics and Porcupine Volcanics along the main monocline of volcanic rocks of the Keweenawan Supergroup. The lowlands immediately to the southeast are underlain by rocks of the Oronto Group in the east-plunging Iron River syncline. Looking east along strike, the stack from the former smelter at White Pine can be seen in the distance. To the north, the interior of the park extends over the rugged topography in the foreground to Lake Superior in the distance. The interior of the park is maintained as a wilderness area with access by hiking only. The park also contains some of the largest stands of virgin timber in Michigan. Most of the park interior is underlain by a thick unit of rhyolite composed of a series of lava flows and domes typified by the rocks seen at this stop (Fig. 7). There are good exposures of coarse rhyolite breccia present along the trail leading to Summit Peak and at the overlook platform west of the summit. This breccia is probably the carapace of a rhyolite dome (stop 3A). Excellent exposures of typical intermediate and felsic units of the Porcupine Volcanics occur on the north side of the hill (549 m elevation, stop 3B) along the Beaver Creek Trail about 0.5 mi from the Summit Peak parking lot. The units dip to the south and include, in stratigraphic order, sparse outcrops of intermediate to mafic rocks as well as massive aphyric rhyolite in the creek bed. Moving up the slope, these rocks are overlain by a coarse rhyolite breccia or debris flow. The breccia contains clasts ranging in size from nearly a meter to less than 1 cm. The breccia is clast-supported and some clasts are subrounded, whereas others are flowbanded. Overlying the breccia is a medium-grained, vesicular basalt flow. Capping the hill, and overlying the basalt, is an aphanitic massive rhyolite that is microspherulitic with some crackle breccia on the easternmost end. The following latitude and longitude readings are given as a guide to locating the different units in the Beaver Creek section. 1) 46.7409ºN -89.77678ºW: massive aphyric rhyolite in stream bed 2) 46.74067ºN -89.77821ºW: rhyolite breccia; some fragments flow-banded 3) 46.74037ºN -89.77795ºW: vesicular basalt flow 4) 46.73981ºN -89.77702ºW: massive to flow-banded microspherulitic rhyolite; eastern end is crackle breccia 5) 46.74067ºN -89.77821ºW: rhyolite breccia; some fragments flow-banded �Figure 7. Location and geologic setting for stops 3A and B. Geology generalized from Cannon and others (1995). Structures are dotted and contacts are thin solid lines. Box indicates the area of multiple units along the Beaver Creek Trail for stop 3B. Stop 4: Historic Nonesuch Mine Site: 46.760ºN -89.620ºW The Nonesuch Mine first opened in 1866, extracting finely disseminated native copper from sandstone and shale near the contact between the Copper Harbor Conglomerate and the Nonesuch Formation (Fig. 8). The mine went through a long history of openings and closings. In the 1880’s the prospects looked quite encouraging, with an operating stamp mill that eventually produced 110 tons of copper, transported by tram to a dock 5 miles away at Union Bay. Four separate shafts on either side of the river extended to a depth of about 460 feet (Butler and Burbank, 1929). A town site with around 300 people sprang up around the operation. However, the very fine-grained nature of the copper made �Figure 8. Location and geologic setting for stop 4. Geology generalized from Cannon and others (1995). Contacts are thin lines and faults are thick lines. extraction difficult. A final attempt with chemical leaching that was successful in small pilot trials proved to be unsuccessful on a large-scale, and the mine closed again late in 1884, with most of the machinery subsequently stripped from the site. Several unsuccessful efforts were made to reopen the mine, one in 1906 and another in 1912, after which the mine closed for good. Total production for the Nonesuch Mine is estimated at 389,000 pounds of copper from 1868-1885 (Butler and Burbank, 1929). The former town site is now listed as a ghost town, with old foundations and lilac and apple trees in a grassy field as the only evidence of the town today. The contact between the Copper Harbor Conglomerate and base of the Nonesuch Formation is well exposed in small rapids on the Little Iron River where beds dip about 30o to the east (Fig. 9A). The upper Copper Harbor is a fluvial sandstone, well cross-bedded, and contains lenses of conglomerate as much as 0.5 m thick with clasts as large as 15 cm. (Fig. 9B). Above a transition zone less than a meter thick, is thinly laminated shale and siltstone of the basal Nonesuch Formation. This is the horizon from which the ore was mined, although mineralization is not evident in this exposure. �A. Laminated gray shale at the base of the Nonesuch Formation a few meters above the contact with the Copper Harbor Conglomerate. B. Conglomerate lens in the uppermost Copper Harbor Conglomerate. The base of the Nonesuch is exposed at the base of the falls (far left). Figure 9. Nonesuch Formation and Copper Harbor Conglomerate exposed during low flow in the Little Iron River near at the Nonesuch mine site. Photographs by Bill Cannon. �Stop 5: Nonesuch Formation at Bonanza Falls: 46.8177ºN -89.5701ºW The most complete exposure of the Nonesuch Formation in the region is along the Big Iron River near Bonanza Falls, although access can be difficult and dangerous at times of high water (Fig. 10). The Nonesuch is exposed nearly continuously in a gently southeast-dipping section from just upstream of Bonanza Falls to the sharp bend in the river near the northeast corner of section 13 (Fig. 11A). A detailed measured section is presented by Suszek (1991). The exposed rocks total 226 m of section, which includes nearly the entire Nonesuch, although neither the upper nor lower contact is directly exposed. Figure 50. Location and geologic setting for stop 5. Geology generalized from Cannon and others (1995). Structures are dotted lines, contacts are thin lines, and faults are thick lines. �A. Exposure of the Nonesuch Formation in the Big Iron River at Bonanza Falls, looking upstream. Beds dip gently upstream. B. Contorted bedding in the Nonesuch at Bonanza Falls, probably generated by seismic liquefaction. Scale card is 85 cm wide. Figure 6. Nonesuch Formation exposed at Bonanza Falls during low flow in the Big Iron River. Photographs by Bill Cannon. �The Nonesuch Formation in the Big Iron River section is dominantly siltstone and fine-grained sandstone with minor shale. Many rocks have trough cross-bedding, symmetrical and asymmetrical ripples, rib and furrow structures, parting lineations, and soft sediment deformational features (Fig. 11B). The finer-grained rocks include well-laminated shale, which is most abundant lower in the section. The shaley units commonly have ball-and-pillow structures and calcareous concretions. The Nonesuch here displays coarsening-upward sequences at scales ranging from a few meters to the entire thickness of the unit. On a smaller scale, upwardly fining sequences are common in units from a few centimeters to a few meters thick. In the lower 10 m of the section, copper mineralization occurs as concentrations of chalcocite, bornite and malachite along bedding planes. The mineralization is cogenetic with the major copper mineralization in the White Pine mine, where the downdip extension of this unit was mined just to the south and east. A good exposure of the mineralized base of the Nonesuch Formation and the top of the Copper Harbor Conglomerate occurs along the Little Iron River near the center of the SW1/4, Section 13, but requires a walk of about 1 mi south from Highway 107. It is an easy walk along an unmaintained trail on the east bank for those who can spend more time in the area. At this location, remains of early mining efforts for native silver occurs there, as well as ‘ore’ specimens from old dumps. Copper was initially introduced to the lower Nonesuch Formation in both the White Pine and Copperwood districts during early diagenesis, probably by upward circulating connate waters which dissolved copper from the underlying redbeds of the Copper Harbor Formation (Swenson and Person, 2000). Chalcocite, largely of microscopic size, formed by the replacement of diagenetic pyrite. A later phase of copper mineralization documented by Mauk and others (1992) in the White Pine mine, introduced native copper mostly along fault zones and adjacent strata. This second stage of native copper mineralization commonly occurs as large thin plates of native copper developed along bedding planes in the Nonesuch. It is locally known as sheet copper and is probably cogenetic with the classic native copper mineralization of the Portage Lake basalts along the Keweenaw Peninsula. Stop 6: Upper part of Copper Harbor Conglomerate at Union Bay Campground: 46.8253ºN -89.6418ºW Good exposures of reddish sandstone containing thin conglomerate beds are abundant along the shore of Lake Superior in the park at the Union Bay Campground (Fig. 12). The Copper Harbor Conglomerate is typically characterized by coarse volcanogenic conglomerate, which forms most of the lower part of the section throughout much of its outcrop belt and which grades up into finer grained sandstone (Elmore, 1984). However, south of the village of Ontonagon, there is a different facies relationship. The lower part of the formation here is mostly sandstone, siltstone, and basalt or andesite lava flows; conglomerate is very subordinate. These rocks underlie the high hills immediately south of Highway 107. A coarse conglomerate facies occurs higher in the section, but forms less than 10 percent of the thickness. The exposures here at Union Bay are near the base of the upper unit and are probably about 1,000 m above the base of the formation. Sandstone layers at Union Bay dip 10-20° to the north. Sandstone is volcanogenic and quartz-poor. There are excellent examples of trough cross-bedding, generally indicating a northeastward current vector. A variety of other sedimentary features including �desiccation cracks, rip-up clasts, oscillation and current ripples, and swash marks are also present (Fig. 13A-D). Figure 7. Location and geologic setting for stop 6. Geology generalized from Cannon and others (1995). Structures are dotted and contacts are thin lines. The exposures of Copper Harbor Conglomerate north of the Porcupine Mountains are the farthest from the source highlands to the south. The relative scarcity of thick coarse conglomerate compared to exposures farther south probably reflects the distal nature of these rocks. These northernmost outcrops are a good representative of much of the Copper Harbor beneath Lake Superior. The rocks at Union Bay show a an irregular coarsening-upward trend as opposed to the fining-upward trend typical of the more proximal parts of the unit. This relationship is consistent with northward prograding alluvial plain deposition. Several large boulders are distributed along the beach and are composed of conglomerate typical of the lower part of the Copper Harbor elsewhere. In the boulders the conglomerate contains, almost exclusively, clasts of Keweenawan Supergroup volcanic rock types common in the region. An �A. Ripple marks B. Cross-bedding C. Mudcracks D. Lineations on bedding surface, possibly indicating current directions. Note variations in direction between various bedding planes. Figure 13. Sedimentary features in red siltstone and sandstone of the Copper Harbor Conglomerate along the Lake Superior shoreline at Union Bay Campground. Scale card is 85 cm in width. Photographs by Bill Cannon. �interesting question is the source of these boulders. Although the Copper Harbor does contain some conglomerate beds nearby (for example, Fig. 9B), none of the streams entering Lake Superior in this area seem capable of transporting such large boulders. Most streams, especially near the lakeshore, have low gradients, and streambeds do not contain such large boulders. These boulders are very likely glacial eratics that have been transported from some distance away. Ice movement was from the northeast, roughly parallel to the present shoreline. The boulders probably came from the Copper Harbor Conglomerate farther east toward the Keweenaw Peninsula where thick conglomerate is common. Stop 7: Basalt flows within the Copper Harbor Conglomerate (Lake of the Clouds overlook): 46.8031ºN -89.7641ºW Along Highway 107 leading to the Lake of the Clouds overlook are several exposures of conglomerate within the Copper Harbor Conglomerate. To the north is a good view of Lake Superior and the lowlands underlain by sedimentary rocks of the Oronto Group. From the overlook parking lot, a short hike leads to the overlook and a spectacular view of Lake of the Clouds and the Porcupine Mountains Wilderness State Park (Figs. 14 and 15). Figure 8. Location and geologic setting for stop 7. Geology generalized from Cannon and others (1995). Contacts are thin lines. �Figure 15. Panoramic view of Lake of the Clouds, Porcupine Mountains Wilderness State Park. Photograph by Bill Cannon. �The overlook is along the south escarpment of a high ridge supported by a series of northdipping lava flows within the Copper Harbor Conglomerate (Fig. 16). These flows are known as the Lake Shore traps. Comparable flows within the Copper Harbor Conglomerate at the tip of the Keweenaw Peninsula have an age of 1087.2 ± 1.6 Ma (Davis and Paces, 1990). The low area south of the ridge, including Lake of the Clouds, is underlain by sandstone and siltstone and a few basalt flows which constitute the lower part of the Copper Harbor Conglomerate. The higher regions farther south are underlain by volcanic rocks, mostly rhyolite of the Porcupine Volcanics. Toward the east end of the overlook area, a large glaciated surface shows a series of thin basalt flows, which average a few meters thick. Individual flows can be readily identified by chilled vesicular bases, in places containing inclusions of older flows, and by rubbly or vesicular tops. Abundant epidote alteration and vesicle fillings impart a distinctive greenish cast to flow margins (Fig. 16). Hubbard (1975) described the flows in the Copper Harbor as mostly andesite with minor basalt, but chemical analyses of two samples from this locality indicate that they are basalt, similar in composition to average basalt from the Porcupine Volcanics. An interesting feature of some of these flows is the incorporation of very fine-grained red sediments both in vesicles near their base and in thin fractures extending for a meter or more up into flows. Thin vestiges of these sediments are also along flow contacts. We interpret these sediments as wind-blown dust that accumulated on flow surfaces shortly after eruption. It was still unconsolidated when the next flow was erupted. The soft sediment was then injected upward by the weight of the overlying flow. Compared to Portage Lake Volcanics, these basalts are enriched in incompatible trace elements and show a distinct negative Nb-Ta anomaly in primitive mantle normalized incompatible trace element patterns similar to the basalts from the Porcupine Volcanics (Fig. 4B). This negative anomaly is a likey indication of crustal contribution to the parent magmas. Figure 16. Contact between two basalt flows at the Lake of the Clouds overlook. The vesicular top of one flow (bottom of photo) has amygdules filled with greenish epidote. The base of the overlying flow has pipe vesicles near the lower contact, grading up to a more massive interior (towards the top of the photo). Reddish fine-grained clastic material occurs in cracks in the upper flow. Scale card is 85 cm wide. Photograph by Bill Cannon. �Stop 8: Porcupine Volcanics: Rhyolite quarry near White Pine: 46.7155ºN -89.4435ºW The rocks exposed here are part of a small rhyolite body near the top of the Porcupine Volcanics (Fig. 17). The body probably does not extend much beyond the hill into which the quarry is cut. The quarry provides a cross section through part of a subaerial agglutinate deposit. Agglutinate deposits form at vents by spatter of erupting magma and buildup of mounds of hot, viscous material. The mound of erupted material eventually flows outward under its own weight, resulting in large flow folds such as seen in this quarry. The near-vent nature of this deposit is deduced from the presence of lithic fragments within the rhyolite, large and small contorted flow folds, and stratification of rhyolite (light and dark units) (Fig. 18A-C). At the edges of agglutinate deposits, flowage typically has homogenized the magma, and, as such, stratification and folding are generally not preserved. Zartman and others (1997) reported an age of 1093.6 +/- 1.8 Ma for rhyolite from this quarry. Figure 17. Location and geologic setting for stop 8. Geology generalized from Cannon and others (1995). Contacts are thin lines and faults are thick lines. �This rhyolite contains feldspar phenocrysts, which are typically aligned parallel to the stratification and foliation. Rhyolite of the Porcupine Volcanics is enriched substantially in such incompatible trace elements as Th, Ba, Zr, Hf, and LREE compared to rhyolite in the Portage Lake Volcanics on the Keweenaw Peninsula. Isotopically, rhyolite at this stop has an initial Nd isotopic signature (εNd(1100Ma) about -14) that reflects a substantial contribution from older crustal sources. The Portage Lake rhyolite from the Keweenawan Peninsula (εNd(1100Ma) about 0) is thought to be derived from partially melted early Keweenawan basalts (Nicholson and others, 1997). A. Lithic fragment showing flow contact between light and dark rhyolite. B. Flattened and drawn-out lithic fragments in rhyolite matrix. C. Contoured flow pattern showing contrasting rhyolite melts. Figure 18. Agglutinate textures in Porcupine Volcanics rhyolite, exposed in quarry blocks. Photographs by Bill Cannon. �References Allen, D.J., Braile, L.W., Hinze, W.J., and Mariano, J., 1995, The Midcontinent rift system, U.S.A.: a major Proterozoic continental rift, in Olsen, K.H. (ed.), Continental rifts: evolution, structure, tectonics: Elsevier, New York, pp. 375-407. Bornhorst, T.J., and Williams, W.C., 2011, The Copperwood sediment-hosted stratiform copper deposit, Upper Peninsula, Michigan, in Hollings, P., MacTavish, A., and Addison, W., (eds.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario: Part 1 – Program and abstracts, v. 58, p. 14. Bornhorst, T.J., Paces, J.B., Grant, N.K., Obradovich, J.D., and Huber, N.K., 1988, Age of native copper mineralization, Keweenaw Peninsula, Michigan: Economic Geology, v. 83, p. 619-625. Butler, B.S., and Burbank, W.S., 1929, The copper deposits of Michigan: U.S. Geological Survey Professional Paper 144, 238 p. Cannon, W.F., 1992, The Midcontinent rift in the Lake Superior region with emphasis on its geodynamic evolution: Tectonophysics, v. 213, p. 41-48. Cannon, W.F., 1994, Closing of the Midcontinent rift – A far-field event of Grenvillian compression: Geology, v. 22, p. 155-158. Cannon, W.F., and Nicholson, S.W., 1992, Revisions of stratigraphic nomenclature within the Keweenawan Supergroup of northern Michigan: U.S. Geological Survey Bulletin 1970A, p. A1-A8. Cannon, W.F., Peterman, Z.E., and Sims, P.K., 1993, Crustal-scale thrusting and origin of the Montreal River monocline – a 34-km-thick cross section of the Midcontinent rift in northern Michigan and Wisconsin: Tectonics, v. 12, p. 728-744. Cannon, W.F., Nicholson, S.W., Hedgman, C.A., Woodruff, L.G., and Schulz, K.J., 1992, Geology of Keweenawan Supergroup rocks near the Porcupine Mountains, Ontonagon and Gogebic Counties, Michigan, in Dickas, A.B., and Brown, B.A., (eds.), Institute on Lake Superior Geology Proceedings, 38th Annual Meeting, Hurley, Wisconsin: Part 2 – Field trip guidebook, v. 38, p. 77-101. Cannon, W.F., Nicholson, S.W., Woodruff, L.G., Hedgman, C.A., and Schulz, K.J., 1995, Geologic map of the Ontonagon and part of the Wakefiled 30′ x 60′ quadrangles, Michigan: U.S. Geological Survey Miscellaneous Investigations Series Map I-2499, scale 1:100,000. Daniels, P.A., Jr., 1982, Upper Precambrian sedimentary rocks: Oronto Group, Michigan-Wisconsin, in Wold, R.J. and Hinze, W.J. (eds.), Geology and tectonics of the Lake Superior basin: Geological Society of America Memoir 156, p. 107-133. Davis, D.W., and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system: Earth and Planetary Science Letters, v. 97, p. 54 -64. Elmore, R.D., 1984, The Copper Harbor Conglomerate: a late Precambrian fining-upward alluvial fan sequence in northern Michigan: Geological Society of America Bulletin 95, p. 610-617. Elmore, R.D., Milavec, G., Imbus, S., Engel, M.H., and Daniels, P., 1988, The Precambrian Nonesuch Formation of the North American Midcontinent Rift: Sedimentary and organic geochemical aspects of lacustrine deposition: Precambrian Research, v. 43, p. 191-213. �Easton, M., Hart, T.R., Hollings, P., Heamon, L.M., MacDonald, C.A., and Smyk, M., 2007, Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon Region, Ontario: Canadian Journal of Earth Sciences, v. 44, p. 1055-1086. Green, J.C., 1982, Geology of Keweenawan extrusive rocks, in Wold, R.J., and Hinze, W.J. (eds.), Geology and tectonics of the Lake Superior basin: Geological Society of America Memoir 156, p. 4755. Hinze, W.J., Braile, L.W., and Chandler, V.W., 1990, A geophysical profile of the southern margin of the Midcontinent Rift system in western Lake Superior: Tectonics, v. 9, p. 303-310. Hubbard, H.A., 1975, Geology of the Porcupine Mountains in Carp River and White Pine quadrangles, Michigan: U.S. Geological Survey Journal of Research, v. 3, p. 519-528. Hutchinson, D.R., White, R.W., Cannon, W.F., and Schulz, K.J., 1990, Keweenaw hot spot: geophysical evidence for a 1.1 Ga mantle plume beneath the Midcontinent Rift System: Journal of Geophysical Research, v. 95, p. 10869-10884. King, E.R., 1987, Aeromagnetic map of the Iron River 1º x 2º quadrangle, Michigan and Wisconsin: U.S. Geological Survey Miscellaneous Investigations Map I-1306F, scale 1:250,000. Klasner, J.S., 1989, Bouger gravity anomaly map and geologic interpretation of the Iron River 1º x 2º quadrangle, Michigan and Wisconsin: U.S. Geological Survey Miscellaneous Investigations Map I1306E, scale 1:250,000. Mauk J.L, Kelly, W.C., van der Pluijm, B.A., and Seasor, R.W., 1992, Relations between deformation and sediment-hosted copper mineralization: Evidence from the White Pine part of the Midcontinent rift system: Geology, v. 20, p. 427-430. Miller, J.D., and Chandler, V.W., 1997, Geology, petrology, and tectonic significance of the Beaver Bay Complex, northeastern Minnesota: Geological Society of America Special Papers 312, p. 73-96. Nicholson, S.W., and Shirey, S.B., 1990, Evidence for a Precambrian mantle plume: a Sr, Nd, and Pb isotopic study of the Midcontinent Rift System in the Lake Superior region: Journal of Geophysical Research, v. 95, p. 10851-10868. Nicholson, S.W., Shirey, S.B., Schulz, K.J., and Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504-520. Ojakangas, R.W., and Morey, G.B., 1982, Keweenawan pre-volcanic quartz sandstones and related rocks of the Lake Superior region, in Wold, R.J., and Hinze, W.J. (eds.), Geology and tectonics of the Lake Superior basin: Geological Society of America Memoir 156, p. 85-96. Paces, J.B., 1988, Magmatic processes, evolution and mantle source characteristics contributing to the petrogenesis of Midcontinent rift basalts: Portage Lake basalts, Keweenaw Peninsula, Michigan: unpublished Ph.D. dissertation, Michigan Technological University, Houghton, Michigan. Paces, J.B., and Bell, K., 1989, Non-depleted sub-continental mantle beneath the Superior Province of the Canadian Shield: Nd-Sr isotopic and trace element evidence from Midcontinent rift basalts: Geochimica et Cosmochimica Acta, v. 53, p. 2023-2035. Ruiz, J., Jones, L.M., and Kelly, W.C., 1984, Rubidium-strontium dating of ore deposits hosted by Rbrich rocks using calcite and other common Sr-bearing minerals: Geology, v. 12, p. 259-262. �Suszek, T.J., 1991, Petrography and sedimentation of the Middle Proterozoic (Keweenawan) Nonesuch Formation, western Lake Superior region, Midcontinent Rift System: unpublished M.S. thesis, University of Minnesota, Duluth, Minnesota. Suszek, T.J., 1997, Petrography and sedimentation of the middle Proterozoic (Keweenawan) Nonesuch Formation, western Lake Superior region, Midcontinent Rift System: Geological Society of America Special Paper 312, p. 195- 210. Swenson, J.B., and Person, M., 2000, The role of basin-scale transgression and sediment compaction in stratiform copper mineralization: implications from White Pine, Michigan, USA: Journal of Geochemical Exploration, v. 69-70, p. 239-243. White, W.S., 1972, The base of the upper Keweenawan, Michigan and Wisconsin: U.S. Geological Survey Bulletin 1354-F, p. F1-F23. White, W.S., and Wright, J.C., 1954, The White Pine copper deposit, Ontonagon County, Michigan: Economic Geology, v. 49, p. 675-716. Zartman, R.E., Nicholson, S.W., Cannon, W.F., and Morey, G.B., 1997, U-Th-Pb zircon ages of some Keweenawan Supergroup rocks from the south shore of Lake Superior: Canadian Journal of Earth Sciences, v. 34, p. 549-161. � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2013 Description An account of the resource Institute on Lake Superior Geology. Houghton, Michigan. May 8-11, 2013. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2013 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/3bcf982a44576e47069c02b68e895c56.pdf 618d1ae16e9ef5c5a6bfbd10589516cd PDF Text Text �Institute on Lake Superior Geology 60TH ANNUAL MEETING May 14-17, 2014 Hibbing, Minnesota Sponsored by PRECAMBRIAN RESEARCH CENTER, UNIVERSITY OF MINNESOTA DULUTH and MINNESOTA GEOLOGICAL SURVEY James D. Miller and Mark A. Jirsa Co-Chairs Proceedings Volume 60 Part 1 – Program and Abstracts Edited by Jim Miller, University of Minnesota Duluth Cover Photo Credit View of mines and the city of Hibbing looking south. Gray area in foreground is the footprint of Hibbing Taconite’s mining; partially flooded, dark red areas in the mid-ground are remnants of historic natural (hematite) ore mines, including the Hull-Rust, Mahoning, Susquehanna, and Scranton; City of Hibbing in background, showing location of meeting hotel (oval). Modified from image provided by Dave Witt—Aero-Environmental Consulting, LLC, Cook, MN i �Table of Contents Institutes on Lake Superior Geology, 1955-2014 iii Sam Goldich and the Goldich Medal vi Goldich Medal Guidelines viii Goldich Medalists and Goldich Medal Committee x Citation for Goldich Medal Award to Laurel Woodruff xi Memorial to Ernest Lehmann xiii Memorial to Jack Everett xiv Eisenbrey Student Travel Awards xv Joe Mancuso Student Research Awards xvi Doug Duskin Student Paper Awards and Award Committee xvii Board of Directors, Local Committee, and Banquet Speaker xviii Session Chairs and Field Trip Leaders xix Corporate and Individual Sponsors of Student Travel Scholarships xxi Report of the Chair of the 59th Annual Meeting xxii Program xxiv Poster Presentations xxix Abstracts 1-130 Reference to material in Part 1 should follow the example below: Field trip authors, date, title: Institute on Lake Superior Geology Proceedings v. 60, Part 1, p. XX. Proceedings Volume 60, Part 1—Program and Abstracts, and Part 2—Field Trip Guidebook are published by the 60th Institute on Lake Superior Geology and distributed by the Institute Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca Some figures in this volume were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the volume when it is available on-line at http://www.lakesuperiorgeology.org. ISSN 1042-99 ii �Institutes on Lake Superior Geology, 1955-2014 95 o o 85 o Wabigoon subprovince90 o 80 48 o Wawa-Abitibi subprovince 48o Wawa-Abitibi subprovince o 45 45o Minnesota River Valley subprovince MEETING LOCATIONS Phanerozoic Mesoproterozoic Map by Mark Jirsa Paleoproterozoic o 90 95o 85o Archean Superior Province # Date Place Chairs 1 1955 Minneapolis, Minnesota C.E. Dutton 2 1956 Houghton, Michigan A.K. Snelgrove 3 1957 East Lansing, Michigan B.T. Sandefur 4 1958 Duluth, Minnesota R.W. Marsden 5 1959 Minneapolis, Minnesota G.M. Schwartz & C. Craddock 6 1960 Madison, Wisconsin E.N. Cameron 7 1961 Port Arthur, Ontario E.G. Pye 8 1962 Houghton, Michigan A.K. Snelgrove 9 1963 Duluth, Minnesota H. Lepp 10 1964 Ishpeming, Michigan A.T. Broderick 11 1965 St. Paul, Minnesota P.K. Sims & R.K. Hogberg 12 1966 Sault Ste. Marie, Michigan R.W. White 13 1967 East Lansing, Michigan W.J. Hinze 14 1968 Superior, Wisconsin A.B. Dickas 15 1969 Oshkosh, Wisconsin G.L. LaBerge 16 1970 Thunder Bay, Ontario M.W. Bartley & E. Mercy iii �# Date Place Chairs 17 1971 Duluth, Minnesota D.M. Davidson 18 1972 Houghton, Michigan J. Kalliokoski 19 1973 Madison, Wisconsin M.E. Ostrom 20 1974 Sault Ste. Marie, Ontario P.E. Giblin 21 1975 Marquette, Michigan J.D. Hughes 22 1976 St. Paul, Minnesota M. Walton 23 1977 Thunder Bay, Ontario M.M. Kehlenbeck 24 1978 Milwaukee, Wisconsin G. Mursky 25 1979 Duluth, Minnesota D.M. Davidson 26 1980 Eau Claire, Wisconsin P.E. Myers 27 1981 East Lansing, Michigan W.C. Cambray 28 1982 International Falls, Minnesota D.L. Southwick 29 1983 Houghton, Michigan T.J. Bornhorst 30 1984 Wausau, Wisconsin G.L. LaBerge 31 1985 Kenora, Ontario C.E. Blackburn 32 1986 Wisconsin Rapids, Wisconsin J.K. Greenberg 33 1987 Wawa, Ontario E.D. Frey & R.P. Sage 34 1988 Marquette, Michigan J. S. Klasner 35 1989 Duluth, Minnesota J.C. Green 36 1990 Thunder Bay, Ontario M.M. Kehlenbeck 37 1991 Eau Claire, Wisconsin P.E. Myers 38 1992 Hurley, Wisconsin A.B. Dickas 39 1993 Eveleth, Minnesota D.L. Southwick 40 1994 Houghton, Michigan T.J. Bornhorst 41 1995 Marathon, Ontario M.C. Smyk 42 1996 Cable, Wisconsin L.G. Woodruff 43 1997 Sudbury, Ontario R.P. Sage & W. Meyer 44 1998 Minneapolis, Minnesota J.D. Miller & M.A. Jirsa 45 1999 Marquette, Michigan T.J. Bornhorst & R.S. Regis 46 2000 Thunder Bay, Ontario S.A. Kissin & P. Fralick 47 2001 Madison, Wisconsin M.G. Mudrey & Jr., B.A. Brown 48 2002 Kenora, Ontario P. Hinz & R.C. Beard 49 2003 Iron Mountain, Michigan L. Woodruff & W.F. Cannon iv �# Date Place Chairs 50 2004 Duluth, Minnesota S. Hauck & M. Severson 51 2005 Nipigon, Ontario M. Smyk & P. Hollings 52 2006 Sault Ste. Marie, Ontario A. Wilson & R. Sage 53 2007 Lutsen, Minnesota L. Woodruff & J. Miller 54 2008 Marquette, Michigan T. Bornhorst & J. Klasner 55 2009 Ely, Minnesota J. Miller, G. Hudak, & D. Peterson 56 2010 International Falls, Minnesota M. Jirsa, P. Hollings, & T. Boerboom, P. Hinz & M.Smyk 57 2011 Ashland, Wisconsin T. Fitz 58 2012 Thunder Bay, Ontario P. Hollings 59 2013 Houghton, Michigan T. Bornhorst & A. Blaske 60 2014 Hibbing, Minnesota J. Miller & M. Jirsa v �Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970’s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 vi �INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL vii �Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. viii �Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. ix �Goldich Medalists 1979 Samuel S. Goldich 1997 Ronald P. Sage 1980 not awarded 1998 Zell Peterman 1981 Carl E. Dutton, Jr. 1999 Tsu-Ming Han 1982 Ralph W. Marsden 2000 John C. Green 1983 Burton Boyum 2001 John S. Klasner 1984 Richard W. Ojakangas 2002 Ernest K. Lehmann 1985 Paul K. Sims 2003 Klaus J. Schulz 1986 G.B. Morey 2004 Paul Weiblen 1987 Henry H. Halls 2005 Mark Smyk 1988 Walter S. White 2006 Michael G. Mudrey 1989 Jorma Kalliokoski 2007 Joseph Mancuso 1990 Kenneth C. Card 2008 Theodore J. Bornhorst 1991 William Hinze 2009 L. Gordon Medaris, Jr 1992 William F. Cannon 2010 William D. Addison & Gregory R. Brumpton 1993 Donald W. Davis 2011 Dean M. Rossell 1994 Cedric Iverson 2012 James D. Miller 1995 Gene La Berge 2013 Tom Waggoner 1996 David L. Southwick 2014 GOLDICH MEDAL RECIPIENT Laurel Woodruff Goldich Medal Committee Serving through the meeting year shown in parentheses. Graham Wilson (2014) Turnstone Consulting Bernhardt Saini-Eidukat (2015) North Dakota State University Mark Smyk (2016) Ontario Geological Survey x �Citation for the Goldich Medal Award to Laurel G. Woodruff It is my pleasure and honor to present the 2014 Goldich Medal to Laurel G. Woodruff. Laurel has been one of the most active and involved members of the Institute for more than 20 years. During that time she has chaired or co-chaired three annual meetings (47th, 49th, 53rd) and served corresponding terms on the board of directors. She was chair of the board of directors in 1995-1996, 2002-2003, and 2006-2007. She has served twice on the student paper award committee, and most recently, from 2010-2013, was a member of the Goldich Award committee, and chaired the committee in 2012-2013. In Laurel in the Brooks Range, Alaska in addition, she has been co-leader of three 2007 during the Alaska Soil Institute fieldtrips and has made numerous Geochemistry Transect. technical presentations at Institute meetings. In case no one has yet noted, this is the first Goldich citation in which the pronoun “she” has been used. Most of Laurel’s career, spanning more than thirty years, has been with the USGS mineral resources research program, with more than 25 of those years in the Lake Superior region. Prior to that Laurel received her formal education at University of Michigan (BS in Geology 1973), Michigan Technological University (MS in Geology 1977), and the University of Chicago (PhD in geology 1989). After completing her MS degree and beginning a PhD at Chicago, Laurel was hired to run the light stable isotope laboratory at the University of Wisconsin-Madison and she participated in a broad variety of stable isotope research. The highlight of this part of her career was research on modern seafloor hydrothermal deposits, which culminated in publication of her Journal of Geophysical Research paper on the stable isotope geochemistry of seafloor hydrothermal vent systems. Laurel joined the USGS in 1983. Her initial assignment was establishing the light stable isotope laboratory in the Branch of Eastern Minerals Resources. When the laboratory became operational, she was responsible for light stable isotope analyses (S, O, and C) of rocks, ores, and mineral samples from a number of locations throughout the world in support of research on seafloor sulfide formation, and precious metal mineralization. In 1986-87 Laurel returned to the University of Chicago to complete her PhD and conducted her dissertation research on diabases of the eastern U.S. Triassic basins as part of a large USGS project on the mineral potential of xi �those basins. Laurel’s scientific contributions to the geology of the Lake Superior region began in the late 1980’s when she became a member of a USGS project team studying the geology and mineral potential of the Midcontinent Rift in Michigan and Wisconsin. Laurel’s contributions included: 1) field work to collect bedrock and mineral deposit samples, 2) preparation of geologic maps and reports, 3) stable isotope analyses to constrain metal sources and characterize regional alteration patterns, and 4) geochemical and 2-D thermal modeling to better understand the origin and distribution of copper mineralization in the rift In the past decade, Laurel has become increasingly involved in environmental research and has been a leader in fostering the incorporation of geology and geochemistry into multidisciplinary studies of the behavior of elements of concern such as arsenic and mercury. Her studies of the effect of forest fires on the mercury content of soils in the Lake Superior region and the cycling of mercury in aquatic ecosystems, conducted in cooperation with colleagues in soil science, hydrogeochemistry, and aquatic biology have provided fundamental new understanding of the mercury cycle. Laurel also has been a key figure in establishing procedures for and conducting geochemical baseline studies from local to national scale. The recently completed soil geochemical survey of the conterminous U.S. has produced a new database and a national geochemical atlas based on 15,000 samples from about 5,000 sites across the country. Laurel was a key member of that project from the earliest planning phase, through pilot studies, and the full survey, to the current activities of producing interpretive research papers. Laurel also continues her research on the Precambrian geology and resources of the Lake Superior region and is the coordinator of a new USGS multidisciplinary project on metallogeny and mineral potential of the St. Croix horst in Wisconsin and Minnesota. On a personal note, Laurel has been a great friend and colleague for more than 25 years as we have wended our way through a kaleidoscope of research from hard rocks, through glacial deposits, and soils, to lake-bottom muck; wanderings across Alaska, the “death marches” on Isle Royale, and hard days of canoeing and portaging through the Boundary Waters and Voyageurs Park. Many of the most vivid and pleasant (at least in hindsight) memories of my career are from those days. Our research has commonly been guided by Laurel’s often expressed philosophy of “Let’s do something even if it’s wrong!” She’s shown over and over through her proclivity for action and her eagerness to plunge into new work, that it is so much easier to make mid-course corrections of something in progress than it is to overcome the inertia of over planning, indecision, and inaction; an attribute that has been unfailingly valuable in so much of what she has done during her career. So, in recognition of her decades of accomplishments and of her dedication to the geology of the Lake Superior region and to the Institute on Lake Superior Geology it is my pleasure to present the 2014 Goldich Medal to Laurel. Bill Cannon, Geologist Emeritus U.S. Geological Survey xii �In Memoriam Ernest K. Lehmann (1929-2013) On December 13, 2013, the Institute on Lake Superior Geology lost one of its industry giants with the passing Ernest (Ernie) K. Lehmann. Ernie was an exploration geologist whose lifelong work in the mining industry took him around the globe. He was awarded the ILSG’s Goldich Medal in 2002 for his pioneering contributions to base and precious metal exploration in the Lake Superior region, especially in Minnesota and Wisconsin. He tirelessly contributed his time and talents to professional organizations such as SME and AIPG, mining advocacy groups such as Mining Minnesota, and minerals outreach programs such as the Minnesota Minerals Education Workshops. Ernie was admired by family, friends and colleagues for his honesty, integrity and perseverance and his ability to tell a tale. He was a quietly generous and caring man who will be greatly missed. Ernie was born in Heidelberg, Germany and emigrated with his parents to the United States in 1935. He was educated in the public primary and secondary schools of New Rochelle, N.Y. and graduated from Williams College, Mass. in 1951 with highest honors in geology. In 1951-52, he did graduate study in geology at Brown University and in 1984 completed the Owners and Presidents Management Program at the Harvard Business School. His career in the mining industry began in 1950 when he worked as a miner and then geologist at a gold mine in Bannack, Montana and then joined Kennecott Copper’s exploration subsidiary, where he was head of a team that discovered the south end of the “New Lead Belt” in Missouri in the 1950s. With the consulting firm he founded in 1958, he undertook a variety of successful exploration projects, including industrial limestone; gold in Montana, the Northwest Territories and Argentina and copper-nickel-platinum group metals in Minnesota. He managed a small fluorspar mine in southern France and undertook valuation of various mining projects in Africa, Indonesia, South and Central America and North America for IFC (International Finance Corporation), the World Bank, major mining companies and metal trading companies. He served as president of North Central Mineral Ventures (NCMV) since its incorporation in February 1986. NCMV has served as Manager of Vermillion Gold since December 2007. Ernie served as a member of the Advisory Board to the Natural Resources Research Institute of the University of Minnesota and the MGS (Minnesota Geological Survey) State Mapping Advisory Committee. He served as a member of the Governor’s Committee of Minnesota’s Mining Future from 2004 to 2007 and as a member of the Minnesota Legislature’s Mineral Coordinating Committee for over 10 years. He was an officer or director of several other private companies, including a Director of Silverthorn Exploration Inc. He was a charter and honorary member of the AIPG, of which he served as national President in 1985; a life member and fellow of the Society of Economic Geologists; a Legion of Honor member of the Society of Mining Engineers; and a member of several other professional and technical societies. Ernie was president of Mining Minnesota, a trade association representing Minnesota’s non-ferrous and precious metals mining industry. xiii �In Memoriam Jack V. Everett (1921-2013) The ILSG lost another long-time supporter with the passing of Jack V. Everett, who slipped away peacefully on August 12, 2013 at his summer home on Ottertail Lake, MN. Jack will be most remembered for his sitting in the front row of ILSG meetings and snapping pictures of slide presentations. Jack lived in Duluth, MN for most of his professional career working as a consulting mining geologist. Jack was born in Roseburg, Oregon, but spent most of his childhood in lower Michigan. He enrolled at Michigan State in the class of 1944 in wildlife management, conservation and zoology, but later chose to major in geology. World War II interrupted his studies and he enlisted on June 6, 1942 in ROTC in field artillery with basic training at Fort Bragg, NC, and was called for active duty on April 16, 1943. He married Eleanor Brown Everett, class of ’44, from Onaway, Michigan at that time. The Army needed infantry officers and on May 6, 1944 sent his entire class to be retained as infantry officers at Fort Benning, GA. On July 18, 1944 he was assigned as a cadre training officer at Fort Meade, MD. After the Japanese surrendered, on September 3, 1945 he received orders to be transferred to Japan and was assigned to serve in the occupation forces of the 77th Division in Hakodate on the northern island of Hokkaido. He was discharged out of the service on September 5, 1946. He went back to MSU and graduated in 1947 with a B.S. degree, cum laude, in Geology. Honors included Phi Kappa Phi for scholastic, Sigma Gamma Epsilon for geologic, and Tau Sigma for scientific. Jack later served in various U. S. Army Reserve and Minnesota National Guard units in Brainerd and Duluth, and retired as a Major in June of 1972. Jack’s professional career started when he was hired as a District Geologist for Pickands Mather & Co. on the Minnesota Cuyuna Iron Range where he discovered four iron – manganese deposits near Emily, MN. These deposits are currently being developed for their manganese ore potential. In 1951 he took a position with W.S. Moore Company as Chief Geologist & Exploration Manager and moved to Duluth, MN. During the 50s and 60s he conducted exploration programs for iron ore deposits in various locations across the United States and Canada, and also Parana, Brazil. In the 1960s he conducted major prospecting programs in unmapped areas of the Northwest Territories exploring for gold deposits from Yellowknife to the Arctic coast and was quoted as saying that for 20 years, he spent 50% of his life living in tents. He also conducted exploration programs in Northern MN and discovered one major copper nickel deposit. Jack started a career as a Certified Professional Geologist in 1971 and worked for more than 100 US and Canadian mining companies as an independent consulting geologist. He conducted exploration programs for copper and gold deposits in Wisconsin. Jack was an avid hunter and fisherman. Although he supported mining, he was a conservationist and supported preservation of unique natural resources and was on the Governor Elmer L. Andersen committee as a consulting geologist and first chair of the Duluth Chapter of the Citizens Committee for the establishment of Voyageurs National Park. In the 1980s he explored and developed placer gold deposits in Alaska. In later years he worked on a variety of geology, geotechnical and hydrology projects, including the tunnel projects on the North Shore of Lake Superior for the MnDOT. In 1995 he became Vice PresidentExploration & Director of Leadville Mining and Milling Corp. where he was involved with the development, geology and exploration of their underground gold mine near Leadville, CO. More recently he worked on the geology and development of El Chanate Gold Project in Sonora, Mexico as a Director for Capital Gold Corporation. And most recently he has been working on the geology and development of Lake Victoria Gold deposit in Tanzania, Africa. He always joked that he planned to retire soon. xiv �Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xv �Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive onehalf of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. In 2013, the ILSG Board of Governors awarded three $500 awards from the Student Research Fund. The winners were: Michael Doyle University of Minnesota-Duluth, Department of Geological Sciences Current degree program: MS Candidate (Advisor: Jim Miller) Geologic and Geochemical Attributes of the Beaver River Diabase and Greenstone Flow: Testing a Possible Intrusive-Volcanic Correlation in the 1.1 Ga Midcontinent Rift Sarah Sauer University of Minnesota-Duluth, Department of Geological Sciences Current degree program: MS Candidate (Advisor: Jim Miller) The Petrology of the DLS “Chill” – Evidence of Venting of Hydrous Magma from the Layered Series at Duluth? Nicholas Fedorchuk University of Wisconsin-Milwaukee, Department of Geosciences Current degree program: MS Candidate (Advisor: John Isbell) Biogenicity of Mesoproterozoic Lacustrine Stromatolites from the Copper Harbor Conglomerate xvi �Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and long-time friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. Student Paper Awards Committee Andrew Ware – PolyMet Mining Prajukti Bhattacharyya – University of Wisconsin-Whitewater Robert Cundari – Ontario Geological Survey xvii �Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Jim Miller (2014-2017) – University of Minnesota Duluth Allan Blaske (2013-2016) – AECOM Peter Hinz (2012-2015) – Ontario Geological Survey Tom Fitz (2011-2014) – Northland College Pete Hollings - Secretary (2013-2016) – Lakehead University Mark Jirsa – Treasurer (2011-2014) – Minnesota Geological Survey Local Committee Chairs Jim Miller – Program Chair Department of Geological Sciences and Precambrian Research Center University of Minnesota Duluth Mark Jirsa – Field Trip Chair Minnesota Geological Survey University of Minnesota Volume Editors Jim Miller – Proceedings Volume Department of Geological Sciences and Precambrian Research Center University of Minnesota Duluth Mark Jirsa and Terry Boerboom – Field Trip Guidebook Minnesota Geological Survey, University of Minnesota Special Projects Amy Radakovich – Minnesota Geological Survey, University of Minnesota Banquet Speaker Dr. Francis M. Carroll University of Manitoba - Winnipeg and St. Johns University "A Line in the Trees: History of the US-Canadian Boundary from Lake Superior to Lake of the Woods" xviii �Session Chairs Al MacTavish – Panoramic Resources, Thunder Bay, ON Joyashish Thakurta – Western Michigan University, Kalamazoo, MI Geoff Pignotta – University of Wisconsin – Eau Claire Marcia Bjornerud – Lawrence University, Appleton, WI Mary Louise Hill – Lakehead University, Thunder Bay, ON Bernie Saini-Eidukat – North Dakota State University, Fargo, ND Field Trip Leaders Field trips have been the mainstay of the ILSG since its inception 60 years ago. We want to give a special thanks to the field trip leaders who volunteered their time and talent in carrying that tradition forward. 1) STRATIGRAPHY, SEDIMENTOLOGY, STRUCTURE AND MINERALIZATION OF THE BIWABIK IRON FORMATION, CENTRAL MESABI IRON RANGE Phil Larson - Duluth Metals Ltd. Marsha Patelke - Natural Resources Research Institute, UMD Jakob Wartman - United Taconite, Cliffs Natural Resources Michael Totenhagen - Arcelor Mittal Mark Jirsa - Minnesota Geological Survey Steven Losh - Minnesota State University – Mankato Peter K. Jongewaard - Cliffs Natural Resources (retired) 2) A WALK IN THE PARK—NEOARCHEAN GEOLOGY OF LAKE VERMILION STATE PARK George J. Hudak - Natural Resources Research Institute – UMD Amy Radakovich - Minnesota Geological Survey Geoff Pignotta - University of Wisconsin - Eau Claire Kelly Schwierske - University of Wisconsin - Eau Claire 3) WESTERN MESABI RANGE MINING OPERATIONS Douglas Halverson - Cliffs Natural Resources, Duluth Daniel Cervin - Cliffs Natural Resources, Hibbing Taconite William Everett – Essar Steel Kevin Kangas - Essar Steel Joseph Nielsen - Magnetation xix �5) VISIONS OF MATURI: THE GEOLOGY OF THE SOUTH KAWISHIWI INTRUSION Dean Peterson - Duluth Metals Ltd. 6) THE ST. LOUIS SUBLOBE AND GLACIAL LAKE UPHAM Phil Larson - Duluth Metals Ltd. Alan Knaeble - Minnesota Geological Survey Howard Mooers - University of Minnesota Duluth Lisa Marlo - Halcon Resources Corporation 7) GEOLOGY AND GOLD MINERALIZATION OF THE VIRGINIA HORN AREA Mark Jirsa - Minnesota Geological Survey William Rowell - Vermillion Gold LLC Richard Sandri - Vermillion Gold LLC Jason Richter - Minnesota Department of Transportation A) STATE DRILL CORE LIBRARY—HIBBING MINNESOTA Minnesota Department of Natural Resources—Division of Lands and Minerals Dave Dahl – Minnesota Department of Natural Resources, Div. of Lands and Minerals Dean Rossell - Kennecott Exploration, Rio Tinto B) HIBBING’S IRON MINING AND CULTURAL HISTORY Henry Djerlev - Superior GEO-Services (retired) Bob Kearney – Hibbing High School (retired) Erica Larson and other Hibbing Historical Society staff C) MINNESOTA DISCOVERY CENTER, CHISHOLM, MN Discovery Center Staff D) COLERAINE MINERALS RESEARCH LABORATORY Natural Resources Research Institute, University of Minnesota-Duluth Dick Kiesel - CMRL Director Dave Hendrickson - Director Strategic Planning Matt Mlinar - Program Coordinator Mineral Processing Basak Anameric - Program Coordinator High Temperature Process) E) MINEVIEW FROM A CANOE : Mark Jirsa - Minnesota Geological Survey Daniel Jordan - Iron Range Resources and Rehabilitation Board Dale Cartwright - Minnesota Dept. of Natural Resources, Div. of Lands and Minerals xx �Sponsors The following organizations and individuals made general contributions to the 60th Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. All of the funds contributed this year go toward travel awards for student registrants. Midwest Institute of Geosciences and Engineering INDIVIDUAL CONTRIBUTORS TO STUDENT TRAVEL SCHOLARSHIPS MARY ARTHUR JOHN BERKLEY KARL EVERETT JOHN GREEN GEORGE HUDAK PETER JONGEWAARD STEVEN LOSH ALLAN MACTAVISH GORDON MEDARIS, JR. MICHAEL MUDREY JILL PETERMAN With an especially generous donation provided by RON SEAVOY xxi �Report of the Chairs of the 59th Annual Meeting Theodore J. Bornhorst and Allan R. Blaske Houghton, Michigan The 59th annual meeting of the Institute on Lake Superior Geology (ILSG) was held May 8 to 11, 2013 in Houghton, Michigan, at the Franklin Square Inn. The meeting was hosted by the A. E. Seaman Mineral Museum of Michigan Technological University and was chaired and organized by Ted Bornhorst (A. E. Seaman Mineral Museum) and Allan Blaske (AECOM). The meeting was attended by a total of 228 delegates from 14 U.S. states (Arizona, Colorado, Illinois, Indiana, Iowa, Massachusetts Michigan, Minnesota, New York, North Dakota, Ohio, Texas, Virginia, Wisconsin) and 4 Canadian provinces (British Columbia, Ontario, Manitoba, Quebec). There were 58 student attendees. The two-day technical session began on Thursday morning with oral presentations on Archean topics and continued on Friday with presentations on Keweenawan and Quaternary geology. There were a total of 25 oral presentations, 10 of which were presented by students. The technical session included a total of 18 poster presentations, 10 of which were presented by students. The meeting offered 5 field trips that highlighted the Keweenawan geology of the western Upper Peninsula of Michigan. Two pre-meeting trips were held on Wednesday: Geologic Overview of the Keweenaw Peninsula, Michigan, led by Ted Bornhorst (A. E Seaman Mineral Museum) and Caledonia Mine, Keweenaw Peninsula Native Copper District, Ontonagon County, MI, led by Bob Barron (Michigan Tech) and Richard Whiteman (Red Metal Minerals). A third scheduled pre-meeting field trip was cancelled because of the unusual lingering snowpack which prevented access to Silver Mountain. Friday afternoon featured a “field trip” open house at the A. E. Seaman Mineral Museum, with guided tours led by Museum Director, Ted Bornhorst. Two post-meeting trips were held on Saturday: Geology of the Keweenawan Supergroup, Porcupine Mountains, Ontonagon and Gogebic Counties, MI, led Laurel Woodruff (USGS), Bill Cannon (USGS), and Robert Wild (Porcupine Mountains Wilderness State Park) and Geology and Environmental Site Conditions of the Copperwood Deposit, Gogebic County, MI, led by Ted Bornhorst (A. E. Seaman Mineral Museum), Allan Blaske (AECOM), Dave Anderson (Orvana Resources US Corp) and Tom Repaal (Orvana Resources US Corp.). The field trips were well-attended with each being at maximum capacity (sold out). The unusual snow cover in the Keweenaw and cold weather on Saturday made some shuffling of field trip logistics necessary. Four Doug Duskin Best Student Paper Awards were given for student paper presentations. Awards were presented for oral and poster presentations, with an award within each category for undergraduate students and for graduate students. The student awardees were Breanne Beh (Lakehead University, graduate student) and Emily Smyk (Lakehead University, undergraduate student) for their oral presentation and Jonathan Dyess (University of Minnesota – Duluth, graduate student) and Brynley Nadziejka (Lawrence University, undergraduate student) for their poster presentation. Eisenbray Student Travel Awards are funded by ILSG and forty students received a travel award. Thanks to very generous support from ILSG corporate sponsors (AECOM, Coleman Engineering Company, Rio Tinto–Eagle Mine, and Superior Copper Corporation) we were able to award ILSG Corporate Student Registration Awards to all students who attended the meeting. The ILSG Corporate Student Registration Award consisted of the meeting registration fee. In addition, students who were presenting papers received additional monetary support in the form of an ILSG Corporate Presenter Award. We are pleased to report that $5,325 were awarded to students. Through the support of corporate sponsors, the ILSG can better promote geologic studies of the Lake Superior region to the next generation of professional geoscientists. The ILSG social and banquet were held at the Franklin Square Inn, Houghton. There were 140 people at the banquet. Jim Ashley of the Lunar Reconnaissance Orbiter Camera Science Operations xxii �Center (LROC) at Arizona State University delivered the banquet address, entitled “Rusty Metal at the Martian Equator: The Search for Life on the Red Planet,” which discussed the occurrence and weathering of iron meteorites on the surface of Mars, and the implications these samples have to the history of water on Mars. The highlight of the banquet was the awarding of the 2013 Goldich Medal to Tom Waggoner (retired chief geologist and lands manager for Cleveland-Cliffs and currently a consulting geologist). Ron Seavoy provided a brief summary of Tom’s contributions to the geology of the Lake Superior region and the ILSG. Tom was greeted with warm applause upon receiving the prestigious ILSG award. The Institute’s Board of Directors met on May 9 to discuss the business of ILSG. The meeting was attended by Ted Bornhorst (Board of Directors meeting Chair), Allan Blaske, Al MacTavish, Tom Fitz, Peter Hinz, Jim Miller, Mark Jirsa and Pete Hollings. ILSG Secretary Hollings took the minutes of the meeting, which are as follows: 1. Accepted report of the Chairs for the 58th ILSG, Thunder Bay, Ontario; as printed in the Proceeding Volume (Hinz), and minutes of last Board meeting, May 17, 2012 (Hollings). 2. Received, discussed, and accepted the 2012-2013 ILSG Financial Summary (Jirsa). 3. Received, discussed, and accepted the 2012-2013 report of the Secretary (Hollings). 4. Approved Allan Blaske as ILSG Board member representing the 2013 meeting. 5. Approved Hibbing as the site for the 60th annual ILSG meeting. The meeting will be hosted by Jim Miller and Mark Jirsa. 6. Discussed and approved renewal of Peter Hollings as Institute Secretary (end of term 2013). This was later approved by a vote of the membership during the technical session. 7. Discussed and approved replacing Laurel Woodruff as the “member from government” on Goldich Committee (end of term 2013) with Mark Smyk. 8. Discussed the possibility of having a short themed session with invited speakers at future meetings. 9. Requested that Secretary Hollings contact past chairs in order to compile statistics on the number of responses to the first circular in relation to attendance at the meeting. 10. Discussed the topic of the level of student participation required to be eligible for an Eisenbrey Student Travel Award. While there was no formal vote, the Board agreed that meeting Chairs should include a statement on the application indicating that full participation in the meeting was required to receive the award. In other words, a student cannot come for ½ day of the technical session or only attend a field trip to pick up the award. The meeting co-Chairs would like to thank all those who assisted with this year’s meeting either by chairing sessions, judging student papers, leading field trips, driving vehicles for field trips, helping with the registration desk, operating the meeting and banquet projectors, and more. These volunteers made the quality of the meeting better and of course, made the job of the meeting chairs far easier than it may otherwise have been. A special thank you goes to Darlene Comfort, who tirelessly and patiently managed the registration and logistics for the meeting. We are gratified by all of the positive comments by participants. While chairing and organizing an ILSG meeting involves time and a bit of stress on occasion, we are happy to have had the opportunity to serve the geological community of the Lake Superior region. We look forward to the 2014 meeting when we can be much more relaxed! Respectfully submitted, Ted Bornhorst and Allan Blaske Co-chairs, 59th Institute on Lake Superior Geology xxiii �PROGRAM WEDNESDAY MAY 14, 2014 All trips leave from the northeast entrance of the Hibbing Park Hotel 8:00am - 5:30pm PRE-MEETING FIELD TRIPS FIELD TRIP 1: STRATIGRAPHY, SEDIMENTOLOGY, STRUCTURE, AND MINERALIZATION OF THE BIWABIK IRON FORMATION, CENTRAL MESABI IRON RANGE Phil Larson, Duluth Metals Ltd. Marsha Patelke, Natural Resources Research Institute, UMD Jakob Wartman, United Taconite, Cliffs Natural Resources Michael Totenhagen, Arcelor Mittal Mark Jirsa, Minnesota Geological Survey Steven Losh, Minnesota State University – Mankato Peter K. Jongewaard, Cliffs Natural Resources (retired) FIELD TRIP 2: A WALK IN THE PARK - NEOARCHEAN GEOLOGY OF LAKE VERMILION STATE PARK George J. Hudak, Natural Resources Research Institute – UMD Amy Radakovich, Minnesota Geological Survey Geoff Pignotta and Kelly Schwierske, University of Wisconsin - Eau Claire FIELD TRIP 3: WESTERN MESABI RANGE MINING OPERATIONS Douglas Halverson, Cliffs Natural Resources—Duluth Daniel Cervin, Cliffs Natural Resources—Hibbing Taconite William Everett and Kevin Kangas, Essar Steel Joseph Nielsen, Magnetation 4:00 pm - 10:00 pm Registration at Hibbing Park Hotel (hallway outside Arrowhead Ballroom) 7:00 pm - 10:00 pm Ice Breaker Social (Arrowhead Ballroom) and Poster Session (Whispering Pines Room) xxiv �THURSDAY MAY 15, 2014 Asterisk * denotes a student eligible for Best Student Paper Award 7:30 am - noon REGISTRATION 8:00 am OPENING REMARKS Jim Miller and Mark Jirsa, Co-Chairs, 2014 ILSG TECHNICAL SESSION I Session Chairs: Al MacTavish – Panoramic Resources Joyashish Thakurta – Western Michigan University 8:10 Peter Hollings and Geoff Heggie Rethinking the Midcontinent Rift – Puncturing the “Plume Paradigm” 8:30 Paul Bedrosian Electrical resistivity structure of the Midwestern United States from EarthScope magnetotelluric data 8:50 Elisa Piispa*, Aleksey Smirnov and Lauri Pesonen Mesoproterozoic Midcontinent Rift intrusives in Thunder Bay area, Ontario, Canada: a paleomagnetic review 9:10 Adam Leu* and Jim Miller Geology and petrology of the Wilder Lake Intrusion, Duluth Complex, northeastern Minnesota 9:30 Klaus Schulz, Laurel Woodruff and Suzanne Nicholson Midcontinent Rift-related satellite mafic-ultramafic intrusions hosting Fe-Ti-V oxide deposits 9:50 COFFEE BREAK AND POSTER SESSION 10:20 Gabe Sweet, Dean Peterson, Phil Larson, Molly Finnegan, Evan Finnes, Charlie Parent, Bob Nowak, Tyler Boley Sulfide highway revisited: New ideas on internal structure and sulfide mineralization of the Nickel Lake Macrodike 10:40 Molly Finnegan and Phil Larson Geochemistry of basalt xenoliths entrained in mineralized troctolitic and anorthositic intrusions, northeastern Minnesota 11:00 Alex Steiner* and Jim Miller Genesis of sulfide mineralization within the footwall granite of the Maturi Cu-NiPGE Deposit of the South Kawishiwi Intrusion, Duluth Complex, NE Minnesota 11:20 Brent Trevisan*, Peter Hollings and Doreen Ames The Thunder mafic to ultramafic intrusion: a PGE and precious metal-bearing early rift conduit system in the Midcontinent Rift 11:40 Jeff Mauk, Laurel Woodruff and Ester Stewart Variable copper mineralization in the lower Nonesuch Formation of the Midcontinent Rift System: Constraints on regional controls xxv �Noon LUNCH BUFFET (free to all registrants) ILSG BOARD MEETING TECHNICAL SESSION II Session Chairs: Geoff Pignotta – University of Wisconsin – Eau Claire Marcia Bjornerud – Lawrence University 1:30 Bill Cannon, Laurel Woodruff, Stacy Saari, and Molly Hagstrom A new occurrence of the Sudbury impact layer in the Gogebic Iron Range of Wisconsin 1:50 Monica Karman*and Philip Fralick Sedimentology and paleogeographic reconstruction of the layers in and adjacent to the Subury Impact Layer in the Lake Superior Basin 2:10 Leif Johnson and Brad Dunn An exploration update and mineralogical study of the Emily District Maganese Deposit, Cuyuna Iron Range, Minnesota 2:30 Adrian Arts* and Philip Fralick Nanoscale features within freshwater lacustrine ferromanganese nodules: Nanospheres, nanotubes and nanowires 2:50 James Walsh Strontium isotope study of Mesabi Iron Range groundwater 3:10 COFFEE BREAK AND POSTER SESSION 3:40 Robert Seal The danger of “Sulfide Mining” in the Lake Superior Region 4:00 Tim McIntyre* and Philip W. Fralick Sedimentology and geochemistry of the Mesoarchean chemical sediments of Wallace Lake and Red Lake 4:20 Rob Cundari, Mark Smyk and Peter Hollings Geology and geochemistry of the Mesoproterozoic Badwater Intrusive Complex, Ontario: Implications for GEON 15 magmatism 4:40 Terry Boerboom, Karl Wirth and Joseph Evers Five newly acquired high-precision U-Pb ages in Minnesota, and their geologic implications 6:00 RECEPTION – CASH BAR 7:00 ANNUAL BANQUET (Arrowhead Ballroom) − Announcement of 61st Annual Meeting Location − 2014 Goldich Award Presentation to Laurel Woodruff − Banquet Presentation by Dr. Francis M. Carroll, Univ. of Manitoba A Line in the Trees: History of the US-Canadian Boundary from Lake Superior to Lake of the Woods xxvi �FRIDAY MAY 16, 2014 Asterisk * denotes a student eligible for Best Student Paper Award 8:00 OPENING REMARKS, UPDATES Jim Miller and Mark Jirsa, Co-Chairs, 2014 ILSG TECHNICAL SESSION III Session Chairs: Mary Louise Hill – Lakehead University Bernie Saini-Eidukat – North Dakota State University 8:10 Jack Berkley Deer Lake Complex redux: Memories and reflections, 1970 - 1972 8:30. Ben Kuzmich*, Peter Hollings and Michel Houle Geochemistry and mineralogy of the Fe-Ti-V-P mineralized ferrogabbroic intrusions of the McFaulds greenstone belt, Superior Province, northern Ontario, Canada 8:50 Jordan Quinn*, Peter Hollings and John Biczok Geochemistry and petrography of a mafic metavolcanic sequence south of Musselwhite Mine 9:10 Lionnel Djon*, Gema Olivo, Jim Miller and Rob Stewart Petrology of the layered North Lac des Iles Intrusion, Ontario: Part I. Stratigraphy and mineral-chemical evidence for multiple magma injection 9:30 Skylar Schmidt* and Mary Louise Hill Structural control of mineralization at Lac des Iles Mine 9:50 COFFEE BREAK AND POSTER SESSION 10:20 Amanda Van Lankvelt*, M. Williams, D. Schneider, S. Seaman, Garnet in the deep crust: The key to linking Archean TTG generation and vertical block motions? 10:40 Jonathan Dyess* and Vicki Hansen Structural and kinematic analysis of the Shagawa Lake Shear Zone: Implications for Archean tectonic processes in the Southern Superior Province 11:00 Simon Dolega* and Mary Louise Hill Strain analysis on the Max Lake polymictic conglomerates in the Wabigoon Subprovince, Ontario 11:15 Leah Clapp* and Mary Louise Hill Evidence of simultaneous brittle and ductile deformation in the Main Break Fault System, Kirkland Lake, Ontario 11:30 Jared Liimu* and Mary Louise Hill The role of brittle-ductile deformation and competency contrast in gold mineralization in the C-zone at Hemlo 11:45 Daniel LaFontaine* and Mary Louise Hill Structural control on the Borden Gold Deposit in Chapleau, Ontario xxvii �Noon LUNCH BUFFET (free to all registrants) 1:20 BEST STUDENT PAPER AWARDS STUDENT TRAVEL AWARDS 2:00-6:00 FRIDAY AFTERNOON FIELD TRIPS FIELD TRIP A: STATE DRILL CORE LIBRARY – HIBBING, MINNESOTA (MINNESOTA DEPARTMENT OF NATURAL RESOURCES –DIVISION O F LANDS AND MINERALS Dave Dahl, Minnesota Department of Natural Resources Dean Rossel, Kennecott Exploration, Rio Tinto FIELD TRIP B: HIBBING’S IRON MINING AND CULTURAL HISTORY Henry Djerlev and Staff from the Hibbing Historical Society FIELD TRIP C: MINNESOTA DISCOVERY CENTER Discovery Center Staff FIELD TRIP D: COLERAINE MINERALS RESEARCH LABORATORY, NATURAL RESOURCES RESEARCH INSTITUTE, UNIVERSITY OF MINNESOTA DULUTH Dick Kiesel, Director CMRL Dave Hendrickson, Director Strategic Planning Matt Mlinar, Program Coordinator Mineral Processing Basak Anameric, Program Coordinator High Temperature Process FIELD TRIP E: MINEVIEW FROM A CANOE Mark Jirsa, Minnesota Geological Survey Dan Jordan, Iron Range Resources and Rehabilitation Board Dale Cartwright, Minnesota Department of Natural Resources SATURDAY MAY 17, 2014 8:00am – 5:00pm POST-MEETING FIELD TRIPS FIELD TRIP 5: VISIONS OF MATURI: THE GEOLOGY OF THE SOUTH KAWISHIWI INTRUSION Dean Peterson, Duluth Metals Ltd. FIELD TRIP 6: THE ST. LOUIS SUBLOBE AND GLACIAL LAKE UPHAM Phil Larson, Duluth Metals Ltd. Alan Knaeble, Minnesota Geological Survey Howard Mooers, University of Minnesota Duluth Lisa Marlow, Halcon Resources Corp. FIELD TRIP 7: GEOLOGY & GOLD MINERALIZATION OF THE VIRGINIA HORN AREA Mark Jirsa, Minnesota Geological Survey Bill Rowell, Vermilion Gold LLC Rick Sandri, Vermilion Gold LLC Jason Richter, Minnesota Department of Transportation xxviii �POSTER PRESENTATIONS Asterisk * denotes a student eligible for Best Student Paper Award Steven D. J. Baumann, Alex B. Cory and David Wilson Fault offsetting in the Proterozoic Lorraine Formations along Government Road, south of Echo Bay, Ontario, Canada Mark Baumgardner, Nathan Brown, Matt Grotte, Alan Jacobson, Jamie Kendall, Claire Ostwald, Nathan Schriner, Justin White, and Dean Peterson Bedrock geologic map of the Gafvert Lake area, St. Louis County, northeastern Minnesota Patrick Belshaw* and Mary Louise Hill Relationship between Microstructure and Rock Mechanics in Shear-Zone-HostedGold Deposits Terry Boerboom and John Green Bedrock geologic map of the Marr Island and Hovland 7.5'quadrangles, North Shore of Lake Superior,Minnesota Anthony Boxleiter*, Joyashish Thakurta, and Thomas Quigley Geochemical investigation of the origin of the Back Forty volcanogenic massive sulfide deposit, Menominee County, Michigan Tom Buchholz, A. Falster, and W. Simmons Zirconium/Hafnium fractionation in some pegmatites of the upper Midwest, USA Val Chandler and Richard Lively Continued work on using the horizontal-to-vertical spectral ratio (HVSR) passive seismic method for determining Quaternary sediment thickness in Minnesota Ben Drenth, Ray Anderson, Klause Schulz, Val Chandler, Bill Cannon, Ben Bloss, Paul Bedrosian, Josh Feinberg, Rob McKay Preliminary interpretation of Precambrian lithology and structure from high-resolution, multi-method geophysics, northeast Iowa and southeast Minnesota Jonathan Dyess* and Vicki Hansen Determination of Vorticity in Archean Tectonites Nicholas Fedorchuk*, Stephen Dornbos, John Isbell, Julie Bowles, Frank Corsetti, Dylan Wilmeth, Victoria Petryshyn Bedrock Geologic Map of the Putnam Lake Area, St. Louis County, NE Minnesota – Precambrian Research Center Capstone Project Kiel Finn* and Julie Bowles Magnetic Mineralogy of Reversely Magnetized Chengwatana Lava Flows of St. Croix Falls, Wisconsin Sidney Firmin* and Julie Bartley An Unusual Mesoproterozoic Carbonate Unit: Relic of a Saline Lake? Paul Fix, Stephen Ginley, Lauren Schraeder, Aaron Summers, Michael Doyle,Terry Boerboom Geology of the Brule River area of the Pine Mtn quadrangle, Minnesota: Capstone mapping project for the Precambrian Research Center’s 2013 field camp xxix �Marine Foucher*, Renee Curganus, Elisa Piispa, Aleksey Smirnov, and Lauri Pesonen Evolution of the Midcontinent Rift system: Paleomagnetic, rock magnetic and anisotropy of magnetic susceptibility study of the Mesoproterozoic Baraga - Marquette dike swarm, MI, USA VJ Grauch, Val Chandler, and Rich Lively Compilation of existing geophysical models in preparation for 3D modeling of the Midcontinent Rift System in the western Lake Superior region, Minnesota, Wisconsin, and Michigan Matt Grotte* and George Hudak, A field and petrographic study of Neoarchean variolitic pillow lavas, Newton Belt, Vermilion District, NE Minnesota. Ivan Guzman* Stratigraphic framework and landsystem correlation for deposits of the Saginaw Lobe, Michigan, USA George Hudak, Stephen Monson Geerts, Larry Zanko, Sara Post, and Bryan Bandli The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter - 2014 Update Darcy Jacobson*, Elisa Piispa, Aleksey Smirnov, and Lauri Pesonen Silica Remobilization in the Biwabik Iron Formation, Minnesota USA Monica Karman* and Phil Fralick Impact ejecta features in the Lake Superior basin from the 1850Ma Sudbury Impact Event Stephen Kissin and Gregory Brumpton PDFs in Sudbury Ejecta in the Gunflint Formation, Ontario: A Comparison of Methods Matthew Lamb* and Prajukit Bhattacharyya Ru-Rh-Pd Mobilization in Flambeau Massive Sulfide Deposit Gordon Medaris, Jr., Tim Flood, Brian Jicha and Bradley Singer Composition and 40Ar/39Ar Age of Pegmatitic Amphibole in the Wausau Syenite Complex, Marathon County, Wisconsin Jim Miller, Sarah Sauer, Jordan Benningfield, Jackson Graham, Sara Kozmor, and Ann Marie Prue Geology of the Lake Three Troctolite, Duluth Complex - 2013 Precambrian Field Camp Capstone Connor Mulcahy, Dan Romanelli, Roger Schulz, Steve Moorhead, Mitchell May, and Mark Jirsa Geologic mapping of Neoarchean and Paleoproterozoic rocks near Hanson Lake, NE Minnesota, by students of the Precambrian Research Center’s 2013 field camp Brynley Nadziejka* and Marcia Bjornerud Petrographic characterization of the Penokean Twelvefoot Falls Shear Zone, Marinette County, Wisconsin: Evidence for coeval ductile and seismic behavior xxx �Ainslee Nolan* and Mary Louise Hill Metamorphism and Deformation at the Wabioon-Quetico subprovince boundary in the Decourcey Lake area Dean Peterson Bedrock geologic map of the Twin Metals Minnesota Project, Northern South Kawishiwi Intrusion and adjacent areas Nadine Piatak, Robert Seal, Perry Jones, and Laurel Woodruff Potential for copper toxicity caused by surface water and stream sediments in unmined mineralized watersheds of the Duluth Complex Patrick Quillen* and Jim Miller Documenting the first lava flows of the Midcontinent Rift by digital mapping and petrographic analysis Amy Radakovich and Howard Hobbs The Arrowhead Pilot Project: Mapping of Precambrian and Quaternary geology in two diverse geologic areas of northeastern Minnesota Bill Rose and Erica Vye Tools for interpreting Keweenaw geoheritage to a broad public Kelly Schwierske*, Geoff Pignotta, and George Hudak The 2.7 billion year old Mt. St. Helens of northern Minnesota: Petrography, geochemistry and economic significance of the Neoarchean Gafvert Lake sequence Laurel Woodruff, Bill Cannon, Federico Solano, and David Smith Geochemistry and Mineralogy of Glacial Soils in the Upper Midwest Chris Yip* and Phil Fralick The evolution of the atmosphere-hydrosphere: A geochemical comparison of two Paleoproterozic Gunflint weathering profiles xxxi �Abstracts ��Nanoscale features within freshwater lacustrine ferromanganese nodules: Nanospheres, nanotubes and nanowires ARTS, Adrian and FRALICK, Philip. Department of Geology, Lakehead University, Oliver Rd. Thunder Bay, ON, P7B 5E1, Canada Iron-hydroxide and manganese-oxide precipitates, often referred to as ferromanganese nodules (FMN), are common occurrences on lake bottoms worldwide (Sozanski and Cronan, 1978). The nodules form at the sediment-water interface, generally on a sandy substrate, at neutral pH (Kindle, 1932). Detailed studies documenting their morphology and geochemistry have been conducted by several authors (Sommers et al., 2002). FMNs can take varying morphological forms. However, they most typically accrete as disk-shaped precipitates with a concentric growth pattern of alternating Fe- and Mnrich bands, around a central pebble or small cobble nucleus (Sozanski and Cronan, 1976) (Fig. 1A, 1B). Bacterially mediated precipitation and changes in redox conditions are believed to be a significant factor in their growth (Dean and Greeson, 1979; Boudreau, 1988). These features are environmentally important as the iron hydroxides composing them adsorb arsenic with concentrations up to 4900 ppm. Despite the considerable amount of work conducted on the nodules, no research has been undertaken to explore FMNs at the micro- and nanoscale to investigate the extreme arsenic uptake abilities. This study was conducted to provide new insights on the micro- and nanoscale features within FMNs and to determine the geochemical composition of these features. The nodules examined were collected from Shebandowan Lake (Ontario), Sowden Lake (Ontario), and Lake Charlotte, (Nova Scotia). The utilization of high resolution field emission scanning electron microscopy (SEM) revealed an intriguing range of nanoscale forms previously undocumented in FMNs. Coccus bacterial forms were commonly found implanted in extracellular polymeric substance (EPS). Figure 1C illustrates the high concentrations of ovoid to round nanospheres (100-200nm diameter) which are embedded within the EPS. Semi-quantitative analysis (SEM-EDX) indicates the areas containing nanospheres are enriched in iron. It has been postulated that similar structures in carbonates provide nucleation sites for biologically induced mineralization (Aloisi et al., 2006) preventing the cellular membrane from being entombed by the precipitates (Bontognali et al., 2008). Nanotubes were also documented and appear to be ubiquitous in the FMN samples. The nanotubes appear as a tangled mass of worm like structures, which range in length, from 2-40 µm, with a diameter range of 50-400 nm. Uwins et al. (1998) reported similar structures in Triassic and Jurassic sandstones. Utilizing three different RNA staining techniques they deduced that the structures are biogenic, contain RNA, and have thus referred to them as nanobes. Finally, nanowires appear to be a common formational constituent of samples from each of the three lakes (Fig. 1E). These wires build together into large sheet-like- masses. SEM-EDX analysis show the wires to be composed of manganese oxides. This study provides new insight as to how FMNs accrete, and how they are able to accumulate high concentrations of toxic metals. Similar to the environmental goals of artificially produced metal nanotubes, the biogenic iron hydroxide nanotubes greatly increase the reactive area allowing far greater arsenic adsorption. 1 �(A) (B) (C) 2.00μm (E) (D) 1.00μm 5.00μm Figure 1. Images of Ferromanganese nodules at different scales. Picture of dorsal (A) and ventral (B) side of a FMN forming around a cobble nucleus, with distinct concentric laminations. (C) SEM image of nanospheres embedded into a smooth layer of extracellular polymeric substance. Mineralization can be seen increasing from the top to bottom of the image. (D) Intertwined mass of nanotubes coated in an iron precipitate. The varying diameters and lengths are evident in the image. (E) Wispy mass of manganese oxide nanowires. They can be seen growing together to form sheet-like layers. References Aloisi, G., Gloter, A., Krüger, M., Wallmann, K., Guyot, F., and Zuddas, P. 2006. Nucleation of calcium carbonate on bacterial nanoglobules. Geology. 34, 1017-1020. Bontognali, T.R., Vasconcelos, C., Warthmann, R.J., Dupraz, C., Bernasconi, S.M., and McKenzie, J.A. 2008. Microbes produce nanobacteria-like structures avoiding cell entombment. Geology. 36, 663-666. Boudreau, B. 1988. Mass transport constraints on the growth of discoidal ferromanganese nodules. Journal of American Science. 288, 777-797. Dean, W.E., and Greeson, P.E. 1979. Influences of algae on the formation of freshwater ferromanganese nodules Oneida Lake, New York. Archiv fur Hydrobiologie. 86, 181-192. Folk, R.L. 1993. SEM imaging of bacteria and nannobacteria in carbonate sediments and rocks. Journal of Sedimentary Petrology. 63, 990-999. Gorham, E, and Swaine, D. J. 1965. The influence of oxidizing and reducing conditions upon the distribution of some elements in lake sediments. Limnology and Oceanography. 10, 268-279. Harriss, R.C., and Troup, A.G. 1969. Freshwater ferromanganese concretions: chemistry and internal structure. Science. 166, 604-606. Kindle, E.M. 1932. Lacustrine concretions of manganese. American Journal of Science. 5(24), 496-504. Sommers, M., Dollhopf, M., and Douglas, S. 2002. Freshwater ferromanganese stromatolites from Lake Vermilion, Minnesota: Microbial culturing and scanning electron microscopy investigations. Geomicrobiology Journal. 19, 207-227. Sozanski, A.G., and Cronan, D.S. 1976. Environmental differentiation of morphology of ferromanganese oxide concretion in Shebandowan Lakes, Ontario. Limnology and Oceanography. 21, 894-898. Sozanski, A.G., and Cronan, D.S. 1978. Ferromanganese concretions in Shebandowan lakes, Ontario. Canadian Journal of Earth Science. 16, 126-140. 2 �FAULT OFFSETTING IN THE PROTEROZOIC LORRAINE AND JACOBSVILLE FORMATIONS, ALONG GOVERNMENT ROAD, SOUTH OF ECHO BAY, ONTARIO, CANADA BAUMANN, Steven D.J.1, DYLKA, Sandra K.1 1 Geology Section, Midwest Institute of Geosciences and Engineering, 2328 W. Touhy Ave. Chicago, IL 60645 Along the east side of Government road (a small road that runs parallel to Trans Canada 17) exists an outcrop about 850 feet long at GPS: 46.46141o -84.05810o. The outcrop is mostly of the Jacobsville Formation, with an approximately 300 foot long outcrop of the Lorraine Formation near the center (see Figure 1). The Lorraine is much more indurated than the Jacobsville. The Lorraine was quarried at this location during sometime in the past. The Lorraine is a nearly white, thinly bedded, crystalline, fine to medium grained, quartz arenite, with minor beds of red jasper and white quartz conglomerate that has been metamorphosed. The Lorraine exposed at the outcrop is probably near the top of the formation. The Jacobsville is dominantly a reddish purple mottled pale yellow brown, cross bedded, fine to medium grained, non-metamorphosed, quartz arenite. Lenses of dark red sandy siltstone with red and green shale breccia are common in the Jacobsville (see Figure 2). The exact stratigraphic position of the Jacobsville is not known at this location. The outcrop displays a section of the Lorraine “poking up” through the Jacobsville (see Figure 2). There are several ways this relationship could have formed. 1) The Lorraine existed as a paleo-high and the Jacobsville was deposited around it, making the contact depositional in nature. A similar situation exists in the Baraboo Range at the Upper Narrows in Wisconsin, where the Precambrian Baraboo Quartzite is in contact with the Cambrian sandstones and conglomerates. 2) The Jacobsville could have been deposited with some initial dip, and as more sediments accumulated the weight created growth faulting along the Lorraine-Jacobsville contact. 3) The outcrop is a horst structure, where the Jacobsville was originally deposited with little to no initial dip, and later extensional tectonic forces lowered the Jacobsville relative to the Lorraine. Due to the field relationships of the outcrop, we believe the exposure to be a horst structure (see Figures 1 and 2) for the following reasons. 1) Perhaps the most compelling line of evidence is that no recognizable clasts of Lorraine exist within the Jacobsville at this outcrop, unlike what is seen at the Upper Narrows in Wisconsin, where clasts of Baraboo Quartzite are commonly seen in the local Cambrian sandstones. 2) The brecciated nature of the green and red shale cobbles within the siltstone facies of the Jacobsville appear jumbled. They were probably deposited as clay in stream beds and later brecciated during faulting. This makes sense that the shale and siltstone would have been more susceptible to deformation than the surrounding quartz arenites. 3) The variation in strike and dip between the Jacobsville surrounding the Lorraine. The Lorraine was not deformed during faulting. However, the Jacobsville north of the north fault has a different strike than it does on the south side of the south fault (the south fault may also have some lateral strike-slip movement). 4) There is complex breccia exposed at the north fault (see Figure 4). 5) Slickensides are present on the faces of the Lorraine at both the north and south faults. References: Baumann, S.D.J., 2013. Contact of the Precambrian, Lorraine and Jacobsville Formations, along Government Road, South of Echo Bay, in an Abandoned Quarry, Ontario. Midwest Institute of Geosciences and Engineering, M-122013-2A Jackson, S.L., 2001. On the Structural Geology of the Southern Province between Sault Ste. Marie and Espanola, Ontario. Ontario Geological Survey, Open File Report 5995 Johns, G.W., Mcllraith S., Muir, T.L., 2003. Precambrian Geology Compilation Series, Sault Ste. Marie-Blind River Area, Ontario Ministry of Northern Development and Mines, MAP 2670 3 �Figure 3: Photo of the Stratigraphic Relationships at the North Fault, U.S. Dollar coin for scale Figure 1: Diagram of the Outcrop and Location Map Figure 2: Conceptual Diagram of the Horst Structure along Government Road 4 �BEDROCK GEOLOGIC MAP OF THE GAFVERT LAKE AREA, ST. LOUIS COUNTY, NORTHEASTERN MINNESOTA Mark Baumgardner1, Nathan Brown2, Matt Grotte3, Alan Jacobson4, Jamie Kendall5, Claire Ostwald6, Nathan Schriner7, Justin White8, and Dean Peterson9 1 Wayne State University, 2Virginia Tech University, 3University of Minnesota Duluth, 4University of Wisconsin Milwaukee. 5Swarthmore College, 6Boston University, 7University of Cincinnati, 8Northwest Missouri State University, 9Duluth Metals Limited and UMD Natural Resources Research Center Each year, students from the Precambrian Research Center (PRC) geology field camp complete “capstone” projects that encompass approximately one week of detailed field mapping followed by one week of mapmaking and map publishing. During the fifth and sixth weeks of the 2013 field camp, eight PRC field camp students, under the direction of PRC Assistant Director Dean Peterson, mapped Neoarchean rocks of the informally named Gafvert Lake Sequence (Peterson and Jirsa, 1999, Peterson, 2001) between eastern Lake Vermilion’s Mud Creek Bay and Armstrong Lake, 6 miles to the east-southeast (Baumgardner et al., 2013). This capstone mapping project sought to: 1) identify the lithologies and determine the detailed stratigraphy within the Neoarchean supracrustal strata in this area; 2) define and characterize the nature of the contacts between various units of the Neoarchean supracrustal strata and intrusive rocks; 3) obtain a better understanding of geological structures and their orientations within the area; 4) produce a detailed geological map of the entire Gafvert Lake sequence stratovolcano. Mapping was carried out over five days by eight students (Figure 1) of PRC field camp and 617 new outcrops were mapped by hiking and lakeshore canoe mapping in the field area. The final map incorporated over 1,500 outcrops from historical work in the area. Emphasis was placed on defining the structure of an Archean stratovolcano within the Vermilion Greenstone Belt. Figure 1. Students of the Gafvert Lake Capstone. Mark Severson, while mapping in the area for US Steel in the early 1980's, first recognized that the volcanic rocks in the area around Gafvert Lake represent a proximal facies dacitic to andesitic volcanic edifice. The morphology of the volcanic complex is best seen at map-scale, 5 �which provides an almost perfect cross section through an Archean stratovolcano of dacitic to andesitic composition. The sequence overlies the Soudan Iron Formation, is overlain by the Upper member of the Ely Greenstone to the east and north, and interfingers with reworked tuff and greywacke of the Lake Vermilion Formation on the west. In simple terms, the complex consists of a core of dacite lava flows that are overlain by coarse fragmental volcanic rocks of dacitic to andesitic composition. The fragmental rocks are in turn overlain by thin- to medium-bedded dacitic lapilli and ash tuffs. The whole complex is cut by multiple intrusions of coarse-grained quartz-feldspar porphyry (with rounded quartz phenocrysts up to 1.5 cm across), which occur as a central plug and thick sills to the east and west. The presence of pumice and scoriaceous clasts in the fragmental rocks indicates that much of the sequence was erupted in extremely shallow water or subaerially. Two large bodies of quartz-feldspar porphyry intrude pillow basalts of the overlying Upper member of the Ely Greenstone and probably represent the last episode of igneous activity associated with the sequence. Tuffaceous greywacke of the Lake Vermilion Formation is inferred to be derived largely from this dacitic complex, and possibly other felsic complexes developed along this stratigraphic horizon. Capping the Gafvert Lake sequence to the east and north is a distinct horizon of multiple-facies iron-formation. References Baumgardner, M., Brown, N, Grotte, M., Jacobson, A., Kendall, J., Ostwald, C., Schriner, N., White, J., and Peterson, D., 2013, Bedrock Geologic Map of the Gafvert Lake Area, St. Louis County, Northeastern Minnesota; Precambrian Research Center, PRC/MAP 2013-04. Peterson, D.M., 2001, Development of Archean Lode-Gold and Massive Sulfide Deposit Exploration Models using Geographic Information System Applications: Targeting Mineral Exploration in Northeastern Minnesota from Analysis of Analog Canadian Mining Camps; University of Minnesota Ph.D. thesis, 503 pages, 12 plates, 1 CD-Rom. Peterson, D. M., and Jirsa, M.A., 1999, Bedrock geologic map and mineral exploration data, western Vermilion district, St. Louis and Lake Counties, northeastern Minnesota: Minnesota Geological Survey, Miscellaneous Map M-98, scale 1:48,000. 6 �ELECTRICAL RESISTIVITY STRUCTURE OF THE MIDWESTERN UNITED STATES FROM EARTHSCOPE MAGNETOTELLURIC DATA BEDROSIAN, Paul, U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 As part of the EarthScope USArray program, long-period magnetotelluric (MT) data have been collected within the Midwestern United States. From 2011-2013, 237 MT stations were collected with 70 km nominal station spacing over all of Minnesota, Wisconsin, Michigan and Iowa as well as parts of Illinois, Indiana, Missouri, Kansas, Nebraska, and Ohio. This data set is unique in its ability to constrain subsurface electrical resistivity at crustal and lithospheric scales over this broad region. Coupled with advances in three-dimensional (3D) MT inversion, the EarthScope MT data can be used to create 3D resistivity models with coverage and resolution comparable to seismic tomography models. I will present a preliminary 3D resistivity model derived from these data. The discretized resistivity inverse model has a horizontal cell size of 10 km, with cell thickness starting at 100 m and increasing logarithmically with depth. The full MT impedance tensor was inverted at 11 periods ranging from 10 – 20,000 sec; vertical magnetic-field transfer functions at these same periods were also inverted. The resulting 3D resistivity model reflects the complex structural collage from the Archean to present. Several zones of high conductivity from the surface to ~5 km depth mimic the distribution of Phanerozoic sediments within the Michigan, Illinois, and Forest City basins. In contrast, structural highs such as the Transcontinental Arch and the Wisconsin and Ozark domes are electrically resistive at these same depths. At upper- to mid-crustal depths, the resistivity model illuminates the first-order structure of the 1.1 Ga Mid-Continental Rift (MCR) system. Highly resistive rocks coincide spatially with high magneticfield anomalies, and are attributed to Keweenawan volcanic rocks along the length of the southwest rift arm. In addition, flanking conductive anomalies trace out thick packages of Keweenawan clastic rocks, some of which appear to extend to ~15 km depth. The structure of the MCR system as imaged within the 3D resistivity model is most striking within the Iowa Horst, but can be traced throughout the known extent of the MCR system, including beneath the Michigan Basin within the southeastern rift arm. Structures predating the MCR system are also reflected in the resistivity model, particularly in the northern half of the model area. One example, the Flambeau anomaly, is imaged as a 300-km long, eastwest trending structure through northern Wisconsin and upper Michigan at 46°N. The southern boundary of this high conductivity zone appears coincident with the Niagara Fault. 7 �Figure 1: Preliminary resistivity structure at 9 km depth as constrained by 3D inversion of EarthScope magnetotelluric data. Background hillshade shows the total magnetic-field anomaly for comparison. 8 �RELATIONSHIP BETWEEN MICROSTRUCTURE AND ROCK MECHANICS IN SHEAR-ZONE-HOSTED-GOLD DEPOSITS Belshaw, Patrick and Hill, Mary-Louise Department of Geology, Lakehead University, 955 Oliver Rd, Thunder Bay, ON P7B 5E1, Canada. Orogenic gold deposits of northern Ontario are often hosted in steeply dipping, mylonitic shear zones. The common challenge of identifying, and mitigating the effects of penetrative planar fabrics with respect to rock mass properties is important in determining the safety of any mining operation. In addition to the relative effect on rock mass behaviour associated with fault planes and joint surfaces, highly ordered flaws in the material can have a significant effect on the orientation, initiation, and propagation of tensile Griffith fractures. In these deposits, where inherent flaws parallel to foliation dip steeply, spalling conditions are often observed in rocks that appear to be adequately competent. Most rocks in these deposits have also undergone grain size reduction during progressive deformation, making the interpretation of penetrative fabrics much more difficult on the outcrop scale. On the microscopic scale however, the morphology and density of grain boundaries and other flaws can be evaluated to develop relationships between the microscopic texture, and macroscopic material properties of the rock mass. 9 �10 �DEER LAKE COMPLEX, REDUX: MEMORIES AND REFLECTIONS, 1970 - 1972 BERKLEY, Jack, Department of Geosciences, Houghton Hall, SUNY Fredonia, Fredonia, NY 14063 USA The Deer Lake Complex (DLC), located between the town of Big Fork and southern tip of Deer Lake in Itasca County, MN, has been the target of precious and industrial metals exploration for over forty years. It was a major Cu-Ni prospect in the early 1970s, but has recently received renewed scrutiny as a gold, PGM, and Cu-Ni prospect. Although written off during the late 60s – early 70s exploration boom as a “dry hole”, application of modern technology and imaginative exploration strategies could conceivably bear positive results in the near future. The DLC consists of at least two gravity-stratified mafic-ultramafic sills, each roughly 250 meters thick, intruded into an Archean greenstone-meta-sedimentary terrain within the Wawa subprovince of the southern Superior province, Northern Minnesota (Berkley, 1972). Each sill differentiated into a lower wherlite / hornblende peridotite, overlain by distinct layers of clinopyroxenite, poikilitic and non-poikilitic gabbro, topped off by a layer of quartz-hornblende diorite (Berkley and Himmelberg; Ripley, 1978). Upper-most diorite units are commonly transected by thin, randomly dispersed veins of granitic assemblages (mostly microcline + quartz), and also display sheaf-like assemblages of acicular augite (Fig. 2a) plus skeletal plagioclase, the result of extreme undercooling at emplacement contacts (Ripley, 1972). Exposed lower contacts of sills (below peridotite) have chilled margins consisting of fine-grained basalt that – along with quench textures noted above – suggest shallow emplacement. Pyroxene reaction rims on peridotite olivine grains (Fig. 2b) indicate emplacement of less than about 6 km depth (e.g., Longhi and Pan, 1987). (a) (b) Figure 1: (a) Location map for DLC exposures. Textured squares represent areas mapped 1970-1973 (Berkley & Himmelberg, Ripley, 1978). (b) DLC geologic map from Ripley, 1978. Initial interest in the area was prompted by 1970 USS Corp. aeromagnetic plots that portrayed pronounced NE-trending, parallel linear patterns. Recent UMD grads, Jack Berkley and David Witt, were dispatched by Sid Iverson to the area to determine the cause of the magnetic anomalies. Upon entering the area using a logging road winding south from highway MN-1, they were eventually rewarded by the discovery of highly sheared black, blue, and blue- 11 �green – magnetite-rich serpentine, instantly accounting for the mag anomalies. What followed was a program using techniques and equipment that might seem archaic by today’s standards, but that nevertheless demonstrate the value of systematic field work leading to discovery. Subsequent mapping during August, 1970 required using a sun compass to compensate for the Brunton compass’ tendency to confuse peridotite exposures for the north magnetic pole. Topographic maps of the area were non-existent (the USGS team arrived the next year), thus geospatial positioning required finding, and using county-installed PLS section posts. Section lines festooned with colorful plastic flagging tape served as base lines for traverses that inevitably crossed insect-infested, soggy high-grass wetlands, waist-high blackberry thickets, and black spruce groves good only for obscuring views of whatever outcrops loomed ahead. By the end of August 1970 our team had completed a crude map and hand-written report, likely representing the first geological report on the DLC ever produced. It reported the existence of possible layered, mafic-ultramafic igneous intrusions, consisting of – at the very least -peridotite, pyroxenite, and gabbroic or diorite components. Berkley returned the next summer (1971) to study and complete a map of the DLC to fulfill the requirements for a master’s degree in Geology at the University of Missouri, Columbia under the tutelage of Dr. Glen R. Himmelberg (PhD, 1965 UM, Twin Cities). As he had in 1970, Berkley resided that summer in the cabin on Deer Lake (Fig. 2c) owned by Dave Witt’s parents, where they were visited at times by various UMD Geology alums and other friends. One very important visitor was UMD’s Dr. Richard Ojakangas, our indefatigable undergraduate instructor who was eager to see the work of his students. We duly escorted him into the depths of the DLC so he could plot the complex on the revised Hibbing Sheet of the Minnesota State Geologic Map. It remains there to this very day! Figure 2. a) Super-cooled pyroxene sheaves, from upper sill contact zone (hand specimen). b) Olivine with augite reaction rim, both enclosed by hornblende (photomicrograph). c) DLC field partners, Dave Witt and Jack Berkley (with “field dogs”) outside the Deer Lake cabin, 1971. References Berkley, J. and Himmelberg, G., 1978, Cumulus mineralogy of the Deer Lake Complex, Itasca County, Minnesota, Report of Investigations 20-A, Minnesota Geological Survey, 18pp. Longhi, J. and Pan, V., 1987. Olivine / low-Ca pyroxene liquidus relations and their bearing on eucrite petrogenesis. Lunar and Planetary Sci. XVIII: 570-571. Ripley, Edward, 1978, Sulfide Minerals in the Layered Sills of the Deer Lake Complex, Report of Investigations 20-B, Minnesota Geological Survey, 32pp. 12 �FIVE NEWLY ACQUIRED HIGH-PRECISION U-PB AGES IN MINNESOTA, AND THEIR GEOLOGIC IMPLICATIONS BOERBOOM, Terrence J., Minnesota Geological Survey, boerb001@umn.edu WIRTH, Karl R., and EVERS, Joseph, F., Macalester College, wirth@macalester.edu In the past year, with much appreciated cooperation from Dr. Mark Schmitz, director of the Boise State University isotope geology laboratory, four new high-precision U-Pb ages (three zircon and one baddeleyite) have been acquired for rocks from various geologic terranes throughout Minnesota. We summarize these results and also include the results from a suite of detrital zircon ages obtained as part of a Macalester College student research project. Funding provided by the USGS Statemap program and the MGS County Atlas program. UTM coordinates given below are in NAD ’83, Zone 15. Midcontinent Rift System – Two samples - a porphyritic rhyolite flow (DG073-AD; Grand Marais rhyolite) and a ferromonzonite intrusion (MH047-AD; Hovland sill) were dated as a followup to detailed 1:24,000 scale bedrock mapping (Boerboom and Green, 2010; Boerboom and Green, 2013. Sample DG073-AD – Six zircon crystals were selected for CA-TIMS (Chemical Abrasion Thermal Ionization Mass Spectrometry) analysis, from which five grains produced concordant isotopic ratios, with a weighted mean 206Pb/238U date of 1095.00±0.33 (MSWD (Mean Square Weighted Deviation) = 0.07) and a weighted mean 207Pb/206Pb age of 1097.26±0.67 (n=5; MSWD 1.47). Using the 207 Pb/206Pb weighted mean date, this age is only slightly younger than the Devil’s Kettle rhyolite (1097.7±1.7; Davis and Green, 1997), which lies roughly 8,000 feet stratigraphically below and is separated by several thick mafic to felsic volcanic units. The nearly identical ages for these two units indicates rapid and voluminous volcanic activity in the upper part of the northeast limb of the North Shore Volcanic Group. Sample from roadcut at the west edge of Grand Marais. (UTM 698364E, 5291847N) Sample MH047-AD – Six baddeleyite crystals selected for dissolution were all variably discordant, but gave equivalent 207Pb/206Pb dates with a weighted mean of 1095.94±0.62 (n=6; MSWD 0.37). This age falls within the range of published ages for various other units of the Beaver Bay Complex, including the Wilson Lake ferrogabbro (1095.75±0.92; Hoaglund, 2010), Sonju Lake intrusion (1096.1±0.8; Paces and Miller, 1993), Silver Bay ferrogabbro (1095.8±1.2; Paces and Miller, 1993), Pine Mountain granophyre (1095.3±3.8; Vervoort and others, 2007), as well as others. The Hovland sill occurs near the base of the upper northeast limb of the NSVG. The sample, a coarse prismatic olivine-pyroxene ferromonzonite that forms the upper differentiated cap to the Hovland sill, was collected from a roadcut on Highway 6, 1.3 miles northeast of the Brule River near Hovland. (UTM 722714E, 5301024N) Yavapai-interval intrusion – Sample 11-BUC-1-459.5. Seven zircon grains were selected for CA-TIMS analysis, and five of these seven analyses were concordant and equivalent, with a weighted mean 206Pb/238U date of 1779.93±0.56 (MSWD = 0.71), and a weighted mean 207Pb/206Pb date of 1782.31±0.64 Ma (MSWD 1.12). This drill core sample is from the heart of the Minnesota River Valley (MRV) subprovince in southern Minnesota, within a prominent north-south elongate magnetic low about 9 x 1.3 miles in dimension. The age of this granite places it in the Yavapai interval, along with the granites that make up the east-central Minnesota Batholith (1772 – 1800 Ma) and other mafic-intermediate intrusions along the southern Minnesota border emplaced into the MRV subprovince, which have been dated at ca. 1792 Ma (Southwick, 1994) and ca. 1760 Ma (Van Schmus, 2006). (UTM 352525E, 4890299N) 13 �Penokean Orogen – Sample HB5716-AD. Three abraded zircon fragments produced concordant isotopic ratios, with a weighted mean 206Pb/238U date of 1882.32±1.21 (MSWD = 0.23), and a weighted mean 207Pb/206Pb date of 1882.96±0.67 Ma (MSWD 1.12). This sample is from a metagabbroic sill encountered in the 7,440’ (2,268.3m) – long ‘Hattenberger’ core (HB-1) located within the Moose Lake-Glen Township panel, in the heart of the Penokean fold-and-thrust belt in eastcentral Minnesota. The drill hole intersected interlayered mafic volcanic and sedimentary rocks metamorphosed under lower amphibolite-grade conditions, as well as lesser proportions of mafic sills thought to be petrologically related to the mafic volcanic rocks (Southwick and others, 2005). The dated sample is from a coarse-grained, feldspathic zone in the interior of a 500-foot thick mafic sill that has chilled upper and lower margins. This sill mostly retains its primary igneous fabric despite the mafic mineral assemblage being composed almost entirely of blue-green metamorphic hornblende. UTM 503780E, 5149030N. The ca. 1,882 Ma age of this sill clearly predates the geon 17 Yavapai-interval ages of the voluminous east-central Minnesota batholith and the ca. 1858-1877 Ma Bradbury Creek granodiorite (Holm and others, 2005); but it postdates the 2,009 Ma Mille Lacs granite (Holm and others, 2005). It is barely older than the 1878.3±1.3 Ma lapilli tuff in the upper Gunflint Iron Formation (Fralick and others, 2002), and the 1,874±9 Ma Hemlock Volcanics interlayered with the Negaunee Iron Formation (Schneider and others, 2002). It overlaps with ages from the Pembine-Wausau portion of the Wisconsin Magmatic Terrane (1,860-1,889; Sims and others, 1989 as reported in Schulz and Cannon, 2007). The sill intruded a carbonate-rich unit tentatively correlated with the Denham Formation, which has a maximum depositional age of 2,072.7±17.9 Ma based on detrital zircons, (Vorhies, 2006). Sample 05BWS001 - Little Falls Formation detrital zircon – A drill core sample of staurolite-garnet schist (Little Falls Formation) yielded abundant detrital zircon grains with U-Pb ages from 1,833 to 2,784 Ma. Most grains have ages from 1,833 - 1,898 Ma; smaller numbers of grains yield clusters (>3 grains each) at 2,420, 2,668, 2,695, and 2,704 Ma. Analysis of the main cluster of ages suggests a separate group with an age of 1,844 Ma, considered to be the maximum depositional age (MDA). The preponderance of ca. 1844 Ma ages (46 of 90 grains analyzed) offers new insights into the long-standing debate about age of the Little Falls Formation. The Little Falls Formation covers a large area of east-central Minnesota, and based on geophysical data is thought to form the upper plate of a thrust sheet that has been ramped west-northwest over rocks correlative with the Penokean Mille Lacs Group. Although it is well known that the Little Falls Formation was affected by the widespread ca. 1760 Ma regional metamorphic event throughout east-central Minnesota (e.g. Holm and others, 2007), the actual depositional age has never gone beyond speculation. If the 1844 Ma maximum depositional age is verified by further studies, this would imply that the Little Falls Formation was deposited synchronous with the much lower grade Animikie basin and that the basin covered a much broader area than previously recognized, necessitating a rethinking of models of the geologic evolution of the Penokean and Yavapai orogens in central Minnesota. Drill hole 05BWS001 UTM 391454E, 5071884N. References: Boerboom, T.J., and Green, J.C., 2010, Minnesota Geological Survey Miscellaneous Map Series Map M-189, scale 1:24,000. Boerboom, T.J., and Green, J.C., 2013, Minnesota Geological Survey Miscellaneous Map Series Map M-195, scale 1:24,000. Fralick, P., Davis, D.W., and Kissin, S.A., 2002, CJES v. 39 no. 7, p. 1085-1091. Hoaglund, S., Miller, J.D., Crowley, J.L., and Schmitz, M.D., 2010, ILSG 56th Annual Meeting; Program and abstracts, p. 25-26. Holm, D.K., D.A. Schneider, D.A., Rose, S., Mancuso, C., McKenzie, M., Foland, K.A., and Hodges, K.V., 2007, Precambrian Research, V. 157, nos. 1-4, p. 106-206. Holm D.K., Anderson, R., Boerboom, T.J., Cannon, W.F., Chandler, V., Jirsa, M., Miller, J., Schneider, D.A., Schulz, K.J., and Van Schmus, W.R., 2007, Precambrian Research V. 157 p.71–79 Schneider, D., Bickford, M., Cannon, W., Shulz, K., and Hamilton, M., 2002, C.J.E.S., v. 39, p. 999–1012. Schulz, K.J., and Cannon, W.F., 2007, Precambrian Research, V. 157, p. 4-25. Sims, P.K., Van Schmus, W.R., Schulz, K.J., Peterman, Z.E., 1989, C.J.E.S., v. 26, 2145–2158. Sims, P.K., Schulz, K.J., Peterman, Z.E., 1992, US Geol. Surv. Prof. Pap. 1517, 65 pp. Southwick, D.L., 1994, Minnesota Geological Survey Report of Investigations 43, p. 1-19. Southwick, D.L., Morey, G.B., Christopher, J.M., McSwiggen, P.L., and Boerboom, T.J., 2005, MN. Geol. Survey R.I. 63, 63 p. Van Schmus, W.R., 2006, University of Kansas, Lawrence; Reported in Jirsa, M.A., Miller, J.D., Jr., Severson, M.J., and Chandler, V.W., 2006, Minnesota Geological Survey Open-File Report OF-06-03, 49 p. Vorhies, S., 2006, B.A. Honors Paper, Smith College, John Brady, Faculty advisor. 14 �BEDROCK GEOLOGIC MAP OF THE MARR ISLAND AND HOVLAND QUADRANGLES, NORTH SHORE OF LAKE SUPERIOR, MINNESOTA BOERBOOM, Terrence J., Minnesota Geological Survey, boerb001@umn.edu GREEN, John C., University of Minnesota-Duluth, jgreen@d.umn.edu The Minnesota Geological Survey has continued ongoing mapping of the bedrock geology of 7.5’ quadrangles adjacent to Lake Superior as part of the USGS STATEMAP program, resulting to date in twenty published 1:24,000 scale maps from Duluth northeast to beyond Hovland, in addition to 10 quadrangles already published under the former USGS COGEOMAP program. The Marr Island and Hovland quadrangles, (Boerboom and Green, 2013), are the most recent of these maps (Fig. 1). These maps are available at the MGS website (www.mngs.umn.edu). Outcrop mapping was augmented by nearly 60 sets of high-quality water well cutting samples, collected at 10 foot intervals by McKeever Well Drilling of Little Marais, Minnesota. This mapping has refined the volcanic stratigraphy of the North Shore Volcanic Group (NSVG) in this area, as well as details and extents of intrusions. In keeping with prior work, the NSVG is subdivided into informal lithostratigraphic packages (listed below), which on this map follow closely those units identified by Green, 2002. Most of the rocks in this quadrangle contain typical zeolite mineral assemblages; however in proximity to mafic intrusions the volcanic rocks contain minor garnet and/or epidote. Volcanic rocks—The volcanic rocks in this map area cross the boundary between the upper (normally-polarized) and lower (reversely-polarized) portions of the northeast limb of the NSVG. The volcanic rocks are intruded by mafic to felsic intrusions inferred to be mostly related to the Beaver Bay Complex in timing. The major lithostratigraphic units are listed below from lowest to highest in the volcanic stratigraphy. Hovland lavas—Predominantly strongly porphyritic trachyandesite, famous for its abundant and large tabular plagioclase phenocrysts. Also includes a poorly-mapped rhyolite known only from water well cuttings. Only the uppermost of the Hovland lavas are present here, most of the unit being to the northeast. Brule River lavas—Predominantly rhyolite, interlayered with variably porphyritic intergranular basalts, pigeonitic ferroandesite, ophitic basalt, and minor sandstone. The upper half of this sequence of lavas is made up of the quartz-and feldspar-phyric Devil’s Kettle rhyolite (1,097.7 ±1.7 Ma: Davis and Green, 1997); other rhyolites contain only feldspar phenocrysts, including the Big Bay rhyolite near the base of the upper normally-polarized lavas, which has a U-Pb zircon age of 1100.2±2 Ma (Davis and Green, 1997). Marr Island lavas—An approximately 1,000 meter thick sequence of dominantly mafic to intermediate lava flows (Green, 2002) that range from ophitic Fe-tholeiite to andesite with minor proportions of icelandite, rhyolite, and sandstone. Pigeonite is present in the basalts and andesites, and some andesites contain fresh glass in the mesostasis. Kimball Creek felsites—Icelandite and rhyolite; the bulk of this unit occurs to the west in the Kadunce River quadrangle (Boerboom and Green, 2011), and only the very base of the section reaches this map area. Intrusive rocks—Multiple intrusions range from coarse-grained gabbroic anorthosite, cumulate- differentiated sills, felsic to intermediate intrusions, and small diabase to ferrodiorite dikes and sills. The major intrusions are summarized below from earliest to latest in timing. Carlson Creek gabbro complex—Anorthositic gabbro with pods of gabbroic pegmatite, and felsic-intermediate rocks that may be related, which locally contain xenoliths of pure anorthosite. Intrusions tentatively assigned to the Brule-Hovland complex—Mainly ophitic gabbro and diabase but including erromonzodiorite hybrids. Ophitic olivine gabbro contains inclusions of volcanic rocks and interflow sandstone. Reservation River diabase—A gently dipping sheet-like body of ophitic olivine diabase that is present mainly to the north and east of the map area; intrudes reversely-polarized flows but is normally-polarized. Horseshoe Bay diabaseOphitic diabase (normally polarized), troctolitic diabase, and ferromonzonite. Troctolitic phase contains Fo62-72 olivine and augite of Mg#70; ophitic diabase contains Fo50-58 olivine and augite of MG#70. Orientations of olivine streaks and sheet joints indicate troctolite is a gently south-dipping sill; this possibly grades into the ophitic diabase. Ophitic diabase contains large inclusions of amygdaloidal basalt, and there are local narrow hybrid melt zones where ophitic diabase intruded rhyolite. Small plug-like bodies of prismatic pyroxene-quartz ferromonzonite may have formed as partial melt segregations from the underlying rhyolite. Chicago Bay ophitic diabasePresumably a sill (normally-polarized) beneath the Hovland sill; may be a marginal phase to it. Augite compositions of Mg#70 and partially olivine average Fo60. 15 �Brule River sillVariably granophyric mafic sill; the lowest part is a cumulate with locally abundant ilmenite and minor poikilitic olivine. Higher parts have layers bearing clots of poikilitic olivine alternating with more coarsegrained granophyric layers. The upper portion may grade into overlying Pine Mountain granodiorite. Lookout sillA south-dipping sill, like the Hovland sill in that it contains cumulate plagioclase, augite, olivine (Fo23 near base, Fo15 higher up), apatite, and abundant ilmenite. The upper part is a miariolitic prismatic ferromonzonite that contains fayalitic olivine (Fo10). Hovland sillAn approximately 15-degree south dipping, subcordant, 300 m-thick sill composed of a basal noncumulate ferrogabbro, a middle cumulate-foliated granophyric ferrogabbro, and an upper coarse-grained felsic cap. Not physically continuous with the Lookout sill but very similar and may be related in timing and paragenesis. Monazite U-Pb age of 1095.94±0.62 (Boerboom and others, this volume). The lower ferrogabbro contains distinguishing, evenly distributed 3-4mm altered olivine clots (5%), intergranular augite (Mg# 59 to 54), and up to 2% pigeonite (Mg# 41). The middle section is strongly cumulatefoliated and typically coarse-grained; the lowest part of this contains abundant cumulate ilmenite plates but higher in the stratigraphy magnetite becomes dominant over ilmenite; also contains cumulate plagioclase, augite, Fe-Ti oxides, olivine (mostly altered), and minor apatite. Olivine content is generally around 2%, but near the top increases to as much as 13%. Pigeonite (Mg# 49) rims augite, for which average Fe/Mg ratios increase from Mg# 57 at the base to 35 at the top (ferroaugite). Mg numbers for olivine range from Fo29 near the base to Fo17 near the top. The coarse-grained felsic cap contains 8-15% prismatic ferroaugite (Mg#35) and 10-15% fayalitic olivine (Fo10) that is mostly altered and varies in form from irregular coarse clots and prismatic grains to acicular trellises up to 30cm in length. Pine Mountain GranophyrePart of a larger body of granophyre and granodiorite located mainly to the west of this map area (e.g. Boerboom and Green, 2011); with a reported U-Pb age of 1,095.3 ± 3.8 Ma (Vervoort and others, 2007). Compositions are gradational from leucogranite into gabbro of the Brule River sill, implying that the Brule River Sill, and by extension, the Lookout sill, may all be close to 1,095 Ma in age. Miscellaneous intrusionsA wide variety of small intrusions are present throughout the map area. These include small hybrid/contaminated ferromonzonitic dikes with intermingled partially melted rhyolite, fine-grained ferrodiorite dikes, ophitic to intergranular olivine diabase, medium-grained pyroxene granodiorite, and ferromonzodiorite. Most of these are dikes, but some appear to be sills. References Boerboom, T.J., and Green, J.C., 2010, Minnesota Geological Survey Miscellaneous Map Series Map M-189, scale 1:24,000. Boerboom, T.J., and Green, J.C., 2013, Minnesota Geological Survey Miscellaneous Map Series Map M-195, scale 1:24,000. Davis, D.W., and Green, J.C., 1997, Canadian Journal of Earth Science, Volume 34, No. 4, April 1997, p. 476-488. Green, J.C., 2002, Minnesota Geological Survey Report of Investigations 58, p. 94-102. Vervoort, J.D., Wirth, K., and Kennedy, B., 2007, Precambrian Research, vol. 157, no. 1-4, p. 235-268. Figure 1. Simplified geologic map of the Marr IslandHovland quadrangles, showing the major lithostratigraphic units discussed in the text. Inset location map shows the extent of the Keweenawan Midcontinent Rift System in Minnesota, and the locations of the Marr IslandHovland quadrangles (dark gray box). 16 �GEOCHEMICAL INVESTIGATION OF THE ORIGIN OF THE BACK FORTY VOLCANOGENIC MASSIVE SULFIDE DEPOSIT IN MENOMINEE COUNTY, MI Anthony Boxleiter1, Joyashish Thakurta1, and Thomas O. Quigley2 1. Department of Geosciences, Western Michigan University, Kalamazoo, MI 49008, anthony.r.boxleiter@wmich.edu; 2 Aquila Resources Inc., Menominee, MI 49858 The Back Forty Volcanogenic Massive Sulfide deposit is located in Menominee County, Michigan. Several Volcanogenic Massive Sulfide (VMS) deposits can be found trending along the Penokean Volcanic Belt in northern Wisconsin and the Michigan Upper Peninsula (Fig. 1). The Back Forty deposit is a Paleoproterozoic ore deposit which formed during the Penokean Orogeny (1874 ± 4 Ma; Schulz et. al., 2007). This deposit is unique because it contains low amounts of copper and high amounts of zinc and gold when compared to other VMS deposits associated with the Penokean Volcanic Belt, such as Crandon and Flambeau. Mineralization of the Back Forty deposit consists of massive, semi-massive, stringer sulfide zones, and sulfide-poor Au and Ag enriched zones (Thakurta and Quigley, 2013). Three chemically distinct varieties of host rhyolite have been identified based upon trace element characteristics, two of which are found to host sulfide mineralization (Quigley et al., 2008). The relationship between rhyolite geochemistry and VMS mineralization has been proposed by Thurston (1981) and Campbell et al. (1982) as an exploration tool for discerning prospective VMS deposits, based on Archean VMS deposits in the Canadian Shield. From this groundwork, Lesher et al. (1986) and Barrie et al. (1993) developed a formal classification scheme for felsic volcanic rocks based on trace element concentrations and they suggested that certain types of rhyolites are more prospective for sulfide mineralization than others. They classified rhyolites associated with VMS deposits into three types: FI, FII, FIIIa, and FIIIb. Conclusions drawn from Lesher et al. (1986), Lentz (1998), and Hart et al. (2004) are that Archean VMS deposits are hosted mainly by FIII rhyolites, whereas most post-Archean VMS deposits are hosted predominantly by FII rhyolites. Under the classification scheme developed by Lesher et al. (1986) and Barrie et al. (1993), FI and FII type rhyolites are least favorable for VMS mineralization while FIIIa/FIIIb types have been proposed as the most prospective. FI type rhyolites appear to be particularly associated with gold-rich VMS deposits, such as the world-class Laronde deposit. The FI rhyolites are alkaline to calc-alkaline, with strongly fractionated REE patterns and strongly negative Ta and Nb anomalies. The FII rhyolites are calc-alkaline to transitional, with moderately fractionated REE patterns and moderate Ta and Nb anomalies and considered more favorable than FI rhyolites. The FIIIa and FIIIb rhyolites are tholeiitic and considered to have the greatest potential for hosting VMS deposits. The FIIIa rhyolites show weakly fractionated REE patterns and weak to nonexistent Nb and Ta anomalies. The FIIIb rhyolites are high-temperature rhyolites with flat REE patterns that lack Ta and Nb anomalies. (Gaboury and Pearson, 2008) Trace element analysis plotted as [La/Yb]CN versus YbCN for the two rhyolites hosting sulfide mineralization in the Back Forty deposit classifies these rhyolites as FI type under the scheme proposed by Lesher et al. (1986) and Barrie et al. (1993). Trace element analysis plotted as Zr/Y versus Y places these two rhyolites under the FII/FIIIa classification. The elements Zr, Y, La, and Yb are most useful for trace element analysis because they are generally immobile during hydrothermal alteration and are representatives of petrogenetic processes (Gaboury and Pearson, 2008). The FI classification of these rhyolites based upon [La/Yb]CN versus YbCN and the high amount of gold associated with the Back Forty is consistent with FI type association under this classification scheme. However, the trace element plot of Zr/Y versus Y places these rhyolites under the FII/FIIIa classification scheme. While this classification scheme has demonstrated the usefulness of rhyolite geochemistry for exploration in some areas, more work is required to characterize each type based on actual mineral deposits. For this reason, Gaboury and Person (2008) suggest that a combination of rhyolite geochemistry, volcanic facies, and the style of sulfide mineralization may be more meaningfully applied in exploration than rhyolite type alone. This is particularly important in the case of FI and FII rhyolites associated with VMS deposits of post-Archean age, such as the Back Forty deposit. 17 �This research will explore the use of not only rhyolite classification, but sulfur isotope analysis and petrographic techniques to characterize the Back Forty VMS deposit. This research will investigate the relationship between the pattern of distribution of sulfur isotopes in sulfide minerals of the Back Forty deposit, the mode of occurrence of the ore body, and textural characteristics of the sulfide ore minerals. Sulfur isotope values will help to characterize the distribution of sulfide minerals in the Back Forty deposit and to model the origin of sources. Sulfur isotope analysis may reveal episodic pulses of hydrothermal fluids as well as the source of sulfur (i.e., magmatic sulfur with δ34S values of 0 ± 2 per mil, biogenic sulfur with negative δ34S values, and surface-derived sulfur with positive δ34S). Sulfur isotope values measured in VMS deposits in other parts of the world, notably the Archean Kidd Creek VMS deposit in Ontario, Canada, indicate isotopic disequilibrium. In the study conducted on Kid Creek by Hannington et al. (2006), δ33S values in conjunction with δ34S values were used to model sulfur isotope systematics in Archean ore deposits. A similar study was conducted by the U.S. Geological Survey and U.S. Department of Interior (Taylor et al., 2010) on the Greens Creek VMS deposit located in Admiralty Island, Southeastern Alaska. Sulfur isotope analysis on the Greens Creek VMS produced δ34S values of 11 to -16‰ and has been interpreted by Taylor et al. (2010) as resulting locally from the organic reduction of seawater sulfate to H2S. Sulfur isotope analysis has never been used on the Back Forty deposit. The relationship of the mode of occurrence of sulfide mineral deposits at Back Forty with sulfur isotope signatures will provide important geochemical constraints on the origin of the deposit. This geochemical dataset will also be useful to model the origins of other VMS deposits in the Penokean Volcanic Belt and to explore for new economic sulfide deposits associated with rhyolitic host rocks. Fig. 1: Locations of VMS deposits along the E-W trend of the Penokean Volcanic Belt in northern Wisconsin. The Back Forty is the easternmost deposit of this trend and is the only VMS deposit found in the Michigan Upper Peninsula. References Barrie, C.T., Ludden, J.N., and Green, T.H., 1993. Geochemistry of volcanic rocks associated with Cu-Zn and Ni-Cu deposits in the Abitibi subprovince: Economic Geology, v. 88, p. 1341-1358. Campbell, I.II., Coad, P., Franklin, J.M., Gorton, M.P., Scott, S.D., Sowa, J., and Thurston, P.C., 1982. Rare earth elements in volcanic rocks associated with Cu-Zn massive sulfide mineralization. A preliminary report: Canadian Journal of Earth Sciences, v. 19, p. 619-623 Gaboury, D. and Pearson, V., 2008, Rhyolite geochemical signatures and association with volcanogenic massive sulfide deposits: Examples from the Abitibi Belt, Canada, Economic Geology, 103, 1531-1562 Hart, T.R., Gibson, H.L. and Lesher, C.M., 2004, Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic massive Cu-Zn-Pb sulfide deposits, Economic Geology, 99, 1003-1013 Hannington, M., Jamieson, J., Wing, B. and Farquhar J., 2006, Evaluating isotopic equilibrium among sulfide mineral pairs in Archean ore deposits: Case study from the Kidd Creek VMS deposit, Ontario. Economic Geology, 101. p. 1055-1061. Lentz, D.R., 1998. Petrogenetic evolution of felsic volcanic sequences associated with Phanerozoic volcanic-hosted massive sulphide systems: the role of extensional geodynamics: Ore Geology Reviews v. 12 p. 289-327. Lesher, C.M. Goodwin, A.M., Campbell, I.II., and Gorton, M.P., 1986. Trace-element geochemistry of ore-associated and barren, felsic metavolcanic rocks in the Superior province. Canada: Canadian Journal of Earth Sciences, v. 23, p. 222-237. Quigley, T., Mahin, B., and Aquila Field Office Geologic Staff, 2008, Back Forty Geology and Mineralization: 54th Annual Institute on Lake Superior Geology, Field Trip #3 Ross, C., Hudak, G., Morton, R., Quigley, T. and Mahin, R., 2011, Preliminary stratigraphy and physical volcanology associated with the Paleoproterozoic Back Forty VMS deposit, Menominee County, Michigan, Institute of Lake Superior Geology Schulz, K.J. and Cannon, W.F., 2007, The Penokean Orogeny in the Lake Superior region. Precambrian Research, 157, 4-25 Taylor, C.D. and Johnson, C.A., 2010, Editors. U.S. Geological Survey Professional Paper 1763, 429 p. Thakurta, J. and Quigley, T.O., 2013. Geochemical characterization of the Back Forty volcanogenic massive sulfide deposit in Menominee County, MI. Western Michigan University; Kalamazoo, MI. Thurston, P.C., 1981. Economic evaluation of Archean felsic volcanic rocks using REE geochemistry: Geological Society of Australia Special Publication 7, p. 439-450. 18 �ZIRCONIUM/HAFNIUM FRACTIONATION IN SOME PEGMATITES OF THE UPPER MIDWEST, USA Buchholz, T. W.1, Falster, A. U.2, and Simmons, W. B. 2 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494, 2Department of Earth and Environmental Sciences, University of New Orleans, New Orleans, Louisiana 70148 Zirconium and hafnium form a coherent pair, being of similar radius and chemistry, and mutually substitute in various Zr minerals. Zircon, nominally ZrSiO4, incorporates Hf levels reflecting the relative Hf contents of the crystalizing melt, which in turn reflects the degree of fractionation or evolution of the melt, rarely culminating (in highly evolved pegmatites) in the very rare Hf dominant analog hafnon. In the course of our studies of pegmatites in Wisconsin, Michigan and Minnesota, considerable data have been developed regarding chemistry of zircons from these pegmatites, in particular regarding Hf, which will be discussed below. HfO2 contents in zircons from source granites are generally low, in the range of 1-2 wt% (Fleischer, 1955, Wang et al, 2000), though higher levels have been reported in evolved granites (e.g. Wang, et al 2000). For local comparisons, a zircon from typical Nine Mile granite, Marathon Co, WI was analyzed, and found to contain 1.0-1.2wt% HfO2, zircons from the Zunker property (site of a former zircon mining attempt) in the Stettin Complex, Marathon Co, WI contained 1.8 to 2.2wt% HfO2, and zircons from the Bell Creek Granite (Marquette Co, MI) contained 3.8 to 4.3wt% HfO2. Results from Midwest pegmatite zircons are summarized below. The lists of associated minerals are not intended to be complete, but rather to suggest the degree of fractionation achieved, with emphasis on Mn/Fe, Nb/Ta, W, Sn and F. Waterloo Quarry, Jefferson Co, WI: (Small fractionated pegmatites emplaced in quartzites/metapelites, associated minerals: columbite-(Mn), tantalite-(Mn), microlite, gahnite, Bi): 8.1 to 10.9 wt% HfO2. Wimmer Pit #3, Marathon Co, WI: (Small fractionated pegmatite emplaced in Nine Mile Granite, associated minerals: cassiterite, tantalian cassiterite, columbite-group minerals, U-rich pyrochlore, and a U-niobate phase): 1.6 to 4.9 wt% HfO2, Hf-rich rims on zircon. Maguire Pit, Marathon Co, WI: (Greisenized pegmatites and aplites emplaced in Nine Mile Granite, associated minerals: huebnerite/ferberite, cassiterite, topaz, W-rich columbite-group minerals, zinnwaldite, fluorite, prosopite, cryolite): 1.8 to 10.4wt% HfO2, Hf enriched cores in and rims on zircon. Pegmatite #22, Koss Pit, Marathon Co, WI: (Small fractionated pegmatite emplaced in aplite body in Nine Mine Granite, associated minerals: columbite-group minerals, tapiolite-(Fe), cassiterite, microlite, monazite, xenotime-(Y), xenotime-(Yb), fluorite): 7.4 to 14.6 wt% HfO2, Hf-rich rims on Urich zircon. Woodland Drive Pegmatite, Marathon Co, WI: (Small unusually fractionated silica-saturated REE & Th-poor pegmatite emplaced in tabular syenite phase of the Stettin Complex, associated minerals: Ta-rich columbite-Fe, Ta-rich pyrochlore, cassiterite, ilmenite, unidentified phases): 3.9 to 8.5 wt% HfO2. Hoskin Lake pegmatite field (Florence Co, WI): (Moderate sized highly fractionated LCT pegmatites, associated minerals: elbaite tourmaline, tantalite-Mn, stibiotantalite, tantite, pegmatite phosphates, pollucite, rynersonite, behierite): 6.1 to 19.8wt% HfO2. Groveland Pegmatite, Dickinson Co, MI: (Small fractionated pegmatite emplaced in metasediments adjacent to the old Groveland Mine, associated minerals: columbite-(Fe), tantalite-(Fe), tapiolite-(Fe), microlite, gahnite, beryl): 1.7 to 6.5wt% HfO2. 19 �Black River Pegmatite, Marquette Co, MI: (Small pegmatite emplaced in Archean Bell Creek Gneiss, associated minerals: columbite-(Fe), late microlite, monazite-(Ce), synchysite-(Ce), synchysite(Y), topaz, fluorite): 3.9 to 5.3wt% HfO2. Orr Pegmatite, St Louis Co, MN: (Moderate sized pegmatite emplaced in biotite schist-rich migmatite, associated minerals: Mn-rich almandine garnet, magnetite, Th-rich monazite, possible chevkinite-(Ce), columbite-Fe, allanite-(Ce), titanite): 2.1 to 5.3wt% HfO2. Substantial enrichment in Hf is evident in many of the above pegmatites and is in accordance with observations made by Fleischer (1955) and Linnen & Kepler (2002). The enrichment present in the Woodland Drive pegmatite, emplaced in the alkalic Stettin Pluton, is particularly interesting as in all other Stettin zircon samples Hf contents are quite low. The highest HfO2 levels, from pegmatites in the Proterozoic Nine Mile Granite, are closely associated with highly fractionated Nb-Ta minerals, as well as high to very high F levels (as evidenced by late fluorite and other F-rich minerals). It is likely that high Hf levels in the evolved zircons are related to the relatively greater stability of Hf-F vs Zr-F complexes (Linnen & Kepler, 2002). Break down of Hf-F complexes triggered by the formation of F-rich minerals released Hf which was then incorporated into the outer zones of growing zircon crystals. The high levels of enrichment in HfO2 observed in the Florence County pegmatites can be attributed to the extreme level of fractionation achieved in these LCT pegmatites. F is present in various phases, but does not approach the high levels observed in the Nine Mile Granite and the Stettin Complex, and is unlikely to be the driving factor in this fractionation. References Fleischer, M. 1955. Hafnium Content and Hafnium/Zirconium Ratio in Minerals and Rocks. US Geological Survey Bulletin 1021-A Linnen, R. L. and Keppler, H. 2002. Melt composition control of Zr/Hf fractionation in magmatic processes. Geochimica et Cosmochimica Acta, 86, no. 18: 3293-3301. Wang, R.C., Zhao, G.T., Lu, J.J., Chien, X.M. and Wang, D.Z. 2000. Chemistry of Hf-rich zircons from the Laoshan I- and A-type granites, Eastern China. Mineralogical Magazine, 64/5: 867-877. 20 �A NEW OCCURRENCE OF THE SUDBURY IMPACT LAYER IN THE GOGEBIC IRON RANGE OF WISCONSIN CANNON, William F.1, WOODRUFF, Laurel G2., SAARI, Stacy M.3, and HAGSTROM, Molly C.3 1 U.S. Geological Survey, MS 954, Reston, VA 20192 wcannon@usgs.gov U.S. Geological Survey, 2280 Woodale Drive, Mounds View, MN 55112 3 Gogebic Taconite, LLC, 402 Silver Street, Hurley, WI 54534 2 Exploration drilling in 2014 by Gogebic Taconite, LLC in the western Gogebic Iron Range in northern Wisconsin provides seven new intersections of the Sudbury impact layer (SIL) (Fig. 1). Together with a hole drilled previously, they reveal features of the SIL along 6 km of strike length. Data presented here are observations of drill core and preliminary microscopic examination. The SIL lies at the expected stratigraphic position- at or very near the contact of the underlying Ironwood Iron-formation and Tyler Formation. The Ironwood-Tyler contact appears transitional as evenly bedded magnetic iron-formation of the Pence Member of the Ironwood grades upward into nonmagnetic laminated argillite of the lower Tyler Formation. The precise location of the contact is generally arbitrary; we have seen no indications of a hiatus in sedimentation between the Ironwood and Tyler. Although the Pence Member is predominantly evenly- and thinly-bedded at this locality, it contains a few interbeds of wavy bedded granular iron-formation suggesting that at the time of deposition of the SIL the area was submerged to depths only slightly below the maximum depth of surface wave agitation. The SIL here consists of ejecta having similarities to many other SIL localities reported previously in the Lake Superior region. The most definitive features are accretionary lapilli (Fig. 2), altered glass spherules and fragments of diverse character (Fig. 3), and a very sparse suite of quartz grains showing relict planar deformation features (Fig. 4). The rocks are somewhat metamorphosed so that biotite, chlorite, and sericite are common in the matrix. Secondary carbonate is widespread and obscures much of the original texture. Fragments of argillite are also common. Together, these lithologies vary in total thickness from about 20 m to only 0.1 m along the 6 km strike length studied to date. In some drill cores, a zone of seismically disrupted sediments underlies the ejecta. Several features suggest that the SIL is largely, or entirely, reworked ejecta mixed with more local sediments. One drill hole contains several clasts of lapillistone at least 5 cm diameter about 15 cm above the base of a 1.2 m thick ejecta layer (Fig. 2). The clasts appear to be original lapilli-bearing material that was lithified (or frozen?), fragmented, and redeposited in its present position. Many clasts of Tyler-like laminated argillite are included within ejecta and vary from thin wisps of apparently soft sediments to much thicker intervals. The thickest of the SIL drilled sections contains four intervals of distorted laminated argillite from 3 m to less than 1 m thick. A second drill core contains a 4 m interval composed of lapilli-bearing ejecta interlayered with five intervals of laminated argillite as much as 1.5 m thick. We interpret the argillite to be clasts of semiconsolidated Tyler Formation that were incorporated into debris flows composed originally of ejecta. This implies that a nearby elevated area existed onto which ejecta originally was deposited and subsequently slumped into the deeper water of this area. Such an elevated area probably existed only a few tens of kilometers to the east. The classic five-fold stratigraphy of the Ironwood, defined in the central and eastern Gogebic Range, includes the Anvil Member, a 21 �shallow water granular iron-formation that overlies the Pence Member in that area. Deposition of the Anvil was probably followed by uplift that raised the Anvil above sea level. The base of the Tyler in that area is a basal conglomerate (Pabst Member of the Tyler Formation) that marks a short-lived erosion interval between the Tyler and Ironwood (see Cannon and others, 2008 for a summary of previous publications on stratigraphic details). Our preliminary model, therefore, begins with deposition of the Sudbury ejecta, in part in a terrestrial setting, to the east of our study area. The ejecta deposit was partly lithified before being remobilized and traveling as debris flows that incorporated newly deposited argillite to the current depositional site. Thus, at least part of the SIL in the western Gogebic Range may not record the instant of impact, but rather is a younger deposit whose deposition was delayed sufficiently to allow partial lithification of the ejecta and deposition of at least a thin layer of Tyler argillite. Cannon, William F., LaBerge, Gene L., Klasner, John S., and Schulz, Klaus, J., 2008, The Gogebic Iron Range-a sample of the northern margin of the Penokean fold and thrust belt: U.S. Geological Survey Professional Paper 1730, 44 p. 22 �CONTINUED WORK ON USING THE HORIZONTAL-TO-VERTICAL SPECTRAL RATIO (HVSR) PASSIVE SEISMIC METHOD FOR DETERMINING QUATERNARY SEDIMENT THICKNESS IN MINNESOTA Val W. Chandler and Richard S. Lively Minnesota Geological Survey, 2642 University Ave., St. Paul, MN 55114 chand004@umn.edu Work has continued on using the horizontal-to-vertical-spectral ratio (HVSR or sometimes H/V) passive seismic method for determining the thickness of Quaternary sediments in Minnesota, which consist chiefly of Pleistocene glacial deposits. The HVSR method is used to estimate the primary resonant frequency (shear wave) of unconsolidated sediments. If the acoustic impedance (density*seismic velocity) at the sediment-bedrock contact differs by a factor of at least 2, and if this surface is reasonably flat, the thickness (z) of the unconsolidated materials can be estimated by the relationship: z=af0b Where f0 is the estimated primary resonant frequency, and a and b are parameters that are calibrated empirically for a given area from control points where a wide range of bedrock depths are known. Once calibrated, the relationship can be used to estimate depths at points lacking control. Due to the pronounced variations in shear-wave velocities for glacial deposits, calibrated parameters are reliable only within fairly localized areas, and multiple calibrations may have to be conducted for larger regions. At control points where z is known, the average shear-wave velocity (Vs) of the unconsolidated sediments can also be estimated. During the spring and summer of 2013 more than 425 passive seismic stations were acquired, bringing the total to over 1100 passive seismic stations in Minnesota and adjacent parts of Wisconsin (Figure 1). The most recent work has been focused in the “Arrowhead” area in northeastern Minnesota, in Kanabec County in east-central Minnesota, and along profiles that traversed parts of Becker, Clay, Hubbard, Todd and Wadena Counties in northwestern Minnesota (Figure 1) Considerable scatter is observed in the f0 vs z relationships at control points in the new study areas, implying that Vs values vary significantly, both laterally and vertically, and more than one depth calibration may be needed for each area. In addition, HVSR results were commonly poor in parts of the Arrowhead area where unconsolidated sediment was thin (<15 meters), most likely reflecting an irregular bedrock surface. In spite of these limitations, the HVSR method was nonetheless useful in mapping the thickness of Quaternary sediments in both the Arrowhead and Kanabec County areas. In Kanabec County the HVSR-results were combined with drillhole and gravity data to produce residual gravity data that further helped in mapping the thickness of Quaternary sediments. Preliminary analysis of HVSR data in northwestern Minnesota indicates good results in Clay County, and in western Becker, northern Hubbard, northern Todd, and southern Wadena Counties, whereas generally poor results were observed elsewhere. Further work is being planned for the northwestern part of the state this summer. 23 �In summary the HVSR passive seismic method continues to be a very useful tool for estimating the thickness of Quaternary sediments in Minnesota and adjacent areas, provided the appropriate cautions are heeded. In some situations the HVSR methods will provide a suitable and much cheaper alternative to conventional seismic sounding, and when not, it will at least help in prioritizing and targeting areas where conventional seismic sounding is necessary. Figure 1. Map showing all passive seismic stations that have been acquired in Minnesota and adjacent parts of Wisconsin through the summer of 2013 (red circles). Stations highlighted in white represent control points where bedrock depth is known through either drillholes or seismic soundings. 24 �EVIDENCE OF SIMULTANEOUS BRITTLE AND DUCTILE DEFORMATION IN THE MAIN BREAK FAULT SYSTEM IN KIRKLAND LAKE, ON L. B. Clapp and M. L. Hill Department of Geology, Lakehead University, 955 Oliver Rd., Thunder Bay, Ontario, P7B 5E1, lbclapp@lakeheadu.ca Microstructural analysis provides evidence of significant ductile deformation concentrated along the main break, in Kirkland Lake, ON. The Main Break is an east-west striking mineralized fault system that has sustained multiple gold mines since its discovery in 1911. It is found in the southern Abitibi gold belt on Kirkland Lake Gold Inc.’s Lakeshore Mine property, which is a structurally controlled orogenic gold deposit. Oriented samples for microstrucural analysis were collected from two new 1-meter long channels across the main break spaced eight meters apart (Fig.1). Samples were examined in transmitted and reflected light microscopy. Evidence of ductile deformation by dislocation creep to produce mylonite includes porphyroclasts of potassium feldspar in an extremely fine-grained matrix, mineral alignment, undulatory extinction and subgrains in potassium feldspar, as well as lenticular aggregates of feldspar and muscovite showing dextral shear sense (Fig.2). Dislocation creep occurs in feldspar during ductile deformation at temperatures in the amphibolite facies of metamorphism or higher which puts deformation of the Main Break at temperatures above 400°C. Sericite aggregates replacing potassium feldspar likely enhanced grain softening during this ductile deformation process. Sericite as well as evidence of pressure solution along grain boundaries indicate the presence of a hydrous fluid during deformation. The most significant evidence of simultaneous brittle-ductile deformation is a potassium feldspar porphyroclast with strong undulatory extinction and subgrains in one half of the grain and microfractures in the other half (Fig.3). We conclude that the main break is a narrow ductile shear zone with minor brittle deformation. Figure 1. Main Break outcrop with channel samples shown in black lines 25 �Figure 2. Lenticular aggregates of feldspar and muscovite showing dextral shear sense Figure 3. K-spar porphyroclast showing mutually overprinting brittle and ductile deformation 26 �GEOLOGY AND GEOCHEMISTRY OF THE MESOPROTEROZOIC BADWATER INTRUSIVE COMPLEX, ONTARIO: IMPLICATIONS FOR GEON 15 MAGMATISM CUNDARI, Robert1, SMYK, Mark1 and HOLLINGS, Peter2 1 Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, 435 James St. S., Suite B002, Thunder Bay, ON, P7E 6S7 Canada 2 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The Mesoproterozoic Badwater Intrusive Complex (a.k.a. Waweig Troctolite Complex; cf. Borradaile and Middleton, 2006) intrudes Archean Wabigoon Subprovince country rocks, 13 km southwest of Armstrong, Ontario. This proposed complex comprises the Badwater Gabbro (BG) and the Badwater Syenite (BS). It is believed to form a multi-phase, intrusive complex which is expressed by a circular magnetic anomaly 12 km in diameter (Borradaile and Bennett, 2008). Initial mapping by MacDonald (2004) identified a variety of intrusive rocks, ranging from gabbro to quartz monzonite and syenite. They are unconformably overlain and largely obscured by Pillar Lake volcanic rocks which were possibly erupted at 1129 ± 4.6 Ma (U-Pb age from titanite) which would then likely represent early Midcontinent Rift magmatism (Heaman et al., 2007; Smyk et al., 2011). The poorly exposed BG was first recognized in 2000 by East West Resources Corp. (Middleton 2004) and was tested for PGE mineralization in drilling campaigns carried out in 2004 and 2008. The BG is described as a layered troctolite-gabbro complex consisting of olivine gabbro, anorthosite, troctolite, glomeroporphyritic rocks and layers of magnetite with sulphides (Middleton and Bennett, 2008). Magmatic layering dips ~45° to ~55° southeast. Modal mineralogy for typical olivine gabbro is listed as: plagioclase (labradorite/bytownite) 55%; clinopyroxene (augite?) 25%; biotite 10%; olivine (partly relict) 3%; talc/sericite (after olivine) 2%; amphibole (secondary actinolite) 2%; opaque (magnetite? pyrrhotite?) 2%; clay?/sericite (after plagioclase) trace (Middleton and Bennett, 2008). The BG is undeformed and displays generally fresh plagioclase and relatively unaltered olivine. The Badwater Syenite likely represents a multi-phase intrusion, characterized by intrusive breccias and hybrid rocks resulting from assimilation and contamination. Syenite dykes crosscut BG and BG xenoliths occur in syenite. High-precision U-Pb dating of baddeleyite yielded an emplacement age of 1598.7 ± 1.1 Ma for the BG and a U-Pb zircon age of 1590.1 ± 0.8 Ma for the BS, which supports observed cross-cutting relationships (Heaman et al., 2007). A possible genetic relationship between the BG and the BS can be tested using geochemistry. The BG is geochemically distinct from the BS on primitive mantle-normalized diagrams. The BG shows a flatter REE pattern with slight LREE enrichment, moderate HREE fractionation (Gd/Ybn = 2.4 to 3.9) and pronounced negative Zr and Hf anomalies (Fig. 1A). The BS shows a steeper REE pattern characterized by strong LREE enrichment, weak HREE fractionation (Gd/Ybn = 1.3 to 2.2) and pronounced negative Eu and Ti anomalies (Fig. 1B). It should be noted that two BG samples (03CAM115 and 03CAM305) were taken from mafic phases within the BS on the shore of Pillar Lake (not from within the main body of the BG) and display lower Gd/Ybn ratios than those taken from the two BG outcrops north of Pillar Lake (BW-01 and BW-02) which display Gd/Ybn ratios of 3.88 and 3.90, respectively. The trace element patterns for the BS show distinct similarities to those for the nearby 1546.5 ± 3.9 Ma (Heaman et. al., 2007) English Bay granite-rhyolite complex (EBC) (Fig. 1C). 27 �Based on similar trace element geochemistry, it would appear that the BS was derived from a similar source to the EBC, despite the 50 m.y. gap between the two units, whereas the BG appears to be sourced from a deeper source region. Hollings et al. (2004) suggested that the anorogenic EBC was derived from a mantle plume and it recorded the northern portion of a Mesoproterozoic plume track which produced anorogenic granites throughout North America. If the BS and the EBC are genetically related, the plume would have been attached to the base of the lithosphere for ~50 m.y. before detaching to create the anorogenic granites to the south in the United States. The BG could represent an early expression of the plume emplaced through a lithospheric-scale structure allowing for the tapping of a deeper-seated source. Alternatively, it may represent an earlier, unrelated plume that exploited the same structures as the EBC. Further work will elucidate intrusive relationships and possible regional associations. References Borradaile, G.J. and Middleton, R.S. 2006. Proterozoic paleomagnetism in the Nipigon Embayment of northern Ontario: Pillar Lake Lava, Waweig Troctolite and Gunflint Formation tuffs. Precambrian Research 144: 6991. Heaman, L.M., Easton, M., Hart, T.R., Hollings, P., MacDonald, C.A. and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario; Canadian Journal of Earth Sciences 44: 1055-1086. Hollings, P., Fralick, P. and Kissin, S. 2004. Geochemistry and geodynamic implications of Mesoproterozoic English Bay granite-rhyolite complex, northwestern Ontario; Canadian Journal of Earth Sciences 41: 13291338. MacDonald, C.A. 2004. Precambrian geology of the south Armstrong-Gull Bay area, Nipigon Embayment, northwestern Ontario; Ontario Geological Survey, Open File Report 6136, 42p. Middleton, R.S. 2004. Diamond drilling on Red Granite property, Pillar Lake sheet, Armstrong, Ontario; unpublished assessment file, Thunder Bay North District, Thunder Bay, 58p. Middleton, R.S. and Bennett, 2008. Drill Report, Armstrong, (Red Granite) property, Pillar Lake area, Thunder Bay Mining Division, Ontario; unpublished assessment file, Thunder Bay North District, Thunder Bay, 125p. Smyk, M., Hollings, P., and Cundari, R. 2011. The Pillar Lake Volcanics: new insights into an enigmatic Mesoproterozoic suite near Armstrong, Ontario. 58th Institute on Lake Superior Geology, Annual Meeting, Ashland, WI, May 18-21, 2011, Proceedings Volume 57, Part 1, p.75-76. 28 �PETROLOGY OF THE LAYERED NORTH LAC DES ILES INTRUSION, ONTARIO; PART I. STRATIGRAPHY AND MINERAL-CHEMICAL EVIDENCE FOR MULTIPLE MAGMA INJECTION Djon, M. L.1, Olivo, G.R.1, Miller, J.D.2 and Stewart, R. D3. 1 Queen's University, Department of Geological Sci. and Geological Engineering, Kingston, Ontario K7L 3N6 Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812 3 North American Palladium Ltd, 10th Avenue, Thunder Bay, Ontario, P7B 2R2. 2 The North Lac Des Iles Intrusion (NLDI-I) is a large, multi-phase, layered ultramafic intrusive component of the Lac des Iles Complex located in the northwest Ontario, which also includes the mafic Mine Block intrusion that has been extensively studied because of the palladium mining over two decades. The NLDII is an extensive (~25km2) well-layered tadpole-shaped complex, characterized by a series of nested bodies emplaced into the Archean tonalitic basement rocks of the Central Wabigoon Subprovince (McCracken et al., 2014). Two major intrusive centres, the Northern Ultramafic Centre (NUC) and the Southern Ultramafic Centre (SUC), are distinguishable based on dominant lithologies and attitudes of layering (Sutcliffe and Sweeney, 1986; Brugmann et al., 1989; Gupta et al., 1990; Brugmann et al., 1997). The focus of this study is the eastern limb of the NUC, which is a nearly circular funnel-shaped body with a mean diameter of approximately 4 km (Stone et al., 2003). Gravity modelling indicates that the NUC body thickens southward from a thickness of 1 km at the north end of NLDI to 3 km at the contact with the SUC (Gupta et al., 1990). North American Palladium’s integration of historical geological data with recent geophysical surveys shows that the magmatic layering is concentric and is preserved in the eastern part of northern centre but, is disrupted in the west by several smaller bodies of irregular to semielliptical shape. Bedrock mapping and reconstruction of 1.5 km long stratigraphic traverse of the eastern flank of the NUC show that it consists of a shallow, west- to northwest-dipping, layered sequences of ultramafic cumulate assemblages composed of olivine, chromite, clinopyroxene (augite) and orthopyroxene (bronzite). These sequences are characterized by regular repetitions of cumulate assemblage that define cyclic units that range from 50 to 250 meters in thickness. The upward progression of a typical cyclic unit is a basal olivine-chromite cumulate (dunite) grading into an olivine-bronzite-augite ± chromite cumulate (olivine websterite) and capped by a locally feldspathic augite-bronzite cumulate (websterite). In some cyclic units, however, the olivine websterite cumulate is absent. Contacts between cumulate assemblages within cyclical units are gradational over 0.5 to 3m to less commonly sharp. Minor plagioclase commonly occurs as an intercumulus phase in the websterite, but locally is abundant (up to 70 modal %) and can display cumulus texture in thin lenticular intervals. Although cyclic units appear throughout the NLDI-I, cyclic units in the lower part of layered sequence are dominated by websterite assemblages and are thus designated as the Pyroxenite Zone. This zone is composed of three, thick (200-250m) cyclic units that have websterite/dunite unit thickness ratios averaging 3:1. The upper cyclic units, in contrast, tend to be thinner (20-175m) and are dominated by olivine-bearing intervals (websterite:dunite unit thickness ratios average 2:3. This sequence is designated as the Peridotite Zone. Electron-microprobe analyses of cumulus olivine, chromite, and pyroxene compositions from drill core NL12-100, which profiles the two zones, are shown in Figure 1. The total ranges of compositions are not significantly different between the two zones, implying that the composition of the parent magma was likely the same for both sequences. Although a cyclical cryptic variation is evident throughout the core and consistent among the different cumulus phases, the breaks in mineral composition (dashed lines in Fig. 1) do not consistently correlate with cyclical boundaries. If this cryptic variation was due to repeated magma recharge pulses followed by fractional crystallization, the mineralogically most primitive dunite cumulate units would be expected to have the highest mg# (=MgO/(MgO+FeO), mole%) and the upper websterites to be the most evolved. However, this type of cryptic variation is clearly evident only in the 29 �upper (third) cycle of the Pyroxenite Zone. The cryptic break at the boundaries of cyclic units 1-2 and 3-4 is displaced upward from the lithologic contact. The density of data in the Peridotite Zone cycle is not sufficient to evaluate a correlation between cumulus phase layering and cryptic layering. The cause of the deviations from expected correlations between cumulus phase layering and cryptic layering and other petrologic aspect of the NLDI-I stratigraphy are still under investigation. Some possible processes being evaluated include: 1) variations in trapped liquid shift wherein primitive cumulus compositions are reset to lower mg#s by reequilibration with intercumulus liquid; 2) differences in the partitioning of MgO and FeO between cumulus olivine, clinopyroxene, and orthopyroxene (MgO is more compatible in pyroxenes than olivine); 3) changes from eutectic to peritectic relations between olivine and orthopyroxene, which could explain why olivine websterite cumulates are not always present;; and 4) contrasting densities of hot, primitive recharging magma and cooler, evolved resident magmas. If the recharging magma is denser than the resident magma, it should intrude beneath the resident magma and produce a sharp phase and cryptic change. If the recharging magma density is lower than the resident magma, it will plume into the chamber and would result an abrupt phase change and a more gradual cryptic shift to more primitive compositions. The relative volumes of recharging and resident magmas will also control the phase and mineral chemical effects. Figure 1: Cryptic variation shown by olivine, chromite, clinopyroxene, and orthopyroxene in the Peridotite and Pyroxenite zones of the Northern Ultramafic Center of the North Lac des Iles Complex. References Brügmann, G.E., Naldrett, A.J., Macdonald, A.J., 1989, Magma Mixing and Constitutional Zone-Refining in the Lac-Des-Iles Complex, Ontario - Genesis of Platinum-Group Element Mineralization: Economic Geology, v. 84, p. 1557-1573. Brugmann, G.E., Reischmann, T., Naldrett, A.J. and Sutcliffe, R.H. 1997. Roots of an Archean volcanic arc complex: The Lac des Iles area in Ontario, Canada; Precambrian Research 81, p. 223-239. Gupta, V.K. and Sutcliffe, R.H. 1990. Mafic–ultramafic intrusives and their gravity field: Lac des Iles area, northern Ontario; Geological Society of America Bulletin, v.102, p.1471-1483. McCracken, T., Kanhai, T., Bridson, P., McBride, W. R., Small, K., Penna, D., Technical Report Lac Des Iles Mine, Ontario, 2014. Stone, D., Lavigne, M.J., Schnieders, B., Scott, J., Wagner, D., 2003, Regional geology of the Lac des Iles area: Ontario Geolgical Survey Open File Rep 6120:15-1, p. 15-25. Sutcliffe, R.H. and Sweeny, J.M., 1986. Precambrian Geology of the Lac des Iles Complex, District of Thunder Bay, Ontario. Ontario Geological Survey, Map 3047, Geological Series-Preliminary Map, scale 1:15840. 30 �STRAIN ANALYSIS ON THE MAX LAKE POLYMICTIC CONGLOMERATES IN THE WABIGOON SUBPROVINCE, ONTARIO, CANADA Simon Dolega and Mary Louise Hill Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario P7B 5E1 Canada (sdolega@lakeheadu.ca) The Max Lake polymictic conglomerates are exposed near Highway 527, about 88 km north of the intersection with Highway 11-17. The conglomerates are part of the Beardmore-Geraldton belt in the Wabigoon Subprovince, Superior Province of Ontario, Canada. The Rf/Phi method for initially elliptical objects was used to estimate the amount strain on the conglomerates. In the Max Lake conglomerates, chlorite-actinolite clasts are more deformed than amphibolite clasts, which are more deformed than granitoid clasts. Heterogeneous strain also occurs among different outcrops. The overall amount of strain is lower where larger, more abundant and more competent clasts are found. Petrographic and microstructural analyses were used to determine the peak metamorphic grade preserved by each clast in the polymictic conglomerate. The matrix of the conglomerate and the chlorite-actinolite clasts preserve a peak metamorphic mineral assemblage stable in the greenschist facies. The amphibolite clasts preserve a peak metamorphic mineral assemblage stable in the amphibolite facies. The preservation of the amphibolite facies metamorphic mineral assemblage in the amphibolite clasts indicates that these clasts were derived from a metamorphic terrane. 31 �32 �PRELIMINARY INTERPRETATION OF PRECAMBRIAN LITHOLOGY AND STRUCTURE FROM HIGH-RESOLUTION, MULTI-METHOD GEOPHYSICS, NORTHEAST IOWA AND SOUTHEAST MINNESOTA DRENTH, Benjamin1, ANDERSON, Raymond2, SCHULZ, Klaus3, CHANDLER, Val4, CANNON, William3, BLOSS, Benjamin1, BEDROSIAN, Paul1, FEINBERG, Joshua M.5, and McKAY, Robert6 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 2 Dept. Earth and Environmental Sciences, Univ. Iowa, Iowa City, IA, 52242 3 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA, 20192-6320 4 Minnesota Geological Survey, 2642 University Avenue W., St. Paul, MN, 55114-1032 5 Dept. Earth Sciences, Univ. Minnesota, 310 Pillsbury Dr. SE, Minneapolis, MN, 55455-0219 6 Iowa Dept. Natural Resources—Iowa Geological and Water Survey, Iowa City, IA, 52242 Large amplitude gravity and magnetic highs over northeast Iowa are interpreted to reflect a buried intrusive complex composed of mafic/ultramafic rocks, the northeast Iowa Intrusive Complex (NEIIC), intruding Yavapai Province (1.8-1.72 Ga) rocks. The age of the complex is unproven, although it has been considered to be Keweenawan (~1.1 Ga). Because only four boreholes reach the complex, which is thought to be covered by 200-700 m of Paleozoic sedimentary rocks, geophysical methods are critical to developing a better understanding of the nature and mineral resource potential of the NEIIC. An initial airborne data collection campaign in the region of Decorah, Iowa, included high-resolution magnetic, gravity gradient (AGG), and time-domain electromagnetic (TDEM) data. Geophysical interpretations are presented in the form of a preliminary geologic map of the basement Precambrian rocks (Fig. 1), largely constructed by interpreting lithologies and cross cutting relationships expressed in magnetic and AGG anomalies. Numerous magnetic anomalies are coincident with AGG highs, indicating widespread strongly magnetized and dense rocks of likely mafic/ultramafic composition. A Yavapai age (Van Schmus et al., 2007) metagabbro unit (Ymg) is interpreted to be part of a layered intrusion with subvertical dip, and is thought to be among the oldest rocks present in the survey area. Another presumed Yavapai-age unit (Ysp) has low density and weak magnetization, observations consistent with granitic plutons. Northeast-trending, linear magnetic lows are interpreted to reflect reversely magnetized diabase dikes, and modeling shows that the anomalies are consistent with Keweenawan magnetization. The interpreted dikes are cut in places by normally magnetized mafic/ultramafic rocks (mi), suggesting that the latter represent younger Keweenawan rocks. Large magnetic highs without coincident AGG highs are interpreted to reflect intermediate or silicic intrusive rocks (ii). Distinctive horseshoe-shaped magnetic and AGG highs correspond with a known gabbro (dg, undated), and surround rocks with weaker magnetization and lower density (dwm). Here called the Decorah Complex, the source body has notable geophysical similarities to Keweenawan alkaline ring complexes, such as the Coldwell and Killala Lake Complexes, and Mesoproterozoic anorogenic complexes, such as the Kiglapait, Hettasch, and Voisey’s Bay intrusions in Labrador. Most units are cut by suspected northwest-trending faults imaged as magnetic lineaments, and one produces apparent sinistral fault separation of a dike in the eastern part of the survey area. The location, trend, and apparent sinistral sense of motion are consistent with the suspected faults being part of the Belle Plaine fault system, a complex transform fault zone within the Midcontinent Rift System. The TDEM data fail to directly image Precambrian rocks, due to their large burial depth. However, most of the overlying sedimentary section is well imaged, and depths to Precambrian rocks are estimated at 400-600 m based on preliminary TDEM interpretations and downward extrapolation of known stratigraphy. This interpretation is consistent with the limited borehole data. Reference Van Schmus, W.R., Schneider, D.A., Holm, D.K., Dodson, S., and Nelson, B.K., 2007, New insights into the southern margin of the Archean-Proterozoic boundary in the north-central United States based on U-Pb, SmNd, and Ar-Ar geochronology: Precambrian Research, v. 157, p. 80-105. 33 �Figure 1: Preliminary geologic map of Precambrian crystalline rocks, interpreted from limited boreholes and high-resolution airborne gravity gradient and magnetic data. 34 �Structural and Kinematic Analysis of the Shagawa Lake Shear Zone: Implications for Archean Tectonic Processes in the Southern Superior Province (Part 2 of 2) DYESS, Jonathan and HANSEN, Vicki, Department of Geological Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth MN 55812 The Archean (3.85-2.5 Ga) Superior Province, to a first approximation, consists of a series of east-west trending subprovinces of supracrustal rocks (greenstone belts) and granitoid rocks (e.g., Percival et al., 2007, and references therein). The Wawa Subprovince, southern Superior Province, is widely interpreted as a transpressional margin with shear zones recording unidirectional dextral strike-slip along the subprovince boundary (Hudleston et al., 1988; Bauer and Bidwell, 1990; Schultz-Ela and Hudleston, 1991), an interpretation held up as fundamental evidence for Archean plate-tectonic processes (Sleep, 1992). Others interpret these shear zones as recording dominantly oblique- to dip-slip shear possibly formed during greenstone sagduction between rising granitoid diapirs (Erickson, 2008, 2010; Wolf, 2006; Goodman, 2008; Karberg, 2009). Differing interpretations invoke different assumptions about non-coaxial shear direction. Due to the proximity of the Shagawa Lake shear zone to the Wawa subprovince boundary, structural and kinematics fabrics recorded within the Shagawa Lake shear zone have direct implications for crustal assembly of the southern Superior Province. If the Shagawa Lake shear zone records significant unidirectional strike-slip, then supported plate-tectonic models for Wawa Subprovince formation will be further constrained. If the Shagawa Lake shear zone does not record significant unidirectional strike-slip, then existing plate-tectonic and structural models of terrane amalgamation along the Southern Superior Province require reevaluation. We conducted a structural and kinematic analysis of the Shagawa Lake shear zone in three phases: 1) analysis of regional tectonic fabrics through Light Detection and Ranging altimetry data; 2) structural analysis of outcrop-scale structures through detailed field mapping; and 3) thin-section kinematic analysis. The Shagawa Lake shear zone contains a regional subvertical metamorphic foliation with an average strike of 065 but varies locally from 065 to 100. Two types of elongation lineation occur within the Shagawa Lake shear zone. These include ridge-in-groove striations on C-foliation surfaces (Lc) and stretching lineations on Ssurfaces (Ls) (Lin and Williams, 1992; Lin et al., 2007). Lc and Ls plunge steeply to obliquely, with local zones of shallow plunge, and non-coaxial shear direction is sub-parallel to elongation lineation (Dyess et al., 2014). Thus non-coaxial shear was dominantly dip- to oblique-slip with localized strike-slip. Microstructures, within foliation-normal, lineation-parallel sections, record both north-side-up and south-side-up shear in different samples. Samples with oblique lineation commonly record an apparent dextral strike-slip shear-sense despite varying lineation orientation. Our data indicate the Shagawa Lake shear zone experienced both N-side-up and Sside-up dip- to oblique-slip with relatively minor apparent dextral strike-slip and does not record significant unidirectional strike-slip as required by accepted plate tectonic models. 35 �References Bauer, R. L and Bidwell, M. E., 1990. Contrasts in the response to dextral transpression across the Quetico-Wawa subprovince boundary in northeastern Minnesota. Canadian Journal of Earth Sciences, 27, 1521-1535. Dyess, J.E., Hansen, V.L., Goscinak, C., 2014. Determination of vorticity in Neoarchean tectonites (Part 1 of 2). Institute on Lake Superior Geology annual meeting, Hibbing, MN. Erickson, E., 2008. Structural and kinematic analysis of the Shagawa Lake shear zone, Superior Province, northeastern Minnesota. M.S. Thesis, University of Minnesota Duluth, MN. Erickson, E., 2010. Structural and kinematic analysis of the Shagawa Lake shear zone, Superior Province, northern Minnesota: implications for the role of vertical versus horizontal tectonics in the Archean. Canadian Journal of Earth Sciences, 47, 1463-1479. Goodman, S., 2008. Structural and Kinematic Analysis of the Kawishiwi Shear Zone, Superior Province. M.S. Thesis, University of Minnesota Duluth, MN. Hudleston, P.J., Schultz-Ela, D., Southwick, D. L., 1988. Transpression in an Archean greenstone belt, northern Minnesota. Canadian Journal of Earth Sciences, vol 25, 10601068. Karberg, S M., 2009. Structural and Kinematic Analysis of the Mud Creek Shear Zone, Northeastern Minnesota. M.S. Thesis, University of Minnesota Duluth, MN. Lin, S., Williams, P.F., 1992. The origin of ridge-in-groove slickenside striae and associated steps in an S-C mylonite. Journal of Structural Geology 14, 315e321. Lin, S., Jiang, D., Williams, P., 2007. Importance of differentiating ductile slickenside striations from stretching lineations and variation of shear direction across a high-strain zone. Journal of Structural Geology, 29, 850-862. Percival, J.A., 2007, Geology and metallogeny of the Superior Province, Canada, in Goodfellow,W.D., ed.,Mineral Deposits of Canada:ASynthesis ofMajor Deposit-Types, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods: Geological Association of Canada, Mineral Deposits Division, Special Publication No. 5, p. 903-928. Schultz-Ela, D.D., Hudelston, P.J., 1991. Strain in an Archean greenstone belt of Minnesota. Tectonophysics, 190, 223-268. Sleep, N., 1992. Archean plate tectonics: what can be learned from continental geology?. Canadian Journal of Earth Sciences, 29, 2066-2071. Wolf, D. E., 2006. The Burntside Lake and Shagawa/Knife Lake shear zones: Deformation kinematics, geochemistry and geochronology; Wawa Subprovince, Ontario, Canada. Masters Thesis, Washington State University. 36 �Determination of Vorticity in Archean Tectonites (Part 1 of 2) DYESS, Jonathan, HANSEN, Vicki, and GOSCINAK, Christopher, Department of Geological Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth MN 55812 There is no consensus about the processes responsible for the formation of Archean crust. The Superior Province of North America is widely interpreted as a series of accreted terranes with subprovinces representing individual terranes (Talbot, 1973; Goodwin and Ridler,1970; Langford and Morin, 1976; Dimroth et al., 1983a, b; Ludden et al., 1986; Sylvester et al., 1987). The Neoarchean (2.7-2.5 Ga) Wawa Subprovince (Fig. Location) forms a NE-trending belt of sub-greenschist to greenschist facies supracrustal rocks cut by multiple shear zones marked by a well-developed metamorphic foliation (Fm) and elongation lineation (Le). The Wawa is interpreted as a transpressional plate-margin with significant dextral strike-slip displacement (Hudleston et al., 1988; Schultz-Ela and Hudleston, 1991). This interpretation is held up as evidence for Archean plate-tectonic processes based on plate-tectonic models that require strike-slip shear zones more than 1200 km of long (Sleep, 1992). Despite apparent broad acceptance of this interpretation, the nature of shear zone deformation within the Wawa remains poorly constrained. Displacement direction and magnitude is genetically linked to vorticity, marked by the vorticity-normal-section (VNS) and vorticity axis (pole to VNS), within L-S tectonites (Passchier, 1998; Xypolias, 2010 and references therein). However, geometric relationships between displacement direction and macroscopic structures, such as Le, can vary depending on shear zone kinematics (Passchier, 1998). Le can form parallel, perpendicular, or oblique to displacement direction during L-S tectonite formation. Therefore determination of L-S tectonite vorticity requires careful consideration before interpretation of displacement direction. A determination of the vorticity axis and observation of the geometric relationships between the VNS and Le can allow for the use of Le as a reference to noncoaxial shear direction. In this contribution, we determine the vorticity for seven samples from two Neoarchean shear zones from the Wawa subprovince. We use a combination of thin-section, X-Ray Computed Tomography, and quartz petrofabric data. We demonstrate that the vorticity axis lies approximately within the Fm and is normal to Le and that the VNS lies approximately within Fm-normal, Le-parallel planes for all seven samples. Thus non-coaxial displacement direction is sub-parallel to Le. Regional Le orientation varies within the two shear zones ranging from steeply to obliquely plunging, with local zones of shallow plunge (Hudleston, 1976; Hudleston et al, 1988; Bauer and Bidwell, 1990; Jirsa, et al., 1992; Goodman, 2008; Erickson, 2010; Johnson, 2009; Karberg, 2009). Data indicate that noncoaxial shear direction is sub-parallel to Le regardless of Le geographic orientation. 37 �References Bauer, R. L and Bidwell, M. E., 1990. Contrasts in the response to dextral transpression across the Quetico-Wawa subprovince boundary in northeastern Minnesota. Canadian Journal of Earth Sciences 27, 1521-1535. Dimroth, E., Imreh, L., Goulet, N., Rocheleau, M., 1983a. Evolution of the south-central segment of the Archean Abitibi Belt, Quebec Part II tectonic evolution and geomechanical model. Can J Earth Sci 20, 1355-1373. Dlmroth, E., Imreh, L., Goulet, N., Rocheleau, M., 1983b. Evolution of the south-central segment of the Archean Abitibi Belt, Quebec Part III plutonic and metamorphic evolution and geotectonic model. Can J Earth Sci 20 1374-l 388. Goodman, S., 2008. Structural and Kinematic Analysis of the Kawishiwi Shear Zone, Superior Province. M.S. Thesis, University of Minnesota Duluth, MN. Goodwln, A.M. and Ridler, R.H., 1970. The Abitibi orogenic belt In Symposium on Basins and Geosynclines of the Canadian Shield. Geol Surv Can Pap 70-40, 1-24. Erickson, E., 2010. Structural and kinematic analysis of the Shagawa Lake shear zone, Superior Province, northern Minnesota: implications for the role of vertical versus horizontal tectonics in the Archean. Canadian Journal of Earth Sciences, 47, 14631479. Hudleston, P.J., 1976. Early deformational history of Archean rocks in the Vermillion district, Northeastern Minnesota. Canadian Journal of Earth Sciences 13, 579-592. Hudleston, P.J., Schultz-Ela, D., Southwick, D. L., 1988. Transpression in an Archean greenstone belt, northern Minnesota. Canadian Journal of Earth Sciences 25, 10601068. Johnson, T., 2009. Structural, Kinematic, and Hydrothermal Fluid Investigation of the GoldBearing Murray Shear Zone, northeastern Minnesota. M.S. Thesis, University of Minnesota Duluth, MN. Karberg, S.M., 2009. Structural and kinematic analysis of the Mud Creek shear zone, northeastern Minnesota; implications for Archean (2.7 Ga) tectonics. M.S. Thesis, University of Minnesota Duluth, MN. Langford, F.F. and Morin, M.A. 1976. The development of the Superior Province of Northwestern Ontario by merging island arcs. Am J Sci 276, 1023-1034. Ludden, J.N., Hubert, C., Gariepy, C., 1986. The tectonic evolution of the Abitibi greenstone belt of Canada. Geol Mag 123, 153-166. Passchier, C.W., 1998. Monoclinic model shear zones. Journal of Structural Geology. 20 (8): 1121-1137. Schultz-Ela, D.D. and Hudleston, P.J., 1991. Strain in an Archean greenstone belt of Minnesota. Tectonophysics 190, 233-268. Sleep, N., 1992. Archean plate tectonics: what can be learned from continental geology?. Canadian Journal of Earth Sciences, 29, 2066-2071. Sylvester, P.J., Attoh, K and Schulz, K.J., 1987. Tectonic setting of late Archean bimodal volcanism in the M1- chipicoten (Wawa) greenstone belt, Ontario. Can J Earth Sci 24, 1120-1134. Talbot, C.J., 1973. A plate tectonic model for the Archean crust. Philos Trans Soc London 273, 413- 427. Xypolias, P., 2010. Vorticity analysis in shear zones: A review of methods and applications. Journal of Structural Geology 32, 2072-2092. 38 �EVALUATING THE BIOGENICITY OF FLUVIAL-LACUSTRINE STROMATOLITES FROM THE MESOPROTEROZOIC COPPER HARBOR CONGLOMERATE, UPPER PENINSULA OF MICHIGAN, USA Nicholas D. Fedorchuka, Stephen Q. Dornbosa,b, John L. Isbella, Julie A. Bowlesa, Frank A. Corsettic, Dylan T. Wilmethc, Victoria A. Petryshynd a Department of Geosciences, University of Wisconsin-Milwaukee, Milwaukee, WI 53211, USA Geology Department, Milwaukee Public Museum, Milwaukee, WI 53232, USA c Department of Earth Sciences, University of Southern California, Los Angeles, CA 90089, USA d Department of Earth, Planetary, and Space Sciences, University of California, Los Angeles, Los Angeles, CA 90095, USA b The Mesoproterozoic (1.09 Ga) Copper Harbor Conglomerate represents alluvial fan, fluvial and lacustrine deposition into the Midcontinent Rift System. The formation outcrops in the Upper Peninsula of Michigan where it contains carbonate stromatolites preserved within both siltstone and conglomerate facies. The purpose of this study is to evaluate the biogenicity of these stromatolites, which lack direct microfossil evidence. The stromatolites were placed into their depositional context, their macro-scale features and thin section microfabrics were analyzed, and growth angles were measured of cobble-draping samples to determine if a phototrophic response existed. A methodology that uses magnetic susceptibility as a biosignature was also performed on these stromatolites. The result of these analyses reveals two distinct types of stromatolites. Stromatolites from the siltstone facies are interpreted as biogenic. They contain detrital laminae, hematite-rich micritic laminae, and fenestral fabrics. The stromatolites formed as microbial mats grew over a mudflat or sandflat with carbonate filled dessication cracks and an eroded topography. Stromatolites from the conglomerate facies are interpreted to have formed by a mix of chemical and biological processes. They are microdigitate and have abiogenic features such as isopachous laminae with radial fibrous calcite fans and botryoids. They also lack a phototrophic response, suggesting that growth was not controlled by cyanobacteria. These stromatolites also have some biogenic signatures such as conical wavy laminae that have been separated by gas build-ups. These stromatolites are interpreted as having formed in a flooded braidplain setting with restricted circulation. Magnetic susceptibility tests yielded inconclusive results in this case because the stromatolites in question contain secondary hematite. This study supports previous studies of these stromatolites, as well as microbial structures and organic-rich paleosols that have suggested freshwater microbial communities were abundant in the Midcontinent Rift during the Mesoproterozoic. It also highlights how variable environmental factors can influence stromatolite growth, even in similar depositional settings and with a consistent microbial presence. 39 �40 �GEOCHEMISTRY OF BASALT XENOLITHS ENTRAINED IN MINERALIZED TROCTOLITIC AND ANORTHOSITIC INTRUSIONS, NORTHEASTERN MINNESOTA FINNEGAN, Molly L1 and LARSON, Phillip C1, 1 Duluth Metals Limited, 306 W. Superior Street, Suite 610, Duluth, MN 55802 Troctolites of the Nickel Lake Macrodike (NLM), South Kawishiwi (SKI) and Partridge River (PRI) Intrusions, as well as the varying lithologies of the Anorthositic Series (An-series), within the 1.1 Ga Duluth Complex, host numerous basaltic xenoliths. These xenoliths provide evidence that these intrusive bodies originally emplaced within the basalts and other accompanying lithologies of the North Shore Volcanic Group (NSVG; Miller and Weiblen, 1990). These xenolith packages, which include Biwabik Iron Formation, Virginia Formation and Colvin Creek metasediment, and An-series material, in addition to both magnetic and non-magnetic basalts, are often associated with sulfide-mineralized troctolitic and anorthositic lithologies. Lithogeochemical data suggests these basalt xenoliths can be grouped into three distinct categories, correlating with different mineralization styles in the host rock lithologies. This study aims to determine the nature of the apparent correlation between the presence of basalt xenoliths and CuNi-PGE mineralization as well as attempt to correlate the basalt compositions with possible sources. Samples were collected from drill holes and outcrop from the NLM, eastern margin of the SKI (East Shore), and the northeast corner of the PRI (Rook)(Fig. 1). Three groups of basalt are distinguished by relative abundances of MgO, Al2O3, Fe2O3, TiO2, Zr, and Mg# (calculated as Mg/(Mg+Fe3+)*100), using factor analysis (which identifies trends between multiple variables) to provide a more robust classification. Type 1 basalt xenoliths are characterized by Mg#s which range from 42-52 accompanied by elevated MgO and Al2O3, and low TiO2 and Zr. Type 2 has a range of Mg#s from 19-33, along with elevated Al2O3, TiO2, and Zr, and low MgO. Type 3 has a mid-range Mg# with respect to the other two types (27-36), accompanied by low MgO and Al2O3, and high TiO2, Zr, and Fe2O3*. An Mg# vs. Zr plot clearly discriminates the three different basalt xenoliths types (Fig. 2). Comparing these basalt xenoliths to basalt compositions from the NSVG using an MgO vs. TiO2 plot demonstrates Type 2 and Type 3 basalts correlate well with compositions of basalts within the NSVG, whereas the composition of the Type 1 basalt is particularly anomalous (Fig. 3). The nearest compositional correlation to Type 1 comes from sample KEW-6, collected from the Larsmont basalts of the NSVG near Knife River, MN (Boerboom et al., 2002). A closer match appears to be the P-Magma of Miller and Weiblen (1990) (Mg# 48), their representative primitive high-Al olivine tholeiitic basalt composition, which plots within the range of the Type 1 basalt xenoliths (Fig. 3). Type 1 basalts are spatially associated with the lower contact of a Cu-Ni mineralized zone at the southwest end of the NLM. Other xenolith lithologies occurring with Type 1 basalts include Virginia Formation and Biwabik Iron Formation. Type 2 basalts are spatially associated with high-grade PGM mineralization along the SKI-An-series contact, occurring with Colvin Creek and An-series xenoliths. Type 3 basalts are spatially associated with high-grade Cu-Ni-PGE mineralization in the PRI. They are the most variable in composition, as well as geographically widespread, occurring along the eastern margin of the SKI and within the NLM as well. Xenolith lithologies associated with Type 3 basalt include Colvin Creek, Virginia Formation, and An-series. 41 �Basalt xenolith populations can potentially be used as an indicator of the prospectivity of heterolithic mafic rocks. Correlating these xenoliths with their source areas potentially allows reconstruction of the sources and pathways of mineralized troctolitic and anorthositic magmas in the Duluth Complex. Figure 2: Plot of Mg# (=Mg/(Mg+Fe3+)) versus Zr (ppm). Type 1 has clustered around higher Mg# values, while Type 2 and 3 have lower and more variable Mg# ranges in conjunction with higher Zr content. Figure 1: Map indicating spatial relationships between the intrusions and showing the xenolith rich areas of the NLM, SKI, and PRI. Figure 3: Plot of TiO2 versus MgO for the three differentiated types of basalt as well as known compositions within the NSVG. Sample KEW-6 (Boerboom et al, 2002) is the sample plotted closest to the Type 1 basalt xenolith trend. The composition of P-Magma (Miller and Weiblen, 1990) is also shown. References Cited Boerboom, T.J., Green, J.C., and Jirsa, M.A., 2002, Bedrock geologic map of the Knife River quadrangle, St. Louis and Lake Counties, Minnesota: Minnesota Geological Survey Miscellaneous Map Series M-129, scale 1:24,000. Miller, J. and Weiblen, P., 1990. Anorthositic rocks of the Duluth Complex: Examples of rocks formed from plagioclase crystal mush. Journal of Petrology, 31, 295-339. 42 �MAGNETIC MINERALOGY OF REVERSELY MAGNETIZED CHENGWATANA LAVA FLOWS OF ST. CROIX FALLS, WISCONSIN FINN, Kiel and BOWLES, Julie Department of Geosciences, University of Wisconsin-Milwaukee, Lapham Hall 366, Milwaukee, WI 53201 Previous paleomagnetic studies on the 1.1 billion year old Chengwatana Volcanic lava flows, located near St. Croix Falls, Wisconsin, found a period of reversed polarity within a predominantly normal sequence that was very similar to that found at another location within the Keeweenawan Rift at Mamainse Point in Ontario. Based on this observation, the two sequences were correlated with each other. However, new research at Mamainse Point (Swanson-Hysell et al., 2011) has discovered that some of the lava flows there carry a self-reversed magnetization. This means that the magnetization is opposite in direction to the Earth’s field in which the rocks originally cooled. Further, data from some of the Chengwatana samples reported by Kean et al. (1997) also showed two components of magnetization approximately antiparallel to each other. The goal of this study was to further investigate the magnetic mineralogy of the Chengwatana Volcanic lava flows in order to test that reversely-magnetized samples from Chengwatana indeed reflect a period of reversed field polarity and not a self-reversed magnetization. Thirty samples were collected in two different locations at Interstate State Park in St. Croix Falls, Wisconsin. The natural remanent magnetization was demagnetized both by alternating field (AF) and thermal demagnetization at the University of Wisconsin – Milwaukee (UWM). Also at UWM, Curie temperature was measured via temperature-dependent susceptibility using an MFK1 Kappabridge susceptibility bridge with furnace insert. Hysteresis measurements to determine grain size were carried out on a vibrating sample magnetometer at the Institute for Rock Magnetism at University of Minnesota. Results indicate that the magnetic remanence held by both normal and reverse polarity flows, as defined previously by Kean et al. (1997), is held by multi-domain to pseudo-single-domain magnetite. Individual samples that carry two anti-parallel components of magnetization are also dominated by magnetite, and the high-coercivity or high-temperature component is similar in direction and polarity to single-component samples from the same flow. It is likely that the antiparallel component was acquired by partial remagnetization during a reheating event during a later period of opposite polarity. It does not completely overprint the primary magnetization, and the polarity sequence identified Kean et al. (1997) remains unchanged. This is in contrast to the self-reversed magnetization found in some of the Mamainse Point basalts by Swanson-Hysell et.al (2011). In those samples, the antiparallel component was controlled by fine grained hematite that acquired its magnetization during the formation of martite within the rocks. There is no indication that the Chengwatana flows share this mineralogy and the conclusions of Kean et al. remain valid. References Kean, W.F., Williams, I., Chan, L., Feeney, J., 1997 Magnetism of the Keweenawan age Chengwatana lava flows, northeast Wisconsin. Geophysical Research Letters, vol. 24, no.12, 1523-1526 Swanson-Hysell, N. L., Feinberg, J. M., Berquo, T.S., Maloof, A.C. 2011 Self-reversed magnetization held by martite in basalt flows from the 1.1-billion-year-old Keweenawan rift, Canada. Earth and Planetary Science Letters, 305, pp.171-184 43 �44 �AN UNUSUAL MESOPROTEROZOIC CARBONATE UNIT: RELIC OF A SALINE LAKE? FIRMIN, Sydney and Bartley, Julie K. Department of Geology, Gustavus Adolphus College, St. Peter, Minnesota, 56082 The Mesoproterozoic Rossport Formation of Ontario, Canada is primarily made up of sandstone and shale. The Rossport Formation is approximately 1.4 billion years old (Franklin et al., 1980) and is generally interpreted to have been deposited in an Figure 1 intracratonic basin, most likely a rift-related lake (Rogala et al., 2005). The Middlebrun Bay Member, in the middle of the formation, consists of cherty limestone containing stromatolites. While examining outcrops of the Middlebrun Bay Member on the Channel Islands of Lake Superior, we discovered an unusual limestone bed on Copper Island. This calcite does not contain stromatolites; it has an unusual bright white color and lacks internal structure (Fig. 1). Previous work on the Rossport Formation suggests that the stromatolites formed when lake levels were low and not much sand was making it to the basin (Rogala et al., 2007). In this model, stromatolites would have formed in a hypersaline lake environment during intervals of low clastic influx. If this interpretation is correct, the non-stromatolitic “white-bed” could have formed as an evaporite bed, now replaced by calcite. In this study, we investigate the hypothesis that the massive white limestone unit is calcitized evaporite. At the outcrop level, the massive white carbonate is approximately 1.1 m thick and occurs between layers of sandstone. The carbonate unit contains sandstone clasts (Fig. 2). This outcrop relationship is consistent with either collapse of sandstone during evaporite dissolution, or upward growth of evaporite rock, causing sandstone to wedge, split, and form clasts surrounded by evaporite. Figure 2 Figure 3 The macroscopic texture of the carbonate unit is both massive and coarsely crystalline, with a texture reminiscent of chicken-wire evaporite (Fig. 3). Chicken-wire texture forms when nodules of gypsum crystals grow and push other material to their edges, forming a coarsely crystalline structure with a network of residuum outlining large crystal domains. Thin-sections of the massive white carbonate were compared to those from Middlebrun Bay Member Stromatolites from Channel Island. The stromatolitic thin sections show relatively small crystals and fine lamination, consistent with their macroscopic texture. In contrast, the “whitebed” rock had large subhedral to euhedral crystals, with zonation apparent by 45 �cathodoluminescence. “White-bed” samples also had a large number of stylolites, indicating substantial dissolution along crystal boundaries. Both stylolites and large crystal edges contain accumulations of insoluble residue (Fig. 4), indicating that dissolution and reprecipitation processes were important in generating the final texture of the white bed. Figure 4 Chemical evidence is consistent with an evaporite origin of the “white-bed” carbonate. Trace element concentrations, measured by ICP-MS, were generally highly elevated in stromatolite samples and moderately elevated in the massive carbonate unit, compared to average Proterozoic carbonate compositions. Taken together, geochemistry suggests that both the stromatolites and the white bed were deposited in a hypersaline lake environment. Trace elements were concentrated in carbonate during precipitation of stromatolites. Primary evaporite phases would also have been highly concentrated in trace elements, but these concentrations would have decreased during dissolution of evaporites and precipitation of secondary calcite. Similar patterns of trace element enrichment are observed in hypothesized calcitized evaporites from the Mesoproterozoic Atar Group, Mauritania (Manning-Berg and Kah, 2013). Based on the evidence collected both in the field and lab, it seems likely that the “white bed” carbonate possesses a unique texture because it was originally precipitated as gypsum. The massive, coarsely crystalline texture indicated pervasive recrystallization, consistent with a primary evaporite miner, like gypsum, which was secondarily replaced by calcite, resulting in coarse, featureless carbonate and collapse of overlying sandstone layers. Geochemical results are consistent with deposition under hypersaline conditions. In further research we will look at sulfate concentrations, both as total S and as carbonate associated sulfate (CAS). Other calcitized evaporites have shown elevated CAS concentrations (Manning-Berg and Kah, 2013). A depositional environment where gypsum formed would indicate a saline lake, consistent with previously proposed environmental conditions for the Rossport Formation. References Rogala, B., Fralick, P.W., Heaman, L.M., and Metsaranta, R., 2007, Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada: Canadian Journal of Earth Sciences, v. 44, p. 1131-1149. Rogala, B., and Fralick, P.W., 2005, Stratigraphy and sedimentology of the Mesoproterozoic Sibley Group and related igneous intrusions, northwestern Ontario: Ontario Geological Survey Open File Report 6174, 128 pp. Franklin, J. M., McIlwaine, W.H., Poulsen, K.H., and Wanless, R.K., 1980, Stratigraphy and depositional setting of the Sibley Group, Thunder Bay district, Ontario, Canada: Canadian Journal of Earth Sciences, v. 17, p. 633-651. Manning-Berg, A.R., and Kah, L.C., 2013, Calcitized Evaporites and the Evolution of Earth’s Early Biosphere: Geological Society of America Abstracts with Programs, v. 45(7), p. 628. 46 �GEOLOGY OF THE BRULE RIVER AREA OF THE PINE MOUNTAIN QUADRANGLE, MINNESOTA: CAPSTONE MAPPING PROJECT FOR THE PRECAMBRIAN RESEARCH CENTER’S 2013 FIELD CAMP Paul J. Fix1, Stephen J. Ginley1, Lauren A. Schraeder1, Aaron J. Summers1, Michael S. Doyle2,Terrence J. Boerboom3 1 2013 Field Camp Participants, Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Hwy., Duluth, MN 55811 2 Dept. of Geological Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 3 Minnesota Geological Survey (MGS), University of Minnesota, 2642 University Ave. West, St. Paul, MN 55114 The Precambrian Research Center at the University of Minnesota Duluth conducted its seventh annual Precambrian Field Camp during the summer of 2013.This presentation is one of a series that detail the results of the 2013 Capstone Mapping Projects which represent the culmination of activities at the field camp. The Capstone projects, conducted during the final two weeks of the field camp, are meant to test the skills obtained by the camp participants by conducting field studies and creating geological maps of areas of poorly understood geology. This Capstone Project involved mapping near Brule River area, in the Pine Mountain 7.5’ Quadrangle, approximately 25 miles north-northwest of Grand Marais, MN. Figure 1. Geology of northeastern Minnesota showing the area of the Pine Mountain capstone mapping project. From Miller and Green, 2002 Prior to this work, the area had only been mapped in ay reconnaissance fashion, and consequently the geologic detail was poorly understood and interpretations were largely derived from only geophysical data. The map area is in a geologically complex Mesoproterozoic terrane comprised largely of mafic to 47 �felsic lavas of the Keweenawan North Shore Volcanic Group (NSVG), and later mafic to felsic intrusions related largely to the Beaver Bay Complex. The map area lies along the western extension of the reversely-polarized Hovland lavas (1107.7 ± 1.9 Ma; Davis and Green, 1997), shown only as a single unit of undivided volcanic rocks on published maps (e.g. Miller and others, 2001). The area is bordered on the south by the Brule-Hovland gabbro complex, and is cut by roughly east-west trending diabase dikes inferred to also be related to the BruleHovland gabbro. Geophysical evidence implies that the latter dikes form a forked dike set which cuts the middle of the mapping area. Our mapping has shown that multiple east-west striking, southeast-dipping, lava flows composed of alternating rhyolite and andesite are present within the map area. The rhyolites are generally quartzand plagioclase-phyric but local aphyric varieties with large (5-10 cm) pumice inclusions were noted. Contorted flow banding and flow-aligned plagioclase laths in the rhyolites give evidence of viscous flow. The andesites are generally plagioclase phyric to glomeroporphyritic, with plagioclase phenocrysts as large as 1-5 cm. Some of the flows contain abundant pyroxene and ilmenite, and may be basaltic in composition, but no thin sections or geochemical analyses were obtained to verify this. The orientations of the lava flows in the andesites were determined by measuring the orientations of oxidation-lamination. Flow contacts were noted via the presence of amygdaloidal flow tops, and amygdules in both the rhyolites and andesites contain quartz, epidote, and chlorite, indicating that they experienced low grade contact metamorphism due to emplacement of the surrounding mafic intrusions. The volcanic rocks are cut by a series of mafic intrusions related in timing to the Beaver Bay Complex (~1096 Ma; Miller and Chandler, 1997). Based on field relationships the oldest intrusive unit is a medium-grained, variably porphyritic, slightly granophyric, poorly- to moderately-foliated anorthositic gabbro. This is intruded by medium-grained ophitic olivine gabbro, which is much more extensive throughout the map area than was recognized prior to this work. Hybrid ferrodioritic rocks are common along the margins of this ophitic gabbro, especially where in contact with felsic volcanic rocks. This hybrid unit contains abundant rhyolite inclusions and felsic stringers mixed with light to dark gray, finegrained chilled mafic phases. The shape and textures of these felsic stringers implies that they were formed from partial melting of the rhyolite inclusions, and that these melts commingled with mafic magma. Sparse inclusions of meter to outcrop scale hornfels basalt are also found in this unit. Two narrow and parallel, high-amplitude linear aeromagnetic anomalies, formerly interpreted as two parallel east-west trending diabase dikes, instead may be caused by the magnetic margins of a single, thick dike of ophitic olivine gabbro. However the ophitic gabbro also covers a large area that is characterized by lower amplitude magnetic anomalies; in order to resolve this further petrographic and rock property studies would need to be completed. In summary, although not all outcrops in the capstone field map area were examined due to time constraints, we have shown that there is much more geologic complexity than had been documented by previous reconnaissance mapping. This small map window should lay the groundwork for and encourage future mapping endeavors in this poorly mapped area. This and other capstone maps produced by the Precambrian Research Center can be viewed at www.d.umn.edu/prc. References: Miller, J.D. Jr., and Chandler, V.W., 1997 , in Ojakangas, R.W., Dickas, A.B., and Green, J.C., eds., Middle Proterozoic to Cambrian rifting, central North America: GSA Special Paper 312, p. 73-96. Miller, J.D., Jr., and Green, J.C. 2002, in Miller,, J.D., and others, 2002, MGS Rept. of Investigations 58, p. 144-163 Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.M., 2001, Minnesota Geological Survey Miscellaneous Map Series Map M-119, scale 1:200,000. Davidson, D.M., Jr., and Burnell, J.R., Jr., 1977, Minnesota Geological Survey Miscellaneous Map M-29, scale 1:24,000 Davis, D.W., and Green, J.C., Canadian Journal of Earth Sciences, v. 34, no.4, p.476-488. 48 �EVOLUTION OF THE MIDCONTINENT RIFT SYSTEM: PALEOMAGNETIC, ROCK MAGNETIC AND ANISOTROPY OF MAGNETIC SUSCEPTIBILITY INVESTIGATION OF THE MESOPROTEROZOIC BARAGA - MARQUETTE DIKE SWARM (MICHIGAN, USA) FOUCHER, Marine, CURGANUS, Renee, PIISPA Elisa J., SMIRNOV, Aleksey V. Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, 630 DOW ESE Building, Houghton, MI 49931-1295, USA and PESONEN, Lauri J. P. Department of Physics, Division of Geophysics and Astronomy, University of Helsinki, Helsinki, Finland The Midcontinent Rift System (MRS) is characterized by multiple diabase dike swarms parallel to subparallel to the rift axis (e.g. Green et al. 1987). The dikes are generally considered to be feeders to now eroded lava flows once deposited on the flanks of the rift. We report the results of a detailed investigation of rock magnetism, paleomagnetism, and anisotropy of magnetic susceptibility (AMS) from 24 dikes of the east-west trending Baraga-Marquette (BM) dike swarm exposed in the Upper Peninsula of Michigan. In addition, preliminary rock magnetic and paleomagnetic data from five dikes of the Central Wisconsin (CW) dike swarm are presented. Thermomagnetic and magnetic hysteresis analyses indicate that the principal magnetic carrier in the studied dikes is single- to pseudo-single domain low-titanium titanomagnetite. Approximately a third of the dikes contain minor amounts of hematite. In addition, several dikes from highly mineralized areas exhibit an additional magnetic phase likely pyrrhotite or maghemite. Twelve of the investigated BM dikes yielded well-defined characteristic remanent magnetization (ChRM) directions similar to the typical directions observed from other reversely magnetized MRS rocks and the directions observed in a prior study of the BM swarm by Pesonen and Halls (1979). The new data from reversely magnetized dikes are combined with the prior study data and the combined dataset (20 dikes) is subjected to paleosecular variation analysis. The results are also compared with the paleomagnetic data obtained from other nearly coeval dike swarms of MRS. Three BM dikes yielded normal ChRM directions with steep inclinations, significantly different from the direction exhibited by other normally magnetized MRS sequences. The normal and reversed polarity dikes are also distinguishable with respect to their magnetic grain size. While the statistical significance of these observations requires further investigation, taken at face value it suggests that the BM swarm may represent at least two emplacement episodes with normally magnetized dikes being older. Three CW dikes yielded stable ChRM directions (two normal and one reversed) typical for the MRS time. The anisotropy of magnetic susceptibility analyses yielded well-constrained magma flow directions in most of the studied dikes. The magma flow directions of the BM dike swarm are discussed in the context of the tectonic evolution of MRS. REFERENCES Green, J.C., Bornhorst, T.J., Chandler, V.W., Mudrey, M.G. Jr., Myers, P.E., Pesonen, L.J., Wilband, J.T., 1987. Keweenawan dikes of the Lake Superior region: evidence for evolution of the middle Proterozoic Midcontinent Rift of North America. In: Halls, H.C., Fahrig, W.F. (Eds.), Geological Association of Canada, Special Paper 34, 289–302. Pesonen, L.J. and Halls, H.C., 1979. The paleomagnetism of Keweenawan dikes from Baraga and Marquette Counties, northern Michigan. Canadian Journal of Earth Sciences 16: 2,136-2,149. 49 �50 �Compilation of existing geophysical models in preparation for 3D modeling of the Midcontinent Rift System in the western Lake Superior region, Minnesota, Wisconsin, and Michigan GRAUCH, V.J.S.1, CHANDLER, Val2 and LIVELY, Richard S.2 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 2 Minnesota Geological Survey, 2642 University Avenue W., St. Paul, MN, 55114-1032 Over the past several decades, both 2D and 3D geophysical models have played a large part in developing our understanding of the subsurface structure and composition of the Midcontinent Rift System (MRS). The overall configuration of the MRS is prominently expressed in regional gravity and magnetic maps. Details of the subsurface configuration of volcanic layers and structures are evident in high-resolution aeromagnetic data and published seismic-reflection sections. Improvements in the resolution and coverage of gravity and magnetic data and technical advances in modeling capabilities in the last decade provide the motivation for new attempts at 3D modeling of the MRS. In particular, the MRS has complex structure in the western Lake Superior region that is appropriate for 3D modeling. In this area, bounding faults of the NNE-trending St. Croix horst abruptly turn more easterly at White’s Ridge and transition into the ENE-striking Lake Owen and Keweenaw faults (Fig. 1). Gravity data west of White’s Ridge suggest that relations between the St. Croix Horst and the Duluth Complex are also 3D in nature (Fig. 2). Thus, we are developing a new 3D model of the MRS for the area surrounding White’s Ridge (Figs. 1 and 2). We start with the 3D gravity models of Allen (1994), who constrained his modeling using seismic-reflection lines (Fig. 1). Refinements are made by incorporating new 2D geophysical models and constraints from analysis of more recent geophysical and geologic data sets. We also plan to re-evaluate the gravity effects of the lower crust and mantle, as interpretations of new deep-looking geophysical data become publically available. As the model develops, the generalized rock units can be subdivided and additional structure added. The first step in the modeling process is to compile existing models and information into the 3D model space so that discrepancies or other issues can be easily recognized. Images captured from published 2D geophysical models, interpreted seismic-reflection sections, and geologic cross-sections (Fig. 1) were input as displays along section lines. The two 3D gravity models of the MRS constructed by Allen (1994) for western Lake Superior and the Minnesota-Wisconsin section (Figure 2) were recently combined by Chandler and Lively (2011) into a 3D visualization of surfaces, where each surface represents the base of a generalized rock package. Grid points from these surfaces were input into the 3D model space, then projected onto the section displays to quickly discover discrepancies between models. The discrepancies found are mostly explained by how rock units are generalized and what rock properties are assigned to which rock units. After consideration of updated rock property information, we chose the following model units and associated densities (in kg/m3) for the 3D modeling, which generally follows those of Allen (1994): Bayfield Group—2,450; Oronto Group—2,650; volcanic rocks—2,950; Duluth Complex—3,000; pre-rift upper crust (<20 km depth)—2,750; and lower crust (>20 km depth)—2,900. References Allen, D. A., 1994, An integrated geophysical investigation of the Midcontinent Rift System: western Lake Superior, Minnesota and Wisconsin [PhD]: Purdue University, 267 p. Cannon, W. F., Woodruff, L. G., Nicholson, S. W., and Hedgman, C. A., 1996, Bedrock geologic map of the Ashland and the northern part of the Ironwood 30' X 60' quadrangles, Wisconsin and Michigan: U.S. Geological Survey Miscellaneous Investigations Map I-2566, scale 1:100,000. Chandler, V. W., and Lively, R. S., 2011, Compilation of Minnesota and western Wisconsin geoscience for the USGS national geologic carbon dioxide sequestration assessment: Enhanced geophysical model for extent and thickness of deep sedimentary rocks: Minnesota Geological Survey Open-File Report 11-03, 37 p. Ferderer, R. J., 1982, Gravity and magnetic modeling of the southern half of the Duluth Complex, northeastern Minnesota [MA]: Indiana University, 86 p. 51 �Figure 1: Rock units of the Midcontinent rift in the western Lake Superior region, area covered by the new 3D model, and locations of previous 2D geophysical models and seismic lines. Geographic boundaries are shown by dashed lines. Figure 2: Color shaded-relief image of Bouguer gravity, showing the new 3D model area (bold black outline). Sections are as in Fig. 1. The white dashed lines outline the two 3D model areas of Allen (1994). 52 �A Field and Petrographic Study of Neoarchean Variolitic Pillow Lavas, Newton Belt, Vermilion District, Northeastern Minnesota GROTTE, M. J.1,2 and HUDAK, G. J.2,3     1 Department of Geological Sciences, University of Minnesota Duluth, grot0133@d.umn.edu Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, 3 Minerals Division, Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN   2 The Newton Lake Formation, located northeast of Ely, Minnesota, comprises a series of tholeiitic to komatiitic lava flows, mafic to ultramafic intrusions, and associated clastic sedimentary rocks. Approximately one-half mile north of CR-88 on the Echo Trail, a sequence of exceptionally well-preserved, steeply dipping, south-topping, lower greenschist-facies metamorphosed variolitic pillow lavas are exposed on the northwest side of the road. This outcrop was recently visited during the 55th Annual Institute of Lake Superior Geology Field Trip 7 and was described as “spherulitic pillow basalt” (Peterson et al., 2009). Although mapped on a regional scale (Peterson et al., 2005), to date no detailed geological sketches or petrographic analyses of these variolitic pillowed flows have been completed for the purpose of understanding the genesis of these variolites. As varioles may be the result of blotchy alteration, magma mingling, or quench crystallization (Arndt and Fowler, 2004; Fowler et al., 2002), a detailed petrographic study was conducted to evaluate the genesis of the variolites that exist at this location. Scanning electron microscopy will be utilized to better understand the compositional characteristics of the variolites. A series of hand samples from this exposure of Newton Lake Formation pillow lavas was collected during fall, 2013. Samples were chosen at different distances from the crusts of individual pillows, as well as from areas where both pillow selveges and associated pillow hyaloclastite occurred. All sample locations were documented using a hand-held GPS unit in the UTM NAD 83 Zone 15 North coordinate system. A series of photographs was taken, and a panorama of these photos was constructed to assist in the mapping of this exceptional outcrop. Samples were prepared into standard and polished thin sections for analysis on a Leica DM EP polarizing microscope and by energy dispersive spectroscopy using a JEOL JSM-6490LV scanning electron microscope, both at the University of Minnesota Duluth. In outcrop, the pillow lavas vary from bun- to mattress-shaped (Dimroth et al., 1978) and range from <1 m to >2.5 m in diameter. A typical cross-section through these pillow lavas is shown in Figure 1. Pillow cores tend to be dark green to pale yellow-green in color depending on the degree of secondary epidote alteration present. Locally, quartz-filled vacuoles are present near the stratigraphic tops of individual pillows. Variolitic textures are most common in the border zone of individual pillows within 0.5 m of the crust of individual pillows, shown in Figure 2. Here, 5-15% individual rounded to spherical variolites up to 1 cm in diameter, as well as nearly massive, globular, coalescing variolites oriented sub-parallel to pillow margins are present. Petrographic studies indicate exceptional preservation of hyaloclastite adjacent to the crusts of individual pillows (Figure 3). The hyaloclastite comprises shard-shaped, jigsaw-puzzle fit lapilli composed of light brown to light green altered glass similar to that comprising the pillow crusts. Small (<5 mm) spherical plagioclase variolites are commonly present in the hyaloclastite shards. Variolites in the border region are zoned moving inward toward the pillow core, with an outer zone comprising of individual, small (2-5 mm), spherical to oval plagioclase spherulites with minor (<20%) fan-shaped to bow-tie shaped inclusions of altered mafic minerals. A secondary zone comprising larger (5-10 mm), round plagioclase spherulites in a matrix composed of fine-grained, chaotically-oriented skeletal amphibole that may be pseudomorphs of original pyroxene. A third zone composed of semi-massive to massive, globular, coalescing spherical plagioclase spherulites that contain 30-55% fan-shaped to bow-tie inclusions of mafic minerals, and a fourth zone containing <10% rounded plagioclase spherulites up to 15 mm in 53 �diameter in a matrix of coarser grained, chaotically oriented skeletal amphibole. The cores of the pillow lavas are composed of fine-grained tabular to acicular plagioclase intergrown with finegrained tabular to prismatic amphibole pseudomorphs of pyroxene. Scanning electron microscopy studies are currently in progress to evaluate compositional differences between the variolites and groundmass minerals in the various variolite zones and pillow cores. Based on the results of this study, variolites in this exceptional exposure of Newton Lake Formation pillow lavas are dominantly composed of rounded to oval, radiating plagioclase spherulites with rare, axiolitic plagioclase spherulites locally present. The presence of needle-like to acicular skeletal plagioclase crystals and absence of phenocrysts suggest that lavas responsible for the pillow lava flows at this location were erupted at temperatures above the liquidus and experienced relatively large degrees of undercooling before undergoing rapid crystallization on the Neoarchean seafloor. Figure 1. Typical pillow lava cross-section. Figure 3. Thin section of pillow margin containing small varioles and well preserved hyaloclastite. Figure 2. Outcrop photo of pillow margin, varioles transitioning into hyaloclastite (from right to left). References Arndt, N. & Fowler, A. D., 2004, Textures in Komatiites and Variolitic Basalts. The Precambrian Earth - Tempos and Events: Elsevier, p. 298-311. Dimroth, E., Cousineau, P., Leduc, M., and Sanschagrin, Y., 1978, Structure and organization of Archean subaqueous basalt flows, Rouyn-Noranda area, Quebec, Canada: Canadian Journal of Earth Sciences, v. 15, p. 902-918. Fowler, A. D., Berger, B., Shore, M., Jones, M. I., and Ropchan, J., 2002, Supercooled rocks: development and significance of varioles, spherulites, dendrites, and spinifex in Archean volcanic rocks, Abitibi Greenstone Belt, Canada: Precambrian Research, v. 115, p. 311-328. Peterson, D. M., Jirsa M. A., and Hudak, G. J., 2009, Field Trip 7 - Architecture of an Archean Greenstone Belt: Stratigraphy, Structure, and Mineralization: Institute on Lake Superior Geology, Proceedings Volume 55 Part 2 – Field Trip Guidebook, p. 178-215. 54 �STRATIGRAPHIC FRAMEWORK AND LANDSYSTEM CORRELATION FOR DEPOSITS OF THE SAGINAW LOBE, MICHIGAN, USA GUZMAN, Ivan R., Department of Geosciences, Western Michigan University, Kalamazoo, MI 49008 Since the time of the Last Glacial Maximum (LGM) the south-central portion of the Lower Michigan Peninsula has been subject to several glacial advances and retreats by the Saginaw lobe. As part of the U.S Geological Survey Great Lakes Geological Mapping Coalition projects, several rotasonic borings were drilled between 2006 and 2013 in Barry, Kalamazoo and Calhoun Counties. Gamma ray logs and textural analyses were completed for each core. Five of these borings were selected according to their diamicton (till) content and correlated using water well logs and surficial geology maps. Glacial deposits such as diamicton serve as evidence of glacial advance/retreat, and are usually present as nearly continuous layers of sediments. Analysis of these layers affords the ability to accurately correlate these types of sediments across an area. Three cores, BA-10-02 and BA-09-02, KA-12-02 were drilled along the Kalamazoo moraine, each one containing 1 to 3 diamicton units separated by lacustrine sediments. The last two cores, CA-11-01 and KA-13-01 were drilled on a drumlinized till plain; both contain 2 to 4 diamicton units separated by outwash sediments. These diamicton units indicate the presence of at least one major and two minor advances/retreats of the Saginaw Lobe. 55 �56 �RETHINKING THE MIDCONTINENT RIFT – PUNCTURING THE “PLUME PARADIGM” HOLLINGS, Peter1, and HEGGIE, Geoff2 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada. peter.hollings@lakeheadu.ca 2 Panoramic Resources Ltd., 1004 Alloy Drive, Thunder Bay, ON P7B 6A5 Canada The mantle plume origin for the Midcontinent Rift (MCR) is widely accepted in the literature (e.g., Hutchinson et al., 1990; Nicholson and Shirey, 1990; Nicholson et al., 1997). However, recent geochronological, geochemical and mineralogical data suggest that the simple plume model should no longer be applied and it is necessary to evaluate alternate models. So, why do we need to rethink the rift? Geochronology – The majority of plume-related Large Igneous Provinces (LIPs) are characterized by a short-duration magmatic pulse or pulses (less than 1–5 My; Ernst et al., 2013a). Recent geochronology has shown that MCR magmatism spans at least 20 million years (Heaman et al., 2007, Hollings et al., 2010 and Dunlop, 2013) and possibly as much as 60 million years. Ultramafic rocks – One argument often put forward in favour of a plume origin for the rift is the presence of ultramafic rocks. Although there are, indeed, ultramafic rocks in the MCR, all of which are hosted in intrusions, mineral chemistry analyses from these intrusions shows that they have maximum olivine forsterite compositions not in equilibrium with the mantle: Seagull (Fo85), Thunder Bay North (Fo82), Eagle (Fo85; Ding et al., 2010) and Tamarack (Fo88; Goldner, 2011). Numerical models of these olivine compositions suggest a parental magma with 8-10 wt% MgO. Even the most primitive Tamarack intrusion suggests a primary magma with a composition of 12wt% MgO and 11wt% FeO (Goldner, 2011). Dyke swarms – The majority of plume-related LIPs are associated with, or even recognized by, the presence of giant, radiating dike swarms up to 3000 km long which project for long distances into cratonic hinterlands and provide evidence for paleo-stress regimes consistent with a central piercing point, likely a plume (Ernst et al., 2013a). To date, no radiating dike swarm has been recognized nor associated with the MCR. Rather, the majority of MCR-related dikes occupy extensional, rift arm-parallel structures (e.g. Hollings et al., 2010). Hanson et al. (1998, 2004, 2006) recognized ~1100 Ma magmatism in the Kalahari Craton, termed the Umkondo event. Ernst et al. (2013b) have recently proposed that this magmatic event may be considerably more widespread, with the recognition of ~1100 Ma magmatism in the Congo, the Amazon and India. They proposed that paleogeographic reconstructions are permissive of these events representing a single LIP that, based on paleomagnetic reconstructions, was distinct from Keweenawan magmatism (Ernst et al., 2013b). The presence of multiple, broadly contemporaneous LIP events suggests that the Mesoproterozoic may have been a period of significant mantle overturn (Stein and Hofmann, 1994) and atypically increased magmatic activity. 57 �The long duration of MCR magmatism, absence of primary ultramafic magmas and lack of a radiating dike swarm all suggest that a passive rifting model may be more appropriate for the rift. According to this model, rifting of the Superior Craton, possibly a response to the Umkondo LIP event, allowed for upwelling of material underplated by earlier plume events thought to have been centered in the vicinity of the present-day Lake Superior (e.g. the Marathon LIP, Halls et al., 2008). References Ding, X., Li, C., Ripley, E.M., 2010. The Eagle and East Eagle sulfide ore-bearing mafic-ultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution. G3 Geochemistry Geophysics Geosystems, 11, 1-22. Dunlop, M., 2013. The Eagle Ni-Cu-PGE Magmatic Sulfide Deposit and Surrounding Mafic Dikes and Intrusions in the Baraga Basin, Upper Michigan: Relationships, Petrogenesis, and Implications for Magmatic Sulfide Exploration. Unpublished MSc thesis, Indiana University, 105p. Ernst, R., Bleeker, W, Soderlund, U. and Kerr, A., 2013a. Large Igneous Provinces and supercontinents: Toward completing the plate tectonic revolution. Lithos, 174, 1-14. Ernst, R., Pereirac, E., Hamilton, M., Pisarevsky, S., Rodriques, J., Tassinari, C., Teixeirah, W., and VanDunemi, V., 2013b. Mesoproterozoic intraplate magmatic ‘barcode’ record of the Angola portion of the Congo Craton: Newly dated magmatic events at 1505 and 1110 Ma and implications for Nuna (Columbia) supercontinent reconstructions. Precambrian Research, 230, 103-118. Goldner, B.D., 2011. Igneous Petrology of the Ni-Cu-PGE-Mineralized Tamarack Intrusion, Aikin and Carlton Counties, Minnesota. Unpublished MSc thesis, University of Minnesota, 155p. Halls, H.C., Davis, D.W., Stott, G.M., Ernst, R.E., Hamilton, M.A., 2008. The Paleoproterozoic Marathon Large Igneous Province: new evidence for a 2.1 Ga long-lived mantle plume event along the southern margin of the North American Superior Province. Precambrian Research 162, 327–353. Hanson, R.E., 2003. Proterozoic geochronology and tectonic evolution of southern Africa. In: Yoshida, M., Windley, B., Dasgupta, S. (Eds.), Proterozoic East Gondwana: Supercontinent Assembly and Breakup, vol. 206. Geological Society of London, Spec. Publ, pp. 428–463. Hanson, R.E., Crowley, J.L., Bowring, S.A., Ramezani, J., Gose, W.A., Dalziel, I.W.D., Pancake, J.A., Seidel, E.K., Blenkinsop, T.G., Mukwakwami, J., 2004. Coeval large-scale magmatism in the Kalahari and Laurentian cratons during Rodinia assembly. Science 304, 1126–1129. Hanson, R.E., Harmer, R.E., Blenkinsop, T.G., Buller, D.S., Dalziel, I.W.D., Gose, W.A., Hall, R.P., Kampunzu, A.B., Key, R.M., Mukwakwami, J., Munyanyiwa, H., Pancake, J.A., Seidel, E.K., Ward, E.K., 2006. Mesoproterozoic intraplate magmatism in the Kalahari craton: a review. Journal of African Earth Sciences 46, 141–167. Heaman, L.M., Easton, R.M., Hart, T.R., Hollings, P., MacDonald, C.A., Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon Region, Ontario. Canadian Journal of Earth Sciences 44, 1055–1086. Hutchinson, D.R., White, R.W., Cannon, W.F., Schulz, K.J., 1990. Keweenaw hot spot: geophysical evidence for a 1.1 Ga mantle plume beneath the Midcontinent Rift System. Journal of Geophysical Research 95, 10,869–10,884. Nicholson, S.W., Shirey, S.B., 1990. Midcontinent rift volcanism in the Lake Superior region: Sr, Nd, and Pb isotopic evidence for a mantle plume origin. Journal of Geophysical Research 95 (10), 10851– 10868. Nicholson, S.W., Shirey, S., Schulz, K., Green, J., 1997. Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development. Canadian Journal of Earth Sciences 34, 504–520. Stein, M. and Hofmann, A.W., 1994. Mantle plumes and episodic crustal growth. Nature. 372. 63–68. 58 �THE MINNESOTA TACONITE WORKERS HEALTH STUDY: ENVIRONMENTAL STUDY OF AIRBORNE PARTICULATE MATTER 2014 UPDATE HUDAK, George1, MONSON GEERTS, Stephen1, ZANKO, Larry1, POST, Sara1, BANDLI, Bryan2 1 2 Natural Resources Research Institute, 5013 Miller Trunk Highway, Duluth, MN, 55811 Department of Geological Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 The Natural Resources Research Institute (NRRI) continues to conduct a detailed characterization of mineral dust in northeastern Minnesota. The purpose of this research is to evaluate the effects of present emissions from taconite mining and processing on air quality throughout the Mesabi Iron Range (MIR) (Figure 1) by characterizing airborne mineral particulate matter within currently operating taconite processing plants, in MIR communities surrounding taconite mining/processing operations, and in population centers in Minnesota not associated with taconite mining. Characterization studies of agedated lake sediments are also being conducted to determine the composition of past particulate matter deposition. NRRI’s sampling and characterization work represents the community/environmental component of the Minnesota Taconite Workers Health Study, a broad University of Minnesota (UM) research effort involving both the NRRI and the School of Public Health. Figure 1. Locations of taconite processing plants on the Mesabi Iron Range being sampled during this study (after Oreskovich and Patelke, 2006) Air sampling was performed within taconite operations, MIR communities, and non-MIR communities by NRRI scientists during both winter and summer seasons from 2009-2012. Sampling was conducted at four process locations within taconite operations, including: 1) secondary crushers; 2) magnetic separators/concentrators; 3) agglomerators/ball drums; and 4) kiln/pellet discharge areas. Sampling within the MIR communities took place on centrally-located rooftops of public buildings, whereas sampling in non-MIR communities occurred on either rooftops, or in remote sampling locations, 59 �so that background air quality away from the MIR could be evaluated. Airborne particles were collected using: 1) a micro orifice uniform deposit impactor (MOUDI) (Marple et al., 1991, 2014), which enables size-fractionated particulate matter collection; and 2) a Total Filter Sampler (TFS). Particulate matter was evaluated via gravimetric analysis and was subsequently subjected to comprehensive particulate matter characterization that included: 1) scanning electron microscopy (SEM) imaging; 2) energy dispersive xray spectroscopy (EDS); 3) electron backscattered diffraction (EBSD); 4) proton induced x-ray emission (PIXE); 5) the Minnesota Department of Health’s 852 Method Transmission Electron Microscopy (TEM) Analysis for Mineral Fibers in Air; and 6) the International Standardization Organization’s Indirect Method 13794 for Ambient air – Determination of Asbestos Fibers. NRRI’s research methods do not produce exposure data, and are not meant to provide data for regulatory purposes. During the past year, the NRRI has been evaluating the physical (gravimetric, morphology, concentration), mineralogical, and chemical characteristics of the particulate matter obtained from sampling at the taconite operations and MIR/non-MIR communities. This includes analysis of samples obtained during 14 sampling events at taconite operations and 79 sampling events at locations within communities and sites in northeastern Minnesota (73) and Minneapolis (6). Lake sediment analysis continues, and will provide important historical data regarding potential mineralogical inputs from iron mining and processing from ~1840 (which pre-dates iron mining on the MIR) to the present, which includes the period where the transition from natural ore mining to taconite mining took place. Community results to date are as follows: • measured particulate matter concentrations for PM2.5 in all MIR communities have been below 12 µg/m3, and for total PM have been below 16µg/m3; • particulate matter concentrations on the MIR are similar to those in the two NE Minnesota background sites (Duluth NRRI, Ely Fernberg site), and are lower than those obtained in Minneapolis (UM Mechanical Engineering Building rooftop); • mineral particulate matter in community air samples reflects the mineralogy of the Biwabik Iron Formation and other Minnesota rock types and geological materials; • elongate mineral particles (EMP) are present in MIR community ambient air samples; however, asbestiform amphiboles were rarely observed (1 asbestiform amphibole EMP in ~22,800m3 of air). In-plant results to date are as follows: • plant environments can be very dusty, with the most dusty environments associated with the agglomerator and kiln discharge areas; • particulate levels (PM1, PM2.5, PM10, and total PM) show a slight increase in the five MIR communities during plant/mine activity, but this increase is not statistically significant compared to when the plants were not operating. The NRRI plans to complete this work in 2014. References ISO 13794 (1999), Ambient air — Determination of asbestos fibres — Indirect-transfer transmission electron microscopy method. Marple, V. A., Rubow, K. L., and Behm, S. M., 1991, A micro orifice uniform deposit impactor (MOUDI): description, calibration, and use: Aerosol Science and Technology, v. 14, p. 434-446. Marple, V., Olson, B., Romay, F., Hudak, G., Monson Geerts, S., and Lundgren, D., 2014, Second Generation Micro-Orifice Uniform Deposit Impactor, 120 MOUDI-II: Design, Evaluation, and Application to LongTerm Ambient Sampling: Aerosol Science and Technology, v. 48-4, p. 427-433. MDH. Method 852 (1999) T.E.M. analysis for mineral fibers in air – 852. Minnesota Department of Health, Microparticulate Unit, St. Paul, MN. 42 pp. Oreskovich, J. A., and Patelke, M. M., 2006, Historical use of taconite byproducts as construction aggregate materials in Minnesota: A Progress Report: Natural Resources Research Institute Report of Investigation NRRI-RI-2006-02, 10 p. 60 �ROCK MAGNETISM AND PALEOMAGNETISM OF THE ~1144 MA LAMPROPHYRE DYKES, THE EASTERN LAKE SUPERIOR REGION, ONTARIO, CANADA JACOBSON, Darcy, M. Department of Physics, Michigan Technological University, Houghton, MI, PIISPA, Elisa J., SMIRNOV, Aleksey V. Department of Geological and Mining Engineering and Sciences, Michigan Technological University, Houghton, MI, and PESONEN, Lauri J. P., Department of Physics, Division of Geophysics and Astronomy, University of Helsinki, Helsinki, Finland Despite several decades of intensive research, the origin of the formation of the ~1.1 Ga Midcontinent Rift system (MRS) remains an open question. The proposed hypotheses vary from active rifting due to a mantle plume or plumes (e.g. Nicholson et al., 1997; Hollings et al., 2010), to passive rifting related to the Grenville Orogeny (Gordon and Hempton, 1986), or separation of the Amazon craton from Laurentia (Stein et al., 2014). Abundant ~1144 Ma lamprophyre dykes in the Eastern Lake Superior region (Ontario, Canada) (Queen et al., 1996) are coeval with the ~1141 Ma Abitibi diabase dyke swarm (Krogh et al. 1987) in the same area. In addition, both dyke suites share similar alcalic composition and appear to fan out from a locus in the present-day Lake Superior. These observations hint that the lamprophyre dykes and Abitibi dykes may form a single radiating dyke swarm representing the earliest magmatic stage of MRS. The existence of such a swarm is consistent with the arrival of a mantle plume, hence supporting the active rifting hypothesis. In order to test this hypothesis, we sampled 173 independently oriented samples from 22 lamprophyre dykes. In addition, at three sites, samples for the baked contact test were collected from the presumably baked and unbaked host rocks. The dependence of low-field magnetic susceptibility versus temperature indicates low titanium titanomagnetite as the dominant magnetic mineral. Subordinate hematite was observed in several samples. Magnetic hysteresis measurements reveal single to pseudo-single domain behavior in most dykes except for four dykes that show multidomain behavior. Both alternating field (AF) and thermal demagnetization were used to determine the paleomagnetic directions. In general, the AF demagnetization technique proved to be more effective in revealing the characteristic paleomagnetic directions. For several dykes, temperature treatment resulted in unstable demagnetization behavior due to alteration. Preliminary measurements of the anisotropy of magnetic susceptibility were also conducted on selected dykes in order to test whether the flow directions are consistent with the potential plume center. The paleomagnetic results of this study will be compared with the results obtained from the Abitibi dykes by Ernst and Buchan (1993) and the possible implications related to the formation of the MRS will be discussed. References Ernst, R.E., and Buchan, K.L. 1993. Paleomagnetism of the Abitibi dyke swarm, southern Superior Province, and implications for the Logan Loop. Canadian Journal of Earth Sciences, 30: 1886- 1897. Gordon, M. B., and M. R. Hempton (1986), Collision-induced rifting: The Grenville Orogeny and the Keweenawan Rift of North America, Tectonophysics, 127(1–2), 1–25, doi:10.1016/00401951(86)90076-4. Hollings, P., Smyk, M., Heaman, L.M., and Halls, H. 2010. The geochemistry, geochronology, and paleomagnetism of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada. Precambrian Research. Precambrian Research, v. 183, iss. 3, p.553-571. 61 �Krogh, T.E., Corfu, F., Davis, D.W., Dunning, G.R., Heaman, L.M., Kamo, S.L., Machado, N., Greenough, J.D., and Nakamura, N.1987. Precise U–Pb isotopic ages of diabase dykes and mafic to ultramafic rocks using trace amounts of baddeleyite and zircon; in Mafic Dyke Swarms, (ed.) H.C. Halls and W.F. Fahrig; Geological Association of Canada, Special Paper 34, p. 147–152. Nicholson, S. W., S. B. Shirey, K. J. Schulz, and J. C. Green (1997), Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: Implications for multiple mantle sources during rift development, Can. J. Earth Sci., 34(4), 504–520. Queen, M., Heaman, L.M., Hanes, J.A., Archibald, D.A., and Farrar, E. 1996. 40Ar/39Ar phlogopite and U–Pb perovskite dating of lamprophyre dikes from the eastern Lake Superior region: evidence for a 1.14 Ga magmatic precursor to Midcontinent Rift volcanism. Canadian Journal of Earth Sciences 33, 958–965. Stein, C. A., S. Stein, M. Merino, G. Randy Keller, L. M. Flesch, and D. M. Jurdy (2014), Was the Midcontinent Rift part of a successful seafloor-spreading episode?, Geophys. Res. Lett., 41, 1465–1470, doi:10.1002/2013GL059176. 62 �AN EXPLORATION UPDATE AND MINERALOGICAL STUDY OF THE EMILY-DISTRICT MANGANESE DEPOSIT, CUYUNA IRON RANGE, MINNESOTA 1 JOHNSON, Leif A. and 2DUNN, Brad M. Barr Engineering Company, 4700 W 77th St. Minneapolis, MN 55435 1 ljohnson@barr.com, 2bdunn@barr.com The Cuyuna Range in east-central Minnesota produced in excess of 100 million tons of manganiferous iron ore from initial discovery in 1904 to the final mine closing in 1984. The presence of higher percentages of manganese (greater than 10 percent) is the main component that distinguishes the Cuyuna Range from other Early Proterozoic iron mining districts in the Lake Superior Region. Similar to the Mesabi Iron Range, The Cuyuna Range has been documented as part of the Animikie Group, containing the Virginia Formation, "Unit A" iron formation (similar to the Biwabik Iron-Formation), and the Pokegma Quartzite (Morey and Southwick, 1993) Numerous works have studied the manganese deposits of the Emily-District. Summaries of the regional structural setting by Southwick et al. (1988) and Morey et al. (1981) show that deformation within the Emily-district is associated the Penokean Orogeny. Regionally, rocks within the Emily district form a broad synclinorium that plunges to the east. Morey and Southwick (1993) presented an in-depth summary of the stratigraphic and sedimentology characteristics that show possible geologic controls of manganese distribution. Dahl et. al. (1994) characterized the mineralogy for the purposes of utilizing insitu mining techniques using two holes drilled in 1990. Other studies have utilized historic drill core obtained from the Minnesota Department of Natural Resources core library in Hibbing. Cooperative Mineral Resources (CMR) controls 80 acres within the Emily-District. Historic resource estimates for this district showed one to two million tons of manganese resource. On their controlled property, thirteen historic drill holes were drilled, which indicated a sizeable manganese resource. In 2011 and 2012, CMR commissioned an expanded exploration drilling program with seven additional holes. This drilling confirmed the historic resource, which is divided into upper and lower manganese-rich zones. Assays from these modern drill holes showed stratigraphically continuous manganese grades of 15 to 20 percent. Further work has continued on the CMR Emily deposit. In 2013, metallurgical studies included a mineral liberation analyses (MLA). The MLA report analyzed fourteen core samples from the 2011 and 2012 drill holes using an automated scanning electron microscope equipped with energy dispersive detectors (SEM-EDS) and MLA software. The MLA technique analyzes X-ray spectrometry (XRF) from individual grains, assigns an elemental composition based on the geometric center of each grain, and assigns the most likely mineralogy based on a database of XRF. The results of the MLA confirmed the mineralogy described by past studies on the Emily-District. Predominant iron mineralogy was assumed to be hematite. Manganese mineralogy occurred in several manganese oxide (MnO) phases. Most notably, MnO was assumed to be the composition of manganite. Several manganese-iron (FeO-Mn) phases were found, which were not assigned a specific mineralogy. Crpytomelane (K12Mn8O16), hollandite (BaMn8O16), stilpnomelane (K(Fe,Mg)8(Si,Al)12(O,OH)27), and calcite_Mn ((Ca,Mn)CO3) were secondary manganese minerals. Quartz is the principal gangue mineral. Past metallurgical work on the manganese resource has shown that primary and secondary grinding, associated with flotation is insufficient in separating individual manganese-oxide from iron-oxide grains 63 �and quartz. The MLA report showed the mineral liberation at various grind meshes. These results will be further refined to develop an alternative processing technique for the manganese resource. References Dahl, L.J., Brink, S.E., Blake, R.L, Tuzinksi, P.A. and Adamson, N.R., 1994. Site characterization of Minnesota Manganese deposits for determining in situ mining potential: Society For Mining, Metallurgy, and Exploration Inc.: Transactions Volume 294, p. 1892-1905. Morey, G. B., Olsen, B. M, and Southwick, D. L., 1981, Geologic Map of Minnesota, east-central Minnesota, bedrock geology: Minneapolis, Minnesota Geological Survey, scale 1:250,000. Morey, G. B and Southwick, D. L., 1993. Stratigraphic and Sedimentological Factors Controlling the Distribution of Epigenetic Manganese Deposits in Iron-Formation of the Emily District, Cuyuna Iron Range, East-Central Minnesota: Economic Geology, v. 88, p. 104-122. Southwick, D. L., Morey, G. B., and McSwiggen, P. L., 1988. Geologic Map (scale 1:250,000) of the Penokean orogeny, central and eastern Minnesota, and accompanying text: Minnesota Geological Survey, Report, Investigation 37, 25p. 64 �SEDIMENTOLOGY AND PALEOGEOGRAPHIC RECONSTRUCTION OF THE LAYERS IN AND ADJACENT TO THE SUDBURY IMPACT LAYER IN THE LAKE SUPERIOR BASIN KARMAN, Monica M.1 and FRALICK, Philip W.1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, mmkarman@lakeheadu.ca, pfralick@lakeheadu.ca Various locations around the area of the Lake Superior Basin reveal stratified layers of the 1878.3±1.3 Ma (Fralick et al., 2002) Gunflint Formation, the 1850 Ma (Krogh et al., 1984) Sudbury Impact Layer, and the overlying 1832±3 Ma (Addison et al., 2005) Rove Formation. Samples were collected and tested, and stratigraphic logs were drawn from these locations to determine the sedimentology of the lithologic units in an attempt to reconstruct the paleogeographic setting at the time of deposition, along with diagenesis. Focus is given to two units, one above and one below the Sudbury Impact Layer (S.I.L), which display time gaps in the stratigraphic record. The units below and above the S.I.L. display a ~28 Ma, and an ~18 Ma year time gap respectively (Addison et al., 2010), indicating a period of subaerial exposure and ultimately, erosion; during this time the Rove Sea had regressed back to the southern edge of the continent. Samples collected below the S.I.L. near the Terry Fox monument and the Harbour Expressway in Thunder Bay, ON, indicate that subaerial exposure was present in this unit. The Terry Fox site displays stalactite-like structures (Figure 1A) composed of silica, indicating growth above the water table. The Harbour Expressway site reveals the same stalactite-like features, except that only a very small, eroded outcrop was found, showing plan-view stalactites termed by the author as silica flowerettes (Figures 1B, 1C). In addition, the Harbour Expressway site reveals botryoidal gypsum-like rosettes (Figure 2), indicative of an arid environment. Although these rosettes were most likely comprised of gypsum at one point in time, XRD and geochemical analysis show that this feature has been overprinted by calcite. Samples collected from the S.I.L. indicate that subaerial exposure affected this ejecta unit, made evident by silica invasion, carbonate crystal and cement growth or replacement. Ejecta features such as sphere-in-sphere structures (Figure 3A), and vesicular glass bubbles (Figure 3B) have been infilled, overprinted/replaced, and/or broken because of subaerial exposure. Samples collected above the S.I.L. display formation in subaerial conditions exemplified by a unit resembling gypsum laths sampled from drill core BDQ (Figure 4), collected near Hwy 588, Thunder Bay, ON. As with the gypsum-like rosettes found near the Harbour Expressway site, these gypsum laths had to have precipitated in an arid, subaerial environment. Another sample taken from drill core BDQ displays a chicken-wire texture of zoned carbonate crystals (Figure 5). SEM-EDX analysis of the crystals displays Mg and Fe zonation, indicating that formation occurred in a sabhka environment with access to meteoric water, rather than being covered by the Rove Sea. Although these zoned crystals were formed through the percolation of meteoric waters, the Rove Sea seems to have abruptly transgressed into the Animikie Basin, as seen by the silt and shale unit situated directly on top of the carbonate unit, seemingly choking it out. 65 �4 1A 1B 1C 2 3A 3B 5 References Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005, Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology, v. 33, n 3, p.193-196. Addison, W.D., Brumpton, G.R., Davis, D.W., Fralick, P.W., and Kissin, S.A., 2010, Debrisites from the Sudbury impact event in Ontario, north of Lake Superior, and a new age constraint: Are they base-surge deposits or tsunami deposits?, in Gibson, R.L., and Reimold, W.U., eds., Large Meteor Impacts and Planetary Evolution IV: Geological Society of America Special Paper 465, p. 245-268. Fralick, P.W., Davis, D.W., and Kissin, S.A., 2002. The age of the Gunflint Formation Ontario, Canada: single zircon U-Pb age determinations from reworked volcanic ash. Canadian Journal of Earth Sciences, v. 39, p. 1085-1091. Krogh, T.E., Davis, D.W., Corfu, F., 1984. Precise U-Pb zircon and baddeleyite ages for the Sudbury area. In, E.G. Pye ed., The Geology and Ore Deposits of Sudbury Structure. Ontario Geological Survey, Special Volume 1, p. 431-446. 66 �IMPACT EJECTA FEATURES IN THE LAKE SUPERIOR BASIN FROM THE 1850 MA SUDBURY IMPACT EVENT KARMAN, Monica M.1 and FRALICK, Philip W.1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, Canada mmkarman@lakeheadu.ca, pfralick@lakeheadu.ca Between the 1878.3±1.3 Ma (Fralick et al., 2002) Gunflint Formation, and the overlying 1832±3 Ma (Addison et al., 2005) Rove Formation, unconformably lies the 1850 Ma (Krogh et al., 1984) impact ejecta unit from the Sudbury Impact Event. These distal ejecta sites that have been discovered in the Lake Superior basin, extend approximately 600-800 kilometers, or ~5-7 crater radii (Spray et al. 2004), from the Sudbury impact crater. The Sudbury Impact Layer (S.I.L.) is composed of two constituents: 1) A chaotic debrisite portion that includes clasts and rip-ups of carbonate grainstone, along with blocks of chert and stromatolite that have been sheared from the underlying Gunflint Formation; 2) The ejected material from the impact event (Addison et al., 2010). Ejecta features of the S.I.L. unit include devitrified glass (Figures 1A, 1B), spherules (Figures 2A, 2B, 2C), planar features (Figure 3), and lapilli (Figure 4). During deposition of the 1850 Ma Sudbury Impact Layer, it is assumed that deposition took place in a subaerial environment (Fralick and Burton, 2008). Samples taken from two locations in the field directly below and above the S.I.L. unit display what seems to be gypsum rosettes (Figures 5A, 5B) indicating an arid environment (Karman and Fralick, 2014). Because of subaerial exposure, many ejecta features have been affected by carbonate replacement/alteration and silicification, either overprinting, or destroying them. Figures: Devitrified vesicular impact glass (DVIG); 1A: Oval DVIG with carbonate-filled vesicles (plane polarized light HEW site), 1B: Spherical DVIG with carbonate-filled void (plane polarized light BC site); 2 - Spherules replaced by chalcedony; 2A: Spherule cluster with irregular intricate banding (crossed polarized light HCP site), 2B: “LP record” spherules (plane polarized light BM site), 2C: Fibrous radial rimmed spherules with carbonate-filled previously hollow centers (crossed polarized light HEW site); 3 - Planar Deformation Feature (PDF) (crossed polarized light JN34 slide); Two sets of PDFs in quartz grain decorated with inclusions; 4 - Lapilli; Accretionary lapilli ~2.5cm width (588 site); 5 Replaced gypsum rosettes; 5A: Cross section of radially banded gypsum rosette replaced by carbonate directly under ejecta (HCP site), 5B: Cluster of botryoidal gypsum rosettes replaced by calcite growing in mat-like form overlying the S.I.L. (HEW site). 1A 1B 67 �2A 2B 2C 3 5B 5A 4 References Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005, Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology, v. 33, n 3, p.193-196. Addison, W.D., Brumpton, G.R., Davis, D.W., Fralick, P.W., and Kissin, S.A., 2010, Debrisites from the Sudbury impact event in Ontario, north of Lake Superior, and a new age constraint: Are they base-surge deposits or tsunami deposits?, in Gibson, R.L., and Reimold, W.U., eds., Large Meteor Impacts and Planetary Evolution IV: Geological Society of America Special Paper 465, p. 245-268. Fralick, P.W., Davis, D.W., and Kissin, S.A., 2002. The age of the Gunflint Formation Ontario, Canada: single zircon U-Pb age determinations from reworked volcanic ash. Canadian Journal of Earth Sciences, v. 39, p. 1085-1091. Fralick, P. W., and Burton, J., 2008, Geochemistry of the Paleoproterozoic Gunflint Formation carbonate: Implications for early hydrosphere-atmosphere evolution: Geochimica et Cosmochimica Acta, special supplement, v. 72, no. 125, p. A280. Karman, M. M., and Fralick, P. W., 2014, Sedimentology and paleogeographic reconstruction of the layers in and adjacent to the Sudbury Impact Layer in the Lake Superior Basin, M.Sc. Thesis (in progress), Lakehead University. Krogh, T.E., Davis, D.W., Corfu, F., 1984. Precise U-Pb zircon and baddeleyite ages for the Sudbury area. In, E.G. Pye ed., The Geology and Ore Deposits of Sudbury Structure. Ontario Geological Survey, Special Volume 1, p. 431-446. 68 �PDFS IN SUDBURY EJECTA IN THE GUNFLINT FORMATION, ONTARIO: A COMPARISON OF METHODS KISSIN, Stephen A. and BRUMPTON, Gregory R. Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Planar deformation features (PDFs) in quartz are considered to be definitive evidence, along with other features, of large terrestrial impact (French 1998). The ejecta from the 1850 Ma Sudbury impact structure has been described in the Animike Group and correlative units in the western Lake Superior area in Michigan, Minnesota and Ontario (Cannon and Addison 2007). PDFs are known to develop in response to shock pressures greater than those developed by terrestrial events (e.g. volcanic eruptions) and form on rational planes in the quartz crystal (von Engelhardt and Bertsch 1969). Indexed planes of PDFs from the Baraga Group and correlative units from Michigan have been reported by Pufahl et al. (2007) and Cannon et al. (2010); however, although PDFs in Gunflint Formation of Ontario have been described (Addison et al. 2005), they have not been indexed. In this study, we have indexed 22 PDFs in 11 thin sections of ejecta from drill holes BP99-2 and MC95-1, which have been described by Addison et al. (2005). The conventional method for indexing of PDFs, as described by von Engelhardt and Bertsch (1969), utilizes the universal stage to determine the pole of a PDF and the orientation of the c-axis of the quartz grain. Based on the polar angle between the c-axis and the pole of the PDF, the results are then compared to a template of the projection of known indices of PDFs in a stereographic projection. The standard template devised by von Engelhardt and Bertsch (1969) allows a margin of error of 5º in the stereographic projection. The conventional method described above suffers from some imprecision in the plotting and manipulations of measurements on a Wulff net, which must be done manually. In order to overcome errors of this sort, Huber et al. (2011) created the Automated Numerical Index Executor program (ANIE), which is based on angular calculations from spherical trigonometry. They demonstrated the apparent superiority of use of the program over manual methods in indexing of PDFs. In our study, we have compared the application of the conventional method with that of the ANIE program, as well as presenting results from the Gunflint Formation, which lies approximately 200 km west of the Baraga Group and hence, more distally from the Sudbury structure. Quartz grains examined in this study in most cases contained only one orientation of PDF in the thin section examined. An example of an exceptional grain with a high density of PDFs in two orientations is shown in the figure. The Miller-Bravais indices of PDFs examined, according to the ANIE program, with numbers determined in parentheses are as follows in order of increasing polar angle: {1014} (1set); {1013} (1 set); {1012} (1 set); {1122} (1 set); {1011} (1 set); {1121} (4 sets); {2131} (2 sets); {2241} (1 set); {3141} (2 sets); {5161} (6 sets); {5160} (1 set); unindexed (1 set). There is good agreement with the planes recorded by Cannon et al. (2010); however, they found most planes to be of low index, whereas the high index planes, e.g. {5161}, of this study are indicative of high shock intensity. Pufahl et al. (2007) found most of the planes seen in this study with a fairly uniform number of occurrences. Using the conventional method with the Wulff net, a number of planes were "near misses" on the template and thus would appear to be unindexed. This a result similar to that found by Huber et al. (2011), which indicates that the ANIE program is superior to the conventional method. As well, the format for input of measurements from the universal stage accepts an error of ± 1º, providing for a more realistic accounting of error of measurement. The program also can provide output in tabular form for ease of recording of data. 69 �Figure 1. A "toasted" quartz grain from section JN34 with two indexed directions of PDFs. References Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W. and Hammond, A.L. 2005. Discovery of distal ejecta from the 1850 Ma Sudbury impact event, Geology 33: 193196. Cannon, W.F., and Addison, W.D. 2007. The Sudbury impact layer in the Lake Superior iron ranges: A time-line from the heavens: Institute on Lake Superior Geology, Proceedings 53, Part 1: 20-21. Cannon, W.F., Schulz, K.J., Horton, J.W. Jr. and Kring, D.A. 2010. The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA, Geological Society of America Bulletin 122: 50-75. French, B.M. 1998. Traces of Catastrophe: A Handbook of Shock-Metamorphic Effects in Terrestrial Meteorite Impact Structures, Lunar and Planetary Institute , Contribution 954. Huber, M./S., Ferriére, L., Losiak, A. and Koeberl, C. 2011. ANIE: A mathematical algorithm for automated indexing of planar deformation features in quartz grains, Meteoritics and Planetary Science 46: 1418-1424. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, G.J. and Edwards, C.T. 2007. Physical and chemical evidence of the 1850 Ma Sudbury impact event in the Baraga Group, Michigan, Geology 35: 827-830. von Engelhardt, W. and Bertsch, W. 1969. Shock induced planar deformation structures in quartz from the Ries crater, Germany, Contributions to Mineralogy and Petrology, 20: 203-23. 70 �GEOCHEMISTRY AND MINERALOGY OF FE-TI-V-P MINERALIZED FERROGABBROIC INTRUSIONS OF THE MCFAULDS GREENSTONE BELT, SUPERIOR PROVINCE, NORTHERN ONTARIO, CANADA KUZMICH, Ben1, HOLLINGS, Pete1, HOULÉ, Michel G.2 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Geological Survey of Canada, GSC-Quebec, 490 rue de la Couronne, Québec, Quebec G1K 9A9 2 The McFaulds Lake area (i.e., Ring of Fire) located in northern Ontario (Canada) has been the site of recent exploration leading to the discovery of several mineralization types including chromite and nickel sulfide deposits. Although the majority of exploration has been focused on chromium, the area also contains significant Fe-Ti-V-P mineralization associated with gabbroic intrusions, of which the Thunderbird and Butler occurrences are the best defined. The study has been focused on the geochemical and petrographic characterization of the intrusions to investigate their petrogenesis. The intrusions are widely distributed throughout the McFaulds Lake area and can be grouped into two main types: (1) large mafic-dominated intrusions and (2) subconcordant to slightly discordant mafic-dominated sills/dikes characteristic of the Thunderbird and the Butler intrusions respectively. Both types are composed of an evolved mafic suite termed the ‘Ferrogabbro’ characterized by the presence of Fe-Ti oxides. Detailed core logging has shown that both intrusions are largely composed of very similar lithologies including iron-rich gabbros, leucogabbros, and anorthosites. Two types of Fe-Ti oxide mineralization occur within these intrusions: (1) Fe-Ti-V and (2) Fe-Ti-P mineralization. Fe-Ti-V mineralization has been intersected within both intrusions, whereas the Fe-Ti-P mineralization has only been identified within the Thunderbird intrusion. The mineralization occurs dominantly as disseminated magnetite and ilmenite (1-10%), but is also present as semi-massive (50-80%), to massive layers (>80%: Fig. 1). These layers typically contain distinct sharp, stratigraphically lower contacts and gradational upper contacts typical of primary igneous layering (Fig. 2). The ilmenite and magnetite occurs as anhedral to subhedral crystals and to a lesser extent, as very fine-grained exsolutions within anhedral magnetite grains. 71 �Figure 1. Massive magnetitite from Butler East intrusion (BP11-V01). Figure 2. Magmatic layering within the Thunderbird intrusion (NOT09-2G25). This research project has also addressed the use of TiO2/V2O5 ratio as a potential vector towards vanadium and/or phosphorous mineralized horizons within the ferrogabbroic intrusions of the McFaulds Lake area. Preliminary data in the Butler intrusions, strongly suggest that the ratio TiO2/V2O5 ratio values are independent of rock type, abundance of magnetite-ilmenite, and/or alteration and could be useful in determining favorable horizons for further vanadium mineralization. Furthermore, the samples from this intrusion that exhibit significant vanadium contents (>0.50 weight % V2O5) are restricted to a narrow range of TiO2/V2O5 values (between 8 and 12). This ratio has the potential to be a significant exploration tool to target magmatic Fe-Ti-V-P mineralization and it has also been an instrumental tool interpretation of stratigraphy of the Butler and Thunderbird intrusions. The ferrogabbroic intrusions may be petrogenetically related to the abundant ultramafic rocks within the McFaulds Lake area, and could possibly represent the late stage end member of a magmatic sequence as has been suggested for the Bushveld complex. However, the rare or absent ultramafic components spatially associated with these ferrogabbroic intrusions, combined with some ultramafic units crosscutting the ferrogabbroic units within the Butler intrusion, may suggest that they could represent two distinct magmatic events rather than a dismembered layered intrusion, as proposed by previous workers. 72 �STRUCTURAL CONTROL ON THE BORDEN GOLD DEPOSIT IN CHAPLEAU, ONTARIO D.J. LaFontaine1 and M.L. Hill1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario, P7B 5E1, djlafont@lakeheadu.ca The Borden gold deposit is located 20 km east of Chapleau, 180 km southwest of Timmins, within the Kapuskasing Structural Zone and Wawa subprovince of the Superior Province. The deposit is a low grade, bulk tonnage style hosting 4.3 million ounces at 1.03 g/t Au in upper amphibolite to granulite facies metamorphic rocks. The metamorphic minerals at Borden include biotite, muscovite, hornblende, sillimanite, garnet, kyanite, cordierite and pyroxene. Based on the abundance of aluminous metamorphic minerals, the protolith is inferred to be pelitic. At such high metamorphic temperatures, deformation is dominantly by ductile mechanisms although microfracturing of competent minerals is also possible. On the macroscopic scale, gold mineralization seems to be controlled by strain heterogeneity related to metamorphic grade and competency. Competent lithons (boudins?) of granulite facies rock appear to be surrounded by more ductile amphibolite facies gneisses and schists, suggesting polymetamorphism with retrograde amphibolite facies metamorphism after granulite facies metamorphism. Gold is typically observed in low strain rocks with weakly developed foliation and also in low strain rocks that are bordered by strongly foliated units. Gold mineralization has been observed at grain boundaries of quartz, within cleavage planes of biotite and associated with euhedral pyrite and anhedral pyrrhotite. This project will provide specific structural and microstructural parameters to guide further exploration and development of the mineralized zone at the Borden gold deposit. 73 �74 �RU-RH-PD MOBILIZATION IN FLAMBEAU MASSIVE SULFIDE DEPOSIT LAMB, Matthew T., and BHATTACHARYYA, Prajukti Department of Geography and Geology, UW-Whitewater, 120 Upham Hall, 800 Main St., Whitewater, WI, 53190 The Flambeau deposit located in Rusk county Wisconsin is a Volcanogenic Massive Sulfide (VMS) deposit that plays host to large amounts of copper (chalcopyrite) and zinc (sphalerite). The Flambeau deposit is located in the Ladysmith-Rhinelander volcanic complex and has a felsic center which is a steeply dipping section of interlayered fragmental quartz-sericite, andalusitebiotite, quartz-eye, chlorite-garnet, actinolite, and chlorite schist (DeMatties, 1994). The deposit has gone through major alteration due to preliminary sulfide mineralization, regional metamorphism, and supergene alteration near the surface (May and Dinkowitz 1996). However, the characteristics of hydrothermal alterations associated with sulfide mineralization at deeper levels of Flambeau deposit have not been well studied. The goal of this research project is to study how hydrothermal fluids at Flambeau deposit concentrated Cu, Zn, and Fe sulfides, and mobilized other metals such as Ru, Rh, Pd, Ti, and Cr in the process. In order to accomplish this goal I am analyzing a core sample from the Flambeau deposit using a Bruker handheld X-Ray Fluorescence (XRF) analyzer. I am collecting data from primary bedrock layers (sericite-rich layers with little or no sulfide), discrete, thin (1-2 cm think) sulfide bands within the host rock, bedrock with visible amounts of sulfide minerals (mixed layers), sharp boundary layers between bedrock and massive sulfide layers, and massive sulfide layers with no visible bedrock. Preliminary data collected from the depths of 620 feet to 640 feet (which is within the hypogene massive sulfide deposit) show that Ru and Pd concentrations are lower within the massive sulfide layers, and progressively increase towards the sericite-rich bedrock layers, while Rh concentration progressively decreases going from massive sulfide layers towards bedrock (Figure 1). Besides Ru, Rh, and Pd, no other PGE, or common “pathfinder” elements like Ba, Co, Ni, etc. are present in the samples in detectable quantities. Ti concentrations are higher in bedrock layers compared to the sulfide layers, but Cr concentrations stay relatively constant in all the layers. The distribution patterns of Ru, Rh, and Pd associated with Flambeau VMS deposit might provide important insights regarding how sulfide mineralization may affect the distribution of trace elements already present in host rocks. Results from this research can potentially help in future explorations for other, similarly formed VMS deposits around the world. References DeMatties, Theodore (1994). Early Proterozoic Volcanogenic Massive Sulfide Deposits in Wisconsin: An Overview. Economic Geology. 1994: 1122-1151 May, Edwarde, and Dinkowitz, Stephen (1996). An Overview of the Flambeau Supergene Enriched Massive Sulfide Deposit: Geology and Mineralogy, Rusk County, Wisconsin. Volcanogenic Massive Sulfide Deposits of Northern Wisconsin: A commemorative volume: (LaBerge, G. L., Ed), Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, Cable, WI, v. 42, part 2, 67-93 75 �Massive sulfide Ru Sulfide bands in bedrock Boundary Bedrock and mixture Rh Pd Figure 1. Change in Ru, Rh, and Pd within Massive Sulfide, Boundary Layers, and Sericite-Rich Bedrock Layers (620-640 Ft). 76 �GEOLOGY AND PETROLOGY OF THE WILDER LAKE INTRUSION, DULUTH COMPLEX, NORTHEASTERN MINNESOTA LEU, Adam, and MILLER, Jim Department of Geological Sciences, University of Minnesota Duluth, Duluth, Minnesota 55812. In a massive forest fire in the autumn of 2011, 160 square miles of dense forest in the Boundary Waters Canoe Area Wilderness (BWCAW) was intensely burned. Underlying what was termed Pagami Creek burn area are mafic intrusive rocks of the 1.1 Ga Duluth Complex - a large, arcuate-shaped, multiple intrusive igneous complex that underlies most of northeastern Minnesota and that constitutes the largest exposed plutonic component of the 1.1 Ga Midcontinent Rift. Situated in the center of the burn area is the Wilder Lake Intrusion (WLI), an incompletely studied, northward-dipping sheet-like mafic layered intrusion known only from reconnaissance mapping of its lake-accessible western extent. The intense burn created a time-sensitive opportunity to access freshly exposed outcrops along the entirety of its 10 km strike length. The WLI is part of the Layered Series – a collection of mafic layered intrusions within the Duluth Complex emplaced near the base of a comagmatic volcanic edifice (North Shore Volcanic Group). While most Layered Series intrusions were emplaced beneath an earlier intrusive suite of anorthositic rocks (Anorthositic Series; Fig. 1), the WLI was intruded entirely within Anorthositic Series rocks (Miller et al., 2002), The overall objective of this study is to document the igneous stratigraphy along its entire strike-length with the goal of better understanding the composition and emplacement and crystallization history of its parental magma(s) that produced its unique petrologic attributes. These attributes, noted by others and confirmed here, include 1) its unique cumulate stratigraphy where Fe-Ti oxide becomes a cumulus phase before augite; 2) the cumulus reversal indicated by the change from a four-phase cumulate of Pl+Cpx+Ox+Ol abruptly giving way up section to a troctolitic (Pl+Ol) cumulate; and 3) its reversed cryptic variation of Fo in olivine and En’ in clinopyroxene (Miller and Ripley, 1996). Detailed field mapping (1:12,000), petrographic observations, and geochemical analyses were conducted to accomplish these goals and objectives. WLI was first discovered by reconnaissance mapping by Phinney (1972) who documented exposures of well-foliated and layered gabbros and troctolites that extend from North Wilder Lake to the west and Arrow Lake to the east, a strike-length of about 10 kilometers. Phinney noted that internal layering and foliation dips to the north to northeast between 15° and 35°, which contrasts with the southerly to easterly (riftward) dip of most layered series intrusions of the Duluth Complex. Reconnaissance mapping and follow-up petrologic studies were conducted by Jim Miller who noted a distinct cumulate reversal (Pl+Cpx+Ox+Ol  Pl + Ol) in the upper section of the western portion of the intrusion, as well as identifying a reversed cryptic variation of upwardly increasing En’ and Fo content of pyroxene and olivine, respectively (Miller, 1986; Miller and Ripley, 1997). Unpublished field, petrographic and geochemical data also collected in the western WLI by Joy Turnbull in 2004 verified the reversed cryptic variation and phase layering in the western part of the WLI. Detailed mapping conducted in 2012 and 2013 for this study shows that most cumulate units of the WLI can be followed along the entire 10km strike length of the WLI, but with some notable exceptions. Remapping in the western part of the WLI has confirmed that the 2km-thick igneous stratigraphy exposed here starts with a basal unit of heterogeneous, intergranular olivine oxide gabbro that is in sharp contact with Anorthositic Series rocks. This marginal gabbro is overlain by a troctolitic unit of Pl+Ol cumulates, which can be subdivided into a heterogeneous subunit, a layered subunit and an anorthosite inclusion subunit. The troctolite unit is overlain by a thin (20-100m thick) oxide troctolite unit of Pl+Ol+ Ox cumulates which abruptly gives way to an olivine oxide gabbro unit of Pl+Cpx+Ox+Ol cumulates. Above the gabbro, an upper troctolitic unit occurs marking a cumulus reversal back to Pl+Ol cumulates. Mapping of the excellent exposures created by the burn reveal that the upper troctolite unit cross-cuts and locally scours out the four-phase gabbro. Thus it is interpreted as a recharge of more primitive magma into the upper part of the WLI chamber rather than a downward crystallizing roof zone unit as proposed by Miller (1986). Detailed mapping by overland traverses in the central and eastern extents of the WLI show it to thin from 2 km in the west to 1 km in the east. Moreover, several units pinch out in the eastern section of the intrusion. The oxide troctolite pinches out just east of center, but swells back to about 20 meters in thickness before pinching out again with the upper gabbro farther east. The lower gabbro also pinches out around the same place and is replaced by a taxitic unit that dominates at the eastern margin; a similar heterogeneous unit also can be found at the western margin. 77 �Petrographic studies of 223 thin sections collected along three profiles across the intrusion at its western, east central and eastern extents helped to confirm and refine the mineralogy and textures described from field observations. In addition, olivine and pyroxenes from many of the thin section samples were analyzed by UMD’s SEM-EDS to document cryptic variation of Mg/Fe ratios. This mineral chemical data was acquired to verify the reversed cryptic variation previously documented in the west and to determine if this variation persists along strike to the east. Reversed cryptic variation of upwardly increasing magnesium number (mg#=MgO/(MgO+FeO), mole%) in olivine (Fo) and pyroxene (En’) was confirmed in the west and in the eastern profiles. However, the data also reveal that the mg# tends to decrease at a particular stratigraphic horizon from west to east. We interpret the reversed cryptic variation up section to be due to a reduced trapped liquid shift within the oxide troctolite and olivine oxide gabbro units. Trapped liquid shift occurs where high mg# cumulus olivine re-equilibrates with low mg# intercumulus liquid. As evidenced by their strong foliation, these rocks are adcumulates with very little intercumulus minerals (i.e., trapped liquid component) and thus retain their high-mg# cumulus compositions. The lateral decrease in mg# to the east as well as the disappearance of the oxide troctolite unit is thought to be caused by the thinning of the intrusion causing the eastern portion of the intrusion to cool more rapidly than the west. This more rapid cooling would have caused more trapping of intercumulus liquid (and thus a stronger trapped liquid shift) and would have promoted oxide and pyroxene to crystallize more synchronously since their liquidus temperatures are not very different. We are currently evaluating whole rock analyses of the basal intergranular gabbro samples to determine if they may be representative of a parental liquid composition. We are using a MELTS-based modeling program, Pele (Boudreau, 2006), to evaluate the phase equilibrium of these compositions and to see if fractional crystallization of these compositions under different conditions of oxygen fugacity and cooling rate can replicate the cumulate stratigraphy observed in the WLI. REFERENCES Boudreau, A. , 2006, Pele. (7.07). Computer modeling program. Duke University. www.nicholas.duke.edu/eos/ Miller, J.D., Jr., 1986, The geology and petrology of anorthositic rocks of the Duluth Complex, Snowbank Lake quadrangle, northeastern Minnesota (PhD thesis) University of Minnesota Miller, J.D. Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.E., and Wahl, T.E., 2002, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota. Minnesota Geological Survey Report of Investigations 58, 207p. w/ CD-ROM Miller, J.D., Jr., and Ripley, E.M., 1996, Layered Intrusions of the Duluth Complex, Minnesota, USA. In: Cawthorne, R.G., Layered Intrusions: Amsterdam, Elsevier Science, p. 257-301 Phinney, W.C., 1972. Northwestern part of Duluth Complex. In: Sims, P.K. & Morey, G.B. (eds.) Geology of Minnesota -A centennial volume. Minnesota Geological Survey, p. 335-345 Wilder Lake Intrusion Figure 1. Generalized geology of NE Minnesota showing the location of the Wilder Lake Intrusion and the Pagami Creek Burn Area. Pagami Creek Burn Area 78 �THE ROLE OF BRITTLE-DUCTILE DEFORMATION AND COMPETENCY CONTRAST IN GOLD MINERALIZATION IN THE C-ZONE AT HEMLO LIIMU, Jared and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Rd., Thunder Bay, ON, P7B 5E1, Canada The Hemlo gold mine is located along the Hemlo shear zone within the Hemlo-Schreiber greenstone belt (Muir, 2002). This greenstone belt lies within the Archean Wawa subprovince of the Superior province (Muir, 2002). The characteristics of an active mineralized stope in the Czone were studied. The area was interpreted to be under amphibolite facies metamorphism, based on ductile deformation of feldspar, and the presence of sillimanite. This is consistent with findings by Powell et al. (1999), who also noted a peak kyanite phase followed by a decrease in pressure but a relatively minor decrease in temperature. Lingering high temperatures are supported by our observation of annealed grain boundaries. Brittle-ductile deformation is present throughout the C-zone, on all scales of measure. Figure 1 shows evidence of this on a microstructural scale, where quartz shows evidence of ductile deformation (undulose extinction and subgrains), as well as fracturing. Different minerals show evidence for different behaviours; for example, clinozoisite deformed in a purely brittle manner. Mutually overprinting brittle and ductile deformation, as well as competency contrasts are also evident on outcrop scale. Competency contrast within the area of study is most obviously seen between the more competent metavolcanics and more ductile biotite schist. Quartz veins within the metavolcanics tend to have relatively straight contacts with the host rock, whereas quartz veins in the biotiteschist tend to exhibit boudinage or pinch and swell textures. Quartz veins in the metavolcanics are localized within this unit and do not extend into the surrounding biotite-schist (Fig. 2). Greenschist facies pressure and temperature are considered by many to be the perfect conditions for simultaneous brittle-ductile deformation due to competency contrasts (Weinberg et al., 2012), which in turn provides an ideal Figure 1: Photomicrograph of a microshear zone with evidence for brittle and ductile deformation within quartz. Figure 2: More competent metavolcanic unit (1) and the less competent biotite-schist (2). 79 �setting for gold mineralization. Results from this project show that similar mineralization can occur under amphibolite facies conditions where ductile deformation dominates. There are three modes of mineralization observed within the C-zone. The first two include mineralization of gold along competence contrast boundaries. The third involves mineralization of gold within metavolcanic hosted fractures. The role competency plays in gold mineralization is two-fold. The first is that competent bodies tend to fail via brittle deformation. This allows pore space for gold-hosting fluids to infiltrate. This may be the basis for mineralization within the metavolcanic hosted fractures. The second role competency plays involves high-temperature diffusion of gold-hosting fluids along boundaries between competent and ductile lithologies. This causes mineralization, a) within the biotite-schist along the boundary of quartz boudins, and b) within the metavolcanics along the boundary with the biotite-schist. References Muir, T., 2002. The Hemlo gold deposit, Ontario, Canada: principal deposit characteristics and constraints on mineralization, Ore Geology Reviews, v. 21, p. 1-66 Powell, W., Pattison, D., Johnston P., 1999. Metamorphic history of the Hemlo gold deposit from Al2SiO5 mineral assemblages, with implications for the timing of mineralization, Canadian Journal of Earth Sciences, v. 36, p. 33-46 Weinberg, R., Groves, D., Hodkiewicz, P., van der Borgh, P., 2012. Controls on gold endowment: Shear Zone Comparison, Hydrothermal Systems, v. 3, p.101-108 80 �VARIABLE COPPER MINERALIZATION IN THE LOWER NONESUCH FORMATION OF THE MIDCONTINENT RIFT SYSTEM: CONSTRAINTS ON REGIONAL CONTROLS MAUK, Jeffrey L.1, WOODRUFF, Laurel G.2, and STEWART, Esther3 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 U.S. Geological Survey, 2280 Woodale Ave, St. Paul, MN 55112 3 Wisconsin Geologic and Natural History Survey, 3817 Mineral Point Road, Madison, Wisconsin 537055100 2 The lower Nonesuch Formation in the Lake Superior portion of the Midcontinent Rift System is a host of major sediment-hosted stratiform Cu mineralization, including the White Pine mine, which produced more than 2 Mt Cu, and the Copperwood deposit, which has measured, indicated, and inferred resources of 0.5 Mt Cu. However, the lower Nonesuch Formation is not uniformly mineralized; instead some areas are relatively well-endowed with Cu mineralization, whereas other areas, such as the Ashland syncline in northern Wisconsin, contain only trace quantities of Cu. These variations in Cu content lead to questions on the possible first-order controls of regional to local mineralization, and we critically evaluate new and preexisting data to help identify possible controls on Cu mineralization. The western Lake Superior portion of the Midcontinent Rift contains clastic sedimentary rocks of the Oronto Group. The basal unit is the conglomerate, sandstone, and siltstone that form red beds of the Copper Harbor Conglomerate. This is overlain by gray siltstone, shale, and fine-grained sandstone of the Nonesuch Formation, which is overlain by reddish brown sandstone of the Freda Formation. Sediment-hosted stratiform Cu deposits such as White Pine and Copperwood occur in the lowest gray beds, which contain organic matter and diagenetic pyrite that can serve as a reductant for cupriferous brines that are introduced from red beds of the underlying stratigraphy. Previous work on the lower Nonesuch has documented that the total organic carbon and sulfur contents are similar in the Ashland syncline and the White Pine-Copperwood area, suggesting that the sedimentary rocks in both places had similar reduction potential for trapping metallic minerals. The hydrothermal fluid that transported metals to the reduction site is widely accepted to be oxidized cupriferous brines. Neither the Ashland syncline nor the White Pine-Copperwood area have abundant evaporite minerals, but both contain evidence for local to minor evaporite minerals in the lower Nonesuch Formation, and the red beds of the Copper Harbor Conglomerate underlie the Nonesuch Formation in both areas, so we infer that the carrying capacity of the diagenetic basinal fluids was similar in both areas. The ultimate source of the Cu is typically interpreted to be the basalt that underlies the Oronto Group, and because both areas are underlain by extensive basalt, we infer that a source of Cu was not the greatest limitation. The sedimentary facies in each of the main formations of the Oronto Group are similar in both areas; this suggests that first-order control of fluid flow by major facies variations was not a likely control on different Cu endowments. However, the detailed stratigraphy of the lower Nonesuch Formation, which shows remarkable continuity in the White Pine-Copperwood area, differs in the Ashland syncline: most of the marker beds in the White Pine-Copperwood area are poorly developed and rarely occur in the Ashland syncline. This observation is consistent with previous interpretations that the Ashland syncline lies within a different subbasin that was either partially or completely separated from the large rift basin that hosts the White Pine and Copperwood deposits. 81 �Taken together, these results suggest that one major constraint on the Cu endowment of favorable strata in the Nonesuch may have been the size of the available mineralizing diagenetic fluid source. The White Pine-Copperwood deposits are on the margin of a large rift basin that would have been able to contribute significant quantities of mineralizing fluid. In contrast, the presumably smaller size and restricted connectivity of the rift sub-basin in the Ashland syncline area would have had a smaller fluid source thereby limiting the potential metal endowment in that area. 82 �SEDIMENTOLOGY AND GEOCHEMISTRY OF THE MESOARCHEAN CHEMICAL SEDIMENTS OF WALLACE LAKE AND RED LAKE MCINTYRE, Tim and FRALICK, Philip, Department of Geology, Lakehead University, Thunder Bay, ON, Canada, P7B 5E1, philip.fralick@lakeheadu.ca In the western Superior Province, Mesoarchean (~2.93Ga) carbonates and iron formation of the Uchi Subprovince represent a large carbonate platform that extended between the Wallace Lake and Red Lake greenstone belts. The aerial extent, sedimentary structures, and geochemistry of the platform indicate that significant changes in oceanic processes were occurring in the Mesoarchean. These changes include the precipitation of aerial extensive platform carbonates and evidence for the addition of free oxygen to limited areas of the hydrosphere by its initial production on semi-restricted platforms. Paleoarchean marine carbonates consisted of thin bedded units precipitating from anoxic water basins (ex. Strelley Pool Chert (Allwood et al., 2010)). Geographically scattered large Neoarchean carbonate platforms show evidence for the gradual build-up of oxygen during this time and leading to a relatively oxygenated atmosphere by 2.4Ga (ex. Steep Rock platform (Fralick and Riding, in press) and the CampbellrandMalmani platform (Kendall et al., 2010)). The occurrence of this large carbonate platform in the earliest transitional period of this change and its similarities to younger Neoarchean platforms is significant in that the information gathered has significant import to processes responsible for the change in carbonate deposition through the Paleoarchean to Paleoproterozoic. The lithofacies of the Wallace Lake and Red Lake carbonate platform represent deposition from peritidal to basinal environments, with many of the structures being present in younger Neoarchean and Proterozoic carbonate platforms. The peritidal lithofacies assemblage consists of herringbone calcite, pseudomorph fans, and tufa. The sub-tidal environment is characterized by large pseudomorph fans (Figure 1A) and laterally linked domal stromatolites (Figure 1B). Upper slope environments consisted of slumps, ribbon rock, and carbonate associated oxide-facies iron formation. Chert-oxide facies iron formation defines the basinal environment. These lithofacies typify younger Neoarchean carbonate platforms contributing to the gradual oxidation of the atmosphere (cf. Sumner and Grotzinger, 2004; Kendall et al., 2010; Fralick and Riding, in press). The rare earth element (REE) geochemistry of the Wallace Lake and Red Lake chemical sediments (Figure 2) can significantly contribute to our understanding the changing early oceans. The PAAS normalised REE patterns (REE(PAAS)) of the basin lithofacies (oxide-facies iron formation) is characterized by LREE depletion and positive Eu anomalies (Figure 2). This pattern mirrors REE geochemistry of Paleoarchean oceans (cf. Allwood et al., 2010). However, the carbonates show very little Figure 1. A) Pseudomorph fans after aragonite. B) Laterally linked domal LREE/HREE fractionation, positive La, Eu, and Y anomalies, stromatolites. and negative Ce anomalies. This implies a significant change A B 83 �REE(PAAS) in ocean chemistry from basin to the shallow carbonate platform. The transitional facies between basin and peritidal platform (upper slope) is characterized by a REE(PAAS) pattern similar to that of the platform carbonates. The significant difference in basin and platform REE(PAAS) patterns and the REE(PAAS) pattern of the upper slope suggests that the platform was semi-restricted and evaporitic. This would lead to density contrasts between the shallow platform and basin waters allowing for seeping and down-welling of platform waters to the basin and imparting the shallow carbonate REE(PAAS) pattern to the upper slope environment. The presence of Ce anomalies in the 1 Average Carbonate (n=28) carbonates and lack thereof in the basin Average Oxide-Facies Iron Formation (n=6) iron formation is indicative of a redox boundary separating the basin and peritidal Carbonate Associated Iron Formation (n=2) environments. The negative Ce anomalies 0.1 are the largest and most consistent in the pseudomorph fan facies. These pseudomorph fans are common structures found in Precambrian carbonate platforms, 0.01 range in depositional environments from La Ce Pr Nd Sm Eu Gd Tb Dy Y Ho Er Yb Lu peritidal to sub-tidal, and are thought to be pseudomorphs after aragonite (Sumner and Grotzinger, 2004). The fans of the Wallace Lake and Red Lake carbonate platform are also characterized by increased Ba, Sr, P, and K concentrations. The presence of the negative Ce anomalies in the fans and lack thereof in most of the rest of the platform indicate that the fans precipitated from relatively oxygenated water. It is suggested here that the fans are not indicative of a particular environment of deposition, but limited to periods of relative oxygenation of the platform and, when occur, are a platform wide occurrence. To summarize: the Wallace Lake and Red Lake carbonate platform is the oldest known carbonate platform in a series of geographically scattered platforms contributing to a gradual change in oceanic processes in the Mesoarchean and Neoarchean that led to an extensively oxygenated atmosphere by 2.4Ga. Figure 2. REE plot normalized to Taylor and McLennen’s (1989) Post Archean Australian Shale (PAAS) (n refers to the number of samples analyzed). References Allwood, A., Kamber, B. S., Walter, M. R., Burch, I. W., and Kanik, I. (2010). Trace elements record depositional history of an Early Archean stromatolitic carbonate platform. Chemical Geology, 270(1), 148-163. Fralick, P. and Riding, R. (in press). Anatomy and geochemistry of an Archean carbonate platform. Kamber, B. (2001). The geochemistry of late Archaean microbial carbonate: Implications for ocean chemistry and continental erosion history. Geochimica Et Cosmochimica Acta, 65(15), 2509-2525. Kendall, B., Reinhard, C.T., Lyons, T.W., Kaufman, A. J., Poulton, S.W., and Anbar, A.D., (2010). Pervasive oxygenation along late Archaean ocean margins. Nature Geoscience (3), 647- 652. Sumner, D. and Grotzinger, J., (2004). Implications for Neoarchaean ocean chemistry from primary carbonate mineralogy of the Campbellrand-Malmani Platform, South Africa. Sedimentology, 51(6), 1273-1299. 84 �COMPOSITION AND 40AR/39AR AGE OF PEGMATITIC AMPHIBOLE IN THE WAUSAU SYENITE COMPLEX, MARATHON COUNTY, WISCONSIN MEDARIS1, Gordon Jr., FLOOD2, Tim, JICHA1, Brian and SINGER1, Bradley 1 Department of Geoscience, University of Wisconsin-Madison, Madison, WI 53706 St. Norbert College, De Pere, WI 54115 medaris@geology.wisc.edu, tim.flood@snc.edu, bjicha@geology.wisc.edu, bsinger@geology.wisc.edu 2 Granitic pegmatites are of great interest for their unusual and exotic minerals. Because they are highly differentiated chemically, they may also contain silicate minerals of endmember compositions, such as arfvedsonite, aegirine, and albite. End-member amphibole crystals ≤ 30 cm in length occur in an NYF (niobium-yttrium-fluorine) granitic pegmatite that intrudes quartz syenite of the Wausau Complex (Fig. 1). The Wausau Complex is one of four, alkaline syenitic to granitic concentric complexes proximate to Wausau in Marathon County. The igneous complexes decrease in age from 1565 Ma in the north (Stettin) to 1506 Ma in the south (Nine Mile) and precede the Wolf River batholith in intrusive age by 30 to 90 m.y. (Van Wyck et al., 1984; Dewayne and Van Schmus, 2007). The granitic pegmatite in the Wausau Complex is ~6m long × ~1m wide and consists of 35% euhedral amphibole, 40% subhedral microcline, 25% anhedral quartz, and accessory pyroxene, albite (Ab 99.7), fluorite, ilmenite, and magnetite. The amphibole is mainly arfvedsonite (Table 1; Fig. 2), with a mean atomic composition of: (Na0.76K0.10)A(Na1.09K0.15Ca0.35Fe2+0.41)B(AlVI0.20Ti0.14Fe3+1.14Fe2+3.22Mn0.13Mg0.17)C (Si7.91AlIV0.09)TO22(OH1.49F0.51) A-site occupancy varies from 0.73 to 0.97, and F contents vary from 0.51 to 1.26 wt.%. In contrast, the rims of amphibole and fine-grained amphibole associated with aegerine are intermediate arfvedsonite-riebeckite, with a mean atomic composition of: (Na0.42K0.02)A(Na1.60K0.06Ca0.01Fe2+0.32)B(AlVI0.10Ti0.03Fe3+1.56Fe2+3.09Mn0.02Mg0.20)C (Si8.27)TO22(OH1.92F0.08) A-site occupancy varies from 0.35 to 0.48, and F contents vary from 0.13 to 0.43 wt.%. The intermediate arfvedsonite-riebeckite contains less TiO2, Al2O3, MnO, CaO, and K20 than the predominant arfvedsonite and is closer to a Na-Fe-Si end-member composition. However, the Si content of arfvedsonite-riebeckite exceeds 8.0 apfu, which may indicate the presence of a pyribole component, which is under further investigation. The pyroxene is very close to end-member aegirine, with a mean atomic composition of: (Na0.956Ca0.004)(AlVI0.038Ti0.007Fe3+0.896Fe2+0.093Mn0.001Mg0.002)(Si2.004)O6 85 �Aegirine occurs interstially and intergrown with arfvedsonite-riebeckite at the rims of arfvedsonite crystals. Such textural and compositional relations are likely the result of increasing oxygen fugacity during crystallization, which stabilizes riebeckite at the expense of arfvedsonite and stabilizes aegirine at the expense of alkali amphibole (Ernst, 1962; Scaillet & MacDonald, 2001). The presence of ilmenite inclusions in arfvedsonite and an association of magnetite with interstitial aegerine are consistent with this interpretation. Ar analysis of a single arfvedsonite crystal was performed by step heating using a CO2 laser, which yielded a well-defined plateau at 1513 ± 5 Ma (Fig. 3), representing ~75% of the 39Ar released. This age is slightly younger than, but within error of, the 1522 ± 6 Ma U-Pb zircon age for syenite in the Wausau Complex (Dewayne and Van Schmus, 2007). The geon 15 igneous intrusions near Wausau are classic examples of alkaline granitic to syenitic concentric complexes, the differentiated parts of which, such as the NYF pegmatite in the Wausau Complex, contain end-member silicate mineral compositions. These igneous intrusions represent an important magmatic event in the Precambrian evolution of the southern Lake Superior region, appearing somewhat earlier than the more voluminous and slightly less alkaline geon 14 Wolf River batholith. References Ernst WG (1962) Journal of Geology, v. 70, 689-736. Myers PE et al. (1984) Institute on Lake Superior Geology, v. 30, Field Trip #3, 58 pp. Scaillet B & MacDonald R (2001) Journal of Petrology, v. 42, 825-845. Van Wyck N et al. (1984) Institute on Lake Superior Geology, v. 30, Part 1 Program and Abstracts, 81-82. 86 �GEOLOGY OF THE LAKE THREE TROCTOLITE, DULUTH COMPLEX - 2013 PRECAMBRIAN FIELD CAMP CAPSTONE MAPPING Jim MILLER, Sarah Sauer, Jordan Benningfield, Jackson Graham, Sara Kozmor, and Ann Marie Prue Precambrian Research Center, University of Minnesota Duluth, Duluth, MN 55812 As a capstone mapping project for the 2013 Precambrian field camp, a crew of four students, a teaching assistant (Sauer) and an instructor (Miller) conducted five days of field mapping bedrock geology in the southern area of Lake Three in the Boundary Waters Canoe Area. The field area was affected by the Fall 2011 Pagami Creek Fire, which burned an area of approximately 93,000 acres. The area immediately south of Lake Three experienced moderate fire intensity such that significant amounts of deadfall still existed on the forest floor. Moreover, in the second summer after the burn, a thick cover of schrubs and vines covered the forest floor creating difficult conditions for inland traverses. Despite these less than ideal conditions, three crews of field partners conducted multiple inland traverses that ultimately extended detailed mapping over a mile south of the Lake Three shoreline. The Lake Three Troctolite (L3T) is a sub-circular body of troctolitic cumulates exposed along the southern shore of Lake Three and intruded into Anorthositic Series rocks. Both the troctolitic and anorthositic rocks are parts of the extensive Duluth Complex, which is the largest exposed intrusive component of the 1.1Ga Midcontinent Rift. Shoreline exposures of the L3T were first mapped by Phinney (1972) and Miller (1986). Additional reconnaissance mapping to south of Lake Three by Miller (1986) and more recent post-fire mapping by Jirsa (2013) suggests that the L3T may extend to as far as 3 miles south of the Lake Three shoreline. This study sought to extend detailed mapping of the L3T south of the shoreline into the area burned by the Pagami Creek Fire. This detailed mapping shows the L3T to have broad asymmetrical synformal structure that trends NNE-SSW with a thinner eastern limb dipping NW and thicker western limb dipping SE. Moreover, the capstone mapping distinguished two distinct map units – an outer (lower) zone of varitextured troctolitic cumulates rich in anorthositic inclusions, and an interior (upper) zone of homogeneous ophitic augite troctolite. More specifically, the outer troctolite unit is composed of light to dark gray, commonly vari-textured (medium- to coarse-grained), locally layered, poorly to well foliated, subophitic to ophitic, Pl-phyric troctolitic rock types that are locally rich in irregular masses of anorthositic lithologies. Modal variants include melatroctolite, leucotroctolite and augite troctolite. The unit typically contains 2%-15% plagioclase phenocrysts typically about 1cm diameter. Modal layering is typically defined by variable olivine:plagioclase concentrations. Contacts with anorthositic series rocks are not exposed, but near inferred contacts, the troctolite becomes vari-textured, generally finer grained, and contains numerous anorthosite inclusions from cm to several meters. Contacts with the interior ophitic augite troctolite are broadly gradational over several to tens of meters. Rocks of the inner augite troctolite unit are typically light gray to dark gray, medium-grained, poor to moderately foliated, homogeneous, ophitic, Pl-phyric augite troctolite. Plagioclase phenocrysts are typically 1-2 cm in diameter, but some are up to 4 cm. Augite oikocryst range from 2-5 cm in diameter and iron oxide commonly occur as subpoikilitic clots less than 1 cm across. Contacts with the troctolite unit are gradational over several to tens of meters. 87 �Anorthositic Series rocks are composed of various anorthositic to leucogabbroic rock types including anorthosite, troctolitic anorthosite, olivine gabbroic anorthosite, leucotroctolite, and augite leucotroctolite.. All lithologies have greater than 75% plagioclase with less than 25% olivine, augite, and Fe-Ti oxides. Augite and oxide always occur as poikilitic to subpoikilitic crystal relative to lathy plagioclase, whereas olivine ranges in habit from subhedral granular to poikililtic. Oikocrysts of olivine up to 20 cm diameter have been observed and commonly recessively weather to create a pocked surface. Most anorthositic lithologies have a well- developed igneous foliation defined by plagioclase alignment, but modal and textural layering are rare. Locally anorthositic series rocks occur as inclusions into both the troctolite and augite troctolite units of the L3T. A previously unrecognized lithology was found along the eastern margin of the L3T and as a large xenolith at the western troctolite-augite troctolite contact during the capstone mapping. It is typically a fine- to medium fine-grained, equigranular (granoblastic) olivine gabbro to augite troctolite. It is typically homogenous, but locally contains abundant coarse-grained clots and stringers rich in oxides and olivine. In some areas, a swirly mixture of medium fine-grained olivine oxide gabbro, fine-grained augite leucotroctolite, and fine grained leuctroctolite with 1cm oikocrysts of olivine is observed. Subtle modal layering with hints of crossbedding have been observed in a few locations. These masses are similar to other areas of the Duluth Complex that have been interpreted to be intensely metamorphosed inclusions of mafic volcanics of the North Shore Volcanic Group. Abundant unburned blowdown trees and thick forest floor vegetation prevented traverses to reach south much more than a mile. Confirmation of the southern extent of the L3T will require additional mapping north of the Isabella River. References Miller, J.D., Jr., 1986, The geology and petrology of anorthositic rocks of the Duluth Complex, Snowbank Lake quadrangle, northeastern Minnesota. Unpublished Ph.D. thesis, University of Minnesota, Minneapolis, MN, 525 p. Phinney, W.C., 1972, Northwestern part of Duluth Complex. In Sims, P.K. & Morey, G.B. (eds.) Geology of Minnesota - A Centennial Volume. Minnesota Geological Survey, p. 335-345 88 �GEOLOGIC MAPPING OF NEOARCHEAN AND PALEOPROTEROZOIC ROCKS NEAR HANSON LAKE, NE MINNESOTA, BY STUDENTS OF THE PRECAMBRIAN RESEARCH CENTER’S 2013 FIELD CAMP MULCAHY, Connor1, ROMANELLI, Dan1, SCHULZ, Roger1, MOORHEAD, Steve1, MAY, Mitchell1, and JIRSA, Mark2 1 2013 Field Camp Students, Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, Duluth, Minnesota 55811 2 Minnesota Geological Survey (MGS), University of Minnesota, 2642 University Avenue W., St. Paul, Minnesota 55114 (jirsa001@umn.edu) The University of Minnesota-Duluth’s Precambrian Research Center conducted its seventh annual field camp in 2013, and this presentation is one of a series that show some of the results. During the fifth and sixth weeks of camp, teams of students participate in “capstone projects” that test student skills by creating new geologic maps in areas of poorly known geology. This capstone project involved mapping an area of the Boundary Waters Canoe Area Wilderness (Fig. 1) accessed by 20 lakes, with Hanson Lake at its center. The map provides details about the complex depositional and deformation history of a Neoarchean, largely metasedimentary terrane that is part of the Wawa subprovince of Superior Province. Figure 1. Generalized bedrock geologic map of northeastern Minnesota showing the Hanson Lake capstone area. The unit labeled “Knife Lake Group” also encloses older volcanic sequences that are not delineated separately at this scale. Dashed line is the border of the Boundary Waters Canoe Area Wilderness. 89 �The Hanson Lake map area lies west of the boundary between the Saganaga Tonalite (ca. 2690 Ma), and superjacent sedimentary strata of the Knife Lake Group that are inferred to have been derived in part from it. Both rock units were variably tilted, folded, faulted, and metamorphosed to low greenschist facies during a regional deformation event at about 2680 Ma, which brackets deposition of the Knife Lake Group between ca. 2690-2680 Ma. Our mapping demonstrated that strata of the Knife Lake Group in this area form a broad, northeast-trending synclinorium, bounded by the Saganaga Tonalite on the east, and an apparently uplifted fault-block of metabasalt on the west. The limbs of this large structure are marked by smaller sympathetic folds, and are dissected by faults and shear zones. Several major faults enclose what appear to be discrete blocks that were uplifted, tilted, and eroded to expose different crustal levels of the stratigraphic succession. As a result, some of the blocks contain older metavolcanic and meta-intrusive rocks that are unconformably overlain by Knife Lake strata. The dominant rock types are interbedded graywacke and slate, which were subdivided into eastern and western sections having somewhat different attributes. The eastern section contains sandstone, graywacke, and mudstone that is locally interlayered with several types of conglomerate and one thin unit of banded iron-formation, and intruded by rare peperite. A conglomerate near Nawaska Lake contains rounded, cobble-to-boulder sized clasts of tonalite, metabasalt, and metagabbro, implying fluvial deposition from streams draining a lithologically diverse terrane. Conglomeratic strata near Gift Lake contain amoeboid and diffuse-edged clasts of dacitic rock, inferred to have been weathered to saprolite before incorporation. Shreds of mafic peperite occur in chaotically bedded gritstone that contains angular grains of feldspar and mafic minerals identical with those in the peperite, indicating synchronous magmatism and sediment deposition. The western section consists of graywacke and slate, together with localized occurrences of more enigmatic strata. At Lake of the Clouds, a trachybasaltic crystal-lapilli tuff and breccias, containing clasts as large as 25cm occurs within an overall package of gray wacke. This implies episodic explosive calc-alkalic volcanism was synchronous with or just predated deposition of graywacke. At least macroscopically, this rock and the peperite appear magmatically-related. In the South Arm of Knife Lake, a unit of metabasalt is capped by basaltic conglomerate having both angular and amoeboid clasts. Presence of the latter implies the basaltic substrate was weathered prior to erosion, which is consistent with the inference that the sedimentary sequence lies unconformably on older basaltic substrate within the fault-bounded block. Taken together, these attributes portray deposition in a largely fault- and unconformitybounded, Timiskaming-type extensional basin. We infer that after the Saganaga Tonalite intruded older basaltic strata, the terrane was uplifted, weathered, and eroded to contribute detritus from both source rocks into the developing Knife Lake basin. The abundant graywacke and slate are interpreted to represent deposition in a lacustrine or marine setting, and the interlayered coarser, polymictic clastic strata may represent braided stream, alluvial fan, or subaqueous fan deposition of sediment shed off the uplifted flanks of the basin. The layered strata exhibit chaotic softsediment deformation features, local growth faults, and abrupt facies changes, suggesting that deposition was synchronous with episodic basin subsidence. Thin layers and lenses of ironformation that are associated with graywacke and slate are inferred to represent chemical precipitation into what may have been a shallow marine environment during periods of relative quiescence. These attributes, together with the association of syn-sedimentation magmatism, are consistent with the model of a Timiskaming-type basin assemblage. Several N20°W-trending, vertically dipping diabasic dikes were also encountered in the area. They are locally as thick as 30m, coarse-grained, subophitic, and have chilled margins. The dikes are inferred to be part of the Paleoproterozoic Kenora-Kabetogama dike swarm on the basis of similarities in trend, thickness, and macromineralogy. This and other capstone mapping projects can be viewed at www.d.umn.edu/prc. 90 �PETROGRAPHIC CHARACTERIZATION OF THE PENOKEAN TWELVEFOOT FALLS SHEAR ZONE, MARINETTE COUNTY, WI: EVIDENCE FOR COEVAL DUCTILE AND SEISMIC BEHAVIOR NADZIEJKA, Brynley and BJØRNERUD, Marcia Geology Department, Lawrence University, Appleton, Wisconsin, 54911 USA The Twelvefoot Falls Shear Zone in northeastern-most Wisconsin lies 25 km south of, and approximately parallel to, the NW-SE-striking Niagara Fault, which is thought to represent the ca. 1.88 Ga Penokean suture between the Archean-Paleoproterozoic Superior Craton and Paleoproterozoic island arc rocks of the Pembine-Wausau terrane (Schulz & Cannon, 2006). The shear zone is well exposed along the Pike River at Twelvefoot and Eightfoot Falls, where it cuts through quartz diorite with a U-Pb zircon crystallization age of 1889 +/- 6 Ma (Schulz & Schneider, 2005). This intrusive body was emplaced into the Quinnesec Formation, a metavolcanic unit interpreted to be part of a precollisional suprasubduction zone ophiolite-island arc complex (LaBerge et al., 2003). The Twelvefoot Falls Quartz Diorite lies on the southern flank of the Dunbar Gneiss Dome, a younger, post-collisional composite intrusion dated at 1862 +/- 5 Ma (Sims et al., 1985; Sims, 1990). Detailed petrographic study of samples from outcrops at Twelvefoot and Eightfoot Falls reveals a complex, multistage deformational history. The rocks have a pervasive, though heterogeneously developed, subvertical NW-striking foliation defined by the preferred orientation of relatively large (2-5 mm) hornblende porphyroclasts and planar quartz-rich domains. The quartz in these bands is typically fine grained (<0.1 mm), with undulose extinction, irregular grain boundaries, and in places ‘core and mantle’ structure. These textures record ductile deformation, dynamic recrystallization and subsequent partial annealing. The hornblende grains have ragged grain boundaries and are commonly dismembered or boudinaged. In many cases, once-joined grain fragments can be recognized through their optical continuity (simultaneous extinction). Alteration of hornblende to chlorite is common. Narrow (<1 cm wide) mylonite zones transect the foliation at oblique angles. Within these bands, both the hornblende and quartz grains are finer than they are in the rest of the rock, but their microstructural characteristics are similar: the hornblende is fragmented in a quasi-brittle manner, while the quartz forms narrow, high-strain bands. This suggests that both the foliation and mylonite zones developed under similar temperature conditions, namely between ca. 300°C (threshold for quartz ductility) and 600°C (onset of hornblende ductility under hydrous conditions; Hacker & Christie, 1999). This is consistent with the upper greenschist- to amphibolite-facies metamorphic conditions in the area surrounding the Dunbar Gneiss Dome (Sims et al., 1985) and also with temperature estimates for rocks with similar compositions and textures from mylonite zones in the Grenville Province (Babaie & LaTour, 1994). At Eightfoot Falls, dark, branching discordant veins 0.3-0.5 cm wide and 10-15 cm long cut across the foliation in the rocks. In thin section, these are found to contain a mesh of fine hornblende crystals with high aspect ratio, arranged with no preferred orientation in a non-crystalline matrix that is dark in plane light. These macro- and micro-scale characteristics suggest that the veins represent devitrified pseudotachylyte – frictional melt glass generated on a fault plane during seismic slip, and injected as ‘hydro’-fractures into the surrounding rock. Significantly, the pseudotachylyte material can be seen in thin section to have been cut by, and in places incorporated into, the mylonitic bands, indicating that brittle seismic failure occurred at least once while the rocks were still at depths and temperatures where crystal plastic deformation was predominant. Such mutually cross-cutting relationships between mylonites and pseudotachylytes have been reported from a small number of sites around the world (e.g., Sibson, 1975; Hobbs et al., 1986) and are interpreted as records of large earthquake ruptures, 91 �usually in convergent tectonic settings, that propagated downward from the fully brittle upper crust into the upper part of the ductile regime. We interpret the fabrics in the rocks from the Twelvefoot Falls Shear Zone to be of mid-late Penokean age, based on the proximity and parallelism of the Zone with the Niagara Fault and on the regional evidence that intrusion, deformation and metamorphism of the Dunbar Gneiss Dome coincided with later Penokean (ca. 1.85 Ga) crustal shortening at the time of the collision of the Marshfield terrane with the Pembine-Wausau terrane (Schulz & Cannon, 2006; Sims, et al., 1985). If so, our results provide insight in the rheology of the middle crust in the heart of a growing mountain belt. References Babaie, H. and LaTour, T., 1994. Semibrittle and cataclastic deformation of hornblende-quartz rocks in a ductile shear zone. Tectonophysics, v. 229, p. 19–30. Hacker, B. and Christie, J., 1990. Brittle/ductile and plastic/cataclastic transitions in experimentally deformed and metamorphosed amphibolite. The Brittle-Ductile Transition in Rocks. American Geophysical Union Geophysical Monograph 56, p. 127-147. Hobbs, B., Ord, A. and Teyssier, C., 1986. Earthquakes in the ductile regime? Pageoph, v. 124, p. 309-336. LaBerge, G., Cannon, W.F., Schulz, K., Klasner, J. and Ojakangas, R., 2003. Paleoproterozoic stratigraphy and tectonics along the Niagara suture zone, Michigan and Wisconsin. In: Cannon, W.F. (ed.), Institute on Lake Superior Geology Field trip Guidebook, v. 49, p. 1-32. Schulz, K. and Cannon, W.F., 2006. The Penokean orogeny in the Lake Superior region. Precambrian Research, v. 157, p. 4-25. Schulz, K. and Schneider, D. 2005. Age constraints on the Paleoproterozoic Pembine ophiolite-island arc complex and implications for the evolution of the Penokean orogen. Geological Society of America Abstracts with Programs, v. 37, no. 5, p. 4 Sibson, R., 1975. Generation of pseudotachylyte by ancient seismic faulting. Geophysical Journal of the Royal Astronomical Society, v. 43, p. 775-794. Sims, P. K., Peterman, Z., and Schulz, K., 1985. The Dunbar Gneiss-granitoid dome: Implications for early Proterozoic tectonic evolution of northern Wisconsin. Geological Society of America Bulletin, v. 96, p. 1101-1112. Sims, P.K. 1990. Geologic Map of Precambrian Rocks of Iron Mountain and Escanaba 1° × 2° Quadrangles, Northeastern Wisconsin and Northwestern Michigan. U.S. Geologic Survey Miscellaneous Investigations Series Map I-2056. 92 �METAMORPHISM AND DEFORMATION AT THE WABIOONQUETICO SUBPROVINCE BOUNDARY IN THE DECOURCEY LAKE AREA A.E., Nolan and M.L., Hill Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario, Canada, P7B 5E1 The Wabigoon and Quetico subprovinces are east-west trending belts of metasedimentary and metavolcanic rocks located within the Superior province, the world’s largest preserved Archean craton. The Wabigoon-Quetico subprovince boundary in the Decourcey Lake area is complex and several kilometers wide. The 130m-long roadcut on Hwy 527 near Decourcey Lake is composed of a garnet-biotite to biotite schist with three intrusions, a pegmatitic intrusion, a felsic intrusion and a mafic dyke. There are also several quartz-carbonate veins that have been folded and boudinaged; the folded veins crosscut the boudinaged veins that are parallel to foliation. In the southern portion of the outcrop, mats of fibrolite in the schist indicate peak metamorphism at amphibolite facies or higher. However, some parts of the outcrop show evidence of overprinting retrograde metamorphism. The amount of retrograde metamorphism increases from south to north in the outcrop. In the southern part of the outcrop there are minimal amounts of chlorite (~1%) and the amount increases to about 30% in the north. The chlorite replaces biotite, the main mica in the southern portion of the roadcut. Chlorite replacing biotite is indicative of retrograde metamorphism. Other evidence for this retrograde metamorphism is the presence of stable garnet in the southern portion of the outcrop and metastable garnet in the northern portion, indicated by euhedral unfractured crystals in the south and anhedral, fractured and separated grains in the north (fig. 1). a b c Figure 1. a) Garnet from southern portion of the outcrop (stable), b) garnet from middle portion of the outcrop (metastable), c) garnet from northern portion of the outcrop (metastable) Throughout the schist, quartz with undulose extinction, irregular grain boundaries and subgrains provides evidence for deformation by dislocation creep. In the southern portion of the outcrop, undulose extinction in feldspar indicates deformation at amphibolite facies temperatures or higher. This higher temperature deformation of feldspar is not evident in the northern portion of the outcrop where retrograde metamorphism is more pervasive. The pegmatite and felsic 93 �intrusions preserve evidence of deformation at the amphibolite facies or higher, including undulose extinction in feldspar and bent mica grains (Fig. 2). Figure 2. Photomicrograph of the felsic intrusion showing evidence for dislocation creep The preservation of evidence for amphibolite facies metamorphism and high temperature deformation in schist in the southern portion of the outcrop as well as in the pegmatite and felsic intrusions suggests that the Decourcey Lake roadcut is composed of metamorphosed Quetico lithologies. Therefore, the Quetico-Wabigoon subprovince boundary lies to the north of this roadcut. The increase in retrograde metamorphism toward the northern portion of the outcrop (in the direction of the subprovince boundary) suggests that this lower temperature metamorphism may be associated with the boundary. 94 �BEDROCK GEOLOGIC MAP OF THE TWIN METALS MINNESOTA PROJECT, NORTHERN SOUTH KAWISHIWI INTRUSION AND ADJACENT AREAS Dean M. Peterson, Senior Vice President, Exploration, Duluth Metals Limited, 306 West Superior Street, Suite 407, Duluth, MN 55802 Twin Metals Minnesota LLC (TMM), is the joint venture company between Duluth Metals Limited (60% ownership interest) and Antofagasta plc (40% ownership interest). TMM is currently in the process of completing a prefeasibility study of the Maturi Cu-Ni-PGE deposit in the northern South Kawishiwi intrusion (SKI) of the Duluth Complex, northeastern Minnesota. Assuming a favorable outcome of the prefeasibility study, TMM will embark on a bankable feasibility study that will include extensive study of groundwater within and around the facilities (the proposed underground mine at Maturi, a concentrator facility, and a tailings storage facility). All of these studies would be incorporated into an environmental review of the project and would be open for comment and review by the general public. As the joint venture looks to the immediate future, Duluth Metals Limited believes that it is imperative that this process be as transparent as can be possible, especially on the publication of confidential geological data that could possibly be linked to water issues in an environmental review process. The bedrock geological map presented in this poster is the result of nearly two decades of geological work by the author integrated with data from seemingly innumerable government (Minnesota Geological Survey, U.S. Geological Survey, Natural Resources Research Institute, Minnesota Department of Natural Resources), academic (Minnesota-Duluth, Indiana, Minnesota, Wisconsin), and industry (Duluth Metals, Twin Metals, Franconia, INCO, Hanna, Bear Creek, Kennecott, Newmont, Duval, Encampment, etc…). 95 �96 �POTENTIAL FOR COPPER TOXICITY CAUSED BY SURFACE WATER AND STREAM SEDIMENTS IN UNMINED MINERALIZED WATERSHEDS OF THE DULUTH COMPLEX. PIATAK, Nadine M.1, SEAL, Robert R. II1, JONES, Perry M.2, WOODRUFF, Laurel G.2, 1 U.S. Geological Survey, Reston, VA 20192, npiatak@usgs.gov, rseal@usgs.gov U.S. Geological Survey, Mounds View, MN 55112, pmjones@usgs.gov, woodruff@usgs.gov 2 The characterization of baseline conditions in unmined mineralized watersheds of the Mesoproterozoic Duluth Complex, northeastern Minnesota, is essential to understanding how to responsibly extract minerals in one of the most prospective mining areas in the United States. Mining could release metals into watersheds that already contain ecologically-significant naturally-occurring amounts of some elements such as Cu and Ni. The potential for metals to be toxic to aquatic organisms is influenced by the amount of organic carbon in the aquatic environment, the cumulative effects of multiple metals, cation competition for biologic binding sites, and speciation of metals. We estimated toxicity in mineralized watersheds using approaches that incorporate these water and sediment quality parameters. Surface-water and streambed-sediment samples were collected from sites along three geologically distinct watersheds in the Duluth Complex: 1. Filson Creek where Cu-Ni-PGM mineralization occurs at the bedrock surface along the basal Duluth Complex; 2. Keeley Creek where Cu-Ni-PGM mineralization occurs only at great depth; and 3. the St. Louis River in the vicinity of Fe-Ti oxide ultramafic intrusions, which occur at the subcrop beneath glacial cover. Samples were collected in September 2012 near base-flow conditions in watersheds dominated by lakes, wetlands, and streams. The geochemistry of the surface waters and stream sediments reflects underlying rock types, glacially transported unconsolidated materials, mineralization style within each watershed, and geochemical processes occurring in the streams. The surface water is oxic, near neutral to slightly acidic (pH 5.9 to 7.6), has low total dissolved solids (41 – 94 mg/L), and is characterized by moderate hardness (18 – 50 mg/L CaCO3), moderate carbonate species concentrations (11 – 38 mg/L CaCO3 as bicarbonate), low sulfate (<0.8 – 3 mg/L), and high dissolved organic carbon (DOC) concentrations (18 – 47 mg/L). The dominant dissolved trace elements are Fe (472 – 3,950 µg/L), Al (54 – 228 µg/L), Cu (0.8 – 8 µg/L), Ni (1 – 5 µg/L), and Co (0.4 – 3 µg/L). Stream sediments contain significant Al (7 – 11 wt. %), Ca (1.5 – 6 wt. %), Fe (1 – 7 wt. %), and Na (2 – 4 wt. %). Sulfur is very low (<0.05 wt. %). Organic carbon reaches 4.7 wt. % in one sample but is ≤1.5 wt. % in all the other samples. Trace metals are dominated by Cr (14 – 346 mg/kg), Cu (10 – 179 mg/kg), Ni (13 – 127 mg/kg), and Zn (23 – 95 mg/kg). On average, Cu and Ni are highest in Filson Creek surface waters and sediments where Ni-Cu-PGM mineralization occurs at the surface. Samples collected from the St. Louis River watershed, where Fe-Ti oxide-bearing ultramafic rocks and a Paleoproterozoic shale/greywacke unit (Virginia Formation) occur, contain the highest average concentrations of As, Fe, and Pb in both surface water and sediments, Cr and Zn in sediment, and sulfate in waters. In water, the toxicity of most metals is assessed on the basis of hardness-based criteria that adjust for the protective effects of Ca and Mg ions, which compete with metal ions for binding sites on organisms. For sediment, consensus-based total-metal guidelines are routinely used and rely on laboratory toxicity tests that document increased toxicity caused by increased metal concentrations (McDonald and others, 2000). However, new guidelines that rely on the Biotic Ligand Model (BLM) utilize a more sophisticated approach incorporating more water and sediment quality parameters including the cumulative effects of multiple metals in sediment, metal speciation in water, and organic carbon complexes in both water and sediment (Di Toro and others, 2005; Paquin and others, 2001). The baseline surface-water and sediment metal concentrations can be compared to aquatic guidelines using the hazard quotient (HQ), which is the ratio of the concentration of a metal in the sample to the guideline. Values above 1 imply toxic conditions, whereas those below do not. In water, HQs for Cu are greater than 1 for several sites in the Filson Creek watershed when calculated using hardness-based criteria (Figure 1). However, as shown in Figure 1, HQs for Cu in water calculated based on the BLM model are significantly less than 1, suggesting a lack of toxicity from all samples. The radically different results from the hardness-based and BLM-based approaches suggest that the former may be inadequate to describe metal toxicity in these watersheds because it is based on a more limited set of parameters (i.e., only hardness). The complexation of Cu with DOC likely significantly affects the bioavailability of dissolved Cu, helping mitigate its toxicity. 97 �The sediment BLM approach also suggests a different level of predicted toxicity from sediments than predicted from consensus-based guidelines. Several HQs for Ni and one HQ for Cu are greater than 1 when calculated using the consensus-based guidelines, which suggests toxic conditions (Figure 2A). In comparison, no toxicity to uncertain toxicity is predicted based on Equilibrium Partitioning Sediment Benchmark (ESB) (USEPA, 2005) (Figure 2B). Specific considerations of this approach include: 1. determining extractable (a proxy for bioaccessible) metal concentrations (i.e., combined simultaneously extracted metals, ƩSEM); 2. adjusting them for potential incorporation into less bioaccessible sulfides (i.e., acid volatile sulfide, AVS); 3. and adjusting for complexation with organic carbon (i.e., fraction of organic carbon, foc). The high organic carbon in some of the sediments could sequester significant amounts of trace elements; however, the low AVS suggest trace elements bound to sulfides are not significant components in these sediments. Applying these more sophisticated and holistic approaches enhances our capability to predict metal toxicity. This improved understanding will be advantageous when developing successful strategies to help minimize future mining impacts and develop appropriate restoration goals. References Di Toro, D.M., McGrath, J.M., Hansen, D.J., Berry, W.J., Paquin, P.R., Mathew, R., Wu, K.B., and Santore, R.C., 2005, Predicting sediment metal toxicity using a sediment biotic ligand model: Methodology and initial application: Environmental Toxicology and Chemistry, v. 24, no. 10, p. 2410-2427. MacDonald, D.D., Ingersoll, C.G., and Berger, T.A., 2000, Development and evaluation of consensus-based sediment quality guidelines for freshwater ecosystems: Archives of Environmental Contamination and Toxicology, v. 39, no.1, p. 20-31. Paquin, P.R., Gorsuch, J.W., Apte, Simon, Batley, G.E., Bowles, K.C., Campbell, P.G.C., Delos, C.G., Di Toro, D.M., Dwyer, R.L., Galvez, Fernando, Gensemer, R.W., Goss, G.G., Hogstrand, Christer, Janssen, C.R., McGeer, J.C., Naddy, R.B., Playle, R.C., Santore, R.C., Schneider, Uwe, Stubblefield, W.A., Wood, C.M., and Wu, K.B., 2002, The biotic ligand model: A historical overview: Comparative Biochemistry and Physiology, v. 133, no. 1-2, p. 3–35. U.S. Environmental Protection Agency, 2005, Procedures for the derivation of equilibrium partitioning sediment benchmarks (ESBs) for the protection of benthic organisms: Metal mixtures (cadmium copper, lead nickel, silver, and zinc): U.S. Environmental Protection Agency 600-R-O2-011, variously paginated. 98 �MESOPROTEROZOIC MIDCONTINENT RIFT INTRUSIVES IN THE THUNDER BAY AREA (ONTARIO, CANADA): A PALEOMAGNETIC REVIEW PIISPA, Elisa J., SMIRNOV, Aleksey V. Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, 630 DOW ESE Building, Houghton, MI 49931-1295, USA and PESONEN, Lauri J. P. Department of Physics, Division of Geophysics and Astronomy, University of Helsinki, Helsinki, Finland Mafic sills and dykes extend over 300 km from south of Thunder Bay to northeast of Lake Nipigon, representing the northern expression of the ~1.1 Ga Midcontinent Rift System (MRS). Recent geochemical and geochronological studies have significantly improved our understanding of the area geology. The Logan sills, south of Thunder Bay, are geochemically similar but not identical to the sills exposed in the vicinity of Lake Nipigon, which in turn can be divided into three separate groups based on their geochemical signatures (e.g. Hollings et al., 2010, 2012). In addition, several discrete mafic/ultramafic intrusions and dyke swarms that represent both the earliest and the latest stages of the Midcontinent Rift magmatism are exposed along the north shore of Lake Superior (Heaman et al. 2007; Hollings et al., 2010). The dykes within the Thunder Bay area are currently grouped into four lithological units based on their orientation, petrology and geochemical differences: the Mount Mollie dyke and the Sibley, Pigeon River and Cloud River dykes (Hollings et al. 2010, 2012; Cundari et al., 2012). In general, the magnetostratigraphy of the MRS can be summarized as follows: a) The earliest rock sequences (~1115-1105 Ma) are reversely magnetized; b) A polarity reversal (or reversals) occurred between 1105-1102 Ma; c) The rocks emplaced after ~1102 Ma are normally magnetized. Withstanding the age uncertainty, this geomagnetic polarity sequence allows for an approximate correlation within and between the MRS rock sequences. When combined with high quality geochronological and petrographical observations as well as detailed geochemical and isotope data, paleomagnetic data can provide valuable information of the development of the MRS. We present the new results of our rock magnetic and paleomagnetic investigation of several MRS intrusives exposed in the vicinity of Thunder Bay (Ontario, Canada). We also re-evaluate the previously published paleomagnetic data based on the newly published geochronological and geochemical data from the sills and dykes of the Thunder Bay area. Finally, we will critically address the observed inconsistencies between the field observations 99 �and paleomagnetic and geochronological data. This study contributes to better understanding of the MCR magnetostratigraphy and further improvement of the late Mesoproterozoic apparent polar wander path for North America. References: Cundari, R.M., Hollings, P., Smyk, M.C., Scott, J.F. and Campbell, D.A. 2012. Whole rock and isotope data from the Midcontinent Rift: implications for crustal contamination history; in Summary of Field Work and Other Activities 2012, Ontario Geological Survey, Open File Report 6280, p.18-1 to 18-10. Heaman, L.M. and Machado, N. 1992. Timing and origin of Midcontinent Rift alkaline magmatism, North America: evidence from the Coldwell Complex; Contributions to Mineralogy and Petrology, v.110, p.289303. Hollings, P., Smyk, M., Heaman, L.M., and Halls, H. 2010. The geochemistry, geochronology, and paleomagnetism of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada. Precambrian Research. Precambrian Research, v. 183, iss. 3, p.553-571. Hollings, P., Smyk, M.C. and Cousens, B. 2012. The radiogenic isotope characteristics of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada; Precambrian Research, v.214-215, p.269-279. 100 �DOCUMENTING THE FIRST LAVA FLOWS OF THE MIDCONTINENT RIFT BY DIGITAL MAPPING AND PETROGRAPHIC ANALYSIS QUILLEN, Patrick, and Miller, Jim Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812 The Ely’s Peak basalts (EPB) located west of Duluth represent the first lava flows to be erupted from the Midcontinent Rift. The EPB lie conformably on siltstones, sandstones, and conglomerates of the Nopeming sandstone, which in turn unconformably overlies the Paleoproterzoic Thomson Formation. This classic outcrop area (Jirsa and Morey, 1979; Green, 1999; Jirsa and Green, 2011) has been a popular field trip location for many years, and today it’s used as a field mapping exercise for a Geologic Maps course at UMD (GEOL 3000) and the Precambrian Research Center’s field camp (GEOL 4500). The area was initially mapped by J.A. Kilburg as part of his MS thesis at UMD (Kilburg, 1972) with a reconnaissance map later produced by Kilburg and Morey (1977). Kilburg (1972) noted that the Duluth Complex thermally metamorphosed the EPB flows and made them difficult to distinguish. The focus of this project is exposure of the EPB north of Interstate 35 (Fig. 1). The research questions for this project were to determine the number of lava flows represented in the area, their petrographic characteristics and their geographical distribution of exposure. The data collected from several years of mapping and new mapping conducted for this study were also compiled and digitized with ArcMAP 10. For this study, 17 new were collected and made into thin sections and combined with 17 samples collected previously. A petrographic analysis of this suite of sections were made, noting textures and mineral assemblages comprising the groundmass, phenocrysts, and amygdules that might be useful for distinguishing individual lava flows and revealing the nature and intensity of thermal metamorphism. Figure 1. Bedrock geology of the field area based on previous mapping by Kilburg (1972). Significantly more outcrops have been discovered by mapping conducted for the UMD geology courses. 101 Prior mapping by Kilburg (1972) and UMD classes concluded that there were two main types of flows. The first is a lower sequence of variably amydaloidal, dense, pyroxene-phyric basalts overlying the Nopeming sandstone. Pyroxene phenocryts up to 1 cm in diameter appear to compose between 10 and 40 vol.% of these flows, but up to 90% pyroxene phenocrysts have been locally observed. Amygdaloidal zones composed of chlorite, quartz and epidote amygdules are locally observed, but flow contacts are difficult to recognize, presumably due to anealling by thermal metamorphism. In the eastern third of the EPB exposure area (Fig. 1), a distinctly aphyric basalt is recognized. It is also distinguishable from the pyroxene-phyric flow by being moderately magnetic. The contact �with the Duluth Complex on the east side of the field area is somewhat ambiguous by being marked by a fine-grained, massive gabbro that has been variably interpreted to be intensely metamorphosed basalt or chilled gabbro. Petrographic observations from this study reveal several interesting and unexpected results. Most notable is that the previous interpretation of pyroxene-phyric basalt overlain by aphyric basalt is a significant oversimplification of the area. Several samples collected from the porphyritic basalt area are actually aphyric. Also, what have been assumed to be pyroxene phenocrysts are actually clusters of pyroxene that may be glomerphenocrysts or perhaps xenoliths of pyroxenite. Curiously, they commonly have amphibolitic reaction rims, suggesting that they may be out of equilibrium with the enclosing basaltic groundmass and thus be xenoliths. The few exposures that have been mapped as comprising an area of aphyric basalt instead appear to be more related to the Duluth Complex. They are subophitic, olivine gabbro and could potentially be the chilled contact zone of the Duluth gabbro. Another surprising observation is that given the density of these basalts, granoblastic recrystallization textures are rarely noted. Instead, original intergranular igneous textures dominate the sections observed. This poster presentation will show a revised geological map and of the area and display photomicrographs of the textures and mineral assemblages observed. References Green, J.C., 1999, Proposal to designate the Grandview Area as a state Scientific and Natural Area for its geological importance. Unpublished report submitted to the MN Dept. of Natural Resources, 6p. Jirsa, M.A., and Morey, G.B., 1979, Jay Cooke State Park and Grandview area: Evidence for a major Early Proterozoic-Middle Proterozoic unconconformity in Minnesota. Geological Society of America Centennial Field Guide – North Central Section, p. 67-72. Jirsa, M.A. and Green, J.C, 2011, Classic Precambrian geology of northeast Minnesota . In Miller, J.D., Hudak, G.H., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 25-45. Kilburg, J.A, 1972, Petrology, structure, and correlation of the upper Precambrian Ely’s Peak basalt. Unpublished MS thesis, University of Minnesota Duluth, 97p. Kilburg, J.A., and Morey, G.B., 1977, Reconnaissance geologic map of the Esko quadrangle, St. Louis and Carlton Counties, Minnesota: Minnesota Geological Survey Miscellaneous Map Series, M-25, scale 1:24,000. 102 �GEOCHEMISTRY AND PETROGRAPHY OF A MAFIC METAVOLCANIC SEQUENCE SOUTH OF MUSSELWHITE MINE QUINN, Jordan1, HOLLINGS, Pete1, BICZOK, John2 1 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Goldcorp Canada Ltd., Musselwhite Mine, P.O. Box 7500, Thunder Bay, Ont., P7B 6S8, Canada The North Caribou Greenstone Belt (NCGB), within the North Caribou Terrane of the Archean Superior Province, is host to multiple ~3.0Ga metavolcanic and metasedimentary assemblages. The assemblages have been metamorphosed from greenschist to upper-amphibolite grade and are bounded by ~2.7-3.0 Ga granitoids and gneisses (Biczok et al., 2012). The study area is located approximately 5km south of Musselwhite Mine within what was previously thought to be the Opapimiskan-Markop metavolcanic assemblage. There were three lithologies identified within the study area (basaltic, komatiitic, and felsic volcanic flows) which have been subdivided into four separate volcanic suites: Volcanic Suite A, Volcanic Suite B, Volcanic Suite C, and Felsic Volcanic Suite. Volcanic Suite A is comprised of a succession of massive and pillowed basaltic flows that were metamorphosed to amphibolites. These flows exhibit a mineral assemblage of amphibole, chlorite, and plagioclase with minor quartz, muscovite, titanite and epidote. Major element geochemistry reveals that this suite is compositionally similar to that of a high-Mg tholeiitic basalt. Primitive mantle normalized plots for this volcanic suite are characterized by a flat rare-earth element (REE) pattern comparable to tholeiites from the South Rim Unit (SRU) (Fig. 1), which have been interpreted to represent oceanic island plateau basalts formed from a mantle plume (Hollings et al., 1999). Figure 1: Comparison of primitive mantle plots from Volcanic Suite A versus tholeiites from the South Rim Unit (Blue trace). Volcanic Suite B was comprised of pillowed and massive basaltic flows that have been metamorphosed to amphibolites. The main mineral assemblage observed in this suite was amphibole and chlorite with minor plagioclase, clinozoisite, quartz, titanite and dolomite. Major element geochemistry indicates that this suite is comprised of high-Mg tholeiitic basalts, komatiitic basalts, and a komatiite. Primitive mantle normalized plots display a relatively flat REE pattern but with a negative Nb anomaly (Fig. 2). The similar 103 �trace element geochemistry of Volcanic Suites A and B suggests that they are both derived from a plume source, however, the negative Nb anomaly in Volcanic Suite B indicates that it has undergone crustal contamination during emplacement. Figure 2: Primitive mantle normalized plot of the samples from Volcanic Suite B. Volcanic Suite C was also comprised of massive and pillowed basaltic flows with a main mineral assemblage of amphibole and chlorite. A single sample was taken from this suite and it was determined to be a high-Fe tholeiitic basalt based on major element geochemistry. Primitive mantle plots of this suite are light rare-earth element (LREE) enriched with a negative Nb anomaly and positive Zr and Hf anomalies. A similar REE pattern was observed in the tholeiitic basalts from the Opapimiskan-Markop Unit (OMU) (Hollings and Kerrich, 1999). The felsic volcanic suite overlies the mafic volcanic suites and is comprised of rhyolitic flows. This suite was LREE enriched with a relatively flat heavy rare-earth element (HREE) pattern and negative Nb and Ti anomalies in conjunction with positive Zr and Hf anomalies. Similar REE patterns were observed in the SRU and interpreted to be derived from a subduction tectonic setting (Hollings et al., 1999). The results of this study are consistent with previous work in the region and suggest that the early history of the area preserved the interaction of a mantle plume with preexisting continental crust. In addition this study has refined the boundaries of the various assemblages within the NCGB. References Biczok, J., Hollings, P., Klipfel, P., Heaman, L., Maas, R., Hamilton, M., Kamo, S., Friedman, R., 2012. Geochronology of the North Caribou greenstone belt, Superior Province Canada; implications for tectonic history and gold mineralization at the Musselwhite Mine. Precambrian research 192:209230. Hollings, P., Wyman, D., Kerrich, R., 1999. Komatiite-basalt-rhyolite volcanic associations in northern Superior Province greenstone belts; significance of plume-arc interaction in the generation of the proto continental Superior Province. Lithos 46.1:137-161. Hollings, P., Kerrich, R., 1999. Trace element systematics of ultramafic and mafic volcanic rocks from the 3Ga North Caribou greenstone belt, northwestern Superior Province. Precambrian research 93.4:257-279. 104 �THE ARROWHEAD PILOT PROJECT: MAPPING OF PRECAMBRIAN AND QUATERNARY GEOLOGY IN TWO DIVERSE GEOLOGIC AREAS OF NORTHEASTERN MINNESOTA RADAKOVICH, A.R.1, and HOBBS, H.C.1 1 Minnesota Geological Survey, St. Paul, MN 55114, rada0042@d.umn.edu, hobbs001@umn.edu The Arrowhead Pilot Project was undertaken by the Minnesota Geological Survey (MGS) to explore the feasibility of County Geologic Atlas-style mapping in NE Minnesota. It integrates new field mapping by the authors with archived data to provide both Precambrian and Quaternary geologic interpretations of two areas of interest in northeastern Minnesota (Fig. 1). Two distinct areas were mapped: the western one contains significant exposure of Archean, Paleoproterozoic and Mesoproterozoic bedrock, and relatively thin and patchy Quaternary deposits; the eastern area has limited bedrock exposure of Mesoproterozoic rocks, and thick Quaternary glacial cover. Together these areas cover part or all of fifteen 7.5’ quadrangles. The intent of this study was multi-faceted: (1) to compile previous maps (both Precambrian and Quaternary) of these areas into single, coherent maps; (2) to augment gaps in data with new mapping; (3) to assess the usefulness of LiDAR (Light Detection And Ranging altimetry) imagery in identifying bedrock outcrop and surficial features; and (4) to assess the costs and feasibility of a mapping project at a similar level of detail over a much larger area, such as that typical of a County Geologic Atlas produced by the MGS. The motivation to produce large scale, comprehensive regional map products is in part driven by increased interest in extracting mineral resources, as well as management of ground and surface water resources. Further, the area’s complex bedrock geology provides insight into more than 1.5 billion years of earth’s history in three distinct terranes: the Archean Giants Range batholith, the Paleoproterozoic Animikie Group, and the Mesoproterozoic Duluth Complex. This study area is also key to understanding the glacial history of the region, as it includes the interlobate junction between the Rainy and Superior lobes, and spans the transition in the provenance of Rainy lobe clasts from the Duluth Complex-dominated eastern portion to the granitedominated western part. Figure 1. Location of study areas showing 1-meter LiDAR (Light Detection and Ranging Altimetry) imagery. 105 �The bedrock portion of the study resulted in creation of modified bedrock geologic and bedrock topographic maps, and a detailed fault and lineament analysis. In the western area, new mapping north of the Mesabi Iron Range helped provide a more complete characterization of the Archean Giants Range batholith. Mapping south of the Iron Range also helped better define layers of the Mesoproterozoic South Kawishiwi Intrusion in the Duluth Complex. In the eastern area, new outcrop data combined with improved aeromagnetic and gravity data helped modify and extend unit contacts within Mesoproterozoic volcanic sequences. LiDAR analysis combined with field mapping and structural analysis identified significant bedrock-controlled linear topographic features which were classified as faults, lineaments, or igneous foliations. Such features are known to potentially play significant roles in the hydrogeologic system. Finally, mapping of the bedrock surface incorporated data from County Well Index (CWI) records, passive seismic data, outcrop elevations, and LiDAR; It revealed that depth to bedrock varies considerably across the area: from zero to as much as 250 feet (~75m). The Quaternary portion of the study revealed the Highland moraine of the Superior lobe to be highly collapsed and strewn with ice-walled lake plains, indicating widespread stagnant ice. The overall texture of the unsorted material of the Highland moraine is rocky sandy loam. The lake plains are composed of sorted material, typically sand and gravel on the raised rims, grading vertically down to fine sand and silt. The center of the plains are presumably composed of silt and clay, but were not investigated. Subglacial meltwater from two large esker systems coalesced and spewed large volumes of meltwater, which deposited outwash that followed the edge of the retreating Rainy lobe. The meltwater was ponded for a time in glacial Lake Dunka, and ultimately flowed through a gap in the Giant’s Range into a lake indirectly connected to glacial Lake Agassiz. The Rainy lobe built several distinct recessional moraines in the mapping area; the Vermilion moraine is the last Rainy lobe moraine south of the International Border. The ice deposited a relatively thin, extremely coarse till between the moraines, and did not discharge copious amounts of meltwater. Rogen moraine ridges are common north of the Vermilion moraine; these ridges formed under the ice, and do not represent ice margins. The Arrowhead Pilot Project demonstrates that successful products can result from regional mapping in areas of northeast Minnesota given appropriate time, funding, and creativity. However, areas of especially sparse data and remote settings will necessarily result in diminished mapping detail, particularly in the depiction of the subsurface distribution of Quaternary materials, compared to other parts of the state where the MGS has produced County Geologic Atlas maps. Nevertheless, even with that limitation, County Geologic Atlas map products for the region would provide a markedly improved geologic framework that would facilitate resource management decisions. Selected References Boerboom, T.J., and Miller, J.D., Jr., 1994., Bedrock geologic map of the Silver Island Lake, Wilson Lake, and western Toohey Lake quadrangles, Lake and Cook Counties, Minnesota: Minnesota Geological Survey Miscellaneous Map M-81, scale 1:24,000. Jirsa, M.A., Chandler, V.W., and Lively, R.S., 2005, Bedrock geology of the Mesabi Iron Range, Minnesota: Minnesota Geological Survey Miscellaneous Map M-163, scale: 1:100,000. Jirsa, M.A., and Miller, James D., Jr., 2004, Bedrock geology of the Ely and Basswood Lake 30' X 60' quadrangles, northeast Minnesota: Minnesota Geological Survey Miscellaneous Map M-148, scale: 1:100,000. Foose, M.P., and Cooper, R.W., 1978, Preliminary geologic report on the Harris Lake area, northeastern Minnesota: U.S. Geological Survey Open-File Report 78-385, 24 p., 1 pl., scale 1:12,000. Friedman, Albert L., 1981, Surficial geology of the Isabella quadrangle, northeastern Minnesota: Unpublished Master's thesis, University of Minnesota. Miller, James D., Jr., and Severson, M.J., 2005, Bedrock geology of the Babbitt quadrangle, St. Louis and Lake Counties, Minnesota: Minnesota Geological Survey Miscellaneous Map M-159, scale 1:24,000. Miller, James D., Jr., and Severson, M.J., 2005, Bedrock geology of the Babbitt Northeast quadrangle, St. Louis and Lake Counties, Minnesota: Minnesota Geological Survey Miscellaneous Map M-160, scale 1:24,000. Miller, James D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.M., 2001, M-119 Geologic map of the Duluth Complex and related rocks, northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map M-119, scale: 1:200,000 and 1:500,000. Miller, J.D., Jr., Boerboom, T.J., and Jerde, E.A., 1994, Geologic map of the Cabin Lake and Cramer 7.5-minute quadrangles, Lake County, Minnesota: Minnesota Geological Survey Miscellaneous Map M-82, scale 1:24,000. Severson, M.J., 1994, Igneous stratigraphy of the South Kawishiwi intrusion, Duluth Complex, northeastern Minnesota: Duluth, University of Minnesota, Natural Resources Research Institute Technical Report NRRI/TR-93/34, 210 p. + plates. Stark, James R., 1977, Surficial geology and ground-water geology of the Babbitt-Kawishiwi area, northeastern Minnesota with planning implications: Unpublished Master's thesis, University of Wisconsin-Madison. 106 �TOOLS FOR INTERPRETING KEWEENAW GEOHERITAGE TO A BROAD PUBLIC ROSE, William I and VYE, Erika, Department of Geological and Mining Engineering and Sciences. Michigan Technological University, Houghton, MI 49931 In the United States, the public is seemingly isolated from geoheritage, perhaps due to a disconnect between the geoscience academic community and how we communicate what we know. Recently retired, and with nearly 45 years in Houghton, a place with a strong geoheritage, the first author has begun to focus on communicating Earth Science to the a broader public with vital help from his co-author. This has generated new interpretive activities and tools paramount for engaging the public in learning about how and why their place has come to be the way it is. Boulder Gardens. Around Lake Superior we have an abundant teaching resource of glacial erratics. We moved some of the most exemplary ones in our region (some with difficulty!) to a public square in the center of campus to serve as an educational and cultural focus (Rose, 2011). The boulders have fresh, glacially polished surfaces and are an assemblage of dozens of outcrops representing all the lithologies of the Keweenaw Rift in one succinct location. The site has drawn educational attention, and is especially useful as an introduction to field trips of the area. GPS, smart phones, QR codes, EarthCaches. We have embraced treasure hunting and technology-based tools to engage people in learning about geolocations, fundamental to 3D visualization and “reading the landscape”. We have identified more than 150 field sites in the Keweenaw and linked them to .kmz file information, Google Maps, and QR codes to be accessed via smartphones. We have also very successfully contributed to GSA’s EarthCache efforts and database (Gochis et al. 2013). Community geotours - bike/walk and trolley. We fashioned a geotour of our local town (Houghton) which can be done on foot or bike. The tour identifies and interprets a variety of features such as mines, large lava flows, faults, veins, aa and pahoehoe flows, amygdaloids, glacial features, river deltas, kame terraces and Anthropocene features in the town (Rose et al. 2013). The Geotour is integrated with other heritage tours and has been successfully used as a fundraiser that uses the local trolley to access a wider range of sites. Grassroots Partnership. We work with local groups who share common goals of conservation and public access to field sites. Partners include the local Chamber of Commerce, Copper Country Trail 107 �National Byway, the Keweenaw Land Trust, the Michigan Nature Association, the Nature Conservancy, national and state parks, local towns and villages, historical societies, museums, and local businesses all wishing to disseminate geoheritage information. A demonstration of geotours with a broad audience (ex. “elderhostal” format) is being done this summer with five two-day geotours which use combined van and boat transport to visit many remote sites of the Keweenaw. The goal is to facilitate public geotours of greater depth which last for several days. We have built our network to engage Earth science teachers, who can replicate geointerpretation efforts and transfer them to their students. This, in turn, helps us find more geosites as they are in everyone’s own yards! International Geoheritage links. We have a strong partnership with colleagues at European Geoheritage sites who are teaching us successful strategies for international geoheritage status; common in Europe, China and other parts of the world, but so far unknown in the US (VanWyk deVries, 2013). All of these efforts are found on a single Keweenaw Geoheritage website (http://www.geo.mtu.edu/ ~raman/SilverI/KeweenawGeoheritage), which broadly links shared technical, geographic and heritage information. We invite public input about this process; how can we improve geoheritage outreach? References: ROSE, WI 2011, KEWEENAW BOULDER GARDEN—A REVITALIZED KAME TERRACE ON CAMPUS, USED AS A TEACHING LABORATORY GSA Abst w Programs 43 (5), p 25 (https://gsa.confex.com/gsa/2011AM/ finalprogram/abstract_195146.htm) GOCHIS, Emily E.1, ROSE, William I.1, VYE, Erika C.1, HUNGWE, Kedmon2, MATTOX, Stephen R.3, and PETCOVIC, Heather4, 2013, INCREASING AWARENESS OF GEOHERITAGE SITES & EARTH SCIENCE LITERACY THROUGH TEACHER-DEVELOPED EARTHCACHES GSA Annual Meeting in Denver: (27-30 October 2013) Paper No. 349-6 (https://gsa.confex.com/gsa/2013AM/finalprogram/abstract_233117.htm) VYE, Erika C.1, ROSE, William I.1, KLAWITER, Mark F.2, and GOCHIS, Emily E. 2013, THE IMPORTANCE OF PARTNERSHIPS FOR IMPROVED EARTH SCIENCE LITERACY AND THE COMMUNICATION OF GEOHERITAGE GSA Annual Meeting in Denver: (27-30 October 2013) Paper No. 349-5 (https://gsa.confex.com/ gsa/2013AM/finalprogram/abstract_232797.htm) ROSE, William I.1, VYE, Erika C.2, KLAWITER, Mark F.2, and GOCHIS, Emily E.2 2013, GEO/BIKE WALK COMMUNICATES GEOHERITAGE IN HOUGHTON, MICHIGAN GSA Annual Meeting in Denver: (27-30 October 2013) Paper No. 318-6 (https://gsa.confex.com/gsa/2013AM/finalprogram/ abstract_226444.htm) VAN WYK DE VRIES, Benjamin, 2013 GEOHERITAGE AND SENSE OF PLACE OF THE CHAîNE DES PUYS AND LIMAGNE FAULT: HOW PEOPLE UNDERSTAND GEOSCIENCE THOUGH BELONGING TO THEIR LANDSCAPE, GSA Annual Meeting in Denver: (27-30 October 2013) Paper No. 318-10 (https://gsa.confex.com/ gsa/2013AM/finalprogram/abstract_223880.htm) 108 �STRUCTURAL CONTROL OF MINERALIZATION AT LAC DES ILES MINE S. SCHMIDT and M.L. HILL Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario, P7B 5E1 SSchmid1@lakeheadu.ca North American Palladium Ltd.’s Lac des Iles mine is located approximately 90km north of Thunder Bay, ON and is the only mine in Canada that is a primary producer of palladium. The mafic Mine Block intrusion that hosts the ore is located in the Wabigoon subprovince north of the regional boundary with the Quetico subprovince, within the Superior province. The property yields evidence for high-temperature deformation in the solid state, indicating that the intrusion is pre- and/or syntectonic, rather than post-tectonic as commonly presumed. The purpose of this MSc thesis project is to discover and examine evidence for high-temperature deformation and assess any controls this deformation may have on the mineralization. Evidence for hightemperature deformation is documented in the North Varitextured rim, Baker zone, Sheriff zone, and Creek zone. From the analysis of structural measurements and field relationships in these areas, two populations of narrow ductile shear zones have been recognized in the North Varitextured rim (NVT), Baker zone, Sheriff zone, and Creek zone. One population has an average orientation of 319/85 (n=30) with a dextral sense of shear, the other population has an average orientation of 058/83 (n=44) with a sinistral shear sense. The NW-striking dextral shear zones are commonly parallel to intrusive features and are variably foliated. The NE-striking sinistral shear zones are discordant to intrusive features. The presence of quartz-carbonate veins within the NE-striking sinistral shear zones may indicate reactivation and/or a local tensile component to stress. The orientation of the intersection lineation between the two populations of ductile shear zones is 81/108. A mutually cross-cutting relationship has been found indicating conjugate formation under the same stress field. 109 �110 �MIDCONTINENT RIFT-RELATED SATELLITE MAFIC-ULTRAMAFIC INTRUSIONS HOSTING FE-TI-V OXIDE DEPOSITS Schulz, K.J., U.S. Geological Survey, 954 National Center, Reston, VA 20192, kschulz@usgs.gov, Woodruff, L.G., U.S. Geological Survey, 2280 Woodale Ave., Mounds View, MN 55112, woodruff@usgs.gov, and Nicholson, S.W., Geological Survey, 954 National Center, Reston, VA 20192, swnich@usgs.gov. The best known Fe-Ti-V oxide deposits in the Midcontinent Rift are in the Duluth Complex, northeastern Minnesota, in two types of deposits: 1) titanomagnetite/ilmenite-rich layers in the early (1107 Ma) Poplar Lake intrusion (formerly Nathan’s Layered Series), and 2) late discordant Oxide-bearing Ultramafic Intrusions (OUI) such as the Longnose and Water Hen intrusions. Less well known are Fe-Ti-V oxide deposits that occur in relatively small (<10 km) mafic-ultramafic intrusions emplaced in country rocks surrounding the Midcontinent Rift. These intrusions are currently known to extend from northwestern Wisconsin into southeastern Minnesota and northeastern Iowa based on geophysics and limited drill core (Fig. 1). The Round Lake intrusion in northwestern Wisconsin is characterized by a large amplitude negative aeromagnetic anomaly; the other intrusions (Clam Lake, WI; Fillmore B1, MN; and Osborne, IA) are characterized by large positive amplitude aeromagnetic anomalies. The Clam Lake intrusion appears to be a plug-like body (Mudrey and others, 2003); the other intrusions appear as linear, northeast-trending dikelike bodies. All the intrusions show variable modal silicate-oxide mineral layering at scales ranging from centimeters to meters; oxide mineral content (Ti-magnetite with variable ilmenite) varies from a few percent to locally massive layers and from intercumulus to cumulus in texture. Strong to moderate igneous flow foliation, defined by aligned plagioclase crystals, is common in all the intrusions. The Clam Lake intrusion is composed of oxide gabbro (plag+cpx+oxide) with some clinopyroxenite layers. Round Lake is dominantly composed of oxide troctolite and melatroctolite (plag+ol+oxide). Both the Clam Lake and Round Lake intrusions are cut by diabase dikes. The Fillmore B1 and Osborne intrusions show greater variability particularly with respect to olivine content and contain oxide dunite and peridotite (ol+oxide) as well as oxide troctolite and melatroctolite (plag+ol+oxide). The Osborne intrusion is oxideand olivine-rich in the upper portion and becomes more plagioclase-rich with depth; it also contains oxide-rich noritic anorthosite layers (plag+opx+oxide). The major element compositions of these intrusions largely reflect their cumulate mineralogy but are dominated by their oxide mineral content. Phosphorous contents are uniformly low (<0.5 wt.%) in all samples and not correlated with TiO2 content. Overall trace element abundances are mostly low as would be expected of dominantly cumulate rocks with low interstitial melt contents. Cobalt, Ni, Sc, and V generally show positive correlations with TiO2 content suggesting that their concentrations are controlled by oxide mineral content. Samples from Round Lake have V contents considerably higher than samples from the other intrusions (up to ~4,500 ppm); the differences in V content may reflect differences in oxygen fugacity between intrusions as V partitioning is strongly dependent on oxygen fugacity. High field strength element (HFSE) covariations within and between intrusions are variable. Samples with high TiO2 mostly show positive Nb-Ta anomalies on primitive mantle-normalized (PMN) trace element plots. However, positive Nb-Ta anomalies are highest in samples from the Clam Lake and Osborne intrusions and weak to absent in samples from Round Lake. In contrast, samples show varying Zr-Hf anomalies on PMN trace element plots ranging from positive anomalies in samples from the Iowa intrusion to no or negative anomalies in samples from Round Lake and Clam Lake. Given the general correlation between HFSE and TiO2 content, it is likely that the HFSE variations are controlled by oxide minerals with different partition coefficients controlled by changing oxygen fugacity. 111 �The REE data show that the intrusions are related to more than one magma type. The Round Lake intrusion has relatively steep REE patterns with enriched light REE and depleted heavy REE. The REE patterns match those of the basal, magnetically reversed basaltic lavas from Pigeon Point and Ely’s Peak. The Fillmore B1 and Osborne intrusions have similar slightly enriched light REE and flat heavy REE patterns. They are likely related to Portage Lake-Chengwatana-equivalent high-TiO2 basalt. The REE patterns for the Clam Lake intrusion have flat light REE and depleted heavy REE; they overlap REE patterns from the BIC intrusion and Eagle Deep dikes in the Baraga basin of Michigan. The Midcontinent Rift-related satellite mafic-ultramafic intrusions and their Fe-Ti-V oxide deposits are very similar to the intrusions hosting Fe-Ti-V oxide deposits in the Permian Emeishan large igneous province of southwest China (Pang and others, 2010). Like the China examples, the Midcontinent Rift intrusions likely formed as conduits experiencing frequent replenishment of fractionated, crystal-rich high-Ti mafic magmas. References cited Mudrey, M.G., Jr., Ervin, C.P., and Olmstead, J.F., 2003, Middle Keweenawan basin evolution inferred from geophysical analysis of a strongly magnetic intrusion, Clam Lake, Wisconsin: Wisconsin Geological and Natural History Survey, Open-File Report 2003-04, 17 p. Pang, K-N, Zhou, M-F, Qi, Liang, Shellnutt, Gregory, Wang, C.Y., and Zhao, Donggao, 2010, Flood basalt-related Fe-Ti oxide deposits in the Emeishan large igneous province, SW China: Lithos, v. 119, p. 123–136. Figure 1. Location of Midcontinent Rift-related satellite mafic-ultramafic intrusions hosting Fe-Ti-V oxide deposits. 112 �THE 2.7 BILLION YEAR OLD MT. ST. HELENS OF NORTHERN MINNESOTA: PETROGRAPHY, GEOCHEMISTRY AND ECONOMIC SIGNIFICANCE OF THE NEOARCHEAN GAFVERT LAKE SEQUENCE SCHWIERSKE, Kelly L.1, PIGNOTTA, Geoffrey S.1 and HUDAK, George J.2 1 University of Wisconsin-Eau Claire, Department of Geology, 105 Garfield Ave., Eau Claire, WI 54701 Precambrian Research Center, Minnesota Natural Resources Research Institute, University of Minnesota-Duluth, 5013 Miller Trunk Hwy, Duluth, MN 55811 2 The Neoachean Gafvert Lake sequence comprises part of the Vermilion District in the Wawa-Abitibi Terrane in northeastern Minnesota and is located in Minnesota’s newest state park, Lake Vermilion State Park (Fig. 1). The Wawa-Abitibi Terrane is the most economically important granite-greenstone belt in the Superior Province, and hosts a wide variety of mineral deposits (including but not limited to shear zone hosted gold deposits, volcanogenic massive sulfide deposits, komatiite-hosted copper-nickelplatinum group element deposits, rare earth element deposits, diamond deposits) in its extents from Minnesota to northeastern Quebec. There has been minimal historic economic mineral exploration in this region despite the striking similarities between the Vermilion District and prolific metal (e.g., Au, Cu, Zn) producing regions across the border in Ontario, Canada. The Gafvert Lake sequence was initially recognized by Peterson and Jirsa (1999) and appeared to represent a stratovolcano complex located immediately up-section from the Soudan Iron Formation, an Algoma-type iron formation unit that hosted Minnesota’s first iron mine, the Soudan Mine. Recent mapping in the Vermilion District, northeast of Ely, MN has documented the regional distribution of rocks associated with the Gafvert Lake sequence which consists of intermediate to felsic volcanic and volcaniclastic rocks intruded by intermediate plutons that are likely age equivalent (Hudak et al., 2004). A dacitic tuff breccia from the Gafvert Lake sequence yielded a 2689.7 ± 0.8 Ma U-Pb age indicating that these deposits lie unconformably on Lower Ely and Soudan members of the Soudan belt (Fig. 1; Lodge et al., 2013). This project examines the petrographic, geochemical and structural characteristics of the Gafvert Lake sequence. In the field, this package of volcanics and associated plutons is strikingly similar to other arc volcano-plutonic complexes found in more recent, Mesozoic and Cenozoic subduction zone related arc systems, like those exposed along the western margin of North America. Field, petrologic, and structural relationships suggest that the Gafvert Lake sequence volcanic rocks are dominantly intermediate in composition and comprised of a series of flows, welded tuffs, and volcaniclastic breccias. Petrographic analyses also show that primary textures are generally well preserved in the volcanics. Preliminary geochemistry indicates that the sequence is dominantly rhyodacite to dacite. Trace element chemistry suggest that the sequence formed in a volcanic arc setting. The volcanics are intruded by a very coarse crystalline to porphyritic tonalite to granite complex called the Gafvert Lake intrusive complex that is geochemically identical to the volcanic package. The volcano-plutonic complex is cut by several steeply dipping, east-west trending, dextral shear zones with stretching lineations that are shallowly east plunging. The tectonic and structural setting of the Gafvert Lake sequence suggests that there is economic potential in this package of rocks due to its strikingly similar characteristics to other economically viable volcano-plutonic systems in the Wawa-Abitibi Terrane. 113 �Figure 1. The Gafvert Lake sequence is exposed in the Vermilion District between Soudan Mine and the Mud Creek shear zone (modified from Lodge et al., 2013). References Hudak, G.J., Heine, J., Jirsa, M. and Peterson, D.M., 2004, Volcanic stratigraphy, hydrothermal alteration, and VMS potential of the lower Ely Greenstone, Fivemile Lake to Sixmile Lake area. 50th Annual Meeting, Institute on Lake Superior Geology, Field Trip Guidebook volume 50. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., Jirsa, M.A. and Hamilton, M.A., 2013, New U–Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince,Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province. Precambrian Research (235), 264-277. Peterson, D.M. and Jirsa, M.A., 1999, Bedrock geological map and mineral exploration data,western Vermilion District, St. Louis and Lake counties, northeastern Minnesota: St. Paul, Minnesota Geological Survey Miscellaneous Map Series M-98. Peterson, D.M., Jirsa, M.A., and Hudak, G.J., 2009, Architecture of an Archean Greenstone Belt: Stratigraphy, structure and mineralization. 55th Annual Meeting, Institute on Lake Superior Geology, Field Trip Guidebook volume 55. 114 �THE DANGER OF “SULFIDE MINING” IN THE LAKE SUPERIOR REGION SEAL, Robert R. II1, PIATAK, Nadine M.1, and WOODRUFF, Laurel G.2 1 2 U.S. Geological Survey, Reston, VA 20192, rseal@usgs.gov, npiatak@usgs.gov U.S. Geological Survey, Mounds View, MN 55112, woodruff@usgs.gov The danger of “sulfide mining” is in the term itself. The term “sulfide mining” undermines meaningful evaluation of environmental risks associated with metal mining because it fails to recognize important influences that geology, hydrology, climate, mining methods, ore-processing methods, and continued evolution of environmental management practices have on environmental risks specific to prospective deposit types. It also places an overemphasis on a single environmental risk – acid generation from the weathering of sulfide minerals – when other risks, such as trace elements in water and solids warrant thorough consideration as well. The mineral deposits in the Lake Superior region, especially the Cu-Ni-PGM ores, highlight the importance of a geologically-based context for assessing environmental risk and designing sound environmental management practices for mining. Acid-mine drainage results from the oxidative weathering of sulfide minerals, principally pyrite or pyrrhotite, in the presence of oxygen or other oxidant, such as dissolved ferric iron. The risk of acid generation from waste rock or mill tailings is directly proportional to the amount of these sulfide minerals present, and inversely proportional to the amount of acid-neutralizing potential available. In the case of the Cu-Ni-PGM ores in the Lake Superior region, the deposits anchor both ends of a wide spectrum in terms of sulfide abundance and therefore acid-generating potential. At one end of the spectrum are the magmatic disseminated sulfide deposits of the Duluth Complex, northern Minnesota; at the other end is the Eagle magmatic massive sulfide deposit in Upper Peninsula of Michigan. Drill core from along the strike of the basal mineralized zone in the Partridge River and South Kawishiwi intrusions of the Duluth Complex, sampled from the Minnesota Department of Natural Resources Core Library, average 0.4 ± 0.7 weight percent S (range <0.05 – 2.6 wt. %), whereas semi-massive to massive sulfide ore at the Eagle mine has sulfur contents ranging from 13 to 36 weight percent (Kennecott Eagle Minerals, 2006). Further, the hydrometallurgical technique likely to be used on the disseminated Duluth Complex ores uses a bulk sulfide concentrate, effectively removing most of the sulfide from the solid waste, whereas the Eagle mine plans to use traditional froth flotation to produce separate copper and nickel concentrates with the pyrrhotite left in the solid waste. These mineralogical, geochemical, and ore-processing features of the proposed ores all affect the acidgenerating potential of solid waste and associated waste management approaches. The acid-generating potential of mine waste is commonly evaluated through a technique known as “acid-base accounting” that compares the acidgenerating potential (AP) of the material to its acid-neutralizing potential (NP). The AP is inferred from the sulfide content of the material and the NP is typically inferred from its carbonate content. The acid-generating potentials of the disseminated and massive sulfide ores vary dramatically, although neither has significant carbonate acidneutralizing potential (Figure 1). From a mine waste management perspective, NP:AP ratios below 1 are considered to be “probably acid-generating waste”, those above 2 are considered to be “non-probably acid-generating waste”, and those between 1 and 2 are considered to have uncertain potential. Silicate minerals are generally not considered because of their limited NP and slow reaction rates; however, two silicate minerals that have some of the highest NP values are olivine and calcic plagioclase (Jambor et al., 2002), which are important constituents of the host rocks of the Lake Superior Cu-Ni-PGM ores, especially the troctolites of the mineralized basal portion of the Duluth Complex. Trace metals, especially Cu and Ni, in mine waste warrant consideration relative to human health and aquatic ecosystem risks. Both Cu and Ni show strong correlations with sulfur concentrations in drill core from the basal mineralized zone (Figure 2), similar to variations found by Ripley (2014). Metallurgical testing on ore from the NorthMet deposit resulted in 88 to 91 percent recovery of Cu and 67 to 73 percent recovery of Ni; this yielded tailings material with Cu concentrations ranging between 320 and 390 mg/kg and Ni concentrations between 280 and 350 mg/kg (Dreisinger, 2009). The difference in recovery between Cu and Ni suggests that approximately 20 percent of the Ni is hosted by a non-sulfide phase such as olivine. Ripley (2014) found Ni concentrations up to 1,800 mg/kg in olivine from the Partridge River intrusion, which is consistent with this interpretation. Nickel in olivine in mill tailings is less likely to be labile and bioavailable than Ni in residual sulfides. The concentrations of Cu and Ni in the experimental tailings are well below both residential and industrial soil screening levels for human health protection (US Environmental Protection Agency, 2013). However, these ranges 115 �are above probable effects concentrations for sediments relative to aquatic ecosystem protection (MacDonald et al., 2000), indicating that waste management practices must be designed to guard against accidental release of tailings to nearby waterways. The potential release of Cu and Ni from mine waste to surface water and groundwater will depend upon hydrologic setting and chemical setting (factors such as the availability of oxygen or other oxidants) of mine waste in the context of waste management practices. Figure 1. Plot of acid-generating potential (AP) and acid-neutralizing potential (NP) of various magmatic NiCu-PGM deposits, modified from Schulz et al. (2010). Figure 2. Plot of bulk sulfur, copper, and nickel concentrations of drill core from the basal mineralized zone of the Duluth Complex. Consideration of the geologic, mineralogical, and ore-processing characteristics of magmatic mineral deposits in the Lake Superior region provides greater insights into environmental challenges associated with mining and mineral extraction than those from the oversimplified perspective of “sulfide mining”. These insights extend beyond acid-generating potential and include the assessment of potential risks to human health and aquatic ecosystems from trace metals. The identification of environmental risks enables effective mine planning for environmental protection. References Dreisinger, D., 2009, Keynote address: hydrometallurgical process development for complex ores and concentrates. In Proceedings of Hydrometallurgy Conference, v. 2009, p. 187-212. Jambor, J.L., Dutrizac, J.E., Groat, L.A., and Raudsepp, M., 2002, Static tests of neutralization potentials of silicate and aluminosilicate minerals: Environmental Geology, v. 43, p. 1-17. Kennecott Eagle Minerals, 2006, Eagle Project Mining Permit Application, Volume I, 126 p. MacDonald, D.D., Ingersoll, C.G., and Berger, T.A., 2000, Development and evaluation of consensus-based sediment quality guidelines for freshwater ecosystems: Archives of Environmental Contamination and Toxicology, v. 39, no. 1, p. 20–31. Ripley, E.M., 2014, Ni-Cu-PGE Mineralization in the Partridge River, South Kawishiwi, and Eagle Intrusions: A Review of Contrasting Styles of Sulfide-Rich Occurrences in the Midcontinent Rift System: Economic Geology, v. 109, p. 309-324. Schulz, K.J., Chandler, V.W., Nicholson, S.W., Piatak, Nadine, Seal, II, R.R., Woodruff, L.G., and Zientek, M.L., 2010, Magmatic sulfide-rich nickel-copper deposits related to picrite and (or) tholeiitic basalt dike-sill complexes—A preliminary deposit model: U.S. Geological Survey Open-File Report 2010–1179, 25 p. (Available at http://pubs.usgs.gov/of/2010/1179/). U.S. Environmental Protection Agency, 2013, Regional Screening Level (RSL) Summary Table (TR=1E-6, HQ=1) November 2013: available only online at http://www.epa.gov/reg3hwmd/risk/human/rbconcentration_table/Generic_Tables/docs/master_sl_table_run_NOV2013.pdf. (Accessed March 26, 2014.) 116 �GENESIS OF SULFIDE MINERALIZATION WITHIN THE FOOTWALL GRANITE OF THE MATURI CU-NI-PGE DEPOSIT OF THE SOUTH KAWISHIWI INTRUSION, DULUTH COMPLEX, NE MINNESOTA STEINER, Ronald Alex and MILLER, Jim Department of Geological Sciences, University of Minnesota Duluth, Duluth MN 55812 The development of the 1.1 Ga Midcontinent Rift (MCR) generated voluminous magmatism resulting in the extensive flood basalts and sub-volcanic intrusions exposed along the flanks of Lake Superior (Miller et al., 2002). In northeastern Minnesota, two intrusions of the Layered Series, the Partridge River Intrusion (PRI) and South Kawishiwi Intrusion (SKI), are known to hosts significant Cu-Ni-PGE sulfide mineralization (Miller et al., 2002). The Maturi Cu-Ni-PGE deposit occurs along the basal zone of the SKI where it is in contact with granitic rocks of the Archean Giants Range Batholith (GRB). Generally Cu-Ni-PGE-enriched sulfides are disseminated throughout a 50-150m-thick basal mineralized zone (BMZ) and locally may be semimassive to massive sulfide (Bonnichsen, 1974). Several researchers (Severson, 1993; Peterson, 1997; Sawyer, 2002; Hovis, 2003) have noted significant sulfide mineralization in the dominantly granitic footwall. Extensive drilling by Twin Metals Minnesota since 2006 has shown that the mineralization within the footwall is typically disseminated sulfide, to locally massive sulfide veins, that is dominantly composed of chalcopyrite, pyrrhotite, pentlandite, as found in the BMZ. The mineralization extends as deep as 100 meters below the basal contact with the SKI (Kevin Boerst, 2013, personal comm.). While sulfur isotope data show that the sulfide in the mineralized granite originated from the same source as that in the overlying gabbro (Ripley, 1986; Molnar, 2009), the mechanism by which footwall mineralization occurred is unconfirmed. The purpose of this study is to evaluate evidence for possible mechanisms by which the Giants Range Batholith may have become mineralized. Two hypotheses will be evaluated: 1. Partial melting of the GRB resulting in buoyant exchange of dense magmatic sulfide fluid and less dense anatectic melts rising from the GRB. 2. Hydrothermal fluids mobilizing sulfide from the mineralized gabbro into the granitoid rocks. These two hypotheses are being tested by acquiring petrographic and geochemical data from four drill cores from the Maturi deposit that penetrate the gabbro-footwall contact and reach below the mineralized zone in the granite. Three of the drill cores represent variations in different styles of mineralization in the gabbro and the granite recognized in recent exploration drilling (Peterson, 2012). A fourth core was selected for the extensive occurrence of mineralization into a large biotite schist enclave within the batholith. Preliminary results suggest a relationship between the degree of partial melting of the GRB and sulfide mineralization. Core logging and subsequent petrographic observations indicate that the footwall experienced pyroxene hornfels grade metamorphism producing orthopyroxene and clinopyroxene at the expense of biotite and amphibole that extends in excess of 30 meters from of the SKI-GRB contact. The volatiles produced by recrystallization of biotite and amphibole likely played a role in promoting anatectic melting of the granite as well. Petrographic evidence of partial melting of the GRB recognized in this and previous studies (Sawyer, 1999 2002; Hovis, 2003) included mylonitic textures, pockets of polygonal quartz-orthoclase-plagioclase aggregates, and lattice-dislocation textures in plagioclase. Leucosome patches have been observed to contain massive to semi-massive sulfide suggesting a relationship between escaping partial melts and sulfide liquid. A retrograde alteration of metamorphic pyroxene to biotite, cummingtonite/actinolite, and chlorite is evidence of post-metamorphic hydrothermal alteration. This late hydrothermal alteration assemblage, which is recognized throughout the granite, typically does not contain significant sulfide. Additionally, where sulfides are present they appear largely unaffected by hydrothermal alteration indicating that this event did not cause significant sulfide remobilization or recrystallization. The presence of rare, late gypsum may indicate that the hydrothermal fluids were strongly oxidized and that any remobilized sulfur was crystallized as sulfate. Petrographic observations implying exchange of anatectic melts and sulfide liquid are also supported by geochemical analyses. All REE became increasingly depleted with increased proximity to the gabbro 117 �contact except for Eu with appears as a peak on the diagrams. During partial melting, Eu is likely being retained in plagioclase whereas other REE will be partioned into partial melts which are able to escape the system. Plotting S concentration against Eu/Ce (Fig. 1A) shows a positive correlation indicating that the amount of anatectic melt escaped generally correlates with an increase in sulfide mineralization. Another proxy of increased escape of anatectic melt is an increase in plagioclase relative to quartz and alkali feldspar. A plot of CIPW norm values of Ab/(Or+Qtz) vs. wt% S (Fig. 1B) shows a similar, though broader, positive correlation. Research is ongoing to further test these hypotheses by evaluating isocon plots (Grant, 1986) of whole rock geochemistry and by acquiring SEM-EDS analyses of pyroxene and amphibole compositions. The isocon method will be applied to determine element mobility through the system in order to better identify the mechanism of mineralization. Partial melting or hydrothermal alteration have distinct elemental signatures that can be identified in isocon modeling. Mineral chemistry acquired by SEM-EDS analyses will be used to trace changes in mineral composition relative to distance from the SKI-GRB contact. References Grant, James A. 1986 "The Isocon Diagram-A Simple solution to Gresens' Equation for Metasomatic Alteration."Economic Geology. Vol. 81, 1976-1982 Hovis, Steven T., 2003,.”Observations on Cu-Ni Mineralization in the Giants Range Batholith Footwall of the South Kawishiwi Intrusion, Duluth Complex, Northeastern Minnesota”., Natural Resources Research Institute; University of Minnesota, NRRI/TR-2003/24 Miller, J.D. Jr., Green, J.C., Severson, M. J., Chandler, V. W., Hauck, S. A., Peterson, D. M., and Wahl, T. E., 2002, “Geology and Mineral potential of the Duluth Complex and related rocks of Northeastern Minnesota”, Minnesota Geological Survey, Report of Investigations 58 Molnar, F., Peterson, D. M., Arehart, G. B., Hauck, S. A., 2009, “Sulfur isotope constraints for a dynamic magmatic sulfide ore deposition model in the sill-like South Kawishiwi Intrusion of the Duluth Complex, Minnesota, USA”, Geological Society of America, Abstract. Peterson D.M. 2012, “Maturi Geological Model” Duluth Metals ltd Presentation to Twin Metals Minnesota LLC Sawyer, E. W., 1999, Criteria for the Recognition of Partial Melting, Physical Chemistry Earth, Vol. 24, No. 2, pp. 269-279 Sawyer, E. W., 2002, “Report on Thin Sections From DDH WM-1, Spruce Road Cu-Ni Deposit, South Kawishiwi Intrusion, Duluth Complex”, Natural Resources Research Institute; University of Minnesota, NRRI/RI-2002/13 Severson, M. J., 1994, “Igneous stratigraphy of the South Kawishiwi Intrusion, Duluth Complex, Northeastern Minnesota”, Natural Resources Research Institute, University of Minnesota, Duluth, NRRI/TR-93/34 Ripley, E. M. & Alawi, J. A., 1986, “Sulfide mineralogy and chemical evolution of the Babbitt Cu—Ni deposit, Duluth Complex, Minnesota”, Canadian Mineralogist Vol. 24, 347-368 118 �SULFIDE HIGHWAY REVISITED: NEW IDEAS ON INTERNAL STRUCTURE AND SULFIDE MINERALIZATION OF THE NICKEL LAKE MACRODIKE SWEET, Gabriel J., PETERSON, Dean M., LARSON, Philip C., FINNEGAN, Molly L., FINNES, Evan, PARENT, Charles, NOWAK, Robert, BOLEY, Tyler D. Duluth Metals, 306 W. Superior St., Suite 610, Duluth, MN 55802 Recent exploratory drilling within the Nickel Lake Macrodike (NLM) by Duluth Metals has facilitated the first subsurface investigations of the magma conduit into the prolifically mineralized South Kawishiwi Intrusion (SKI). Seventeen holes were drilled in late 2012 and early 2013 along 6500’ feet of strike length of the NLM, with 4 holes reaching depths of over 4000’. Intercepts of the primary surface lithologies suggests that the youngest units of the NLM (the variably pegmatoidal oxide gabbro (N-xG) and layered troctolite (N-Tl)) are dipping irregularly (~30⁰) towards the northwest anorthositic sidewall of the NLM. Based on limited pierce points through this internal stratigraphy, the oxide gabbro appears to extend into the anorthosite series wall rocks beyond the surface-mapped northwestern sidewall contact of the NLM. At depth, the top of the N-xG truncates the xenolithbearing (dominantly hornfelsed North Shore Volcanic Basalt and Biwabik Iron Formation) heterogeneous troctolite (N-Th) and sulfide-bearing troctolite (N-Ts) packages emplaced along the southeast-dipping (60⁰) northwest margin of the NLM. Below the N-xG and N-Tl units is a second package of heterogeneous troctolite (Th). Unlike the heavily xenolith-bearing N-Th unit present at and near surface, this troctolite is sparsely populated by small (~10’) hornfelsed basalt xenoliths. This same heterogeneous troctolite hosts a series of large (100’ thick) rafts of Virginia Formation argillite and greywacke at depth (~3000’) in the south-central portion of the NLM. Sulfide mineralization was encountered along the northwestern margin of the NLM both as the down-dip extension of outcropping and subcropping mineralization, as well as at greater depths. Three distinct types of mineralization were defined with respect to overall sample grades, interval thicknesses and lithological associations: Type 1 - long intervals (~50’-200’+) of moderate grade disseminated to blebby chalcopyrite and pyrrhotite (broadly 0.4%Cu, 0.1%Ni and 0.15g/t Pt+Pd+Au) associated with variably hornfelsed basalt xenolith-bearing Th, Type 2 - short intervals (~5’-35’) of higher grade disseminated to blebby chalcopyrite and pyrrhotite (upwards of 0.55%Cu, 0.11%Ni and from 0.25g/t to over 2.0g/t Pt+Pd+Au) generally intercepted deeper than the larger, moderate grade intervals, and Type 3 - variable-length intervals (~25’-400’) of low grade, disseminated to coarse-grained pyrrhotite and minor chalcopyrite (generally < 0.25%Cu, <0.10%Ni and <0.10g/t Pt+Pd+Au) hosted by a variably oxide-rich pyroxenite. Mineralization Type 1 and Type 2 tend to occur at shallow depths in the majority of drill holes along the western margin of the NLM. However, the Type 3 mineralization is confined to the shallow southwestern NLM, where it is found in close proximity to a large iron-formation xenolith (~1300’ strike length as mapped at surface). Comparison of the geochemical signature of NLM sulfide mineralization types to basal SKI mineralization suggests a distinctly different fractional history for the NLM sulfide populations. Copper-nickel ratios for Types 1 and 2 tend to fall between 4:1 to 5:1, and 5:1 to 6:1, respectively. Coupled with low to moderate TPM tenors, NLM mineralization is distinct from the broadly 3:1 Cu:Ni ratio and propensity towards elevated precious metal tenors of the basal SKI mineralization. At this time, no direct analogue to basal SKI mineralization has been identified in the NLM. The deviations in precious metal tenor between Type1 and Type 2 mineralization, and the SKI may speak more directly to processes operating “downstream” of the NLM. With Cu:Ni ratios of 1:1 to 3:1, Type 3 mineralization is distinctly more Ni-rich than Types 1 and 2, and the basal SKI mineralization. Type 3 mineralization is further distinguished by its unique pyroxenite host rock, and highly elevated P (up to 1%) and Zn (generally >175ppm), respectively up 119 �to two orders of magnitude and double that of the vast majority of basal SKI mineralization. However, the Type 3 lithological association and distinct geochemical signature shows affinity with reported rock types and whole rock compositions of apatite-bearing, mineralized oxide-rich ultramafic intrusions (OUIs; Ripley et al., 1998). The spatial association with a large xenolith of metamorphosed iron formation is also in line with the observations of Severson (1995), who noted that OUIs in the basal central Duluth Complex occur in close proximity to metamorphosed Biwabik Iron Formation. The lithological relationships noted in the drilling confirm the intrusive sequence of the NLM as suggested by Peterson et al. (2006), but imply a slightly different geometry for the youngest intrusive phases internal to the NLM. The difference between the xenolith populations in the near surface N-Th unit and the Th at depth may indicate origination from different pulses of troctolitic magma through the conduit. The long intervals of xenolith-poor Th (up to 2500’+) at depth within the central NLM may represent areas of high magma flow through the conduit. The existence of multipe types of mineralization along the margins of the NLM magma conduit indicates sulfide mineralized magmas passed through the conduit. Variation of the NLM mineralization grade and tenor from that of basal SKI mineralization may be the result of fractionation processes that occurred down-stream of the NLM conduit (e.g., sulfide dissolution upgrading; Kerr and Leitch, 2005), or it may ultimately correlate with undiscovered mineralization with the SKI-NLM system. Mineralization Type Type 1 Type 2 Type 3 Boulder Lake Figure 1 – Geological map of the southern Nickel Lake Macrodike with major lithological units and Duluth Metals’ drill pads (black dots). Figure 2 – Zn vs P2O5 comparison of NLM mineralization types with oxideapatite- rich samples from DDH IV-2 from the Boulder Lake Intrusion in the southwestern Duluth Complex, from Ripley et al., 1998. REFERENCES Kerr, A. and Leitch, A., 2005, Self-Destructive Sulfide Segregation Systems and the Formation of High-Grade Magmatic Ore Deposits: Economic Geology, Vol. 100, p. 311-332. Peterson, D.M, Albers, P.B., and White, C.R., 2006, Bedrock Geology of the Nickel Lake Macrodike and Adjacent Areas, Lake County, northeastern Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Map Series NRRI/MAP-2006-04, scale 1:10,000. Ripley, E.M., Severson, M.J. and Hauck, S.A., 1998, Evidence for Sulfide and Fe-Ti-P-Rich Liquid Immiscibility in the Duluth Complex, Minnesota: Economic Geology, Vol. 93, p. 1052-1062. Severson, M.J., 1995, Geology of the Southern Portion of the Duluth Complex: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-95/26, 185p. 120 �THE THUNDER MAFIC TO ULTRAMAFIC INTRUSION: A PGE AND PRECIOUS METAL BEARING EARLY-RIFT CONDUIT SYSTEM IN THE MIDCONTINENT RIFT TREVISAN, Brent1, HOLLINGS, Pete1, and AMES, DOREEN2 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario, P7B 5E1 Geological Survey of Canada, Central Canada Division, 750-601 Booth St., Ottawa, Ontario, K1A 0E8 2 In 2002, high grade massive Ni-Cu-PGE sulphide mineralization at the Eagle deposit near Marquette, Michigan was discovered stimulating exploration programs in search of small mafic to ultramafic intrusions hosting “conduit-type” magmatic sulphide mineralization associated with the early stages of the Mesoproterozoic Midcontinent Rift (MCR; Miller and Nicholson, 2013; Ripley, 2014). Since the Eagle discovery over a half-dozen poorly exposed mineralized early-rift mafic to ultramafic intrusions have been discovered within the Lake Superior region, prompting active petrological research (e.g., Ding et al., 2010) and re-evaluation of the current MCR tectono-magmatic model (e.g., Miller and Nicholson, 2013). However, from an exploration stand point the small size of these buried mineralized mafic to ultramafic intrusions makes them difficult to locate both on the ground and on regional magnetic survey maps (Ames et al., 2012). This study is a collaborative project between the Geological Survey of Canada, the Ontario Geological Survey, and Lakehead University as part of the Ni-Cu-PGE-Cr project, Targeted Geoscience Initiative-4 (TGI-4; Ames et al., 2012). The objective is to characterise the petrology, mineralization, and alteration footprint of the Thunder intrusion within the context of the MCR as a whole, in order to identify criteria for targeting buried mineralization. The Thunder intrusion is a small, layered mafic to ultramafic intrusion located on the outskirts of Thunder Bay, ON, which has been explored by Rio Tinto (formerly Kennecott Canada Exploration Inc.) in 2005 and 2007 (Bidwell and Marino, 2007). The intrusion is interpreted to be associated with the early magmatic stages of the MCR based on geochemical similarities to mafic and ultramafic rocks of the Nipigon Embayment (Hollings et al., 2007) and an unpublished 207Pb/206Pb baddeleyite age of 1110.33±0.92 Ma (Ames, pers. comm., 2014). This intrusion is distinct from the other known mineralized early-rift intrusions as it is the only known occurrence hosted by the Archean Shebandowan greenstone belt. The intrusion is approximately 800 by 1000m by < 500 m thick and dips steeply to the south. Major textural and geochemical differences can be used to divide the lithostratigraphy into a lower mafic to ultramafic basal unit and an upper gabbroic unit, however, similar trace and rare earth element ratios of the two units suggests they formed from a single magmatic pulse that has undergone subsequent fractionation. Ni-Cu-PGE mineralization is hosted by clinopyroxenite in the lower mafic to ultramafic unit adjacent to the basal wall rock, including 20 m of 0.22% Cu, 0.06% Ni, 0.25ppm Pt, 0.29ppm Pd (Bidwell and Marino, 2007). Sulphides rarely comprise up to 30% by volume but more typically 1-5%, with textures ranging from medium- to fine-grained disseminated, globular and rarely net-textured. Pyrrhotite, chalcopyrite and rare pentlandite with common secondary marcasite-pyrite replacement are present along with trace kotultskite, naldrettite, merenskyite, sperrylite, electrum and native silver. The δ34S values of sulphide minerals from the Thunder intrusion are similar to the adjacent wall rock forming a tight range between +3.8 and -3.1‰. Although δ34S values are broadly consistent with a 121 �mantle origin (0 ± 2‰) the involvement of crustal sulphur during the mineralization process remains a possibility. Radiogenic isotopes were measured from select samples to investigate possible contamination of the Thunder intrusion. The εNd values from the intrusion range between -0.74 and +0.99, with no trends towards wall rock compositions, whereas the 87Sr/86Sr values range from 0.7031 and 0.7061 and trend towards wall rock values of 0.7071 and 0.7087. The decoupling of the two radiogenic isotope signatures is consistent with crustal contamination at depth and local contamination during the emplacement of the Thunder intrusion. References Ames, D.E. et al. 2012. Update on Research Activities in the Targeted Geoscience Initiative 4 MagmaticHydrothermal Nickel-Copper-Platinum Group Elements Ore System Subproject: System Fertility and Ore Vectors. Summary of Field Work and Other Activities 2012. Ontario Geological Survey, Open File Report 6280. Bidwell, G. E., and Marino, F. 2007. 2007 drilling assessment report for the Geoinformatics Exploration Canada Ltd Thunder Project; Thunder Bay South District, Assessment Files, AFRO report number 2.34638, 112p. Ding, X., Li, C., Ripley, E.M., Rossell, D., and Kamo, S. 2010. The Eagle and East Eagle sulfide ore‐bearing maficultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution. Geochemisty, Geophysics, Geosystems, v. 11, p. 1-22. Hollings, P., Hart, T., Richardson, A., and MacDonald, C.A. 2007. Geochemistry of the Mesoproterozoic intrusive rocks of the Nipigon Embayment, northwestern Ontario: evaluating the earliest phases of rift development. Canadian Journal of Earth Sciences, v. 44, p. 1087–1110. Miller, J., and Nicholson, S. 2013. Geology and Mineral Deposits of the1.1Ga Midcontinent Rift in the Lake Superior Region: An Overview, in Field Guide to Copper-Nickel-Platinum Group Element Deposits of the Lake Superior Region. Precambrian Research Center Guidebook, v. 13-01, p. 1-50. Ripley, E. M. 2014. Ni-Cu-PGE Mineralization in the Partridge River, South Kawishiwi, and Eagle Intrusions: A Review of Contrasting Styles of Sulfide-Rich Occurrences in the Midcontinent Rift System. Economic Geology, v. 100, p. 309-324. 122 �GARNET IN THE DEEP CRUST: THE KEY TO LINKING ARCHEAN TTG GENERATION AND VERTICAL BLOCK MOTIONS? VAN LANKVELT, A1, WILLIAMS, ML1, SCHNEIDER, DA2, SEAMAN, SJ1 1 Department of Geosciences, University of Massachusetts-Amherst, 611 North Pleasant St, Amherst, MA 01003 USA 2 Department of Earth Sciences, University of Ottawa, 140 Louis-Pasteur Pvt. Ottawa, ON K1N 6N5 Canada There is substantial debate about the differences between geologic processes that operated during the Archean and those operating today (e.g. Percival et al., 2006; Van Kranendonk, 2010). Two notable differences between modern terranes and Archean cratons are the presence of very large volumes of TTGs (tonalite-trondhjemite-granodiorite) and tectonic structures suggesting that locally, TTG blocks moved up relative to adjacent mafic greenstone belts. Explanations for the structural differences between modern and early Earth vary between two endmember models: density-driven (vertical) and modern-style (horizontal) tectonics. Many field-based structural studies of Archean rocks contain evidence for both vertical and horizontal tectonics (e.g. Lin, 2006). When the spatial scales of vertical and horizontal structures are considered, horizontal structures seem to dominate the large-scale province framework and are more prevalent in high-grade rocks, whereas low-grade and local (greenstone belt-scale) structures exhibit more evidence for vertical processes (Van Kranendonk, 2010). Current interpretations for the origins of TTGs favor partial melting of a mafic, hydrous, garnetbearing source rock. TTGs are chemically similar to modern arc rocks, so the suprasubduction-zone setting of modern arcs is commonly invoked for the generation of TTGs as well (see Moyen & Martin, 2012, and references therein). Other suggestions invoke plume-related process for generating larger volumes of TTGs than are found in modern arc settings (e.g. Moyen & Martin, 2012). Below, we attempt to integrate geodynamic and petrogenetic models for Archean tectonics. One approach to better understand geodynamic and petrogenetic processes operating in the Archean is to compare rocks from different crustal levels. The North Caribou Terrane is located in the central Superior Province, and it is dominated by Meso- to Neoarchean TTGs. Lin (2006) studied the structures within and adjacent to greenstone belts in the North Caribou and concluded that these rocks preserve structures consistent with synchronous vertical and horizontal tectonism. Several studies of the TTGs in the North Caribou show that their compositions are consistent with typical Archean TTGs (e.g. Wyman et al., 2011). The thermobarometric and geochemical data we present from TTGs in the North Caribou are consistent with Wyman et al.’s (2013) data and indicate that the TTGs were emplaced in the mid-crust and later metamorphosed at shallower levels. The Athabasca granulite triangle is an exposure of deep-crustal rocks that straddles the Snowbird Tectonic Zone north of Lake Athabasca. The granulite-facies (0.9-1.9 GPa, 700-950 ºC; Baldwin et al., 2003) terrane contains several generations of mafic rocks, two of which, Neoarchean gabbros and Paleoproterozoic mafic dikes, are possible analogues to TTG source rocks, as both preserve primary hornblende and contain garnet. Although the source for the older gabbros is not fully understood, the dikes are not associated with arc magmatic rocks (Flowers et al., 2006). Some of the mafic dikes have undergone anatexis due to dehydration melting of hornblende, resulting in tonalitic melt and peritectic garnet (Williams et al., 1995). Implications The rocks in the Athabasca provide an interesting option for the generation of TTGs. The mafic dikes, which are not related to subduction, indicate the potential for long-term storage of water in the 123 �lithospheric mantle, so concurrent subduction and melting are not necessarily required. Their anatectic textures also suggest the possibility for melting through underplating or mantle upwelling (Williams et al., 1995). Regardless of the melting mechanism, extraction of significant volumes of TTG parent magma from mafic rocks would leave a garnet-rich restite, similar to what has been postulated by Saleeby et al. (2003) to exist below the Sierra Nevada batholith. Foundering of this gravitationally unstable lithospheric root may be the cause of the uplift of the Sierra Nevada (Saleeby et al., 2003), and a similar delamination scenario in the Archean may explain relative uplift of TTGs compared to adjacent greenstone belts. This is observed in the North Caribou, and a higher frequency of delamination events in the Archean due to widespread extraction of melt could explain the evidence for vertical tectonics. Delamination of dense restite and subsequent mantle upwelling could also trigger additional melting, creating a positive feedback mechanism that could produce significant amounts of tonalitic magma, like the large batholiths that are common in the North Caribou. This scenario would require a wet mantle, either through the release of primordial water or an earlier introduction of volatiles. Structural evidence for horizontal tectonics preserved in large-scale structures suggests that subduction may have operated during the Archean, but not all TTG-type magmas need to have been derived at subduction zones. Instead, TTGs could be generated in several non-unique tectonic settings, and garnet-driven delamination of the lower crust can explain both evidence for vertical tectonics and large volumes of TTGs. References Baldwin, JA, Bowring, SA, Williams, ML. 2003. Petrological and geochronological constraints on high pressure, high temperature metamorphism in the Snowbird tectonic zone, Canada. J Metamorphic Geol 21, 81-98. Flowers, RM, Bowring, SA, Williams ML. 2006. Timescales and significance of high-pressure, hightempperature metamorphism and mafic dike anatexis, Snowbird tectonic zone, Canada. Contrib Min Petrol 151, 558-581. Lin, S. 2006. Synchronous vertical and horizontal tectonism in the Neoarchean: Kinematic evidence from a synclinal keel in the northwestern Superior Craton, Canada. Precam Res 139, 181-194. Moyen, J-F, Martin, H. 2012. Forty years of TTG research. Lithos 148, 312-336. Percival, JA, Sanborn-Barrie, M, Skulski, T, Stott, GM, Helmstaedt, H, White, DJ. 2006. Tectonic evolution of the western Superior Province from NATMAP and Lithoprobe. Can J Earth Sci 43, 1085-1117. Saleeby, J, Ducea, M, Clemens-Knott, D. 2003. Production and loss of high-density batholithic root, southern Sierra Nevada, California. Tectonic 22, 1064-1087. Van Kranendonk, MJ. 2010. Two types of Archean continental crust: plume and plate tectonics on early Earth. Am J Sci 310, 1187-1209. Williams, ML, Hanmer, S, Kopf, C, Darrach, M. 1995. Syntectonic generation and segregation of tonalitic melts from amphibolite dikes in the lower crust, Striding-Athabasca mylonite zone, northern Saskatchewan. J Geophys Res 100, 15717-15734. Wyman, DA, Hollings, P, Biczok, J. 2011. Crustal evolution in a cratonic nucleus: Granitoids and felsic volcanic rocks of the North Caribou Terrane, Superior Province, Canada. Lithos 123, 37-49. 124 �STRONTIUM ISOTOPE STUDY OF MESABI IRON RANGE GROUNDWATER Walsh, James F. Minnesota Department of Health, St. Paul, MN 55164 On the Mesabi Iron Range, significant differences in 87Sr/86Sr exist between long residence time groundwater from wells completed in the Biwabik Iron Formation, especially where covered by the Virginia Formation, and short residence time groundwater from wells completed in overlying glacial aquifers and in surface waters. The relatively low 87Sr/86Sr observed at the iron formation wells falls within the range commonly observed for weathering of Phanerozoic marine carbonates, whereas the higher values observed at the drift wells and surface waters are more characteristic of weathering of Archean silicate minerals. Short residence time water from Biwabik Iron Formation wells situated in the subcrop of the formation span a wide range of 87Sr/86Sr and in some cases is more radiogenic than that observed at the glacial drift wells and surface water bodies. These results likely reflect the impact of glacial provenance on the distribution of strontium-bearing minerals within their groundwater flow pathways. The glacial deposits on the Mesabi Range are dominated by northeast-sourced glaciers of the Rainy Lobe, whose sediments are characterized by an abundance of Archean granitic material and scarcity of Phanerozoic marine sediments. However, northwest-sourced glacial sediments are recognized locally and have contributed sediments relatively rich in Phanerozoic marine carbonate and shale, especially along the west-central Mesabi Range. It is likely that water samples with high 87Sr/86Sr and low strontium concentrations are predominantly influenced by recharge through Rainy Lobe glacial sediments. In contrast, those that are relatively low in 87Sr/86Sr but high in strontium concentration are predominantly reflecting dissolution of carbonate minerals from northwest-sourced glacial deposits or from the iron formation itself. 125 �126 �GEOCHEMISTRY AND MINERALOGY OF GLACIAL SOILS IN THE UPPER MIDWEST WOODRUFF, Laurel G., U.S. Geological Survey, St. Paul, MN 55112 (woodruff@usgs.gov) CANNON, William F. and SOLANO, Federico, U.S. Geological Survey, Reston, VA 20192 SMITH, David B., U.S. Geological Survey, Denver, CO 80225 The U.S. Geological Survey has recently completed a low-density (1 site per 1,600 square kilometers, 4,857 sites) geochemical and mineralogical survey of soils of the conterminous United States (Smith et al., 2013). Three samples were collected, if possible, from each site; (1) a sample from a depth of 0 to 5 centimeters, (2) a composite of the soil A horizon, and (3) a deeper sample from the soil C horizon or, if the top of the C horizon was at a depth greater than 1 meter, from a depth of approximately 80–100 centimeters. The <2-millimeter fraction of each sample was analyzed by a combined inductively coupled plasma-atomic emission spectrometry/mass spectrometry method for a suite of 45 major and trace elements following near-total multiacid digestion. The major mineralogical components in samples from the soil A and C horizons were determined by a quantitative X-ray diffraction method. Regional- and national-scale element and mineral patterns can be related to (1) soil parent materials, (2) climate factors, (3) soil age, and (4) possible anthropogenic loading to surface soils. This presentation will describe the influence of source provenance and soil age factors on the geochemistry and mineralogy of the soil A and C horizons in the upper Midwest. In the upper Midwest, melting of glacial ice left the region mantled with a blanket of mixed, immature sediments from which present day soils developed. Individual ice lobes of the late Wisconsinan glaciation created distinct patterns in soil geochemistry and mineralogy because of varying provenance and transport paths. Carbonate- and shale-rich ‘gray’ tills in Minnesota, North Dakota, South Dakota, and Iowa, deposited by the Des Moines and James lobes were derived from Cretaceous sedimentary rocks (dolostone, limestone, shale); glaciolacustrine sediments of Glacial Lake Agassiz along the North Dakota/Minnesota border have a similar provenance (Wright, 1972). Gray tills were transported significant distances to the south and southeast from their source and deposited on Precambrian bedrock that is largely devoid of carbonate minerals. ‘Red’ tills were deposited in northeastern Minnesota and northern Michigan and Wisconsin by the Rainy and Superior lobes. The Rainy lobe provenance is mainly Precambrian crystalline rocks of the Canadian Shield and the Superior lobe provenance is mainly basalts and sediments of the Precambrian Keweenawan Supergroup (Wright, 1972). In the lower Great Lakes region, carbonate- and shale-bearing tills sourced from the Cambrian to Devonian sedimentary bedrock units that rim the Michigan basin were deposited by the Green Bay and Lake Michigan lobes in western Wisconsin and northern Illinois, by the Saginaw lobe in central Michigan, and by the Huron-Erie lobe in eastern Indiana and western Ohio (Johnson, 1986; Hofer and Szabo, 1993). Soils developed on glacial sediments are relatively young and often retain easily weathered minerals and mobile elements, such as carbonates and related elements (e.g., Ca and Mg), typically leached from older, more mature soils beyond the southern extent of the last glaciation. As expected from their differing provenance, soils developed on red tills have much lower clay contents and much higher quartz and feldspar contents compared to soils developed on gray tills. This divergent mineralogy creates striking contrasts in element concentrations. Soils on gray tills have higher Ca contents from carbonate as well as higher As, Cd, Mo, Sb, and U concentrations, likely contributed by the shale component, compared to soils on red tills, which have higher Na and K contents from the higher feldspar content. Soils developed on the James lobe have somewhat higher Mn contents than soils developed on the Des Moines lobe, perhaps related to local redox conditions. Soils developed on Lake Agassiz clays have relatively higher Li, Sc, Ti, V, and Zn contents compared to soils developed on surrounding gray tills. One of the more dramatic characteristics of some glacial soils in the upper Midwest is a high concentration of primary dolomite. Tills of the Green Bay and Lake Michigan lobes are characterized by an especially high content of dolomite relative to calcite, and as a consequence, these soils have some of the 127 �highest soil Mg concentrations in the conterminous United States. The Green Bay and adjacent Lake Michigan lobes, as well as the Saginaw and adjacent Huron-Erie lobes are all largely sourced from similar rocks (dominantly dolostone, limestone, and black shale). However, a strong contrast in Mo contents in the soil C horizon between the Green Bay and Lake Michigan lobes and between the Saginaw and Huron-Erie lobes is an indication of the proportion of black shale incorporated into their respective glacial sediments by the individual lobes (Figure 1). The higher the percentage of black shale, the higher Mo content of soils, as well as a number of other elements such as As, Cd, Co, K, Sb, Tl, U, and Zn, all of which may be enriched in black shale. Because of this shale influence, large areas of northern Ohio and Indiana have some of the higher soil Mo concentrations in the conterminous United States. Thus, glacial dispersal of materials sourced from different bedrock sources, especially relatively thin shale units, had a widespread effect on soil geochemistry and mineralogy throughout the glaciated upper Midwest. Figure 1. Interpolated concentration map depicting molybdenum (Mo) in the soil C horizon in the lower Great Lakes region. The hachured black line is the maximum southern extent of Wisconsinian glaciation; dotted lines are the approximate margins of named individual ice lobes, with arrows indicating major ice flow direction (after Grimley, 2000; Hofer and Szabo, 1993). References Grimley, D.A., 2000. Glacial and nonglacial sediment contributions to Wisconsin Episode loess in the central United States. Geological Society of America Bulletin 112, 1475-1495. Hofer, J.W., and Szabo, J.P., 1993. Port Bruce ice-flow directions based on heavy-mineral assemblages in tills from the south shore of Lake Erie in Ohio. Canadian Journal of Earth Sciences 30, 1236-1241. Johnson, W.H., 1986. Stratigraphy and correlation of the glacial deposits of the Lake Michigan lobe prior to 14 ka BP. Quaternary Science Reviews 5, 17-22. Smith, D.B., Cannon, W.F., Woodruff, L.G., Solano, Federico, Kilburn, James E., and Fey, David L., 2013. Geochemical and mineralogical data for soils of the conterminous United States. U.S. Geological Survey Data Series 801, 19 p. Wright, H.E., 1972. Quaternary history of Minnesota, in Sims, P.K., and Morey, G.B., eds., Geology of Minnesota: a centennial volume. Minnesota Geological Survey, 515-547. 128 �THE EVOLUTION OF THE ATMOSPHERE-HYDROSPHERE: A GEOCHEMICAL COMPARISON OF TWO PALEOPROTEROZIC GUNFLINT WEATHERING PROFILES YIP, Christopher and FRALICK, Philip, Department of Geology, Lakehead University, Thunder Bay, ON, P7B 5E1, cyip@lakeheadu.ca, philip.fralick@lakeheadu.ca The 1878±1 Ma year old Gunflint Iron Formation is a chemical sedimentary unit that forms one of the members of the Animike Group. It is well known that the Gunflint Formation is made up of a transgressive-regressive-transgressive sequence, which represent the advance and retreat of the ancient sea that filled the Animike basin. This sequence traps a unique point in the evolution of the atmospherehydrosphere during the Precambrian. This interaction with the atmosphere should be seen in the rock record at the point where the water depth was at its shallowest, such as the initial transgression or the regressive surfaces. If oxygen was present, the rocks underlying the transgression, as well as the initial transgressive strata that were precipitated in the shallow ocean should contain geochemical markers such as Ce anomalies. The newly created two-lane Highway 11/17 outside of Thunder Bay, shows a clean example of the basal section of the Gunflint Formation (Figure 1a). This section of the Gunflint overlies an Archean granodiorite unit. The overlying Gunflint carbonate grainstones show no unique features. What is important is the underlying granodiorite unit. The granodiorite shows a very clear example of spheroidal weathering that should occur if joints in the bedrock were the site of intense chemical weathering. Samples were collected starting from the base of the granodiorite where it is the freshest and least weathered up through the increasingly altered portions of the weathering profile into the Gunflint grainstone. Samples were prepared and analysed for major oxides as well as trace and rare earth elements. Another prime example of the basal section of the Gunflint Formation can be seen at Schreiber Beach outside of Schreiber ON, which was studied by Polat et al. (2012) (Figure 1b). This area differs from the 11/17 site in that the Gunflint strata sits above a Neoarchean pillow basalt sequence that shows the well preserved basaltic pillows overlain by hyaloclastites made up of shattered pillow breccias and flows. Directly below the contact with the Gunflint, the pillow sequence has been weathered and exfoliated creating red to brown highly fractured pillows topped off by brown to green pillow basalt soils. A geochemical comparison of these two sample locations was performed. When plotted on Nesbitt (2003) A-CN-K feldspar diagrams the 11/17 outcrop shows an enrichment in Al2O3 (Figure 2a), whereas the ACN-K diagram plotted by Polat et al. (2012) shows that the weathered layer is enriched in K2O and Al2O3 (Figure 2b). The difference in the parent material of the two weathered profiles and possibly potassic metasomatism in the basaltic material is controlling these weathering trends. The spheroidal weathering granodiorite also has an intense enrichment in Fe, Mn and Mg, probably the result of interactions with Gunflint derived fluids, which overprinted the effects of weathering. This period of alteration by Gunflint fluids also resulted in intense leaching of rare earth elements. 129 �A B Figure 1.a) the outcrop on the side of Highway 11/17, showing the weathering profile starting with the granodiorite and working up through the weathered section eventually capped by Gunflint grainstone. B) The stratigraphic sequence of the Schreiber beach outcrop modified from Polat et al., (2012). B A Figure 2. a) A-CN-K diagram for the data collected from the Highway 11/17 outcrop outside of Thunder Bay, ON. The fresh granodiorite samples plot in the middle and the weathered samples plot at the top showing CaO, Na2O and K2O depletion. B) The A-CN-K diagram from Polat et al.(2012) showing the enrichment of K from the unweathered pillows to the weathered brown to green basalts. References: Nesbitt, H.W., 2003, Petrogenesis of silicalstic sediments and sedimentary rocks, in Lentz, D.R., ed., Geochemistry of sediments and sedimentary rocks: Evolutionary Considerations to Mineral Deposit-Forming Environments: Geological association of Canada, GeoText4, p. 39-51 Polat, A., 2012, Extreme element mobility during transformation of Neoarchean (ca. 2.7 Ga) pillow basalts to a Paleoproterozic (ca. 1.9 Ga) paleosol, Schreiber Beach, Ontario, Canada. Chemical Geology, 326-327, 145173. 130 � https://digitalcollections.lakeheadu.ca/files/original/cfaebe60172260fedcb902ba9425270e.pdf a55faff9c466cceda748bca918036e0e PDF Text Text �INSTITUTE ON LAKE SUPERIOR GEOLOGY 60TH ANNUAL MEETING May 14-17, 2014 Hibbing, Minnesota Sponsored by PRECAMBRIAN RESEARCH CENTER, UNIVERSITY OF MINNESOTA DULUTH, and MINNESOTA GEOLOGICAL SURVEY James D. Miller and Mark A. Jirsa Co-Chairs Proceedings Volume 60 Part 2 – Field Trip Guidebook Edited by Mark Jirsa, Terrence Boerboom, and Amy Radakovich, Minnesota Geological Survey Cover Photos Historic photo of steam shovel was acquired from the archives of Minnesota Discovery Center Research Library; all other photos were taken by Jirsa. Photos vaguely represent the subject matter of some field trips described in this guidebook (relevant trip numbers are shown in parentheses). Photos in large “6” include conjugate faults (1) in oxidized Biwabik Iron Formation at Susquehanna Mine; glacial till in Albany Mine (6); taconite pellets (3); folded Soudan Iron Formation (2); gray-green stromatolites photographed from polished sample in collection of Dan England, Eveleth Fee Office (1); and jointed, oxidized Biwabik Iron Formation from Glenn Mine (E). Photos in large “0” include taconite-bearing drill core (A); red stromatolites from Dan England collection (1); pillowed metabasalt from near Gilbert (7); and historic photo of steam shovel, location and date unknown (B, C). i �Generalized geologic map showing locations of field trips in this guidebook Note that no Trip 4 is shown, as it was cancelled Geology simplified from Minnesota Geological Survey Map S-21 statewide bedrock geology ii �TABLE OF CONTENTS PROCEEDINGS VOLUME 60 PART 2— FIELD TRIP GUIDEBOOK Pre-meeting field trips, Wednesday, May 14 TRIP 1: STRATIGRAPHY, SEDIMENTOLOGY, STRUCTURE, AND MINERALIZATION OF THE BIWABIK IRON FORMATION, CENTRAL MESABI IRON RANGE ............................................ 1 TRIP 2: A WALK IN THE PARK – NEOARCHEAN GEOLOGY OF LAKE VERMILION STATE PARK .............................. 37 TRIP 3: WESTERN MESABI RANGE MINING OPERATIONS .................................................. 76 TRIP 4: LAURENTIAN VISION RECLAMATION Cancelled Post-meeting field trips, Saturday, May 17 TRIP 5: VISIONS OF MATURI: THE GEOLOGY OF THE SOUTH KAWISHIWI INTRUSION 86 TRIP 6: THE ST. LOUIS SUBLOBE AND GLACIAL LAKE UPHAM 102 TRIP 7: GEOLOGY AND GOLD MINERALIZATION OF THE VIRGINIA HORN AREA ………….119 Syn-meeting field trips, Friday afternoon, May 16 TRIP A: STATE DRILL CORE LIBRARY – HIBBING, MINNESOTA ...................................... 137 TRIP B: TRIP C: TRIP D: TRIP E: HIBBING’S IRON MINING AND CULTURAL HISTORY ........................................... 140 MINNESOTA DISCOVERY CENTER ..................................................................... 146 COLERAINE MINERALS RESEARCH LABORATORY .............................................. 147 MINEVIEW FROM A CANOE................................................................................ 148 The editors extend sincere thanks to all who contributed to this field trip guidebook. The time and effort expended to prepare field trip descriptions are greatly appreciated. Special thanks to Minnesota Coaches Voyageur Bus Company in Duluth for substantially discounting transport costs, and to Greyhound Bus Museum in Hibbing for providing a vintage bus for field trip B. Reference to material in Part 2 should follow the example below: Field trip authors, date, title: Institute on Lake Superior Geology Proceedings v. 60, Part 2, p. XX. Proceedings Volume 60, Part 1—Program and Abstracts, and Part 2—Field Trip Guidebook are published by the 60th Institute on Lake Superior Geology and distributed by the Institute Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca Some figures in this field trip guidebook were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the guide when it is available on-line at http://www.lakesuperiorgeology.org. ISSN 1042-9964 iii ��FIELD TRIP 1 Wednesday, May 14, 2014 STRATIGRAPHY, SEDIMENTOLOGY, STRUCTURE AND MINERALIZATION OF THE BIWABIK IRON FORMATION, CENTRAL MESABI IRON RANGE LEADERS: Phil Larson (Duluth Metals) Marsha Patelke (Natural Resources Research Institute – Duluth) Jakob Wartman (United Taconite, Cliffs Natural Resources) Michael Totenhagen (Arcelor Mittal) Mark Jirsa (Minnesota Geological Survey) Steven Losh (Minnesota State University – Mankato) Peter K. Jongewaard (Cliffs Natural Resources, recently retired) Figure 1. Bedrock geologic map of the central Mesabi Iron Range showing 3 main field trip localities. Stops 1-4 are located on United Taconite’s Thunderbird North and South Mines near Eveleth; stop 5 is the Mary Ellen Mine near Biwabik. Archean metavolcanic, metasedimentary, and granitic rocks are shades of green, blue, and pink, respectively; Paleoproterozoic Pokegama Quartzite is yellow; Biwabik Iron Formation is red; and Virginia Formation is gray (modified from Jirsa and others, 1998). 1 �INTRODUCTION This field trip explores the geology of the Paleoproterozoic Biwabik Iron Formation (BIF) in the central Mesabi Iron Range of northeastern Minnesota. The formation hosts iron-ore deposits that have been mined continuously for nearly 125 years, constituting the most economically significant mining district in the United States. This trip will visit exposures in 3 localities (Fig. 1): the Cliffs Natural Resources - United Taconite LLC Thunderbird North Mine in Eveleth, an active magnetite taconite mine; the inactive satellite Thunderbird South Mine; and the Mary Ellen Mine near Biwabik, a closed directshipping (hematite) ore mine. Because this field trip guide was compiled by a large number of co-leaders having different experiences and perspectives, it may contain some content that is repetitive or has minor inconsistencies. The iron-bearing strata of the Biwabik Iron Formation were first noted in 1866 by Henry Eames, on what was to become the eastern end of the Mesabi Iron Range. Sporadic exploration for iron ore deposits began soon after, notably by Peter Mitchell and the Ontonagon Syndicate. Lack of infrastructure hampered exploration and development until 1884, when the Duluth and Iron Range Railroad to the Soudan Mine on the Vermilion Range was completed. The resulting exploration boom initially focused on the well-exposed, metamorphosed iron-formation near the contact with the Duluth Complex. It was not until 1890 that the Merritt Brothers re-directed their exploration focus to an area farther SE along the Mesabi range, correctly surmising that iron ore float located south of the Giants Range had been transported some distance by glaciation. Their discovery of soft, friable, high-grade iron ore at the future Mountain Iron Mine on November 16, 1890 revolutionized the global iron ore and steel industries. The first shipment of 4,245 tons in 1892 came at the crest of a wave of exploration that was to quickly discover and develop billions of tons of iron ore along the 175 km strike-length of the Biwabik Iron Formation. The Mesabi Iron Range rapidly developed into the largest iron mining district in the United States, a status it continues to hold. For decades after its discovery, it was the largest iron ore district on earth, accounting for nearly half of global production in the late 1940s. The Biwabik Iron Formation ranges from 0.5 to 5.0 km width (0.25 to 3 miles) along a strike length of 175 km (100 miles) (Fig. 2). The formation, as much as 220 meters (750 feet) thick, generally dips gently to the southeast at angles of about 7° to 15°. Unweathered, unaltered iron-formation, colloquially known as “taconite,” contains about 30 percent iron and 45 percent silica, with the balance (2-10%) composed principally of MgO, CaO, and MnO. In numerous places along the length of the range, typically along joints, fractures, folds, and other structurally prepared zones, silica and other elements were leached under tropical weathering conditions, locally enriching the iron content to as much as 64 percent. So-called “natural” or direct-shipping ores dominated production through the Second World War. Depletion of higher grade reserves, combined with increasingly stringent quality requirements, led to a rapid conversion of the industry during the 1950s. Perfection of fine grinding, magnetic separation, and pelletizing technology (the “taconite process”) allowed for economic exploitation of pristine ironformation. Between 1956 and 1977, eight taconite facilities with a combined capacity in excess of 54 million tons per year (mtpy) were brought to production. Production of beneficiated iron ores (gravity and taconite concentrates) exceeded that of direct-shipping ores in 1958, and by 1967, taconite concentrates accounted for over half of production. At this writing, six taconite (40 mtpy capacity), and three tailings recovery (3 mtpy capacity) facilities are in production (Fig. 3). 2 �Figure 2. Location map of the Mesabi Iron Range (maroon). Note the Duluth Complex (Keweenawan, 1.1 Ga) on the east side). Figure 3. A) Aerial distribution of taconite pits and cities. B) A longitudinal section of the Biwabik Iron Formation showing: average thickness of the iron-formation at each taconite operation (along with the thickness of the various members at each operation), and mined taconite intervals (as black columns adjacent to the sections). From Severson and others (2009). 3 �REGIONAL GEOLOGY Basement rocks in the central Mesabi area consist of 2.7 Ga Archean granitic and greenstone terrains of the Superior craton. These were intruded by the 2.1 Ga Kenora-Kabetogama dike swarm, but were apparently eroded to a peneplained surface by about 1.9 Ga. The peneplained Superior craton formed the platform upon which a Paleoproterozoic continental shelf and margin assemblage was deposited. 1.88 Ga iron-formation and associated clastic sediments are preserved along 3000 km of strike length near the margin of the craton, and were once much more extensive. Available evidence indicates iron-formation accumulation occurred at all margins of the craton simultaneously, and is a reflection of global ocean chemical conditions rather than local basin geometry. Most of the iron-formation is preserved in mobile belts at the craton margin, and are thus variably deformed and metamorphosed (Trommald Iron Formation of east-central Minnesota, Gogebic Iron Formation of Wisconsin-Michigan, Negaunee Iron Formation of Michigan, Sokoman Iron Formation of Labrador and Quebec, Cape Smith belt of Quebec, Kipalu Iron Formation of Hudson Bay, Sutton Inlier iron-formation of Ontario). Relatively flat-lying, undeformed, and un- or weakly-metamorphosed iron-formation is locally preserved inboard of the craton margin (Temiscamie Iron Formation of Quebec, the Gunflint Iron Formation of Minnesota-Ontario, and the Biwabik Iron Formation). These were likely deposited on a clastic-starved platform, and indicate that an epeiric sea may have nearly or completely inundated the peneplained craton during the peak of ironformation accumulation. Paleoproterozoic sedimentation and iron-formation accumulation followed a general sequence of depositional events throughout the Superior craton. Nearshore, tidally influenced clastic sedimentation was succeeded by chemically-precipitated iron-formation. The transition from clastic to iron-formation is typically abrupt; however the presence of iron-rich, chert-cemented epiclastic strata indicate that iron and silica precipitation was occurring prior to significant accumulation of the epiclastic-poor iron-formation. Significant iron-formation accumulation was apparently triggered by a lack of epiclastic input, rather than abrupt onset of favorable iron-precipitating geochemical conditions. Iron-formation accumulation proceeded at very low rates until resumption of clastic deposition in the basin. Available evidence indicate that the Biwabik Iron Formation records accumulation over as many as 15 million years within a significantly larger time span (Larson, 2013). Iron-formation accumulation across the craton was likely diachronous within this time span, as significant internal disconformities are evident within individual iron-formations, and correlations between individual iron-formations are problematic. Iron-formation accumulation proceeded until resumption of clastic deposition in the basin; similar to the basal epiclastic strata, significant amounts strata containing iron-rich precipitates are found in the overlying epiclastic units. In the Animikie Basin, the basal, near-shore, epiclastic sequence is represented by the Pokegama Formation, which was succeeded by deposition of the Biwabik Iron Formation. In excess of 200m of iron-formation accumulation was abruptly terminated at the contact with the overlying argillites, siltstones, and greywackes of the Virginia Formation. Within the regional Paleoproterozoic depositional system, this contact is also marked by chaotic (paleoseismic) deformation and deposition of ejecta related to the 1.85 Ga Sudbury meteorite impact event. (Jirsa and others, 2011). In the central Mesabi area, a conglomeratic bed containing angular argillite, chert, and carbonate clasts within the Upper Slaty member Dolomite/Limestone unit (submember US-2; see Stratigraphy section below) (Severson and others, 2009) has been correlated with the Sudbury meteorite impact event (Addison et al., 2005). Turbiditic sediment of the overlying Virginia Formation was ultimately sourced from 1.87-1.83 Ga volcanic rocks, and represents the collision of an island arc with the southern margin of the Superior craton during the Penokean orogen. Collision and sedimentation at the continental margin led to development of a foreland basin and the thick turbidite sequence. Paleoproterozoic sedimentary rocks (including the Animikie Group) along the southern margin of the Superior craton were bisected by the 1.1 Ga Midcontinent Rift, a 2000km (1200 mile)-long rift system extending in an arcuate fashion from northeastern Kansas to southeastern Michigan. The rift consists predominantly of mafic flows and intrusions overlain by rift-fill sedimentary strata. In addition, numerous 4 �mafic dikes, chonoliths, and other small intrusions were emplaced locally into rocks of the Animikie Group rocks at significant distances from the rift axis. These include mafic dikes cross-cutting the Biwabik Iron Formation near Keewatin, and a series of sills emplaced into iron-formation in the vicinity of Aurora. However, no such intrusions are known from the Central Mesabi area. Significant thermal metamorphism of iron-formation is limited to the area generally east of Aurora, within the aureole of the Duluth Complex. By the end of the Cretaceous, peneplaination produced topography similar to that of modern day. The central Mesabi area lay close to the eastern extent of the Cretaceous Interior Seaway, and the Biwabik Iron Formation locally is overlain by near-shore shale and sandstone. A basal iron-ore-bearing conglomerate is locally present, indicating extensive formation of supergene-enriched, direct-shipping ores predated Cretaceous sedimentation. The roughly coeval formation of secondary iron oxides and hydroxides implies that supergene enrichment may have occurred under a tropical climate during the Cretaceous (Symons, 1966; Purucker, 1973). During this and subsequent weathering, it is likely that the iron-formation served as a geochemically resistant “cap” that protected the underlying Giants Range Batholith, and is indirectly responsible for formation of the Missabe Wachu (“Big Man Hills” from the classic David D. Owen’s 1852 Report of a Geological Survey of Wisconsin, Iowa, and Minnesota). The area is now known physiographically as the Giants Range. Ice sheets advanced across the Biwabik Iron Formation multiple times during the course of the Pleistocene, primarily during the last 2 million years. Unconsolidated saprolite, including the supergene enriched direct-shipping ore (DSO), was preferentially eroded, leaving only remnants in deep, structurehosted, trough-shaped bodies and stratiform layers variably protected by resistant cap rocks. Locally, large blocks of weathered iron-formation and even DSO were eroded by glaciotectonic activity. A variable thickness of till, outwash, and glaciofluvial sediment was deposited over the iron-formation during the final glacial cycle (See related Field Trip in this guidebook). Archean Rocks The Neoarchean bedrock of the central Mesabi Iron Range lies near the southern edge of the Wawa subprovince of the Superior Province, and constitutes the southwestern-most exposures of the terrane. The Archean supracrustal rocks on the Mesabi Range are separated from the well-known Vermilion district to the north by the Giants Range batholith, a large, composite body consisting of granitoid rocks of several generations and compositions. The Archean rocks are covered to the south, east, and west by Paleoproterozoic strata, including iron-formation of the Mesabi Iron Range. The Archean supracrustal rocks are subdivided into northern and southern panels on the basis of contrasting metamorphic grade and deformation style (Fig. 4). The northern panel, adjacent to the Giants Range batholith, contains intensely lineated, amphibolite-grade schist of volcanic, intrusive, and clastic protolith. The southern panel contains a similar stratigraphic sequence, but has minerals that indicate it underwent metamorphism to much lower grades, ranging from prehnite-pumpellyite to low greenschist facies. The two panels are separated by the east-trending, post-metamorphic, Laurentian fault. Amphibolite-grade rocks of the northern panel (north of the Laurentian fault) comprise the Minntac sequence that contains locally strongly layered schists having geochemical and outcrop-scale characteristics of volcanic, intrusive, and turbiditic protoliths. The lower grade strata within the southern panel are subdivided into the Mud Lake and Midway sequences. The Mud Lake sequence forms a broad, southwest plunging syncline (the Mud Lake syncline) defined by outer limbs of calc-alkalic and tholeiitic strata, and cored by graywacke, slate, and minor felsic tuff. The Mud Lake strata are unconformably overlain by, and locally lie in fault contact with, fluvial and alluvial conglomerate, subaerially deposited trachyandesitic flows, and pyroclastic rocks that comprise the Midway sequence. 5 �Figure 4. Geologic map of the central Mesabi range area, illustrating features of the Archean bedrock (From Jirsa and Boerboom, 2003). The Z-shaped fold/fault structure apparent in the strike of iron-formation is known locally as the “Virginia horn.” 6 �Paleoproterozoic Animikie Group The Paleoproterozoic Animikie Group unconformably overlies the Mille Lacs Group to the south in central Minnesota, and the Archean basement on the Mesabi Range to the north (Southwick and Morey, 1991). The Animikie Group consists of three major formations on both the Mesabi and equivalent Gunflint Iron Ranges. The respective units are the Pokegama Formation and the Kakabeka Quartzite (the lowest units), the Biwabik and Gunflint Iron Formations (the middle units) and the Virginia and Rove Formations (the upper units, composed of graywacke and shale). The Thomson Formation in the northern part of east-central Minnesota is inferred to be correlative with the Virginia and Rove Formations. The Biwabik and Gunflint Iron Formations are on strike with each other and were probably continuous prior to intrusion of the Duluth Complex at about 1.1 Ga. Age The age of the Animikie Group is relatively poorly constrained in the central Mesabi area due to a paucity of datable cross-cutting or intercalated units. A minimum age of deposition for the Pokegama Formation is 1,930 ± 25 Ma (Pb/Pb), which was obtained from quartz veins that cut the Pokegama Formation (Hemming and others, 1990). An age of 1,878 ± 2 Ma (U/Pb on euhedral zircons) has been obtained from an ash layer in the upper Gunflint Iron Formation (Fralick and others, 2002); this horizon may correlate with the Intermediate Slate of the Lower Slaty Member of the Biwabik Iron Formation (LS-1 submember). The ejecta layer correlated with the 1,850 Ma Sudbury impact event dates the stratigraphic top of iron-formation (Addison and others, 2005; Jirsa 2010). Zircon ages from an ash layer at the very base of the overlying Virginia Formation are dated at 1,832 ± 3 Ma (Addison and others, 2005). The latter sample was collected a few inches above the base of the Virginia Formation in drill hole VHD-00-1, located immediately to the west of the Thunderbird Mine. Vallini and others (2007) dated a metamorphic xenotime overgrowth on detrital zircon from the Pokegama Formation at 1763 ± 14 Ma, attributing this to a ~1786 Ma regional basin-wide, subtle thermal pulse. Stratigraphy In the central Mesabi area, the Animikie Group is composed of three formally defined formations: the Pokegama Formation, Biwabik Iron Formation, and Virginia Formation; and an informally named unit of breccia and ejecta related to the Sudbury meteorite impact event. Pokegama Formation The Pokegama Formation consists of up to 300’+ of shale, siltstone, and chert-cemented quartz arenite (quartzite). The formation consists mostly of siltstone and shale. Silica-cemented quartz arenite is confined to an interval a few meters thick beneath the overlying Biwabik Iron Formation. Ojakangas (1983) used gross stratigraphic relationships to subdivide the Pokegama Quartzite into three informal members: a basal member consisting dominantly of thinly bedded to laminated shale and lesser amounts of siltstone; a middle member consisting of shale and siltstone and scattered thin beds of quartz arenite; and an upper member consisting mostly of quartz arenite. The formation was deposited on an irregular Archean bedrock surface. A basal conglomerate contains a poorly sorted array of clasts derived from the underlying bedrock set in a matrix of fine-grained sandstone to siltstone. Basal strata are marked by a second conglomerate that has angular to sub-rounded, pebble- and granule-size clasts of chert, jasper, algal fragments, and vein quartz. This implies that Pokegama-like clastic sedimentation and Biwabik-like chemical precipitation were, for a time, contemporaneous (Jirsa and Morey, 2003). At most localities, the contact between the Pokegama Quartzite and the overlying Biwabik Iron Formation is conformable and gradational; the presence of beds of chert 6 to 12 meters beneath the Biwabik–Pokegama contact has been cited as evidence for a gradational sedimentary regime between the two formations (Ojakangas, 1983). Biwabik Iron Formation The Biwabik Iron Formation ranges from about 175-300 feet thick at the extreme eastern end of the Mesabi Iron Range (Dunka Pit) (Bonnichsen, 1968), to 730-780 feet thick in the Central Mesabi area near 7 �Eveleth, decreasing to around 500 feet thick on the western Mesabi Iron Range near Coleraine, and eventually exhibits a “nebulous ending about 15 miles southwest of Grand Rapids” (Marsden and others, 1968). The formation is subdivided into four informal members (from bottom to top): Lower Cherty, Lower Slaty, Upper Cherty, and Upper Slaty (Wolff, 1917). Individual beds can be described as either sand-textured granular iron-formation (gif), composed predominantly of rounded oolitic grains and intraclasts, or mud-textured and laminated banded iron-formation (bif). Although interlayering of these two lithotypes occurs on all scales, the “cherty” members are composed predominantly of medium- to thick-bedded gif; the ‘slaty’ members are composed predominantly of thin-bedded bif. The terms “slaty” and “cherty” were originally used by miners, and are not indicative of metamorphism or slaty cleavage, or a predominance of silica as chert. The cherty gif members are largely composed of iron oxides and chertcemented granules of iron silicates and carbonates. The slaty members are generally composed of laminated iron silicates and iron carbonates. The slaty bif members are envisioned to have been deposited below storm wave base. Granules comprising the cherty gif members formed in high-energy environments. Two models have been applied to explain formation of granules in the Paleoproterozoic iron-formations: direct precipitation, and reworking of intraclasts shoreward during storm events where they are reworked into granules in shallower water. Some granules appear to be the product of reworking of laminated bif material, however the self-similarity of granule sizes, lack of apparent intraclast source material, and geochemical dissimilarity between gif and bif material suggest the gif formed due to direct precipitation as oolites in shallow water. A few diagnostic marker units within the formation allow basin-scale correlation. Two stromatolitebearing intervals several feet thick are present, one at the base of the Lower Cherty member and the other in the middle of the Upper Cherty member (UC-6 submember). For submember terminology, see Detailed Stratigraphy of Biwabik Iron Formation discussion below. The black Intermediate Slate (LS-1 submember) at the base of the Lower Slaty member reportedly contains ash-fall tuff, with up to 5.5% Al2O3 (Morey, 1992). The top of the Upper Slaty member (US-2 submember) contains several feet of limestone and dolomite. Most of these marker units, which are prominent in the eastern and central parts of the range, pinch out in the vicinity of Hibbing, about 60 miles from the west end of the range (Severson and others, 2009). The Lower Slaty member is not present at the far western end of the range. Sudbury Impact Layer The contact between iron-formation and overlying slate of the Virginia Formation is marked by a thin layer of deformed substrate and overlying ejecta formed during the 1850 Ma Sudbury meteorite impact event. The horizon can only be seen in drill core on the Mesabi Iron Range; however, it is well exposed in the equivalent Gunflint Formation to the northeast. There, the deformed layer consists of ≤ 7m of chaotically folded and locally brecciated substrate iron-formation. At least some of the iron-formation beds were ductily deformed, implying they were not yet fully lithified at the time of deformation. The rheologic behavior of siliceous layers was more brittle, and they were dislodged and shattered. In the context of a major meteorite impact event, the deformation is inferred to have occurred by impact-induced seismicity (Jirsa and others, 2011). The deformed strata are draped by ejecta (≤ 1m-thick) containing abundant petrographic evidence of impact origin, including the presence of zoned spherules and quartz fragments displaying multiple planar deformation features. The impact-related horizon, known as the Sudbury Impact Layer, is well exposed in several locations in the Lake Superior region (Fig. 5). It occurs in the Gunflint Lake area of northeast Minnesota (op. cit.), in the Thunder Bay area of adjacent Ontario (Addison and others, 2005), and in Michigan (Cannon and others, 2010; Pufahl and other, 2007). At exposures near Thunder Bay and Gunflint Lake, the contact between the impact layer and the overlying Rove Formation (Virginia equivalent) is inferred to be a disconformity that may represent a significant depositional or erosional hiatus, perhaps as long as 15-40 million years (Jirsa and others, 2011). Whether this is also true for the contact on the Mesabi range is currently unclear. It is noteworthy that the depositional environment in which the Virginia and Rove Formations were deposited is nearly identical with that of the underlying iron-formation—the primary difference being a paucity of chemical precipitates (iron and silica) in the former. 8 �Figure 5. Generalized correlation diagram of Paleoproterozoic strata in the Lake Superior region. The position of the Sudbury Impact Layer is from Cannon and Addison (2007). Virginia Formation The Virginia Formation overlies the Biwabik Iron Formation and Sudbury Impact Layer on the south side of the Mesabi Iron Range, and is inferred to extend southward beneath glacial cover for an unknown distance. Presumably, it reappears in east-central Minnesota as a folded and metamorphosed sequence called the Thomson Formation. The Virginia Formation is a thick turbidite sequence composed of interbedded argillite, graywacke, and volcaniclastic rocks (Fig. 5). The 450 meters of Virginia Formation strata penetrated at a site south of Biwabik have been described in considerable detail in Lucente and Morey (1983). The lower part of the formation is composed almost entirely of alternating beds of darkgray, silty mudstone and black carbonaceous shale. Quartz-rich siltstone and very fine- to fine-grained lithic graywacke become more abundant stratigraphically higher. The basal part of the Virginia Formation contains several beds of coarse-grained feldspathic graywacke and volcaniclastic rocks, as well as many lenses and irregular beds of limestone and dolomite. Dolomite-rich concretions of various sizes and shapes, also characterize the lower several hundred meters of the formation. Sandstone beds become coarser-grained and more abundant up-section, and are composed of angular quartz and feldspar grains in a matrix of muscovite and chlorite. Much of the matrix consists of diagenetically altered lithic fragments. Hemming and others (1995) documented that shales of the Virginia Formation have Nd-depleted, mantle model ages of 2.35 to 2.14 Ga. Interbedded volcaniclastic sediments have younger model ages of 1.99 to 1.86 Ga. Craddock and others (2013) showed a dominant spectrum of Penokean orogeny ages (1.85-1.8 Ga) for detrital zircon populations from the correlative Thompson and Rove Formations. These data indicate Virginia Formation sediment was sourced from a comparatively young, differentiated volcanic arc, most likely from the Wisconsin Magmatic Terrane and equivalent rocks in Minnesota to the south. Depositional Environments The Animikie Group records a sedimentalogical transition from nearshore, tidally influenced, allochthonous clastic deposition, through shallow, autochthonous, chemical platform deposition, interrupted by impact-dominated deformation and deposition, and followed by deep water basinal turbidite sedimentation (Fig. 6). The Pokegama Formation is interpreted to have been deposited in a tidally influenced, shallow marine setting near the shoreline, having received clastic detritus from the Archean basement to the north (Ojakangas, 1983; Craddock and others, 2013). In the central Mesabi area, the Pokegama Formation consists of a lower (argillaceous) member, a middle member of intercalated argillaceous and silty sedimentary strata, and an upper member of quartz sandstone. This succession is interpreted to represent a 9 �transition from upper tidal flat to lower tidal flat/subtidal depositional environments (Ojakangas, 1983). Elsewhere along its strike length, the Pokegama contains pebble conglomerate that may represent a transgressive lag. The transition from near-shore, clastic-dominated sedimentation to autochthonous chemical sedimentation is recorded by an abrupt gradation into iron-formation. The abrupt decrease in clastic input is consistent with non-accretionary transgression across the peneplained craton, whereby relatively small rises in eustatic sea level translate into dramatic shifts in shoreline position. In general, iron-formation can be geochemically divided into two components: an autochthonous chemically precipitated component, and an allochthonous clastic component. The elements comprising the autochthonous component (Fe, Si, Mg, Ca, Mn, P) are precipitated directly from seawater, while the allochthonous component (Al, Ti, K, Na) is derived from terrestrial sources (dust or suspended sediment). In general, the allochthonous component of the BIF averages 2.5%. The close correlation between lithofacies and mineralogy indicates that iron-formation mineralogy is a sensitive recorder of redox conditions at the sediment-water interface and, by extension, the depositional environment. The Lower Cherty Member (LC) of the Biwabik Iron Formation is interpreted to have been deposited on a shallow marine shelf. The LC is dominated by granular iron-formation (gif), characterized by rounded oolitic grains, cross-bedding, and other features indicative of a high-energy environment. Accumulation was driven nearly entirely by autochthonous chemical precipitation, with very little epiclastic input. Ferric hydroxide apparently co-precipitated with carbonates, evidenced by an abundance of ankerite. The common presence of stylolites indicates significant volume loss of both iron- and carbonate minerals by chemical dissolution during compaction. The presence of herringbone crossstratification in the lowermost LC indicates deposition in a tidally-influenced environment. The presence of thin laminae of slaty banded iron-formation interbedded with medium-bedded gif farther up-section within the LC suggests a transition to a deeper shelf environment. There, steady deposition of mudtextured banded iron-formation (bif) below the fair-weather wave base was periodically punctuated by rapid accumulation of gif, perhaps by shelf-to-basin transport during storm events. The upper portion of the LC contains a thick-bedded, coarse-textured (with intraformational conglomerate), pink (oxidized) gif overlain by a medium-bedded, sand-textured, green (reduced) gif (LC6/LC-7 contact) with abundant interbedded bif layers. This major lithostratigraphic break represents a significant disconformity within the overall BIF sequence. Trace element data show an abrupt transition from a high Al:Ti clastic source for LC strata below the break, to a low Al:Ti clastic source for all BIF strata above the break, indicating a change in sediment provenance. The LC sequence above the LC6-LC-7 disconformity records a transition to deep water pelagic bif sedimentation, culminating in the Lower Slaty (LS) member Intermediate Slate (LS-1) submember. The Intermediate Slate/LS-1 is composed of iron silicates and sulfides in a laminated bif. The relatively large clastic content indicates a decrease in the rate of autochthonous chemical sediment accumulation, driven in part by redissolution of ferric hydroxide precipitates in anoxic bottom water. The LS to Upper Cherty (UC) (LS-2 to UC-4) sequence is dominated by laminated bif, punctuated by lenticular or channel-shaped bodies of gif (Interbedded Cherts or IBCs). The bifs record deepwater sedimentation; the gifs consist of sand-textured granules that were exported from a high-energy environment and deposited in bypass channels on the shelf (channels) or deepwater fans (lenses). The upper portion of the UC (submembers UC-5 to UC-7) is characterized by gifs deposited in a high-energy environment, as evidenced by abundant rounded oolitic grains, stromatolites, and cross bedding. Mineralogically, these submembers are characterized by the presence of significant amounts of primary hematite relative to the Biwabik Iron Formation as a whole, and the Algal Chert Horizon/UC-6 submember is significantly enriched in Mn, indicating deposition in oxic conditions. Combined, these features suggest deposition in shallow, tidally-influenced or subtidal environment. The high-energy environment, combined with a paucity of clastic sediment input, suggest this may have been analogous to a modern carbonate platform environment. Chauvel and Dimroth (1974), noted similar features (chiefly oolitic and intraclastic sand) in the Sokoman Iron Formation, and their corresponding textural similarity to sediments in modern carbonate environments. They applied a carbonate facies model to the iron-formation, and attributed gif deposition 10 �to a lagoonal platform depositional environment, characterized by rapidly shifting banks of oolitic and intraclastic sand, mud banks, and small, lagoonal basins ringed by oolite shoals. Sommers and others (2000) concluded that ooids and stromatolites from the Lower Algal Chert member of the Gunflint Formation (stratigraphic equivalent of the UC-6) were originally deposited as carbonates and later replaced by silica. Rankey and Reeders (2011) described how the interaction of wind, waves, and tides with channel morphology on platforms to create “Spin Cycle” currents which control formation, transport, grain size, sorting, and deposition of chemically precipitated oolites. Such currents are a viable mechanism for keeping precipitating granules frequently in suspension until they reached a critical size; this in turn provides an explanation for the remarkable self-similarity in granule size observed in gifs within the BIF (Patelke, in progress). The sequence from UC to Upper Slaty (UC-8 to US-2) is dominated by laminated bifs, indicating a return to a deeper-water shelf environment. The US-2 is a limestone of enigmatic origin. The BIF is overlain by argillite and graywacke of the Virginia Formation, deposited by northwardadvancing basinal turbidite sedimentation. Some volcanic ash evidently settled into the basin, locally forming graded beds with a totally volcanic composition. The dominance of black, fissile shale suggests the "raining out" of clay and deposition in deep, anoxic water below the wave base. Minor, thin, sandstone lenses were deposited by bottom currents (Lucente and Morey, 1983) . Figure 6: Possible siliciclastic environments for deposition of portions of the Biwabik Iron Formation in cross-sectional view (facies division in a shallow sea). Compiled, and drawn, by Marsha Patelke (in Severson and others, 2009) from sources including Boggs (2001), Leeder (1999), Nichols (1999), and Pratt and others (1992). 11 �Mesoproterozoic Despite being regionally extensive, no known Mesoproterozoic intrusive rocks are known in the Central Mesabi area. Similarly, the entire area of this field trip lies beyond the extent of thermal metamorphism of the Biwabik Iron Formation by the Duluth Complex. Cretaceous Cretaceous rocks are thought to be more or less continuous beneath glacial drift throughout the western half of the State, and form numerous outliers in the eastern half. On the Mesabi Range, Cretaceous rocks have been exposed by mining, and recognized in drill core and well cuttings (Fig. 7). The Cretaceous Coleraine Formation in the Mesabi district comprises iron-ore conglomerate, shale, and sandstone that form a thin irregular mantle over an uneven surface on the underlying bedrock (Sloan, 1964). The rocks are dominantly of marine origin in the west, but grade eastward in the central Mesabi to continental (fluvial and deltaic) sediments. In the western part of the district, fine conglomerate grades upward into a ferruginous grit and sandstone, including bluish-green shale in the vicinity of the Hill Annex mine near Calumet. The rocks grade eastward into continental sediments, including a widespread basal conglomerate composed of fragments of iron ore. The Cretaceous strata are virtually horizontal except locally, where they are interpreted to have slumped or to have compacted differentially over depressions in the underlying bedrock (Owens, 1956). Depressions and slumps were common topographic expressions of supergene-enriched natural iron ores, enabling the preservation of Cretaceous remnants post-glaciation. In at least two notable central Mesabi locations, the Coleraine Formation conglomerates were mined as ore. At the Judson Mine near Buhl, eroded iron ore was transported short distances southward, and conglomerates were deposited in northwestward-trending channels cut in the underlying bedrock (Everett, 1956). Owens (1956) describes a 60-ft section of iron ore conglomerate, red and white shales, and lignite in the Enterprise Mine northeast of Virginia. The occurrence of Cretaceous iron-ore conglomerates have been confirmed just east of Gilbert (Sloan, 1964). Figure 7. Panoramic view of the Enterprise Mine near Virginia; slump area in foreground and light-colored unit in background are Cretaceous sandstone (Owens, 1956). 12 �Quaternary The central Mesabi area was glaciated repeatedly during the course of the Pleistocene. Unconsolidated saprolite, including a significant amount of supergene enriched direct-shipping ore, was preferentially eroded, leaving only remnants in deep, structure-hosted trough-shaped bodies and in stratiform layers variably protected by resistant cap rock. Final ice retreat occurred about 13 C 14 thousand years before present, when the margin of the Rainy Lobe retreated north of the Giant’s Range, depositing a sandy-textured, boulder till in its wake. The St. Louis Sublobe advanced in a surge across the glacial lake to the south, reaching the toe of the Giant’s Range, depositing a silty, boulder-poor till. A later glacial lake capped the low lying areas with silty glaciolacustrine sediment. Glaciotectonism played a significant role in glacial erosion of direct-shipping ores. Large blocks of loose, porous oxidized and weathered iron-formation were frozen en masse onto the toe of the glacier, and thrust into the debris load. In the vicinity of the Fayal Mine in Eveleth, stripping operations in advance of a ‘milling’ mine encountered a block of direct-shipping ore entirely encased by glacial drift. A similar occurrence can be seen in the pit wall of the Glen Mine near Chisholm (Field Trip E, this guide book). The so-called Moose Track Mine produced in excess of 30,000 tons of ore, indicating the block contained over 10,000 m3 of material (Leith, 1903). BIWABIK IRON FORMATION: STRATIGRAPHY, STRUCTURE, MINERALOGY, AND ORE DEPOSITS Stratigraphy of the Biwabik Iron Formation The four-fold stratigraphy of Lower Cherty, Lower Slaty, Upper Cherty, and Upper Slaty members (Wolff, 1917) is still used at each of the currently operating (and inactive) iron mines on the Mesabi Iron Range. However, each of the mining companies further subdivides the Biwabik Iron Formation into several submembers based on lithostratigraphic units and mineralogical assemblages observed in drill core and mine exposures (Fig. 8). District-wide correlation of individual mine stratigraphy is problematic because lithostratigraphic or mineralogically defined units important at the mine scale are not necessarily extensive at the district scale. Severson and others (2009) defined and correlated 25 laterally extensive submembers within the BIF based on examination of more than 380 drill holes along 75 miles strike length. Units were named for their characteristic bedding types. Definition of these 25 “Rosetta” units serve as a starting point for more detailed sequence stratigraphic studies and basin analysis (Fig. 9). There clearly are variations in some of the units and the four main iron-formation members along strike that are related to facies changes. Figure 8. Textural characteristics of the Biwabik Iron Formation (from Severson and others, 2009 – modified from Pfleider and others, 1968). Dark bands represent magnetite-rich rock. 13 �Figure 9. Graphic summary of the 25 major “Rosetta” units in the Biwabik Iron Formation that were identified and described in Severson and others (2009). Most of these units have corresponding submember designations at each of the taconite mines. 14 �In addition to the member subdivisions recognized in the BIF throughout the Mesabi Iron Range, Thunderbird Mine geological staff further subdivide the iron-formation into 23 submembers based on lithologic, metallurgical, and mineralogical characteristics. These submembers form the basis for both resource estimation and grade control at the mine, and are described in detail below. For reference, the corresponding “Rosetta” unit name of Severson and others (2009) is listed in parentheses after the Thunderbird Mine submember name. The drilled thickness of the Biwabik Iron Formation in the vicinity of the Thunderbird North Deposit is approximately 686’ (Table 1). Member Thickness Upper Slaty 51 feet (15.5 meters) Upper Cherty 347 feet (121 meters) Lower Slaty 52 feet (16 meters) Lower Cherty 236 feet (72 meters) Table 1. Average thicknesses of the four members, as recognized at Thunderbird North Lower Cherty The Lower Cherty member is approximately 236 feet thick in the Thunderbird North deposit. It has been subdivided into the following eight subunits: LC-1 (Basal Red Unit) The submember is a pink-green-gray heterogeneous unit comprised of interbedded thin- bedded slaty and thin-bedded cherty carbonate-silicate(minnesotaite-talc-stilpnomelane) iron-formation. LC-1 comprises the basal 64 feet of the iron-formation. It is defined as the footwall thickness of the iron formation, the magnetite grade of which is subeconomic. It is in general poorly described since the majority of exploration and development drilling terminates in the upper few feet of this unit. LC-2 (Regular-Bedded Unit) The submember is a gray thin-bedded cherty carbonate silicate(minnesotaite-talc) iron-formation. Magnetite occurs as disseminated and diffuse idiomorphic granules and as replacement of thin slaty laminae and early burial stylolites. Magnetite (slaty) laminae often have thin stringers of white talc. LC-2 averages 16 feet in thickness, but varies across the extent of the Thunderbird North Deposit, being thinner in the southwest extent of the deposit and thicker in the northeast extent. In the northeast portion of the deposit, the unit is of sufficient thickness and grade to warrant mining despite dilution by the overlying LC-3 waste unit. LC-3 (Regular-Bedded Unit) Rocks of the submember are characterized by interbedded greenish-gray thin-bedded cherty and green medium-laminated slaty iron-formation. The unit is weakly magnetic, with the cherty beds conspicuously low in magnetite. The unit averages 13 feet thick, but varies across the deposit. In the southwestern extent of the deposit, the unit is up to 30 feet thick and predominantly composed of slaty iron formation. In the northeastern extent of the deposit, the unit is consistently 10 feet thick and composed predominantly of thin bedded non-magnetic granular chert. LC-4 (Wavy-Bedded Unit) The submember is composed of gray medium-bedded cherty oxide-carbonate(ankerite)silicate(minnesotaite-talc) iron-formation with minor thin irregular thin beds of slaty (magnetite) iron formation. Magnetite occurs as disseminated idiomorphic granules, patchy haloes cored by coarse slaty intraclasts, and replacement of thin slaty laminae. LC-4 varies from 40-50 feet thick at Thunderbird North, thickening to the southwest. Notable features of LC-4 are magnetite haloes or reaction rims around 15 �small intraclasts, and wispy laminae of magnetite, likely a later diagenetic overprint of early burial stylolites. The LC- 4 and its equivalents are widespread across the Mesabi District, and perhaps the most economically significant subunit, having a high weight recovery (36%) and being capable of producing a low silica concentrate (~2.0%). LC-5 (Wavy-Bedded Unit) The submember is composed of pink-gray medium- to thick-bedded cherty oxide-chertcarbonate(ankerite/kutnahorite) iron-formation. Magnetite occurs as disseminated grains and in mottles. The unit is notable for its high carbonate content, containing up to 3.0% CaO in ankerite. LC-5 varies from 40-50 feet thick at Thunderbird North, thickening to the southwest. LC-5 contains a small but variable amount of ‘primary’ (e.g. pre-supergene oxidation) hematite. LC-5 has appreciably more matrix chert than the underlying LC-4, and produces a significantly higher silica concentrate (~6.0%). LC-6 (Variably-Bedded and/or Mottled Unit) The submember averages seven feet thick, and consists of a pink massive- to thick-bedded cherty oxide-chert-carbonate(kutnahorite) iron-formation with conspicuous pink carbonate mottles. The unit is composed principally of coarse grained intraclasts, reflecting a relatively high energy depositional environment. The unit also contains an appreciable content of “primary” hematite, and has relatively low magnetite recovery. The unit is remarkably tough, and poses a challenge to mining in that it resists fragmentation during blasting and tends to produce large chunks. LC-7 (Bold Striped Unit) The submember is composed of interbedded thick irregular magnetite-carbonate-silicate slaty and green thin- to medium-bedded cherty carbonate(siderite)-silicate(greenalite) iron-formation. The unit averages 13 feet thick. The unit is remarkable in that magnetite occurs predominantly in the thick slaty laminae, resulting in a boldly-striped appearance in drill core. Green LC-7 sharply overlies the pink LC-6, and the contact is a highly visible stratigraphic marker throughout the Virginia Horn area. The transition upward from thick-bedded coarse-grained to thin-bedded fine-grained iron-formation, as well as the contrasting mineralogical assemblages at the LC-6/LC-7 contact, suggests an abrupt transition in the depositional environment. Very fine-grained magnetite (25-45 µm) and intimate association with very fine-grained chert and siderite contribute to grinding and processing difficulties with the LC-7. LC-8 (Mesabi Select Unit) The submember is visually similar to LC-7, consisting of a interbedded green medium- to thicklaminar massive slaty and greenish-gray thin-bedded granular cherty carbonate(siderite)silicate(greenalite) iron-formation. However, the unit contains little or no magnetite and is a waste product that makes excellent aggregate material. The LC-8 averages 21 feet in thickness. LC-8 from the Thunderbird North Mine is the type material for “Mesabi Select” crushed aggregate currently being marketed regionally for road construction, noted for its high specific gravity and angular fragmentation. Lower Slaty The Lower Slaty member, as defined at Thunderbird North, averages 52 feet thick and is characterized by non-magnetic and thin-bedded waste rock between the Lower Cherty and Upper Cherty member ore horizons. Other interpretations (Severson and others, 2009) of the Lower Slaty in the Virginia Horn area extend it to up to the top of the UC-4 submember, and up to 309 feet thick. LS-1 (Intermediate Slate) The submember is composed of predominantly black massive- to thinly-laminated slaty carbonate(siderite)- silicate (stilpnomelane-minnesotaite)-sulfide iron formation. The LS-1 averages 17 feet in total thickness, and is divisible into a lower half composed of a composed of thick bedded massive intraformational debris flow breccias and an upper half composed of thinly-laminated planar- bedded slaty iron-formation. Locally, thin to medium bedded black flinty cherts are present in the lower portion; these flinty cherts typically occur in pod-like bodies extending a few 100s of feet along strike. The upper portion of LS-1 has undergone extensive bedding-parallel deformation; essentially the entire horizon served as a low-angle fault plane. Small-scale folds are common, as are bedding-parallel syntectonic 16 �quartz-carbonate(ankerite-siderite) veins. The thinly-laminated planar-bedded slaty iron-formation in the upper portion is the so-called Intermediate Slate, a district-scale marker horizon. LS-1 is notable in that it contains a very high percentage of Al2O3 (1.86%) and other elements indicative of clastic input, suggesting the basin was experiencing either an influx of clastic detritus, or a sharp reduction in the rate of iron-formation deposition. LS-2 (Lowermost Thin-Bedded Unit) The submember is composed of a green to greenish-gray well-cemented very thinly-laminated slaty carbonate-silicate(minnesotaite) iron-formation. The unit averages 35 feet thick. The top of the LS-2 is defined by the appearance of significant magnetic slaty iron-formation. Commonly, this corresponds to the first appearance of thin-bedded intraclast breccias. These breccias commonly have a magnetite-rich matrix. Upper Cherty The Upper Cherty member at Thunderbird North contains all potential ore horizons situated above the Lower Slaty waste horizons. This comprises 347 feet in thickness and the remainder of the thickness of the iron-formation exposed in the present workings. The lowermost 257 feet of the Upper Cherty, as defined at Thunderbird North, consists of alternating horizons of dominantly slaty- and cherty-ironformation; these horizons have been included in the Lower Slaty by other workers (Severson and others, 2009) including US Steel. The Upper Cherty has been subdivided into eleven subunits at Thunderbird North. LUC-1 (Ore Zone of Lowermost Thin-Bedded Unit) The submember is composed of gray laminar thin-bedded slaty chert-silicate(stilpnomelane) magnetite iron- formation. The unit averages 18 feet in thickness and is notable for producing a very high silica concentrate (up to ~10% SiO2). This unit, in common with the other slaty iron-formation horizons in the Upper Cherty member, has a relatively high Al 2O3 content (~0.56%). LUC-2 (Lower IBC Unit) The submember is a heterogeneous unit, composed variously of green-gray thin-bedded slaty ironformation, interbedded green-gray thin-bedded slaty iron-formation, thin-bedded cherty iron- formation, and gray thick-bedded cherty iron-formation. The unit as a whole averages 46 feet thick. For the unit as a whole, thin-bedded granular cherty horizons predominate over thin- to medium-laminated shales. The abundance and frequency of cherty horizons generally increases up-section within the unit. Locally, pink to green-grey massive- to thick-bedded, coarse-grained, magnetite- bearing granular cherts, upwards of 20 feet thick, are present within the unit. These beds are characterized by significantly higher weight recovery, and significantly lower concentrate silica grades than the unit as a whole. LUC-3 ( Middle Thin-Bedded Unit) The submember is composed of dark reddish-brown thin/planar-bedded slaty chert-silicate ironformation. The unit averages 19 feet thick. Nodules and beds of chert are increasingly abundant upsection, culminating in the presence of a 1-2 foot thick horizon containing thin bedded flinty chert, an important marker horizon in the mine. UC-1 (Middle IBC Unit) The submember is composed of pinkish-gray thick-bedded cherty oxide-chert-silicate iron-formation. The unit averages 29 feet thick; however, it is not extensive through the deposit. The overall aspect of UC-1 is of a lenticular body on the order of several km in extent. UC-1 is notable in that it contains an appreciable content of ‘primary’ hematite; this hematite is intimately intergrown with magnetite, and thus, is recovered in the Fairlane concentrator circuit. UC-2 ( Middle Thin-Bedded Unit) The submember is a dark reddish-brown thin-bedded slaty chert-silicate iron-formation, averaging 46 feet in thickness. UC-2 is generally characterized by low weight recovery and a high concentrate silica, and thus, is marginal ore at best. 17 �UC-3 and UC-3A (Upper IBC Unit) The submembers are gray thick-bedded cherty iron-formation. Combined, they average 91 feet in thickness. Similar to UC-1, these units are not laterally extensive, and have the overall aspect of lenticular bodies on the order of several km in extent. The two units comprise a single depositional package; however, the lower half (UC-3) is generally characterized by low weight recovery, while the upper half (UC-3A) is characterized by higher carbonate and magnetite content. UC-3A was historically one of the primary ore units at the Thunderbird North Mine; however, it is not now being mined and is only poorly exposed in the pit. UC-3A is notable for an abundance of coarse-grained jasper intraclasts; the vivid colors of these intraclasts have resulted in the name ‘confetti ore’ being attached to UC-3A. UC-4 (Uppermost Thin-Bedded Unit) The submember is a dark reddish-brown thin-bedded slaty silicate iron-formation, averaging 18 feet in thickness. In areas where UC-1, UC-3 and UC-3A are absent, UC-4 is the upward continuation of UC2. The unit typically has a very low weight recovery. The uppermost 1-5 feet of the subunit is commonly a black, thin-bedded non-magnetic slaty silicate iron-formation. This has been recognized as an important marker horizon, and for some workers (Fig. 9), marks the top of the Lower Slaty member. UC-5 (Alternating-Bedded Unit) The submember consists of interbedded reddish-brown thin-bedded slaty silicate iron-formation and thin-bedded cherty iron-formation. The unit averages 15 feet thick. The thin cherty beds commonly contain abundant coarse-grained jasper intraclasts. UC-6 (Algal/Conglomerate Unit) The submember is very distinct in that it is composed of red medium- to thick-bedded coarse-grained intraclast conglomerates. Clasts in the conglomerate are composed predominantly of resedimented cherty algal stromatolites (oncolites). The conglomeratic matrix is commonly composed predominantly of manganiferous carbonates., including rhodochrosite (Zeilinski and others, 1994). UC-6 is notable in having the highest manganese content in the Biwabik Iron Formation, averaging ~6.0% Mn (7.8% MnO). UC-7 (Regular/Medium-Bedded Unit) The submember is composed of gray to red thick-bedded oolitic cherty oxide-chert-carbonate ironformation. The unit averages 37 feet thick in the Thunderbird North Mine area, and is known mostly from oxidized drill hole intercepts. The unit appears to consist of a lower red hematitic oolitic cherty ironformation and an upper gray magnetite-bearing oolitic chert- carbonate (ankerite) cherty iron-formation. The upper gray horizon contains abundant coarse poikiloblasts of ankerite; commonly these are weathered away, leaving vesicle-like vugs in the oolitic cherts. UC-8 (Thin-bedded unit) Similar to UC-5, the UC-8 consists of interbedded green-red thin-bedded slaty silicate iron- formation and thin-bedded cherty iron-formation. The unit averages 28 feet thick. UC-8 is known only from (commonly) oxidized drill hole intercepts. The thin cherty beds commonly contain abundant coarsegrained jasper intraclasts. The contact between UC-8 and the overlying US-1 is poorly defined. Upper Slaty The Upper Slaty member in the vicinity of the Thunderbird North Mine is only known from oxidized intercepts in a few drill holes, and is not exposed in outcrop. US-1 The submember is comprised predominantly of reddish-brown thin-bedded slaty iron-formation, and is about 50 feet thick. US-2 The submember is not exposed in outcrop or mine workings in the Thunderbird Mine, nor has it been intercepted in drilling. Regional drilling data indicates the unit is composed of grey thin-bedded micritic calcareous carbonate, and is about 19 feet thick. 18 �Structure of the Biwabik Iron Formation Recent studies of bedrock structure along the Mesabi Iron Range (Jirsa and others, 1998; 2002; 2005a; 2005b; Jirsa, 2006) reveal that a protracted history of deformation affected the Biwabik Iron Formation. Much of the formation forms a south-dipping homocline that contains little evidence of tectonic disruption, with the exception of locally well-developed deformation structures. A general sequence of deformation events can be inferred from those localized structures. The precise ages of events on the Mesabi range are unknown; however, a relative chronology for various structural elements can be established from cross-cutting relationships. Assigning deformation events to specific structures is speculative; nevertheless, the "D 0, D1, D2…" nomenclature is applied here to refer to suites of apparently related structures. The oldest are those presumably related to soft-sediment deformation (D0), including slumps, sedimentary breccias, and structures that appear to be the result of differential compaction and localized faulting synchronous with deposition. The earliest "regional" deformation (D 1) is manifest in localized, small-scale rotational structures, bedding-parallel slickensides, and larger nappe and sheath folds. The structures commonly lie along boundaries between units having strong rheologic contrast, such as the contacts between packages dominated by mudstone vs. those composed of siliceous, intraclastic grainstone. Nearly all of these structures display a sense of asymmetry that indicates southover-north tectonism. This northward vergence, and the apparent timing relative to later structures, is consistent with compressional deformation—potentially related to the Penokean orogen. One of the longstanding controversies in iron-ore genesis is the question of whether oxidation and leaching of ironformation to form the high-grade hematite ores occurred by supergene or hypogene processes. Although not conclusive, the observation of several early-formed, south-dipping thrust faults with folded, mineralized wall rocks, and bedding-parallel slickensides that host abundant secondary iron and silica implies that at least some of the mineralization was coincident with compressional deformation, perhaps during Penokean orogenesis. This is consistent with the hypogene model proposed by Morey (1999) that attributes oxidation and leaching to groundwater flow driven northward from uplift in the Penokean fold and thrust belt. A second regional suite of structures (D2) is largely extensional. These structures are monoclines and normal faults that are mutually transgressive; that is, faults that have sympathetically folded wall rocks, and folds that pass gradationally into faults along the trend of axial planes. These are some of the major structures along which oxidation and leaching has occurred, and the focus of most direct-shipping (hematite) ore mining. Veins, vugs, and other secondary mineralization features are abundant. D2 structures likely formed as localized responses to regional tilting. The most recent deformation effects (D3) are trough-like collapse structures, presumably related to post-leaching subsidence. The collapse, and associated oxidation and weathering, are best developed in the uppermost subcrop of iron-formation, implying supergene alteration played a significant role in their development. Thus, the answer to the supergene vs. hypogene debate appears to be that both processes were significant, perhaps at different times. Lacking finite ages, the structures can only be inferred to record components of Penokean (Geon 18), Yavapai (Geon 17), Mazatzal (Geon 16), and/or Keweenawan (Geon 11) deformation events. The overall shallowly-dipping, northeast-striking homoclinal trend of Paleoproterozoic strata along the Mesabi Iron Range is interrupted near the city of Virginia, where strata are warped around an apparent anticline-syncline pair to form the structure known locally as the Virginia horn. In this area, dips as steep as 25° occur, and strikes trend N, NE, and NW. The origin of this structure has been variously ascribed to faulting and folding associated with uplift of Archean bedrock that now cores the anticlinal portion of the Z-shaped horn structure. Intuitively, the granitoid basement rocks were too competent to accommodate ductile compression, and it is therefore unlikely that the horn formed by simple flexural folding. The Alpena fault and several others are marked by differences in the thickness of internal units across them, indicating some deformation was synchronous with deposition (i.e., growth faults). Several faults are essentially continuous along strike with those in the Archean basement. Where displacement sense or magnitude differs significantly between these faults in the two ages of bedrock, the portions affecting Paleoproterozoic strata are inferred to have been reactivated along faults of Archean parentage. The conceptual model shown below (Fig. 10) depicts an interpretation of the structural development in the 19 �horn that invokes some combination of faulting and folding, and addresses the reactivation of what were likely Archean faults reactivated during the Paleoproterozoic. The development of direct shipping (hematite-goethite) ores appears to have been localized to varying degrees by fault and fold structures within iron-formation. Regionally, natural ore bodies extend from the bedrock surface to depths as great as 120 m. These ores formed by oxidation, hydration, and subsequent Figure 10. Schematic model showing structural evolution of the Virginia horn structure (from Morey, 2003. 20 �leaching by through-flowing solutions after lithification. Bedding plane fractures, folds, and faults presumably acted as hydraulic conduits and traps for descending and/or ascending solutions that selectively altered certain lithologic units. The direction of fluid movement and the possibility of multiple episodes of alteration are currently unclear, but work to more fully understand these questions is underway (e.g., Losh and Rague, 2013; see Diagenesis, Alteration, and Fluid Flow discussion below). Zones of oxidation along structures are apparent in derivative aeromagnetic imagery (Fig. 11). Linear zones of less magnetic, presumably oxidized iron-formation typically cross the strike of iron-formation at varied angles. Where it can be verified on the ground, many of these zones are coincident with mapped faults, axial planes of minor folds, and major joint networks. Some are also coincident with mined natural ore bodies; though to be clear, most magnetic lows depicted do not represent mineable deposits of natural ore. However, they do represent local zones having variably decreased overall magnetite content. Figure 11. Derivative aeromagnetic map of the central Mesabi Range. Image was created from total field magnetic data, which was regridded from flight lines and band-pass filtered to remove broad wave-length (low frequency) anomalies and reveal contrasts in the short wave-length (near-bedrock surface) anomalies. Magnetic highs are light; lows are dark. In the north-trending limb of the Virginia horn structure, most of the linear magnetic lows that cross the strike of the overall high correspond with folds and faults that have been mapped within iron-formation (White, 1954; and field work by the authors). North-south striping is an artifact of gridding flight line data. 21 �Origin of Iron-Formation Iron-formation formed by chemical precipitation of dissolved ferrous (Fe2+) iron as a solid phase, most likely a ferric (Fe3+) bearing species. A reduced or low-oxygen atmosphere relative to modern conditions was necessary to allow accumulation of high concentrations of dissolved ferrous iron in seawater. Mineralogical and geochemical evidence indicates co-precipitation of variable amounts of Mg, Ca, Mn, P, Si, and CO2 in addition to Fe. Silica precipitation may have occurred by adsorption onto ferric iron species settling through the water column (Fischer and Knoll, 2009), by diagenetic reaction with ferric iron precipitates at the sediment-water interface, or direct precipitation on the seafloor (Maliva and others, 2005). It is likely that all of these mechanisms may have played a role. Geochemistry of Iron Deposition A number of mechanisms have been proposed to explain iron precipitation and deposition, including direct oxidation as a byproduct of oxygenic photosynthesis, anoxygenic photosynthesis utilizing iron as an electron acceptor, and abiotic photochemical oxidation. Regardless of mechanism, the reactions can be generalized as: Fe2+ + O2 + H2O → Fe(OH)3 + H+ (1), or 2+ 4Fe + 11H2O + CO2 → 4Fe(OH)3 + CH2O + 8H+ (2) In each of these models, iron oxidation is placed close to the surface within the photic zone. The photosynthetic models intimately associate iron precipitation with biological activity, and presume that iron precipitates are raining out of the water column along with organic material. Each of these models also presumes a reservoir of dissolved iron in anoxic waters lying beneath a chemocline. In the case of shallow water iron precipitation, this implies a current- or tidal-driven flux of anoxic bottom waters into shallower water environs. Fixation of Ferric Iron into Ferrous Iron species Ferric iron hydroxide (Fe(OH)3) precipitates formed at or near the top of the water column and settled to the bottom. Hydrous ferric iron oxides have not been recognized as primary minerals in iron-formation. All iron-bearing minerals in iron-formation may have been produced by diagenetic reactions at or near the sediment-water interface. Basic iron-fixing reactions include: 2Fe(OH)3 → Fe2O3 + 3H2O (Hematite) (3) Fe(OH)3 +CO2 + H+ → FeCO3 + H2O (Siderite) (4) 3Fe(OH)3 + 2SiO2(aq) + 3H+ → Fe3Si2O5(OH)4 + 4H2O (Greenalite) (5) Fe(OH)3 + 2HS- → FeS2 + 3H2O (Pyrite) (6) With the exception of hematite, all these iron-bearing minerals contain ferrous (Fe2+) iron, indicating diagenetic iron-fixation was accompanied by iron reduction. Iron reduction was driven by a combination of settling of precipitates through the chemocline into anoxic bottom water, or respiration of organic material at the sediment-water interface, or a combination of both processes. Absent carbonate, silica, or sulfur with which to react and form a stable mineral species, the transformation of insoluble ferric iron precipitate to highly soluble ferrous iron would return dissolved iron back into the water column. Formation of geologically stable iron-formation is not a function of deposition of ferric iron precipitates, rather that of fixation of ferric iron into stable, dominantly ferrous iron mineral species. It is more appropriate to speak in terms of iron-formation accumulation and accumulation rates than ironformation deposition and deposition rates. Viewed in this context, apparent decreases in “deposition” rate are actually decreases in accumulation rates, and may reflect a lack of suitable fixative at the sedimentwater interface rather than a decrease in iron precipitation rates at the top of the water column. Iron-Formation Mineralogy The iron-bearing minerals in iron-formation consist of oxides, carbonates, silicates, and sulfides. James (1954) recognized that the iron mineralogy varied systematically, and reflected distinct lithostratigraphic facies, at least in part. His iron-formation facies concept (oxide, silicate, carbonate, and sulfide facies) continues to provide a compelling framework within which to interpret iron-formation sedimentology and mineralogy. The diversity of iron minerals found in iron-formation (Table 2) is a 22 �direct reflection of the diversity of the sedimentological and geochemical environments in which the ironformation formed. The likelihood of extensive replacement of primary iron precipitates has resulted in significant controversy regarding the precise nature of the primary precipitate, and the precise reaction pathways responsible for formation of the observed mineral assemblages (Simonson, 2003). Eh and pH are major controls on the stability of the iron minerals in both the depositional and diagenetic environments (Ojakangas and others, 2005). Klein (2005) has suggested that the original precipitate materials were probably hydrous Fe-silicate gels of a greenalite-type composition; Na-, K- and Al-containing gels approximating stilpnomelane compositions; SiO2 gels; Fe(OH)2 and Fe(OH)3 precipitates; and very finegrained carbonate oozes. A variety of other primary chemical precipitates for iron-formation in general have also been postulated by an assortment of authors and include siderite, iron hydroxides, iron silicates (Konhauser and others, 2002; Rajan and others, 1996), and colloidal iron silicates (Lascelles, 2007). “Clastic” components such as Al2O3, TiO2, K2O, and Na2O were likely deposited as eolian dust, and reflect a far-travel clay mineral component eroded from exposed cratons. Nevertheless, within the ironformation these elements are typically found in the iron silicate mineral stilpnomelane, suggesting that even the clastic component participated in diagenetic reactions. Mineral Oxides Magnetite Hematite Goethite Silicates Chert Chalcedony Microcrystalline Quartz Stilpnomelane Minnesotaite Talc Greenalite Chamosite (Al-rich Fe-chlorite) Carbonates Siderite Ankerite Kutnohorite - Ferroan Dolomite Kutnohorite Calcite Sulfides Pyrite Pyrrhotite Formula Fe3O4 Fe2O3 FeO(OH) SiO2 SiO2 SiO2 K(Mg, Fe+2, Fe+3)8(Si,Al)12(O,OH)27 Fe3Si4O10(OH)2 Mg3Si4O10(OH)2 Fe3Si2O5(OH)4 Fe3(Al,Si)2O5(OH)4 FeCO3 Ca(Fe,Mg)(CO3)2 (Ca,Mn)(CO3)2 - Ca(Mn,Mg,Fe)(CO3)2 CaMg(CO3)2 CaCO3 FeS2 Fe(1-x)S Table 2. Common mineral names and formulas associated with the Biwabik Iron Formation (excluding the more highly metamorphosed eastern Mesabi Iron Range in proximity to the Duluth Complex). 23 �Oxides Hematite is the iron-bearing mineral most commonly associated with iron-formation. However, primary hematite is a relatively rare component of the BIF, occurring most prominently in oxide facies iron-formation of the UC member. Magnetite is common throughout the BIF sequence. The relatively coarse-grained idiomorphic magnetite characteristic of gifs are late diagenetic in origin (LaBerge, 1964; LaBerge and others, 1987; Zanko and others, 2003) and form by the replacement of pre-existing iron silicates and iron carbonates (French, 1973). Fine-grained magnetite in ‘slaty’ bif layers likewise formed by diagenetic reaction of iron silicates and carbonates. Earlier work on the oxidized taconites of the western Mesabi Iron Range was accomplished by Bleifuss (1964). He showed that late hematite was developed by the oxidation and pseudomorphic replacement of magnetite octahedra, that layers of goethite were precipitated from solutions likely derived from the oxidation of siderite, and that some goethite formed by the oxidation of acicular iron silicate minerals. Silicates Greenalite is considered to most closely reflect the composition of an initial ferric hydroxide/silica gel precipitate in that it exhibits no detectable replacement of any pre-existing phase (French, 1973; LaBerge and others, 1987; Simonson, 1987; Klein, 2005). Within gifs, greenalite most often occurs as roundshaped granules that are <1 mm in diameter. Stilpnomelane is a secondary mineral that commonly replaces early iron silicates (greenalite) (French, 1973). The presence of alumina and potassium suggests reaction with the detrital dust component found in the iron-formation. French (1973) suggests that stilpnomelane formed under conditions ranging from diagenesis to low-grade metamorphism. Minnesotaite is a common component of throughout the BIF sequence, and the type locality for this mineral was located in the north end of the Thunderbird North mine. Stoichiometrically analogous to magnesian talc, but structurally dissimilar, it generally occurs as sheaves or needles replacing greenalite granules (French, 1973). True talc, including ferroan talc, locally comprises a significant amount of the silicate fraction within gifs. McSwiggen and Morey (2008) show that both chamosite and talc are common throughout portions of the Biwabik Iron Formation. Carbonates Ankerite and siderite are common early diagenetic minerals. Siderite commonly occurs within laminated bifs, while ankerite commonly occurs as idiomorphic replacement of primary granules within gifs. Locally, coarse-grained poikiloblastic aggregates of ankerite (mottles) are found in gifs. These mottles are clearly late diagenetic replacements. McSwiggen and Morey (2008) report manganese substitution for iron in the dolomite-ankerite series, leading to kutnohorite and ferroan kutnohorite composition in the Lower Cherty. Both Mg and Mn substitute for Fe in siderite in the Lower Cherty, with some samples containing as much as 20 to 25 mole percent MnCO3. The average composition of siderite from the Lower Cherty [(Fe.60-.78Mg.15-.19Ca.02-.03Mn.04-22)CO3] differs considerably from that of the remainder of the iron-formation, where siderite compositions averages [(Fe.88-.84Mg.03-.11Ca.02-.03 Mn.02.07)CO3] (McSwiggen and Morey, 2008). Sulfides Sulfide minerals are ubiquitous throughout the BIF sequence. Pyrite and pyrrhotite are the most common, with minor amounts of arsenopyrite, cobaltite and galena (Theriault, 2011). Pyrite occurs ubiquitously as in trace amounts as idiomorphic cubic or dodecahedral crystals, framboids, and spheroids. Less commonly, sulfide occurs in larger blebs. The Intermediate Slate (LS-1) contains the greatest abundance of sulfide, apparently associated with elevated mercury and arsenic concentrations (Morey and Lively, 1999). 24 �Alteration and Regional Fluid Flow Common features within the Biwabik Iron Formation are quartz-carbonate±iron silicate veins that occupy vertical fractures and bedding-parallel slip planes. These veins postdate magnetite formation and diagenesis of the iron-formation, but display textures indicative of syntectonic growth, suggesting they may be related to far-field deformation, perhaps associated with the Penokean orogen. In the vicinity of direct-shipping ore deposits, these veins are commonly overprinted by: 1. complete or partial dissolution of carbonate minerals; 2. brecciation of the quartz, perhaps associated with volume loss collapse of the iron formation; and 3. recementation by secondary iron oxides and silica. Recent fieldwork in the Hibbing Taconite, Thunderbird North, and Thunderbird South/Fayal Mines, combined with petrographic, SEM, fluid inclusion, and geochemical techniques, have elucidated oxidation by deep, saline, hydrothermal-diagenetic waters at relatively low water/rock ratios (e.g., Losh and Rague 2013). Fluid inclusions in fault breccia and low-angle and high-angle veins containing secondary minerals (quartz, calcite, minnesotaite, stilpnomelane, hematite) have average homogenization temperature of 155°C ± 17°C (n=278), and salinity of 9.5 ± 5.3 wt% NaCl equivalent (n=160). Temperature correction due to pressure is on the order of 50°C. There is no significant difference in fluid inclusion homogenization temperature or salinity between fault breccias (including quartz cement) and veins of diagenetic affinity. These results agree well with oxygen isotopic temperatures of 150°–200°C for diagenesis determined by Perry and others (1973). The oxidizing fluids, a mixture of diagenetic and meteoric fluids, infiltrated along high-angle faults that contain vein quartz cemented by quartz ± iron oxides (typically goethite), and brought about oxidation of magnetite to hematite and Fe-silicates to goethite (the latter reaction also yielding silica as a reaction product), accompanied by quartz recrystallization. Silica liberated from this oxidation filled microfractures, typically only a few microns in width, and pits within altered magnetite grains. This contributes to the general observation of high silica in magnetite concentrates from oxidized ores. As silica was not dissolved but rather only remobilized during this oxidation event, ore was not significantly upgraded; in fact, the introduction of quartz into microfractures in magnetite locally diminished magnetic taconite ore quality. Quartz-filled microfractures in magnetite are also observed in unweathered ‘slaty’ iron formation near bedding-parallel faults, as in the Thunderbird North mine. Ore Deposits The Biwabik Iron Formation contains about 30% iron throughout its thickness, irrespective of lithofacies, mineralogy, and grain size. However, only a fraction of the formation hosts recoverable iron mineralization. Historically, two deposits types have proven economically feasible to mine: directshipping (natural) ores (DSO) and magnetic taconite. Direct-shipping ores are composed of hematite and goethite enriched by supergene leaching of silica from the pristine iron-formation. Early in the development of the Mesabi Range, DSO were shipped directly from the mine. In later years, gravity concentration was used to upgrade the iron content of the ores. DSOs formed from leaching of any lithofacies ranging from thin-bedded slaty band iron-formation to thick-bedded, coarse-grained granular iron-formation. Because of this, DSO quality was quite variable in terms of deleterious elements, including phosphorus, alumina, manganese, and structural water (from goethite). This necessitated an elaborate and extensive system of ore grading and blending of railcar size shipments at rail yards and ore docks at the shipping ports to maintain consistent quality blast furnace feed. Magnetic taconite ores are composed predominantly of coarser grained magnetite found in granular iron-formation. These rocks are capable of producing a high-grade magnetite concentrate after fine grinding by magnetic separation. The magnetic concentrates are agglomerated and fired into the 3/8”-1/2” taconite pellets familiar to would-be slingshot assassins throughout the Great Lakes region. The resulting pellets have a significantly higher iron grade, significantly lower deleterious element content, and superior smelting properties relative to the DSO production they have replaced. Furthermore, the product is easily transportable. 25 �Direct-shipping Ores The direct-shipping ores of the Mesabi Iron Range fall under the Soft Iron Ore category of Marsden and others (1968). They are generally porous masses of hematite, goethite, and minor magnetite and manganese oxides. Gangue minerals consist of quartz, clay, and minor carbonate. Fundamentally, soft ore formation is the product of preferential leaching of silicate and carbonate components from the ironformation, and alteration of the primary iron oxide, silicate, and carbonate minerals to secondary hematite and hydrous iron oxides (Marsden and others, 1968). The iron content of the iron-formation is increased by loss of gangue material (primarily silica and carbonate), rather than enrichment or replacement by supergene iron minerals. On the Mesabi Iron Range, soft ore bodies occur in trough, fissure, and irregular ore bodies, reflecting variable degrees of ore formation along faults, folds, or zones of fracturing. They also occur as stratiform ore bodies, reflecting ore formation along favorable horizons. Generally, soft ore bodies extend from the bedrock surface to depths of 400 or 500 feet (Marsden and others, 1968). Ore formation was evidently a multi-stage process. Early desilicification of the iron-formation was accompanied by alteration of primary magnetite to hematite, and alteration of primary iron silicates and iron carbonates to goethite. Much direct-shipping ore exhibits textures indicative of a second stage of enrichment. Secondary porosity induced during desilicification is commonly filled by paragenetically late iron hydroxides and hydrous iron oxides. Dripstone textures indicate that at least some of these secondary iron hydroxides were precipitated in the vadose zone. Leith (1903) noted that hydrous minerals were more abundant in the shallower portions of the deposits, suggesting the presence of a supergene enrichment zone, perhaps coincident with a paleo-water table. Ore formation and desilicification were accompanied by mass loss (as much as 50% by weight) and, to a variable extent, volume loss. Unaltered iron-formation has a specific gravity of 3.3-3.4; Leith (1903) reported typical direct-shipping ore specific gravities in the range of 2.6-3.1, with some ores as low as 2.0-2.1. Mass loss was typically accompanied by structural collapse and formation of a synclinal structure in the ore body (D 3 deformation; see discussion above in Structure of the Biwabik Iron Formation). Commonly, the ore retained bedding and geochemical traits inherited from the precursor iron-formation, with the steepest dips adjacent to the margins of the deposits. The nature and timing of ore formation has been the subject of much debate. The clear association of many deposits with fault and fracture zones, as well as the sharp wall contacts, has been cited as evidence in favor of a hydrothermal origin (Morey, 1999). In contrast, the clear association of the stratiform bodies with the paleosurface argues strongly in favor of a supergene origin. The complex paragenesis of the ores suggests that multiple events may ultimately have been responsible for development of the ores. Oxidation of iron-formation in the Mesabi Range has long been thought to have been solely the result of near-surface interaction with meteoric water, most intensely during saprolite formation during the Cretaceous (Leith, 1903; Sloan, 1964; Marsden and others, 1968), or during other time periods with a tropical climate. Gruner (1956) and Morey (1999) proposed that at least some if not all of the intense oxidation was of hydrothermal origin but did not characterize the effects, nature, or ultimate source of the fluids responsible for this alteration. Hydrothermal oxidation, accompanied by dissolution of non-oxide minerals, has been implicated in the upgrading of iron ores in Australia (Thorne and others, 2008) and Brazil (Rosiere and others, 2008). On the Mesabi Iron Range, oxidized magnetic taconite ore has been locally characterized by high Davis Tube concentrate silica values, particularly adjacent to faults. Hydrothermal oxidation may have taken place during the Paleoproterozoic, when the currently-exposed Biwabik Iron Formation was the most deeply buried and was undergoing diagenesis. Later lateritic weathering, perhaps during the Cretaceous, dissolved silica and all other non-iron oxides, resulting in the natural ores as found in the Fayal Mine. Geochemically, these ores are characterized by pronounced cerium anomalies, which can result from intense oxidation near the surface, consistent with a lateritic interpretation for these ores. The older, fault-related hydrothermal oxidation did not produce cerium anomalies, consistent with its deep-seated setting. 26 �Magnetic Taconite Ores The taconite reserves of the Mesabi Range are comprised of magnetite-rich horizons in the Biwabik Iron Formation. Although the iron content of the Biwabik Iron Formation is relatively uniform, the proportion contained in magnetically recoverable magnetite is highly variable, ranging from less than 10% (typically in slaty banded iron-formation) to greater than 25% (typically in cherty granular ironformation). Mineable horizons exist throughout the entire iron-formation thickness in the central Mesabi district. Principal ore units include the middle 100’ of the Lower Cherty (UTAC LC-4, LC-5), and in variable-thickness Upper Cherty submembers (e.g. LUC-2, UC-3 & 3A). Ores are classified on their concentrate weight recovery (typically >25%), crude magnetic iron (~1725%), and concentrate Fe and SiO 2 grades (product averages 66% Fe, 4.5-5% SiO2). Certain cherty ores (LC-4) can produce concentrate silica grades as low as 2% with standard grinding and separating techniques (75-80% -325 mesh, or P80 45 µm), with most cherty ores averaging ~5-6% concentrate SiO2. Slaty magnetite taconite ores produce concentrates of higher concentrate SiO 2, reflecting finer magnetite grain size and textural intergrowth with gangue minerals. Minor contaminants (CaO, MgO, MnO from carbonates, K2O and Al2O3 from silicates) are related to very specific stratigraphic horizons, allowing accurate mine-to-mill blending reconciliations. Three-position blending and maximizing the use of a single high-silica ore source are required for stable processing operations. The magnetite in magnetite taconite ores formed as a result of low-temperature diagenetic recrystallization, likely from reaction of oxide and/or carbonate precursors: Fe2O3 + FeCO3 → Fe3O4 + CO2 (7) (Burt, 1972) Textural relationships also suggest formation directly from a carbonate or silicate precursor: 3FeCO3 + 3H+ → Fe3O4 + CO2+ 3 H2O (8) Fe3Si2O5(OH)4 → Fe3O4 + 2SiO2 + 3H2O + H+ (9) Magnetite-rich iron-formation is typically enriched in iron relative to non-magnetite rich iron-formation (Fig. 12), suggesting reactions with dissolved ferrous iron may play a role in magnetite formation: Fe2O3 + Fe2+ + H2O → Fe3O4 + 2H+ (10) (Ohmoto, 2003) 2Fe(OH)3 + Fe2+ → Fe3O4 + 2H2O + H+ (11) 2FeCO3 + Fe2+ → Fe3O4 + CO2 (12) Reaction of precursor ferric oxide with organic material has also been proposed as a mechanism: 3Fe2O3 + CH2O → 2Fe3O4 + CO2 + 2H+ (13) (modified from Perry and others, 1973) Overall, magnetite formation shows a clear affinity for horizons with sedimentological, mineralogical, and geochemical evidence for a significant carbonate component in the primary precipitate. The association of carbonate with subsequent magnetite formation suggests that a buffered pH was as significant a control as Eh. Textural relationships indicate that magnetite formation occurred after the onset of burial stylolitization and significant chemical compaction. However, it was apparently complete prior to post-Penokean deformation and fluid flow, as it is cross cut by hydrothermal quartzcarbonate veins associated with this event. Figure 12. Total iron versus percent of iron in magnetic fraction of unweathered, unmetamorphosed Biwabik Iron Formation. Note that the highest total iron contents (>35%) are associated with the highest magnetic iron fraction; this suggests magnetite formation was accompanied by iron enrichment. 27 �Production Annual production of direct-shipped ore and taconite pellets produced on the Mesabi Iron Range are shown in Figure 13. Production of direct-shipping ore started in 1892 and rose steadily until 1942, when a record 54 million tons were produced. Gravity concentrate production rose steadily thereafter, until a record 77 million tons of direct-shipping and gravity concentrate ore was produced in 1953. Reserve Mining Company initiated the first large scale taconite operation in 1955, and by 1967 taconite production from six taconite facilities accounted for more than half of iron ore production. The Mesabi Range iron ore industry weathered the global resource recession of the 1980s largely intact, accounting for over 75% of US iron ore production by the end of the decade. The industry continues to evolve, with six taconite facilities (40 mtpy capacity), three tailings recovery facilities (3 mtpy capacity), and a valueadded direct reduced iron facility (0.5 mtpy capacity) in production. Figure 13. Annual production of direct-shipping ore, gravity concentrates, and taconite concentrates from the Mesabi Iron Range for the period 1892-2012. 28 �DESCRIPTION OF FIELD STOPS STOP 1—Stratigraphic section of the Biwabik Iron Formation, Thunderbird North Mine 535000E/5257910N (UTM Zone 15 coordinates, NAD83 datum) Eveleth 7.5’ USGS Quadrangle; SWSW, Section 29, T58N, R17W *NOTE: THIS SITE IS LOCATED ON AN ACTIVE MINE SITE. DO NOT ATTEMPT TO ENTER WITHOUT FIRST OBTAINING PERMISSION. Directions: Beginning in Hibbing, proceed east on US Highway 169 to the interchange with Highway 53 in Virginia (22 miles). Turn south (right) to merge onto US Highway 53, and drive 4.1 miles to the stoplight intersection with Grant Avenue in Eveleth. Turn west (right) onto Grant Avenue, and drive south 0.5 miles to the Cliffs Natural Resources Thunderbird Mine entrance. Historical Overview: Material mined at the Thunderbird North Mine consists of taconite ore horizons from the Lower Cherty, Lower Slaty, and Upper Cherty members. Direct-shipping ore, also referred to a natural ore, was originally mined in the immediate vicinity from the Auburn (1894-2002), Virginia (1910-1953), and Gross-Nelson (1944-1977) deposits. Exploration for magnetic taconite at this site began in earnest in 1960, after the opening of pioneering taconite operations at the Reserve Mining Company (now Northshore Mining) and Erie Mining Company (now Cliffs Erie site) in the mid-1950s. Drilling by Oglebay Norton Company identified a substantial magnetic taconite deposit in the area and the property was jointly developed with the Ford Motor Company – groundbreaking occurred in June, 1964. The Thunderbird North mine and Fairlane plant began producing in November, 1965, with an initial rated capacity of 1.6 million tons of iron ore pellets per year. In 1977, addition of a second concentrating and pelletizing line, and the opening of the adjacent Thunderbird South mine, increased rated capacity to 6.0 million tons of pellets. The Thunderbird South mine closed in 1992, and in 1996, ownership of the operation was transferred to Eveleth Mines LLC. Eveleth Mines closed the concentrating and pelletizing Line 1 in May, 1999, reducing capacity to 4.2 million tons of pellets. The remaining operation was idled in May, 2003. The idled facility was purchased and re-opened by United Taconite LLC in December, 2003 (now owned 100% by Cliffs Natural Resources). They subsequently refurbished and reactivated Line 1 in December, 2004, which increased the annual rated capacity to 5.2 million tons of pellets. Description: Depending on access, one or more sites with mine exposures of the Lower Cherty, Lower Slaty, and Upper Cherty members will be visited. Refer to the detailed stratigraphy section for more information regarding the submembers visited. STOP 2—Fault and Associated Quartz Veining and Alteration/Oxidation, Thunderbird South Mine 534100E/5255520N (UTM Zone 15 coordinates, NAD83 datum) Eveleth 7.5’ USGS Quadrangle SWSW, Section 5, T57N, R17W *NOTE: THIS SITE IS LOCATED ON AN ACTIVE MINE SITE. DO NOT ATTEMPT TO ENTER WITHOUT FIRST OBTAINING PERMISSION. Directions: Proceed southward on company roads through the Thunderbird North Mine. Cross County Highway 101 (through two remote operated gates on either side of the highway) and continue southeast to the north side of Thunderbird South. 29 �Description: The site contains multiple exposures of Lower Slaty units on several benches, with the conspicuous fault/quartz vein area trending SW into the flooded pit. The complete Lower Cherty section remains as reserve in Thunderbird South The exposure in the Thunderbird South pit contains a N45E-trending high-angle fault in the LS2/LUC1 units. The fault is approximately 30 cm wide, and contains quartz veins that have been brecciated and cemented by very fine-grained quartz intergrown with goethite. Fluid inclusions in the quartz cement, which is inferred to have precipitated during the hydrothermal oxidation event (hence its intergrowth with goethite), have average homogenization temperatures of 155° C (n=22) and salinity of 7.3 wt% NaCl equivalent, clearly indicating the involvement of saline hydrothermal fluids in goethiteforming oxidation, and furthermore implicating more widely-distributed diagenetic fluid in that alteration. Breccia clasts of quartz vein from this fault zone have essentially the same homogenization temperatures and salinities as the quartz cement. In terms of trace element geochemistry, the fault breccia is remarkably similar to iron-formation, implying it was largely buffered by iron -formation in a rockdominated system. Notably, the fault breccia displays a distinctive positive Europium anomaly, as does the iron-formation (see also Planavsky and others, 2009). Similar oxidized high-angle faults are known throughout the Iron Range. Adjacent to these faults, iron-formation is oxidized, with iron silicates altered to goethite + quartz, chert textures are overprinted by recrystallized quartz, and magnetite is oxidized to martite. The hydrothermal oxidation is commonly overprinted by late red hematite (+/- goethite) that formed near Earth’s surface: it coats fractures and is associated with silica dissolution. STOP 3—Security Reserve/Fayal Complex Direct-shipping Ore 535070E/5255050N (UTM Zone 15 coordinates, NAD83 datum) Eveleth 7.5’ USGS Quadrangle NWSE, Section 6, T57N, R17W *NOTE: THIS SITE IS LOCATED ON AN ACTIVE MINE SITE. DO NOT ATTEMPT TO ENTER WITHOUT FIRST OBTAINING PERMISSION. Directions: Drive around the western and southern sides of the Thunderbird South pit, past the crusher, and east to ramp into the Fayal Pit. The site is an approximately 200 meters (650 feet) long exposure of ironformation and direct-shipping ore exposed along a northeast-trending access ramp into the flooded Fayal Mine complex. Historical Background: The Fayal Mine (1895-1965; total production 44.5 million tons) was discovered in November, 1893 by David T Adams of Duluth. The mine site was initially explored by the McInnis Mining Company and was sold to the Minnesota Iron Company (a component of the 1901 United States Steel merger), after which the mine was operated by the Oliver Iron Mining Company. Production of direct-shipping ore began in 1895 and was initially extracted by shaft from underground operations. Open pit operations facilitated a rapid increase in production, reaching1.9 million tons in 1902. Through the end of 1919, the complex had yielded an aggregate of 29.9 million tons – more than a million tons per year since 1895 (and two-thirds the ultimate production). The Fayal complex was closed in 1933, but was reopened on a smaller scale, as an open-pit truck operation, to recover lower grades of ore between 1944 and 1965. Final development plans included recovering approximately 794,000 tons of Lower Slaty- and lower Upper Cherty-hosted ores along the south side of the deposit (Security Reserve). However, by the time final mining was contemplated by Auburn Minerals LLC (ca. 2000), the sulfur content of the reserve was deemed unacceptably high. 30 �Description: Included in the Security Reserve is an access ramp to the flooded Fayal Mine. Along this ramp, direct-shipping ore is exposed in both the floor and wall of the ramp. This site is one of the few remaining locations on the Mesabi Iron Range to view in situ direct-shipping ore. All direct-shipping ore in the Fayal deposit falls under the Soft Iron Ore Classification of Marsden (1968). The Fayal ore consists predominantly of hematite and goethite, with minor magnetite and manganese oxides, as is common with the other soft ore deposits of the Mesabi Iron Range. Silica and clay minerals are the predominant gangue minerals. In 1901, Fayal direct-shipping ore was reported to averaged 63.8% iron, 0.037% phosphorus, and 2.95% silica (dry basis; Leith, 1903). The direct-shipping ore visible in the Fayal ramp occurs along the margin of the deposit and likely has lower iron and higher silica content than the typical higher grade ore shipped from the Fayal deposit for most of its life. Iron-formation exposed along the east side of the Fayal ramp parallels the contact of the directshipping ore deposit. The ore is formed from predominantly slaty proto-ore, and displays varying degrees of desilicification (leaching) and oxidation. Bedding in direct-shipping ore on the north end of the ramp clearly displays a relatively steep dip to the west and lies near the center of the trough. The west side of the ramp parallels the northeast-trending Fayal fault, a high-angle, west-dipping normal fault, and one of the larger structures cross-cutting the Biwabik Iron Formation. The fault is occupied by a thick, brecciated, and re-cemented quartz ± carbonate vein. Visible immediately adjacent to the large vein is drag folding in the footwall iron-formation, indicating a hangingwall- (westside-) down sense of motion. STOP 4—Drill Core Display 535000E, 5257910N (UTM Zone 15 coordinates, NAD83 datum) Eveleth 7.5’ USGS Quadrangle SWSW, Section 29, T58N, R17W *NOTE: THIS SITE IS LOCATED ON AN ACTIVE MINE SITE. DO NOT ATTEMPT TO ENTER WITHOUT FIRST OBTAINING PERMISSION. Directions: Proceed back north along the same route to the core shack within the Thunderbird North Mine. Description: A drill hole cored through most of the Biwabik Iron Formation, and a portion of the upper Pokegama Formation, will be on display inside the core shack or, weather permitting, will be laid outside. Ironformation submembers will be labeled according to the Thunderbird North classification scheme. STOP 5—Algal/Conglomerate unit of the Upper Cherty member, Mary Ellen Mine 548260E, 5264380N (UTM Zone 15 coordinates, NAD83 datum) Biwabik 7.5’ USGS Quadrangle NENW, Section 10, T58N, R16W *NOTE: THIS SITE IS LOCATED ON AN ACTIVE MINE SITE. DO NOT ATTEMPT TO ENTER WITHOUT FIRST OBTAINING PERMISSION. Directions: From the Thunderbird Mine entrance, turn left (north) on Grant Avenue. Proceed to the .intersection with Highway 53 (0.5 miles), and turn left (north). Proceed 1.5 miles to the intersection with Highway 135, and turn right (east) to merge onto MN Highway 135. Drive east 10 miles to the intersection with County Road 715, located just outside the western outskirts of Biwabik. Turn left (north) on 715, and proceed 0.2 miles. The entrance to the Mary Ellen Mine will be on the south (left) side of the road. 31 �Historical Background: The Mary Ellen Mine was first opened in 1924 by Pioneer Mining (Stanley Mining, operator), and saw regular production of what was termed ‘hard, bluish-red siliceous hematite’ through 1928. Stanley Mining reopened the Mary Ellen in 1948, and it experienced sporadic production through to final topography in 1962 (last operated by Pittsburgh-Pacific). Total shipments in the period 1924-1962 were 4.6 million tons of gravity concentrates. Description: The Mary Ellen mine is perhaps most notable for its exposures of the algal submember of the Upper Cherty (equivalent to the UC-6 submember at the Thunderbird Mine, and the I submember of Gundersen and Schwartz (1962)). Here, stromatolites occur as mounds of fossilized algal colonies separated by intraformational jasper conglomerates. The algal and conglomeratic units exhibit a combined thickness of 2-20 feet. This horizon occurs only in the eastern half of the Mesabi Iron Range, pinching out in the vicinity of Hibbing. Planavsky and others (2009) attribute the stromatolites to Fe-oxidizing bacteria present in the Animikie Basin and in similar settings world-wide, where microbial communities proliferated at specific shallow-water redox boundaries in late Paleoproterozoic oceans (see Fig. 14, below). The Mary Ellen mine is noted for the abundance of colonies of finely-laminated, small (~1cm diameter) digitate, columnar stromatolites. They occur in mound-like aggregations that appear to have been buried in-situ on the seafloor, in contrast to the largely resedimented oncoliths comprising the algal chert submember farther to the west. Discussion: The presence of stromatolites and intraformational conglomerate at this stratigraphic horizon within the iron-formation is consistent with extremely shallow water, and perhaps even emergent (subaerial) conditions. This likely represents maximum marine regression during the transgressive-regressivetransgressive cycle that characterizes deposition of the Biwabik Iron Formation. The carbonate rocks that comprise the uppermost Upper Slaty member of the iron-formation (submember US-2), though enigmatic, may relate to development of a second regression. Figure 14. Model of stromatolite depositional environment (Planavsky and others, 2009). 32 �REFERENCES Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005, Discovery of distal ejecta from the 1840 Ma Sudbury impact event: Geology 33, p. 193-196. Bleifuss, R.L., 1964, Mineralogy of oxidized taconites of the western Mesabi and its influence on metallurgical processes: Transactions of AIME, v. 229, p. 235-244. Bonnichsen, B., 1968, General geology and petrology of the metamorphosed Biwabik Iron Formation, Dunka River area, Minnesota: Unpublished Ph.D. dissertation, University of Minnesota, Minneapolis, 240 p. Boggs, S., 2001, Principles of Sedimentology and Stratigraphy (third edition): Upper Saddle River, New Jersey, Prentice Hall, p. 321-447. Cannon, W.F. and Addison, W.D., 2007, The Sudbury Impact layer in the Lake Superior iron ranges: A time-line from the heavens: Institute of Lake Superior Geology, 53rd Annual Meeting, May 8-13, 2007, Lutsen, Minnesota, v. 53, Part 1-Proceedings and Abstracts, p. 20-21. Cannon, W.F., Schultz, K.J., Horton, W. Jr., and Kring, D.A., 2010, The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA: Geological Society of America Bulletin, v. 122, p. 50-75. Dimroth, E. and Chauvel, J.J., 1974, Facies Types and Depositional Environmenta of the Sokoman Iron Formation in Part of the Central Labrador Trough, Quebec, Canada, Journal of Sedimentary Petrology, v. 44, p. 299-327. Craddock, J.P., Rainbird, R.H., Davis, W.J., Davidson, C., Vervoort, J.D., Konstantinou, A., Boerboom, T.J., Kerber, L., and Lundquist, B., 2013, Detrital zircon geochronology and provenance of the Paleoproterozoic Huron (~2.4-2.2 Ga) and Animikie (~2.2-1.8 Ga) basins, southern Superior Province: Journal of Geology, v. 121, p. 623-644. Everett, J. V., 1956, Judson mine: Geol. Soc. America Guidebook Series, Field Trip #1, Minneapolis, Minn. mtgs., p. 216-218. Fischer, W.W., and Knoll, A.H., 2009, An iron shuttle for deepwater silica in Late Archean and Early Paleoproterozoic iron formation: Geological Society of America Bulletin, v. 121, p. 222-235. Fralick, P.W., Davis, D.W., and Kissin, S.A., 2002, The age of the Gunflint Formation, Ontario, Canada: Single zircon U-Pb age determinations from reworked volcanic ash: Canadian Journal of Earth Sciences, v. 39, p. 1089-1091. French, B.M., 1973, Diagenetic and low-grade metamorphic iron-formation: Economic Geology, v. 68, no.7, p. 1063-1074. Gruner, J., 1956, The Mesabi Range: Precambrian of northeastern Minnesota; Geological Society of America Annual Meeting Guidebook, Field Trip 1, pp. 182-215 Gundersen, J.N., and Schwartz, G.M., 1962, The Geology of the metamorphosed Biwabik Iron-formation, eastern Mesabi District, Minnesota: Minnesota Geological Survey, Bulletin 43, 139 p. Hemming, S.R., McLennan, S.M., Hanson, G.N., and Krogstad, K.M., 1990, Pb isotope systematics in quartz [abs]: Eos, v. 71, no. 17, p. 654-655. Hemming, S.R., Hanson, G.N., and McLennan, S.M., 1995, Precambrian crustal blocks in Minnesota: Neodymium isotope evidence from basement and metasedimentary rocks: U.S. Geological Survey, Bulletin 1904-U, 15 p. James, H.L, 1954, Sedimentary facies of iron-formation: Economic Geology v. 49, no.3, p. 235-293 Jirsa, M.A., 2010, Stratigraphy of Sudbury “impactite” near Gunflint Lake, NE Minnesota: Institute on Lake Superior Geology Proceedings, 56th Annual Meeting, Proceedings and Abstracts, v. 56, part 1, p. 31-32. Jirsa, M.A., and Boerboom, T.J., 2003, Geology and mineralization of Archean bedrock in the Virginia Horn, p.10-73, in Jirsa, M.A., and Morey, G.B., eds., Contributions to the geology of the Virginia horn area, St. Louis County, Minnesota: Minnesota Geological Survey Report of Investigations 53. 33 �Jirsa, M.A., Boerboom, T.J., and Morey, G.B., 1998, Bedrock geologic map of the Virginia horn, Mesabi Iron Range, St. Louis County, Minnesota: Minnesota Geological Survey Miscellaneous Map M-85, scale 1:48,000. Jirsa, M.A., and Chandler, V.W., 2006, Structure of the Biwabik Iron Formation, Mesabi Iron Range, Minnesota: Proceedings, Institute on Lake Superior Geology, v.52. Part 1, p.32-34. Jirsa, M.A., Chandler, V.W., and Lively, R.S., 2005a, Bedrock geology of the Mesabi Iron Range, Minnesota: Minnesota Geological Survey Miscellaneous Map M-163, scale 1:100,000. Jirsa, M.A., Fralick, P.W., Weiblen, P.W., and Anderson, J.L.B., 2011, Sudbury impact layer in the western Lake Superior region, in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 147–169, doi:10.1130/2011.0024(08). Jirsa, M.A., and Morey, G.B., eds., 2003, Contributions to the geology of the Virginia horn area, St. Louis County, Minnesota: Minnesota Geological Survey Report of Investigations 53, 135 p. Jirsa, M.A., Setterholm, D.R., Bloomgren, B.A., and Lively, R.S., 2002 , Bedrock topographic and depth to bedrock maps of the western half of the Mesabi Iron Range, northern Minnesota: Minnesota Geological Survey Miscellaneous Map M-126, scale 1:100,000. Jirsa, M.A., Setterholm, D.L., Bloomgren, B.A., Bauer, E.J., and Lively, R.S., 2005b, Bedrock geology, database, depth to bedrock, and bedrock topography of the eastern half of the Mesabi Iron Range, Northern Minnesota: Minnesota Geological Survey Miscellaneous Map M-158, scale 1:100,000. Klein, C., 2005, Some Precambrian banded iron-formations (BIFs) from around the world: Their age, geologic setting, mineralogy, metamorphism, geochemistry, and origin: American Mineralogist, v. 90, p. 1473-1499. Konhauser, K. O., Hamade, T., Raiswell, R., Morris, R. C., Ferris, F. G., Southam, G., and Canfield, D. E., 2002, Could bacteria have formed the Precambrian banded iron-formations? Geology, v. 30, p. 1079- 1082. LaBerge, G.L., 1964, Development of magnetite in iron-formations of the Lake Superior region: Economic Geology, v. 59, no. 7, p. 1313-1342. LaBerge, G.L., Robbins, E.I., and Han, T.M., 1987, A model for the biological precipitation of Precambrian iron-formations, A: Geological evidence, in Appel, P.W.U., and LaBerge, G.L. eds., Precambrian iron-formations: Athens, Greece, Theophrastus Publications, p. 69-96. Larson, P., 2013, Chemostratigraphy of the Biwabik Iron Formation: Implications for basin longevity and evolution: Proceedings, Institute on Lake Superior Geology, v. 59, Part 1, p. 50-51. Lascelles, D. F., 2007, Black smokers and density currents: A uniformitarian model for the genesis of banded iron-formations. Ore Geol. Reviews, v. 32, p. 381-411. Leeder, M., 1999, Sedimentology and Sedimentary Basins, from Turbulence to Tectonics: Blackwell Science Ltd., p 444-464. Leith, C.K., 1903, The Mesabi iron-bearing district of Minnesota: U.S. Geological Survey Monograph 43, 316 p. Losh, S., Rague, R., 2013, Silica remobilization in the Biwabik Iron Formation, Minnesota;Proc. and Abstr. Instit. On Lake Superior Geol., v 59, part 1, p. 60. Lucente, M.E., and Morey, G.B., 1983, Stratigraphy and sedimentology of the Lower Proterozoic Virginia Formation, northern Minnesota: Minnesota Geological Survey Report of Investigations 28, 28 p. Maliva, R.G., Knoll, A.H., and Simonson, B.M., 2005, Secular change in the Precambrian silica cycle: Insights from chert petrology: Geological Society of America Bulletin, v. 117, p. 835–845 Marsden, R.W., 1968, Geology of the iron ores of the Lake Superior region, in v.1, of Ore deposits of the United States, 1933-1967 (Graton-Sales volume): New York, American Institute of Mining, Metallurgical, and Petroleum Engineers, p. 489-506. Marsden, R.W., Emanuelson, J.W., Owens, J.S., Walker, N.E., and Werner, R.F., 1968, The Mesabi Iron Range, Minnesota, in v.1 of Ore Deposits of the United States, 1933-1967 (Graton-Sales volume): New York, American Institute of Mining, Metallurgical, and Petroleum Engineers, v. 1, p. 518-537. 34 �McSwiggen, P.L., and Morey, G.B., 2008, Overview of the mineralogy of the Biwabik Iron Formation, Mesabi Iron Range, northern Minnesota: Regulatory Toxicology and Pharmacology, v. 52, no. 1, supplement 1, p, S11-25. Morey, G.B.,1992, Chemical composition of the eastern Biwabik Iron Formation (Early Proterozoic), Mesabi range, Minnesota: Economic Geology, v. 87, p. 1649-1658. ———1999, High-grade iron ore deposits of the Mesabi range, Minnesota—product of a continentalscale Proterozoic ground-water flow system: Economic Geology, v. 94, p. 133-142. Morey, G.B., 2003, Paleoproterozoic Animikie Group, related rocks, and associated iron-ore deposits in the Virginia Horn: in Jirsa, M.A., and Morey, G.B., eds., Contributions to the geology of the Virginia Horn area, St. Louis County, Minnesota: Minnesota Geological Survey Report of Investigations 53, p. 74-102. Morey, B.B., and Lively, R.S., 1999, Background levels of mercury and arsenic in Paleoproterozoic rocks of the Mesabi Iron Range, northern Minnesota: Minnesota Geological Survey Information Circular 43, 14p. Nichols, G., 1999, Sedimentology and stratigraphy: Oxford, United Kingdom, Blackwell Science Limited, p 135 -193, p 264-289, Chap 11, 14, 21. Ohmoto, H., 2003, Nonredox transformations of magnetite-hematite in hydrothermal systems: Economic Geology, v. 98, p. 157-161. Ojakangas, R.W., 1983, Tidal deposits in the early Proterozoic basin of the Lake Superior region—the Palms and the Pokegama Formations: Evidence for subtidal-shelf deposition of Superior-type banded iron-formation, in Medaris, L.D., Jr., ed., Early Proterozoic geology of the Great Lakes region: Geological Society of America Memoir 160, p. 49-66. Ojakangas, R.W., Severson, M.J., Jongewaard, P.K., Arola, J.L., Evers, J.T., Halverson, D.G., Morey, G.B., and Holst, T.B., 2005, Geology and sedimentology of the Paleoproterozoic Animikie Group: The Pokegama Formation, the Biwabik Iron Formation, and Virginia Formation of the eastern Mesabi iron range, and the Thomson Formation near Duluth, northeastern Minnesota, in Robinson, L., ed., Field trip guidebook for selected geology in Minnesota and Wisconsin: St. Paul, Minnesota Geological Survey Guidebook 21, p. 208-237. Owens, J. S., 1956, Enterprise and Alpena mines: Geol. Soc. America Guidebook series, Field Trip #1, Minneapolis, Minn. mtgs., p. 168-174. Patelke, M.M,, in progress, Mechanisms for the Formation of Granules in the Biwabik Iron Formation, Mesabi Iron Range, Northeastern Minnesota, Master’s Thesis, University of Minnesota Duluth. Perry, E.C., Tan, F.C., and Morey, G.B., 1973, Geology and stable isotope geochemistry of the Biwabik Iron Formation, northern Minnesota; in Precambrian iron-formations of the world: Economic Geology, v. 68, p. 1110-1125. Pfleider, E.C., Morey, G.B., and Bleifuss, R.L., 1968, Mesabi deep drilling project: Progress report no. 1, in Mining Symposium, 29th Annual, and American Institute of Mining and Metallurgical Engineers, Minnesota Section, 41st Annual Meeting, Duluth, [Proceedings]: University of Minnesota, p. 52-92. Pietrzak, N J; Duke, N; Scott, G; Lukey, H., 2009; Ore textures and mineral chemistry within the oxidecarbonate-silicate flotation ores at the Cliffs Natural Resources' Tilden Mine, Michigan:Proceedings, Institute on Lake Superior Geology. Meeting, v. 55, Part 1, p. 66-67. Planavsky, N., Rouxel, O., Bekker, A., Shapiro, R., Fralick, P., Knudsen, A., 2009, Iron-oxidizing microbial ecosystems thrived in late Paleoproterozoic redox-stratified oceans; Earth and Planetary Science Letters, v. 286, pp. 230-242. Pratt, B.R., James, N.P, and Cowan C.A., 1992, Peritidal Carbonates, in Walker, R.G., and James, N.P. ed., 1992, Facies Models Response to Sea Level Change: Geologic Association of Canada, p. 303-322. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, G.J., and Edwards, C.T., 2007, Physical and chemical evidence of the 1850 Ma Sudbury impact event in the Baraga Group, Michigan: Geology, v. 35, p. 827-830. Purucker, M., 1983, Time of formation of soft iron ore on the Gunflint and Mesabi Ranges (Ontario, Canada and Minnesota, USA). Economic Geology, v. 78, p. 502-506. 35 �Rajan, S., Mackenzie, F. T., and Glenn, C. R., A thermodynamic model for water column precipitation of siderite in the Plio-Pleistocene Black Sea. Amer. Jour. Sci., v. 296, p. 506-548. Rankey, E.C. and Reeder, S.L., 2011, Holocene oolitic marine sand complexes of the Bahamas, Journal of Sedimentary Research, v. 81, p. 97-117. Rosiere, C.; Spier, C., Rios, F, and Suckau, V., 2008, The itabirites of the Quadrilatero Ferrifero and related high-grade iron ore deposits: An overview; in Economic Geology Reviews, v. 15, pp. 223-254 Severson, M.J., Heine, J.J. and Patelke, M.M., 2009, Geologic and stratigraphic controls of the Biwabik Iron Formation and the aggregate potential of the Mesabi Iron Range: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-2009/09, 179 p. Severson, M.J., Ojakangas, R.W., Larson, P.C., and Jongewaard, P.K., 2010, Geology and Stratigraphy of the Central Mesabi Iron Range, in Field Guide to the Geology of Precambrian Iron Formations in the Western Lake Superior Region, Minnesota and Michigan, Precambrian Research Center Professional Workshop Series, PRC Guidebook 10-01, pp. 15-52. Simonson, B.M., 1987, Early silica cementation and subsequent diagenesis in arenites from four early Proterozoic iron formations in North America, Journal of Sedimentary Petrology, vol. 57, no. 3, p. 494-511. Simonson, B.M., 2003, Origin and evolution of large Precambrian iron formations: Geological Society of America Special Paper 370, p. 231-244. Sloan, R., 1964, The Cretaceous system in Minnesota; Minnesota Geological Survey Report of Investigations, v. 5, 64 pp. Sommers, M.G, Awramik, S.M., and Woo, K.S. Evidence for initial calcite-aragonite composition of Lower Algal Chert Member ooids and stromatolites, Paleoproterozoic Gunflint Formation, Ontario, Canada, Canadian Journal of Earth Science, v. 37, p. 1229-1243. Southwick, D.L., and Morey, G.B., 1991, Tectonic imbrication and foredeep development in the Penokean orogen, east-central Minnesota: Minnesota Geological Survey Report of Investigations 37, 25 p. Symons, D.T.A., 1966, A paleomagnetic study on the Gunflint, Mesabi, and Cuyuna iron ranges in the Lake Superior region; Economic Geology, v. 61, p. 1336-1361. Theriault, S.A., 2011, Mineralogy, spatial distribution, and isotope geochemistry of sulfide minerals in the Biwabik Iron Formation: unpublished MS thesis, University of Minnesota-Duluth, 165 p. Thorne, W., Hagemann, S; Webb, A.; Clout, J.; 2008, Banded iron formation –related iron ore deposits of the Hamersley Province, Western Australia; Economic Geology Reviews, v. 15, pp. 197-221. Vallini, D.A., Cannon, W.F., Schulz, K.J., and McNaughton, N.J., 2007, Thermal history of low metamorphic grade Paleoproterozoic sedimentary rocks of the Penokean orogeny, Lake Superior region: Evidence for a 1786 Ma overprint: Precambrian Research, v. 157, p. 169-187. White, D.A., 1954, Stratigraphy and structure of the Mesabi range, Minnesota: unpublished PhD Dissertation, University of Minnesota, 146 p., 4 plates. Wolff, J.F., 1917, Recent geologic developments on the Mesabi Iron Range, Minnesota: American Institute of Mining and Metallurgical Engineers, Transactions, v. 56, p. 229-257. Zanko, L.M., Severson, M.J., Oreskovich, J.A., Heine, J.H., Hauck, S.A., and Ojakangas, R.W., 2003, Oxidized taconite geological resources for a portion of the western Mesabi Range (west half of the Arcturus Mine to the east half of the Canisteo Mine), Itasca County, Minnesota—a GIS-based resource analysis for land-use planning: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-2001/40, 85 p. Zielinski, A.M., Mancuso, J.J., Frizado, J.P., and Waidler, R.J., 1994, Manganese-rich oncolites in the Biwabik Iron Formation, Eveleth Mine, Mesabi Iron Range, Minnesota: in D.L. Southwick, ed., Short Contributions to the Geology of Minnesota 1994, Minnesota Geological Survey Report of Investigations 43, pp. 48-59. 36 �FIELD TRIP 2 Wednesday, May 14, 2014 A WALK IN THE PARK—NEOARCHEAN GEOLOGY OF LAKE VERMILION STATE PARK LEADERS: George J. Hudak (Natural Resources Research Institute – Duluth) Amy Radakovich (Minnesota Geological Survey) Geoff Pignotta and Kelly Schwierske (Geology Dept., University of Wisconsin - Eau Claire) INTRODUCTION The Vermilion District of northeastern Minnesota contains one of the classic greenstone belts in the United States. The district comprises the southwestern part of the Wawa Abitibi Terrane (Stott et al., 2007; Stott and Mueller, 2009) which encompasses Neoarchean metavolcanic, metasedimentary, and meta-intrusive rocks that extend northeastward through northwestern Ontario and Quebec. In Canada, this terrane hosts numerous volcanogenic massive sulfide deposits (e.g. Winston Lake, Geco, Noranda), goldrich volcanogenic massive sulfide deposits (Horne (Noranda camp), Bousquet 2 – LaRonde 1, LaRondePenna; Mercier-Langevin et al., 2010), as well as a large number of lode (orogenic) gold deposits (for example, in the Hemlo, Timmins, and Kirkland Lake camps). The Vermilion District is known for its numerous, previously mined massive hematitic iron ore deposits (including the Pioneer Mine in Ely and the Soudan Mine in Soudan) which locally occur within regional extensive Algoma-type banded iron formations. To date, no volcanogenic massive sulfide, gold-rich volcanogenic massive sulfide, or lode gold deposits have been discovered in the Vermilion District, although several studies (Peterson and Jirsa, 1999; Peterson, 2001; Hudak et al., 2002a; Peterson and Patelke, 2003; Hoffman, 2007; Hudak et al., 2007; Hudak et al., 2012; Lodge et al., 2013) have indicated that evidence for volcanic, hydrothermal, and structural processes associated with these types of mineral deposits is present throughout the Vermilion District. Since the late 1990’s considerable geological research has been conducted in the region between Tower, MN (in the west) to Ely, MN (in the east) within the Vermilion District. Much of this research has been conducted to better understand the stratigraphy, structural geology, and economic geology of the belt, and the results of these studies have provided a solid foundation for the geological research that has, and currently is, taking place in Lake Vermilion State Park. Several recent Institute on Lake Superior Geology (ILSG) field trips (Hudak et al., 2004; Jirsa et al., 2004; Peterson and Patelke, 2004; Larson and Mooers, 2009; Peterson et al., 2009a; Jirsa and Hillman, 2009; Peterson et al., 2009b) describe these finding for specific areas in and around the Vermilion District. The Vermilion District’s iron ore mining heritage is currently preserved at two state parks located near Soudan, Minnesota. With the donation of land and infrastructure associated with the former Oliver Iron Mining Division’s Soudan Mine by United States Steel to the State of Minnesota in 1965, Soudan Underground Mine State Park was established. This state park currently preserves the historical surface and underground workings from, as well as the wilderness adjacent to, Minnesota’s oldest iron ore mine, the Soudan Mine. This mine operated from 1882 until December, 1962 and produced approximately 15.5 tons of hematic iron ore. This popular tourist site continues to be the focus of a wide variety of research spanning geology, geochemistry, hydrogeology, biology, biochemistry and physics. Lake Vermilion State Park is Minnesota’s newest state park, and comprises over 3,000 acres of land, including over five miles of undeveloped shoreline on Lake Vermilion (Bakst, 2013). In 2008, Minnesota State Legislature set aside $20 million in bonding authority to buy, plan, and develop the park, which is located immediately east of Soudan Underground Mine State Park. The park was established in June 2010 after land was 37 �purchased from U. S. Steel Corporation. At the present time, the park is undergoing considerable development, including establishment of trails, roads, and campsites. The park boasts a rich natural and human history, including a wide variety of ~2.7 billion year old rocks that were formed by a wide variety of genetic process, abundant wildlife, as well as archaeological evidence for human habitation dating back over 6,000 years. Additionally, considerable evidence for recent (within the past 140 years) mineral exploration efforts can be readily identified in the park. In 2010 and 2011, students and faculty from the Precambrian Research Center at the University of Minnesota Duluth had the opportunity to conduct detailed (1:5000 scale) geological mapping in both Soudan Underground Mine State Park (Vallowe et al., 2010) and Lake Vermilion State Park (Radakovich et al., 2010; Heim et al., 2011). Twelve students (Nick Heim, Robert Kilduff, Chris Mahr, Charlie Parent, Molly Partridge, Rita Pierce, Amy Radakovich Andrew Ritts, Christine Rahtz, Heather Scott, Andrew Vial, Spencer Young) and instructor George Hudak performed geological mapping in the northwestern (2010) and northeastern (2011) parts of Lake Vermilion State Park. Recently, Geoff Pignotta and Kelly Schwierske of the Geology Department at the University of Wisconsin Eau Claire compiled these geological maps and conducted lithogeochemical evaluations of several lithological units in the park (Schwierske et al., in press). This trip builds on these findings, and will be the first formal geology field trip in Lake Vermilion State Park. It will include a walk up-section through the stratigraphic sequence exposed along a single two-track trail that traverses the park. As well, two outcrops outside the state park boundary will be investigated, as they comprise classic outcrops that will add context to the geological story developed through observing rocks in the park. REGIONAL GEOLOGIC SETTING Supracrustal rocks in the Vermilion district consist of volcanic-dominated stratigraphic sequences of the Wawa Abitibi Terrane within the Superior Province of the Canadian Shield. Rocks of the Wawa Abitibi Terrane in northern Minnesota are divided on the basis of stratigraphic and structural setting into: (1) the Soudan belt, to the south, and (2) the Newton belt, to the north (Jirsa et al., 1992; Southwick et al., 1998). The boundary between these contrasting structural panels can be traced geophysically across the width of Minnesota, and was informally designated the Leech Lake structural discontinuity (Jirsa et al., 1992). In the region west and north of the Lake Vermilion State Park, the Leech Lake structural discontinuity occurs along the Mud Creek shear zone (Hudleston et al., 1988), small segments of the Vermilion and Wolf Lake faults (Sims and Southwick, 1985), and the Bear River fault (Jirsa et al., 1992). A simplified regional geological map of the Neo-Archean terranes of northeastern Minnesota and adjacent Ontario is presented in Figure 1. The Soudan belt (Figure 2) contains large, broad folds involving calc-alkalic and tholeiitic volcanic strata overlain by, and locally interdigitated with, turbiditic rocks. In contrast, the Newton belt consists of elongate, northeast-trending, and mostly northward-younging volcanic and volcaniclastic sequences. Volcanic rocks of the Newton belt differ from those of the Soudan belt in containing locally abundant komatiitic flows and peridotitic sills. The two belts are fault- bounded, and the relationships between stratigraphic units within each belt are largely conformable (although faults obscure contacts locally). In its eastern extension, the Soudan belt is continuous with the Saganagons assemblage in Ontario and terminates against the Saganaga pluton and Northern Light Gneiss. The Newton belt extends discontinuously eastward into the Shebandowan District of Ontario to form the Greenwater and Burchell assemblages. Intrusive rocks in both belts vary from gabbroic and felsic porphyries demonstrably related to volcanism, to large plutons emplaced post-tectonically. Both districts contain unconformable, Timiskaming-type sequences composed of calc-alkalic volcanic rocks, conglomerates, and finer grained sedimentary rocks. Lithostratigraphic units in the western Vermilion district (Table 1) include: (1) the Lower member, Soudan Iron-Formation member, and Upper member (Upper Ely) of the Ely Greenstone Formation, the Lake Vermilion Formation (including the informally named Britt and Gafvert Lake sequences), and the Knife Lake Group of the Soudan belt; (2) the Bass Lake sequence (Peterson and Jirsa, 1999) and the 38 �Newton Lake Formation of the Newton belt; and, (3) syn- to post-tectonic granitoid intrusions of the Giants Range batholith, and a suite of post-tectonic alkalic stocks and plutons. Contacts between the different units are typically conformable, although considerable overlap in time and space is documented between volcanic and sedimentary sequences (Southwick, 1993). Geochronological information for supracrustal and intrusive lithologies in the Vermilion District is relatively sparse (Figure 2). Peterson et al. (2001) obtained a U-Pb zircon age date of 2722 ± 0.9 Ma from a quartz-phyric rhyolite dome in the Fivemile Lake Sequence of the Lower Member of the Ely Greenstone Formation. Lodge et al. (2013) obtained a U-Pb zircon date of 2689.7 ± 0.8 Ma for a Gafvert Lake Sequence dacitic tuff breccia that occurs approximately 2m north of the contact with the Soudan Iron-Formation member of the Ely Greenstone Formation. As well, Lodge et al. (2013) obtained detrital zircon dates ranging from 2680-2690 Ma from greywackes that comprise the Lake Vermilion formation. This date confirms the source of the detritus in the Lake Vermilion Formation was derived locally from the volcaniclastic rocks comprising the Gafvert Lake Sequence. Jirsa et al. (2012) obtained a U-Pb age of 2690.7 ± 0.6 Ma for synvolcanic intrusions that cross-cut volcaniclastic rocks that comprise the Knife Lake Group. The upper part of the Knife Lake Group includes conglomerates which contain clasts derived from the Saganaga Tonalite, which has been dated by Driese et al. (2011) at 2690.83 ± 0.26 Ma. Peterson et al. (2001) also dated a non-foliated feldspar porphyry intruded into Newton Belt strata at 2683.1 +1/-4 Ma. This date provides a minimum age for the regional D2 deformation event that is described below. Wawa-Abitibi Terrane Figure 1. Simplified correlation map of Neoarchean assemblages in Minnesota and northwestern Ontario (after Peterson et al., 2001). Inset map illustrates location of the Wawa-Abitibi Terrane in Minnesota and northwestern Ontario (Stott et al., 2007). The Leach Lake structural discontinuity is illustrated in red. The red star symbols indicate location of Lake Vermilion State Park. 39 �Figure 2. Generalized geology of the Vermilion District in the vicinity of the Tower-Soudan anticline (modified after Peterson, 2001). Locations, ages, and sources of U-Pb ages dates within the district are noted in the callout boxes. Generalized lithologies for each of the groups, formations or sequences are also noted. 40 �Intrusive Rocks Late Intrusions Vermilion Granitic Complex Giants Range Batholith Supracrustal Rocks Newton Belt Newton Lake Formation Bass Lake Sequence Soudan Belt Knife Lake Group Lake Vermilion Formation Gafvert Lake Sequence Britt Sequence Upper Member – Ely Greenstone Soudan Member – Ely Greenstone Lower Member – Ely Greenstone Central Basalt Sequence Fivemile Lake Sequence Plutons and stocks of syenite, monzonite, diorite, and lamprophyre. A U-Pb zircon age date of a non-foliated feldspar porphyry intrusion in the Newton belt is 2683 ± 1.4 Ma (Peterson et al., 2001). Granite, schist, amphibolite, and schist-rich migmatite Granite, granodiorite, monzodiorite, and schist-rich migmatite Tholeiitic and komatiitic basalt lava flows, intrusions, and clastic strata Tholeiitic basalt lava flows, iron-formation, and felsic porphyries Graywacke, slate, conglomerate, and sheared equivalents Graywacke, slate, dacitic tuff, minor conglomerate. Detrital zircons from planar bedded, normal-graded resedimented volcaniclastic rocks have U-Pb age dates of 2680-2690 Ma (Lodge et al., 2013) Dacitic to rhyodacitic tuff, lapilli-tuff, tuff-breccia, and ironformation. Basal dacite tuff-breccia deposits in Lake Vermilion State Park have U-Pb age date of 2689.7 ± 0.8 Ma (Lodge et al., 2013) Tholeiitic basalt lava flows Tholeiitic basalt lava flows and iron-formation Oxide-facies iron formation with intercollated basalt lava flows and felsic volcaniclastic rocks Calc-alkaline and tholeiitic basalt-rhyolite lava flows, tuffs, epiclastic rocks, and minor iron-formation Calc-alkaline to tholeiitic sparsely amygdaloidal basalt and minor basaltic andesite lava flows with MORB-like or back arc basin-like chemical affinities within 100-200 meters of the overlying Soudan Member iron-formation; FII- and FIIIa-type felsic volcanic and volcaniclastic rocks Calc-alkaline to transitional moderately to highly vesicular basalt and andesite lava flows and volcaniclastic rocks with arc-like chemical affinities: FI-, FII-, and FIV-type felsic volcanic and volcaniclastic rocks. Rhyolite dome at near Fivemile Lake has U-Pb age date of 2722.6 ± 0.9 Ma (Peterson et al., 2001). Epithermal-like zinc stringer mineralization is present near Fivemile Lake (Hudak et al., 2002a). Table 1. Lithostratigraphic units within the western Vermilion District (modified after Peterson and Jirsa, 1999; Peterson et al., 2009; Hudak et al., 2012). 41 �STRUCTURAL GEOLOGY The structural geology of the Vermilion District has been well described by Peterson et al. (2009), and is reproduced below. Periods of generally N-S directed compression resulted in three major regional deformation events in the Neoarchean terranes of northern Minnesota. The earliest deformation event (D1) produced broad, locally recumbent folds within the Soudan belt and major fault zones throughout the region. In the Newton belt, D1 was accommodated by thrust imbrication of large crustal blocks, resulting in mainly northward stratigraphic facing. Field relationships indicate that uplift, faulting, and the deposition of Timiskaming-type clastic sedimentary sequences in local fault-bounded basins occurred late in D1 deformation (Jirsa, 2000). A large, map-scale structure related to D1 deformation in the western Vermilion District is the Tower-Soudan Anticline, which is a west-plunging anticline within which the axis and plunge changes orientation along strike from nearly vertical in basalts to shallow NE plunging in the western sedimentary rocks. Axial-planar cleavage associated with this early fold typically is lacking, although Bauer (1985), Hooper and Ojakangas (1971), Hudleston (1976), and Jirsa et al. (1992) have described early cleavage (S1) locally. A second deformation event (D2) associated with synchronous regional metamorphism resulted in foliation development and structures having largely dextral asymmetry. D2 is constrained in the Vermilion District to the time period 2674 to 2685 Ma (Boerboom and Zartman, 1993), and between about 2680 and 2685 Ma in the Shebandowan (Corfu and Stott, 1998). Because D2 deformation affected all of the supracrustal rocks in the area and is reasonably constrained by geochronology, the regional foliation (S2) can be used in the field to temporally relate other structural, intrusive, and deformation events. The relationship between S2 fabric and shear structures indicates that most shearing occurred relatively late in the D2 event. Major shearing that produced the Mud Creek and related shear zones is attributed to the late stages of D2 dextral transpression. Structures related to the third deformation event (D3), which led to juxtaposition of the Wawa Abitibi and Quetico terranes (Peterson and Patelke, 2003), include abundant NE- and NW-trending faults that dissect the stratigraphic assemblages. Named structures related to D3 include the NE-trending Waasa and Camp Rivard faults east of the Soudan Mine area, and the WNW-trending, crustal-scale Vermilion and related faults that form the Wawa-Quetico Subprovince boundary. 42 �Figure 3. Geologic map of Lake Vermilion State Park (after Peterson and Patelke, 2003; Radakovich et al., 2010; Heim et al., 2011; Schwierske et al., in press). 43 �GEOLOGY OF LAKE VERMILION STATE PARK Lake Vermilion State Park contains a variety of supracrustal and intrusive lithological units. Supracrustal rocks that can be observed in the park (Figure 3) include the Lower Member of the Ely Greenstone Formation (both the Fivemile Lake and Central Basalt Sequences), the Soudan Member of the Ely Greenstone Formation, and the Gafvert Lake Sequence of the Lake Vermilion Formation. As well, a wide variety of syn- and post-volcanic intrusive rocks crop out within the park, including diabase, gabbro, diorite, granodiorite, various types of quartz-feldspar porphyries, feldspar-porphyries, and lamprophyre (Peterson and Patelke, 2003; Radakovich et al., 2010; Heim et al., 2011; Schwierske et al., in press). Two northwest-trending faults (which based on detailed mapping (Peterson and Patelke, 2003; Radakovich et al., 2010; Heim et al., 2011) possess higher concentrations of synvolcanic hydrothermal alteration mineral assemblages proximal to the structures) appear to be reactivated synvolcanic structures that offset stratigraphic units in the central and northwestern part of the park. As well, Peterson and Patelke (2003) have identified four major, more-or-less east-west trending shear zones that displace stratigraphy in the southern one-third of the park. The northernmost two shear zones represent the northern and southern limits of the Mine Trend Shear Zone, which extends westward into Soudan Underground Mine State Park, and appears to have played a key role in the development of hematite-rich iron orebodies that were historically mined there. The Mine Trend Shear Zone displaces lithological units higher in the stratigraphic sequence to the east. The southern two shear zones represent the northern and southern edges of the Murray Shear Zone (Peterson and Patelke, 2003). This fault system also displaces rocks higher in the stratigraphic sequence to the east. Rocks sandwiched between the southern edge of the Mine Trend Shear Zone and the northern edge of the Murray Shear Zone are in a structural domain known as the Linking Zone (Peterson and Patelke, 2003). According to Peterson and Patelke (2003), the net slip along the Mine Trend Shear Zone may have been as much as 7 km, whereas the net slip along the Murray Shear Zone may have been as much as 13.8 km (Table 2). Table 2. Calculated displacements among the Mine Trend and Murray Shear zones (Peterson and Patelke, 2003). Ranges of values were calculated geometrically by using the average plunges of lineations associated with the shear zones, and two measured lines of possible correlative stratigraphy offset by the bounding shear zones. See Peterson and Patelke (2003) for further details. 44 �45 Figure 4. Regional stratigraphic correlations across the Vermilion District (after Hudak et al., 2012). �During our field trip, we will be observing exposures of various lithologic units that occur within the Vermilion District. Overall, lithological units observed in Lake Vermilion State Park correlate well with lithological units mapped regionally in the Vermilion District. A diagram (Hudak et al., 2012) illustrating stratigraphic columns from the Soudan Mine Area (in the west), the Fivemile Lake area, and the Twin Lakes area (in the east) is provided in Figure 4. Stratigraphic and intrusive units that occur within Lake Vermilion State Park are described below in order from oldest to youngest units. The Fivemile Lake Sequence, Lower Member of the Ely Greenstone Formation The Fivemile Lake Sequence comprises the lowermost mafic to intermediate and felsic volcanic and volcaniclastic lithologies associated with the Lower Member of the Ely Greenstone Formation. This generally east-west striking, north-topping sequence is dominated by well-vesiculated basaltic to andesitic pillow lavas (Hudak et al., 2007; Peterson et al., 2009; Hudak et al., 2012) that display bun, mattress, and lobe morphologies using the nomenclature of Dimroth et al. (1978). Locally, these pillow lavas display exceptional multiple selveges (Hudak et al., 2002b). Multiple pillow selvedge morphologies have been interpreted as an indication of eruption in shallow water active volcanic settings (Kawachi and Pringle, 1988). Within Lake Vermilion State Park, horizons of Fivemile Lake pillow lavas are up to 1100m thick. Subordinate massive sheet lava flows associated with the Fivemile Sequence have been identified by Hoffman (2007) in one locale near Soudan. As well, numerous relatively thin horizons of massive to bedded basalt tuff, lapilli-tuff, and lapillistone deposits are present in the southwest ¼ of Section 25 in the south-central part of Lake Vermilion State Park. According to Hoffman (2007), these deposits vary from 50-150 meters thick, have a strike length of up to 350 meters, and comprise poorly-sorted and poorlygraded, matrix-supported, thickly to very-thickly bedded mafic pyroclastic deposits containing up to 65% lapilli- (64mm) to block-sized (>64mm) scoria fragments. Hudak et al (2007, 2012) have evaluated the lithogeochemical characteristics of Fivemile Lake Sequence mafic to intermediate lava flows and have found them to be dominantly calc-alkaline to transitional basalt and andesite/basalt using the classification schemes of Ross and Bedard (2009) and Pearce (1996). These rocks also are characterized by significant negative niobium (Nb) anomalies when plotted on primitive mantle-normalized spider diagrams. This suggests derivation of the magmas associated with the Fivemile Lake sequence in an arc-like volcanic terrane. Felsic volcanic and volcaniclastic rocks also occur within the Fivemile Lake Sequence, and crop out within the southern one-third of Lake Vermilion State Park based on mapping completed by Peterson and Patelke (2003). These include coherent and volcaniclastic facies dacitic to rhyolitic lithologies including lava flows, monolithic and heterolithic breccia deposits, and tuff and lapilli- tuff deposits. Lithogeochemically, Hudak (2007) and Hudak et al. (2012) have shown felsic rocks in the Fivemile Lake Sequence to be calc-alkaline to transitional andesites to rhyolites using the classification schemes of Pearce (1996) and Ross and Bedard (2009), respectively. As well, these felsic rocks have trace element characteristics of FI, FII, and FIV rhyolites based on classifications from Hart et al. (2004). One relatively thin horizon (<20 meters thick) of oxide-facies iron-formation identified as being within the Fivemile Lake Sequence has been observed proximal to the northern margin of the Murray Shear Zone approximately 600 meters east of the Lake Vermilion State Park boundary. This iron formation typically occurs as localized, thin (<3m thick) horizons interbedded with Fivemile Lake Sequence pillowed lava flows (Peterson and Patelke, 2003). The various lithofacies comprising the Fivemile Lake Sequence of the Lower Member of the Ely Greenstone Formation are summarized in Table 3. 46 �Lithofacies Associated with the Fivemile Lake Sequence, Lower Member, Ely Greenstone Formation Unit Symbol (Figure 3) Lithofacies FM1a Massive Basalt - Andesite FM1b Pillow Basalt - Andesite FM1h Scoriaceous Basalt –Andesite Tuff and Lapilli-Tuff FM1i Foliated Basalt-Andesite FM2a Coherent Dacite – Rhyolite (lava flows and lava domes) FM2c Felsic Polymict Breccia FM2d Felsic Monolithic Breccia FM2e Dacite-Rhyolite Tuff FM4a Oxide-facies Banded Iron-Formation Table 3. Lithofacies and map symbols associated with lithologies comprising the Fivemile Lake Sequence of the Lower Member of the Ely Greenstone Formation. The Central Basalt Sequence, Lower Member of the Ely Greenstone Formation The Central Basalt Sequence crops out in the east-central part of Lake Vermilion State Park, and is composed of massive and pillowed basalt lava flows, structurally deformed foliated basalt, and local thin (up to 3 meters thick) horizons of Algoma-type banded iron-formation. Within Lake Vermilion State Park, the Central Basalt Sequence varies from approximately 300-1000 meters in stratigraphic thickness. The Central Basalt Sequence mafic lava flows appear to be regionally correlative with basaltic lava flows comprising the Armstrong Lake Sequence in the northernmost parts of the Eagles Nest Quadrangle mapped by Jirsa et al., 2001 (Peterson and Patelke, 2003). The Central Basalt Sequence mafic lava flows can be distinguished from the Fivemile Lake Sequence mafic-intermediate lava flows using the following criteria: 1) the Central Basalt Sequence mafic lava flows commonly comprise exceptionally well-preserved primary volcanic textures - such textures are generally not present in the Fivemile Lake Sequence mafic flows due to destruction of these textures from a combination of synvolcanic hydrothermal alteration combined with recrystallization during greenschistfacies regional metamorphism; 2) the Central Basalt Sequence mafic volcanic rocks tend to be dark green to green colored, whereas the Fivemile Lake Sequence mafic-intermediate volcanic rocks typically vary from gray green to blueish green in color; 3) the Central Basalt sequence mafic volcanic rocks tend to lack amygdules or be sparsely amygdaloidal, whereas the Fivemile Lake Sequence mafic-intermediate volcanic rocks tend to contain abundant amygdules; and 4) to date, multiple selvege pillow lavas have not been identified in the Central Basalt Sequence, whereas they are locally abundant within the Fivemile Lake Sequence. Within Lake Vermilion State Park, the Central Basalt Sequence is composed primarily of pillowed basalt lava flows. These mafic volcanic rocks are medium green to dark green in color and tend to be sparsely amygdaloidal (<5% 2mm-1cm rounded to oval gray quartz-filled amygdules). Dark green, locally exceptionally well-preserved interpillow hyaloclastite deposits, ranging from 1-5cm wide, separate individual pillow structures. Locally these rocks are moderately- to strongly quartz- and epidote-altered. As well, in the southeastern part of the park, hydrothermally altered interpillow hyaloclastite deposits containing abundant andradite garnets have been identified. Massive basalt lava flows (interpreted as sheet flow facies lava flows) are also quite common, and comprise green to dark green, aphyric to sparsely-plagioclase phyric basalt. Foliated basalts locally occur in close proximity to shear zones present in the park. Hudak et al (2007, 2012) have evaluated the lithogeochemical characteristics of mafic lava flows in the Central Basalt Sequence in the vicinities of Needleboy and Sixmile Lakes, which are located 47 �approximately 4-5 kilometers east of the eastern boundary of Lake Vermilion State Park. These researchers have found them to range from calc-alkaline to tholeiitic basalt and andesite/basalt using the classification schemes of Ross and Bedard (2009) and Pearce (1996), respectively. Hudak et al. (2007) first observed that Central Basalt Sequence mafic flows could be divided into two types based on rareearth element characteristics. The first of these types is characterized by calc-alkaline to transitional compositions with arc-like chondrite- and primitive-mantle-normalized rare earth element spider diagrams. The second type comprises tholeiitic basalt characterized by flat chondrite-normalized and primitive mantle-normalized rare earth spider diagrams that are characteristic of mafic volcanic rocks erupted within mid-ocean ridge (MORB) or back-arc basin (BABB) extensional tectonic environments. Detailed mapping indicates that these tholeiitic, MORB/BABB compositions only occur within 200 meters of the lower contact with the overlying Soudan Member Iron Formation. Hudak et al. (2007, 2012) have used both these lithogeochemical results, and results from detailed regional mapping at Lake Vermilion State Park, the Needleboy Lake-Sixmile Lake area, the Twin Lakes area (located approximately 14km east of the eastern boundary of Lake Vermilion State Park), and the Purvis Lake area (on the southern limb of the Tower-Soudan anticline approximately 5km south-southeast of the southern boundary of Lake Vermilion State Park) to indicate that the major hydrothermal event that led to the formation of the Soudan Member Algoma-type iron-formation occurred immediately following the opening of a nascent rift or back-arc basin environment during the youngest part of Central Basalt Sequence mafic volcanism. Hudak et al. (2002b) and Hoffman (2007) have identified several localized occurrences of felsic coherent and volcaniclastic strata within the Central Basalt Sequence to the south and east of Lake Vermilion State Park. Hudak et al. (2007, 2012) have evaluated the lithogeochemical characteristics of these rocks, and have found them to be calc-alkaline andesite to rhyolite/dacite using the classification schemes of Ross and Bedard (2009) and Pearce (1996), respectively. As well, these felsic rocks have trace element characteristics of FII and FIIIa rhyolites based on classifications from Hart et al. (2004). The various lithofacies comprising the Central Basalt Sequence of the Lower Member of the Ely Greenstone Formation are summarized in Table 4. Lithofacies Associated with the Central Basalt Sequence, Lower Member, Ely Greenstone Formation Unit Symbol (Figure 3) Lithofacies Cb1a Massive Basalt Cb1b Pillow Basalt Cb1i Foliated Basalt Cb1u Undivided Basalt Cb2eh Polymict Dacite-Rhyodacite Tuff and Lapilli-Tuff Cb2e Dacitic-Rhyodacite Tuff and Lapilli-Tuff Cb2f Felsic Epiclastic Deposits Cb4a Interbedded Oxide-facies Banded Iron-Formation and Basalt Table 4. Lithofacies and map symbols associated with lithologies comprising the Central Basalt Sequence of the Lower Member of the Ely Greenstone Formation. The Soudan Member of the Ely Greenstone Formation The Soudan Member of the Ely Greenstone formation is dominantly composed of laminated to thinly bedded Algoma-type oxide facies banded iron-formation, with subordinate, locally interstratified, sparsely amygdaloidal massive to pillowed basalt lava flows and resedimented felsic tuff deposits. Regionally, the stratigraphic thickness of the Soudan Member of the Ely Greenstone Formation varies 48 �from 50-3,000 meters, with an average stratigraphic thickness of approximately 700 meters (Peterson et al., 2009). Within Lake Vermilion State Park, the Soudan Member ranges in stratigraphic thickness from approximately 300 – 680 meters in thickness. Individual horizons of oxide-facies iron formation range from approximately 70-345 meters thick, whereas the Soudan basalt lava flow units range from approximately 60-300 meters in thickness. A gradational contact over several tens of meters to two hundred meters occurs between the underlying Central Basalt Sequence rocks and the overlying oxide facies iron-formations of the Soudan Member. This transitional zone is characterized by a decrease in abundance of basalt lava flows and associated volcaniclastic rocks, and an increase in the abundance and thickness of oxide-facies ironformation horizons, as one moves toward the basal contact of the Soudan Member (Hudak et al., 2002b; Peterson and Patelke, 2003; Hudak et al., 2007; Hoffman, 2007; Hudak et al., 2012). Several characteristics suggest that the Soudan Member was deposited in relatively quiet water in a relative deep subaqueous environment (>200m, probably greater than 1400 m). This evidence includes: 1) a lack of primary mafic or felsic pyroclastic deposits within the stratigraphic sequence; 2) a lack of multipleselvege pillow lavas in the stratigraphic sequence; 3) planar laminations and bedding combined with an absence of any wave-associated sedimentary bedforms within both the chemical and clastic sedimentary rocks within the sequence; and 4) lithological and geochemical evidence for the development of an extensional tectonic environment that resulted in deepening of the depositional environment in the uppermost sections of the stratigraphically underlying Central Basalt Sequence. Within Lake Vermilion State Park, the Soudan Member oxide-facies banded iron-formation is planar laminated to medium-bedded, with black magnetite-rich horizons, light gray to black chert horizons, red to blueish-black hematite-rich horizons, and red jasper horizons defining the bedding. Locally, very tight, chaotically oriented folds, resulting from syn-depositional soft sediment deformation and subsequent tectonic deformation, are present. These iron formation deposits can be intimately interbedded with basalt lava flows such that mapping individual iron-formation and basalt horizons is impossible at 1:5000 scale. Where this occurs, these rocks have been mapped as a stratigraphic unit called “Basalt and IronFormation” by Peterson and Patelke (2003). Basalt lava flows associated with the Soudan Member of the Lower Ely Greenstone are characterized by a medium green to dark green color. They are aphyric to sparsely plagioclase ± pyroxene (now actinolite)-phyric. Plagioclase phenocrysts are present in abundances up to 3%, are typically less than or equal to 1mm in length, and vary from subhedral to euhedral tabular in morphology. Locally, 5-7% dark green actinolite pseudomorphs of pyroxene phenocrysts may be present. Where amygdaloidal, the unit contains up to 7% oval to round, light gray quartz-filled amygdules ranging from <1-4mm in diameter. The various lithofacies comprising the Soudan Member of the Ely Greenstone Formation are summarized in Table 5. Lithofacies Associated with the Soudan Member, Ely Greenstone Formation Unit Symbol (Figure 3) Lithofacies S1a Massive Basalt S2eq Aphyric- to Quartz-phyric Rhyodacite Tuff S4a Oxide Facies Banded Iron-Formation Table 5. Lithofacies and map symbols associated with lithologies comprising the Soudan Member of the Ely Greenstone Formation. The Gafvert Lake Sequence, Lake Vermilion Formation The Gafvert Lake Sequence (mapped as the “Upper Sequence” by Peterson and Patelke, 2003; Radakovich et al., 2010: and Heim et al., 2011) comprises dacitic to rhyodacitic volcaniclastic and epiclastic rocks that are locally interbedded with Algoma-type banded iron-formation and chert deposits. 49 �This sequence, which is part of the Lake Vermilion Formation, has been found to unconformably overlie the Soudan Member of the Ely Greenstone Formation in the north-central and northwestern parts of Lake Vermilion State Park based on recent mapping and geochronological evidence reported by Lodge et al. (2013). Within Lake Vermilion State Park, the overall stratigraphic thickness of the Gafvert Lake Sequence is up to approximately 1300 meters thick, with individual felsic volcaniclastic deposits having stratigraphic thicknesses ranging from approximately 75 – 400 meters thick, and individual Algoma-type oxide facies banded iron formations and associated massive- to bedded chert deposits ranging from 25250 meters and up to 175 meters in stratigraphic thickness, respectively. To the wes,t in Soudan Underground Mine State Park, the Gafvert Lake Sequence is locally interlayered with, and overlain by, greywacke deposits associated with the Lake Vermilion Formation. Within Lake Vermilion State Park, several lithofacies comprise the Gafvert Lake Sequence. The basal member of this sequence comprises massive, very-thickly bedded, quartz- and plagioclase-phyric polymict dacitic to rhyodacitic tuff, lapilli-tuff, and tuff-breccia deposits. These light gray, non-sorted, non-graded, matrix-supported deposits contain 3-8% 1-2mm (rare 3mm) pale gray anhedral to subhedral quartz phenocrysts, 10-15% <1-2mm subhedral to euhedral tabular plagioclase phenocrysts, and a wide variety of lapilli- to block-sized clasts including: 1) 10-20% 1-10 cm quartz- and plagioclase-phyric coherent dacite to rhyodacite lapilli and blocks; 2) 5-7% <3cm diameter pale gray-green lens-shaped, locally quartz- and plagioclase-phyric pumice lapilli; 3) up to 1% dark gray to light gray angular chert lapilli ranging from 0.5-3cm in diameter; and 4) 1-3% 0.5-5cm dark gray to black to red magnetite-rich, hematite-rich, or jasper-rich banded iron formation lapilli. These deposits are overlain by, and interbedded with, light gray, matrix-supported, non-sorted and non-graded quartz- and plagioclase-phyric dacitic to rhyodacitic tuff deposits which contain 10-25% 1-3mm subhedral to euhedral tabular plagioclase phenocrysts, 1-3% 1-3mm subhedral to anhedral, commonly broken, quartz phenocrysts, as well as 10-15% subangular quartz- and plagioclase-phyric coherent dacite to rhyodacite lapilli and up to 5% locally quartz- and plagioclase-phyric pumice lapilli. Spectacular felsic epiclastic deposits comprising polymict volcaniclastic conglomerates and lithic sandstones are also present in the Gafvert Lake Sequence and crop out west of Lake Vermilion State Park in Stunz Bay (Radakovich et al., 2010). In addition to felsic volcaniclastic and epiclastic rocks, two types of chemical sedimentary rocks have also been identified in the Gafvert Lake Sequence. These include laminated to medium-bedded Algomatype banded iron formation that varies from red (hematite- and jasper-rich) to dark gray (magnetite-rich) to light gray (chert-rich) in color. Immediately west of the Lake Vermilion State Park boundary, light gray to black laminated to very thickly bedded black chert deposits are present. These chert deposits may represent the distal facies equivalent of Algoma-type banded iron-formation horizons that are present in the northeast part of Lake Vermilion State Park south of Cobble Bay. A limited number of Gafvert Lake Succession felsic volcaniclastic rocks have been studied by lithogeochemical means by Geoff Pignotta and Kelly Schwierske at the University of Wisconsin Eau Claire (Figure 5). These researchers (Schwierske et al., in press) have found that the volcaniclastic and epiclastic deposits associated with the Gafvert Lake Sequence consistently plot as rhyodacite/dacite in composition when using the immobile element lithological classification scheme of Winchester and Floyd (1977). These compositions are very similar to the composition of a quartz- ± plagioclase-phyric rhyodacite sill that crops out in the northeastern part of Lake Vermilion State Park (see Field Trip Stop 9 below), although the sill has consistently higher Nb/Y ratios than the Gafvert Lake volcaniclastic and epiclastic rocks. This sill may represent a synvolcanic intrusion that is genetically related to the evolution of the Gafvert Lake sequence based on this lithogeochemical evidence, as well as field evidence from Peterson and Jirsa (1999), which indicates that this intrusion is most commonly observed within the Gafvert Lake Sequence where it is thickest, in proximity to Gafvert Lake west of the Mud Creek Road. Peterson (2001) has hypothesized that this intrusion represents feeder intrusions to a Gafvert Lake Sequence stratovolcano located in this area. The various lithofacies comprising the Gafvert Lake Member of the Lake Vermilion Formation are summarized in Table 6. 50 �Figure 5. Chemical classification of various lithologies within Lake Vermilion State Park (Schwierske et al., in press) using the classification scheme of Winchester and Floyd (1977). Open triangles represent samples from a quartz- ± plagioclase-phyric rhyodacite/dacite sill in the northeastern part of Lake Vermilion State Park. The black squares, large black diamonds, and small black diamonds represent various Gafvert Lake Succession volcaniclastic and epiclastic rock units. Lithofacies Associated with the Gafvert Lake Sequence, Lake Vermilion Formation Unit Symbol (Figure 2-3) Lithofacies US1,4 Interbedded Basalt and Oxide-facies Banded Iron-Formation US2b Dacite-Rhyodacite Tuff-breccia US2cf/US2f Dacite-Rhyodacite Epiclastic Deposits US2e Quartz- + Plagioclase-phyric Dacite-Rhyodacite Tuff/Lapilli-tuff US2eh Polymict Dacite-Rhyodacite Tuff/Lapilli-tuff US4a Oxide-facies Banded Iron-Formation Table 6. Lithofacies and map symbols associated with lithologies comprising the Gafvert Lake Sequence of the Lake Vermilion Formation. 51 �Intrusive Rocks Eight types of intrusive bodies have been mapped within the boundaries of Lake Vermilion State Park (Peterson and Patelke, 2003; Hoffman, 2007; Radakovich et al., 2010; Heim et al., 2011; Figure 3). From oldest to youngest, these intrusions include:  Gabbro (mapped by Peterson and Patelke, 2003; Hoffman, 2007; Radakovich et al., 2010; Heim et al., 2011, unit Gb) – Identified as sills throughout the central one-third of Lake Vermilion State Park, this unit is characterized by grayish-green to black, medium-grained equigranular gabbro that is locally highly magnetic and displays ophitic texture.  Diabase (mapped by Peterson and Patelke, 2003; Hoffman, 2007, unit Db) – Identified in the southern one-third of Lake Vermilion State Park, this unit comprises black to dark green, finegrained plagioclase-phyric diabase dikes and sills that have been interpreted to represent feeder dikes to mafic volcanic rocks located stratigraphically up-section.  Coarsely porphyritic Quartz-Feldspar Porphyry (mapped by Heim et al., 2011, unit GLIC) – Identified in the northeastern part of Lake Vermilion State Park, this intrusion comprises light gray, massive, quartz ± plagioclase-phyric coherent rhyodacite. The light gray aphanitic groundmass contains 3-7% gray to light blue subhedral rounded to euhedral square quartz phenocrysts that range from 3-10mm in diameter, and 10% pale gray to tan, subhedral to euhedral tabular plagioclase phenocrysts ranging from 1-4mm in length. Similar intrusive rocks have been mapped in the vicinity of Needleboy and Sixmile Lakes by Hudak et al. (2002b), and near Gafvert Lake by Peterson and Jirsa (1999) and Peterson (2001).  Diorite (mapped by Peterson and Patelke, 2003; Hoffman, 2007, unit D) – Occurs as a generally east-west striking sill in the southern one-third of Lake Vermilion State Park. Composed of gray to gray-green, fine- to medium-grained, equigranular diorite. This unit was informally named the “Sugar Mountain Diorite” by Peterson and Patelke (2003), and is notable for its massive, indurated nature and lack of prominent joints, veins, and alteration.  Granodiorite (mapped by Peterson and Patelke, 2003; Hoffman, 2007, unit Gd) – Identified in the central part of Lake Vermilion State Park, and composed of whitish-pink to gray-green, fine- to medium-grained, commonly xenolith-rich granodiorite and locally hornblende granodiorite.  Quartz-Feldspar Porphyry (mapped by Peterson and Patelke, 2003; Hoffman, 2007; Radakovich et al., 2010; Heim et al., 2011, unit Qfp) - Found locally throughout Lake Vermilion State Park, this intrusion comprises a light gray to pale green-gray groundmass that contains 20-25% 1-3mm (locally up to 5mm) subhedral to euhedral tabular plagioclase phenocrysts and 7-12% 1-3mm (locally up to 5mm) subhedral to euhedral gray-blue quartz phenocrysts  Feldspar Porphry (mapped by Peterson and Patelke, 2003; Hoffman, 2007; Radakovich et al., 2010; Heim et al., 2011, unit Fp) – Identified in the south and central parts of Lake Vermilion State Park, this intrusion is white to pink in color, and contains subhedral rounded to euhedral tabular 4mm feldspar phenocrysts and locally, subhedral to euhedral prismatic to tabular actinolite pseudomorphs of hornblende phenocrysts.  Lamprophyre (mapped by Peterson and Patelke, 2003, unit L) – Located in the southwestern part of Lake Vermilion State Park, this intrusion is characterized by black, fine-grained, massive hornblende-feldspar rock that contains 10-15% fine hornblende needles in a gray-black to red matrix, as well as large (>25cm) rounded granite and supracrustal rock xenoliths. 52 �TERMINOLOGY OF VOLCANICLASTIC ROCKS It is important to note the terminology utilized in this field trip guide for: 1) volcaniclastic rocks; and 2) bedding characteristics. Use of consistent terminology is required in order to accurately describe these geological features. Volcaniclastic rocks contain abundant volcanic material irrespective of their origin or depositional environment. Such rocks can be formed directly from volcanic eruptions (whether subaerial or subaqueous), result from resedimentation of non-lithified volcanic deposits (for example, resedimentation of pyroclasts prior to lithification), or result from weathering and resedimentation of pre-existing lithified volcanic rocks. Primary (juvenile) volcaniclastic particles result directly from eruptive processes, and are of three types:  Pyroclasts, which form by explosive fragmentation of magma into particles (including ash, highly vesiculated glass (pumice, scoria), crystals and crystal fragments, and lithic fragments);  Hydroclasts, which form by explosive interaction with external water (via phreatic (steam only) and/or phreatomagmatic (steam and magma) explosions) or by non-explosive quenching and granulation of lava (for example, the formation of hyaloclastite fragments on the margins of submarine lava flows or intrusions into wet sediments); and  Autoclasts, which form by frictional breakage of moving viscous lava flows (for example, to form carapace breccias on the margins of subaerial lava flows). Based on these different types of fragmentation, four types of primary volcaniclastic deposits have been identified by White and Houghton (2006):  Pyroclastic deposits, which are generated from volcanic plumes and jets or pyroclastic density currents as particles first come to rest. Deposition mechanisms associated with these processes include suspension settling, traction, or en masse freezing;  Autoclastic deposits, which are generated during effusive volcanism when lava cools and fragments as a result of thermal processes, or recently cooled lava breaks during flow. Deposition for these types of rocks is under the influence of continued lava flowage;  Hyaloclastite deposits, which are generated during effusive volcanism when magma or flowing lava is chilled and fragmented as a result of contact with water. Deposition of such deposits is under the influence of the continued emplacement of the lava in the presence of water; and  Peperite deposits, which are generated when magma intrudes into unconsolidated clastic material and mingles with (generally wet) debris to form a volcaniclastic deposit. Deposition of peperite deposits takes place essentially in-situ. Secondary volcaniclastic particles are known as epiclasts:  Epiclasts are lithic clasts and/or crystals derived from physical weathering and erosion of preexisting rocks. Epiclasts are volcaniclasts when the pre-existing rocks are volcanic. In recent years, the terminology for volcaniclastic rocks has become increasingly confusing because different classification schemes (for example Fisher, 1961; Fisher 1966; Schmid, 1981; Cas and Wright, 1987; McPhie et al., 1993; White and Houghton, 2006) are preferentially used in different parts of the world, and terminology relating to volcaniclastic rocks is commonly misused. Four classification schemes have been used most commonly in the recent geological literature:  Fisher (1961, 1966) – Classification based on particle size, particle formation, or particle fragmentation mechanism;  Schmid (1981) – Particle type within the deposit;  Cas and Wright (1987) – Mode of fragmentation and deposition; and  McPhie et al. (1993) – Transport and deposition mechanisms. According to R. V. Fisher (1998), the difficulties with volcaniclastic rock classification can be understood because “volcaniclastic rocks are essentially igneous on the way up and sedimentary on the way down”. 53 �In fact, Fisher’s thesis advisor, when observing the volcaniclastic rocks that were the focus of his thesis studies, indicated that they were “the ugliest and most undistinguished rocks I’ve seen in my 30 years of petrology!” As well, classification is especially difficult in ancient volcaniclastic rocks because key aspects of classification can be obscured by subsequent metamorphism and/or structural deformation (e.g. particle type, particle size) or because genetic processes cannot be ascertained unambiguously (e.g. transport and deposition mechanism, fragmentation mechanisms). Figure 6. Volcaniclastic rock classification schemes of Fisher (1966) and White and Houghton (2006). This field trip guidebook will classify volcaniclastic rocks using Fisher’s (1966) classification scheme. For this field trip guidebook, we will utilize Fisher’s (1966) classification (Figure 6) for volcaniclastic rocks. This classification scheme is based on the relative proportions of ash-sized material (< 2mm), lapilli-sized material (64mm), and blocks/bomb sized material (>64mm) in the rock. Both Gibson et al. (1999) and Mueller and White (2004) suggest that this classification be used for field-based rock classification (mapping, diamond drill core logging, petrography) of ancient volcaniclastic deposits for the following reasons:  The classification scheme is “field-user friendly” because it accommodates both the historically important pyroclastic rock names and enables comparison at both the hand sample and thin section scale (Mueller and White, 2004);  It is a Wentworth-based scale, and thus enables comparison of volcaniclastic deposits to sedimentary deposits; and  Rock classification does not require knowledge of the specific transport mechanism or depositional processes involved with the genesis of the deposit. More recently, White and Houghton have developed a modified version of Fisher’s (1966) volcaniclastic classification scheme (Figure 6). The scheme is essentially equivalent to the Fisher (1966) scheme, with the exception that the lapill-tuff field in the White and Houghton (2006) classification comprises the lapilli-tuff and lapillistone fields of Fisher’s (1966). 54 �Specific terms for bedding thicknesses are also used in this guidebook The terms used, and their bedding thickness characteristics, have been adopted from McPhie et al. (1993) and include:  Laminated <1 centimeters thick  Very thinly bedded 1-3 centimeters thick  Thinly bedded 3-10 centimeters thick  Medium bedded 10-30 centimeters thick  Thickly bedded 30-100 centimeters thick  Very thickly bedded >100 centimeters thick ROAD LOG AND FIELD TRIP STOPS All stop locations for this field trip are given in Universal Transverse Mercator (UTM) coordinates, Zone 15N, using the North American Datum of 1983 (NAD83). Section subdivisions read from smallest to largest quarter; e.g., “NW, SE” should be read “NW quarter of the SE quarter.” The small topographic map insets illustrating field trip stop locations have been taken from the Tower and Soudan USGS 7.5minute quadrangle maps. A selected number of field trip stops will take place outside Lake Vermilion State Park, with the majority of the stops taking place along a 4.5 mile traverse through the state park. More detailed geological maps with stop locations are given in Figures 7 and 8 later in this guidebook. From Hibbing, our field trip route will proceed north and east from the Hibbing Park Hotel along Minnesota Highway 169 North. We will make one stop at a spectacular outcrop displaying various flow facies of Central Basalt flows that are located south of Lake Vermilion State Park on the south side of Highway 169 North (Figure 7). We will then proceed into Lake Vermilion State Park (Figure 8) where we will observe an outcrop of Neoarchean gabbro just southeast of the Old Ely Road. After a coffee break, we will strap on our hiking boots and make several field trip stops at outcrops along an approximately 4.5 mile traverse through Lake Vermilion State Park. All major stratigraphic units in the park will be observed during this traverse. We will then leave Lake Vermilion State Park, and start our return to Hibbing via Highway 169 South. We will make one final field trip stop on the south side of Highway 169 South just west of Tower to investigate a recent road cut comprising Gafvert Lake Sequence rhyodacite tuffs, lapilli tuffs, and tuff breccias (Figure 7) prior to returning to the Hibbing Park Hotel via Highway 169 South. Mileage for this roadlog starts at the intersection of East Howard Street and Highway 73/169 North. Roadlog (vehicle) mileage will be denoted in bold italic text. Mileage for the traverse through the park will be denoted in italic text. Bus Log 0.0 miles 22.0 miles 26.6 miles 54.1 miles Turn north on Highways 73/169 North and proceed to Virginia, Minnesota. Turn left on to Highways 53/169 North. Veer right at the exit for Ely, Minnesota on Highways 1/169 North. You will pass the “Y” store at approximately 44.2 miles, and the intersection for Highway 135 on the west side of Tower at approximately 48.3 miles. Continue on Highway 1/169 North through Tower. At approximately 50.2 miles you will see the intersection of Main Street, Soudan, Minnesota, and a sign for Soudan Underground Mine State Park. Continue on Highway 1/169 past Soudan. At approximately 53.4 miles you encounter the intersection of Highway 1/169 North and the Murray Forestry Road (on the south side of Highway 1/169). Get prepared to turn south off of Highway 1/169 on to a dirt road in approximately 0.6 - 0.7 miles. Turn right (south) on to the dirt road and immediately park in the open area at the base of the hill. Hike 0.15 miles (approximately 240 meters) up the hill on the dirt road to Field Trip Stop 1. 55 �Figure 7. Map illustrating regional geology in the vicinity of Lake Vermilion State Park. Field trip stops 1 and 2 are outside the state park boundaries and are illustrated. The location of Figure 8, a more detailed map illustrating field trip stops in the state park, is illustrated by the bold black box. 56 �Figure 8. Detailed geologic map of Lake Vermilion State Park (after Peterson and Patelke (2003), Radakovich et al. (2010) and Heim et al. (2011)). Field trip stops within the park are labeled. 57 �Stop 1: Central Basalt Sequence Sheet Flows, Pillow Lavas, and Perlitic Hyaloclastite Location: T. 62N, R.14W, sec. 19, SE, SW, Soudan 7.5-minute quadrangle UTM: 562,000E / 5,297,805N This classic outcrop has been visited during field trips associated with both the 2004 and 2009 ILSG conferences (Hudak et al., 2004; Peterson et al., 2009). This is a no-hammer outcrop, as the preservation of the delicate textures here rivals those observed in other classic Neoarchean camps in the Superior Province containing well-preserved volcanic textures such as Noranda, Quebec and Timmins, Ontario. The description and figure below has been taken from Peterson et al. (2009). The Central Basalt sequence (Peterson and Patelke, 2003) comprises a steeply north-dipping (75°vertical), north-facing sequence of sparsely amygdaloidal pillowed and massive lava flows of basalt andesite to basalt composition that are believed to be correlative with the tholeiitic Armstrong Lake volcanic sequence mapped in the Eagles Nest quadrangle (Jirsa et al., 2001), approximately 11km to the east. Hudak et al. (2007), Jansen et al. (2007), and Hudak et al. (2012) have shown that the lowermost sections of the Central Basalt Sequence are composed of submarine basaltic andesite to basalt lava flows that have rare earth element lithogeochemical patterns similar to mafic rocks in oceanic volcanic arcs. However, locally, submarine basalt lava flows that occur within 50-200m stratigraphically below the contact between the Central Basalt Sequence and the overlying Soudan Member of the Ely Greenstone Formation illustrate MORB-like or back-arc basin-like lithogeochemical patterns. This change in rare earth element characteristics may be interpreted to indicate a change from an oceanic arc to back-arc environment immediately prior to the deposition of the Soudan Member. Relative to massive and pillowed basalt and andesite flows in the Fivemile Lake sequence, Central Basalt sequence lava flows are notably less amygdaloidal, and lack multiple pillow rind structures. In addition, the Central Basalt sequence lacks the thick sequences of scoriaceous basalt-andesite lapilli tuffs that are commonly interstratified with lava flows in the Fivemile Lake sequence. These characteristics of the Central Basalt sequence indicate eruption and deposition in a deeper submarine environment than the stratigraphically older Fivemile Lake sequence, and suggest overall increasing water depth during the temporal development of the Lower Ely. Deepening of the water column could be accommodated by extensional tectonics and normal faulting associated with the development of the proposed back-arc environment. The outcrop comprises two east-southeast striking massive basalt flows, ranging from at least five to nine meters in thickness, that are separated by a ten meter thick flow unit comprising pillows and pillow lobes (Fig. 9). All three lava flows at this vicinity illustrate tholeiitic, MORB-like lithogeochemistries (Hudak et al., 2007). 58 �Figure 9. Detailed geological map of sheet flows, pillow lavas, and associated hyaloclastite deposits at field trip Stop 1. Flow 1, at the southern part of the outcrop, is composed of a pale- to dark green, faintly feldspar-phyric (~10% 0.5-1 mm laths), sparsely amygdaloidal, basalt sheet flow that locally exhibits tortoise-shell jointing formed in response to contraction during cooling. The uppermost 10-40 cm of the coherent part of Flow 1 is generally silicified and epidotized. Petrographic observations indicate that this section of the flow also contains up to 70% <0.1 cm round spherulites. An irregular contact occurs between the coherent basalt flow and an overlying one- to two meter thick unit of dark green, exceptionally well-preserved perlitic in-situ hyaloclastite and associated self-peperite (c.f. Batiza and White, 2000). The hyaloclastite formed from non-explosive fracturing of the basalt glass developed on the flow top due to quenching by 59 �water, whereas the perlite formed following deposition by hydration of volcanic glass. An irregular contact occurs between the hyaloclastite and Flow 2, which is composed of north-facing mattress- to bunshaped pillow lavas and pillow lobes with numerous “neck and knob” structures. Individual Pillow structures have well developed perlitic hyaloclastite margins that range from 1-4 cm in width. Pillow buds indicate propagation from east to west, suggesting the volcanic vent was located east of this location. The coherent pillows and lobes are overlain by up to 2.5 meters of hyaloclastite breccia that contains 20-40% subround to subangular pale gray green basalt lapilli in a jigsaw puzzle-fit dark green perlitic hyaloclastite matrix. The upper contact of Flow 2 and the overlying basalt sheet flow (Flow 3) is irregular, and is marked by thin (1-8 cm thick), sheet- like basalt fragments that are up to 1.6 meters in length. These fragments locally appear to be isoclinally folded about an east-west-trending fold hinge. Although the genesis of this structure is currently not well understood, it may be due to syneruptive deformation of either thin slabs of hot, basal flow margin crust from the overlying flow, or thin injections of basalt magma into the hyaloclastite from either the underlying pillows or the overlying sheet flow. Flow 3 comprises an at least ten-meter thick pale green-gray, slightly feldspar-phyric, sparsely amygdaloidal sheet flow. Steep, NNE-trending west dipping D3 joints are well developed in this unit, as are lens-shaped pseudo-pillows that are up to 50 cm in diameter. Return to the bus by walking back down the hill. 54.1 miles 54.7 miles 54.8 miles 56.1 miles 56.4 miles 56.9 miles Turn left and follow Highway 1/169 South approximately 0.6-0.7 miles to the west. You will see the Murray Forestry Road on your left (south side of road) Turn right (north) on the dirt road immediately north of the intersection with Highway 1/169 South and the Murray Forestry Road. Proceed north approximately 0.1 miles. Park the vehicle on the Old Ely Road immediately outside the gate to Lake Vermilion State Park. Walk approximately 0.3 miles northeast on the Old Ely Road, then approximately 0.04 miles southwest along the trail to Field Trip Stop 2. Drop off field trip participants on Old Ely Road immediately east of the gate to Lake Vermilion State Park. After dropping off field trip participants, the bus will drive northeast up the Old Ely Road for 1.32 miles. Make sharp left turn on dirt road farthest to the south. Continue west on dirt road approximately 0.28 miles to the gate at Lake Vermilion State Park. Enter Lake Vermilion State Park through gate. Proceed 0.48 miles west, where you will see a two-track trail on the south of the road immediately before the well-maintained road turns right sharply to the north. Park the bus in the grass on the south side of the road at the intersection of the well-maintained road and the two-track trail. Park Bus. 60 �Stop 2: Neoarchean Gabbro Sill Location: T. 62N, R. 14W, sec. 19, SW, SW, Soudan 7.5-minute quadrangle UTM: 561,385E / 5,297,735N Detailed geologic mapping in the Vermilion District (Peterson, 2001; Hudak et al., 2002a, Hudak et al., 2002b; Hudak et al., 2006) has indicated the presence of several gabbro/diabase dikes and sills within both the Fivemile Lake and Central Basalt sequences of the Lower Member of the Ely Greenstone Formation. These intrusive rocks vary from fine-grained diabase with well-developed trachytic textures, to medium- to coarse-grained gabbro and quartz gabbro that locally display well developed ophitic and sub-ophitic textures. Petrographic observations indicate the presence of relatively unaltered (minor sericite ± carbonate alteration) euhedral to subhedral plagioclase and subhedral to anhedral actinolite pseudomorphs of original clinopyroxene. At this location, we will observe dark green to dark greenish-black medium- to coarse-grained gabbro that locally displays exceptional sub-ophitic and ophitic textures. Return to the bus for our morning coffee break and a brief explanation of our traverse through the park We will now begin our traverse through Lake Vermilion State Park. Make sure to have proper field gear (hat, rain gear, etc), your lunch, and drinks along with you, as we will be out on the trails in the park for nearly the remainder of the trip. Per state park rules, no hammering on the outcrops, or taking of samples, will be allowed while we are on the traverse. Traverse Log 0.0 miles 0.35 miles Leave the bus and go through the gate at the entrance to Lake Vermilion State Park for 0.26 miles. Turn south along the dirt road/trail and proceed approximately 0.9 miles along the trail to Stop 3. Stop at outcrop on ridge of hillside. Stop 3: Garnet-altered Central Basalt Sequence Pillow Lavas Location: T. 62N, R. 15W, sec. 25, SW, NE, Soudan 7.5-minute quadrangle UTM: 560,670E / 5,297,210N In several locations in the vicinities of Sixmile Lake (Hudak et al., 2006) and Twin Lakes (Moosavi et al., 2007), mafic volcanic and volcaniclastic rocks in the Central Basalt Sequence have been intensely altered to form mineral assemblages comprising quartz, epidote (both pistacite and zoisite/clinozoisite), actinolite, sericite, and/or chlorite (both Mg-rich and Fe-rich compositions). Locally, these altered rocks also 61 �contain minor to moderate abundances of subhedral to euhedral, dark reddish-brown garnets that have been identified as andradite via x-ray diffraction analysis (andradite chemical formula is Ca 3Fe2Si3O12, a member of the ugrandite garnet series (Phillips and Griffen, 1981, p. 117)). At this location, you will observe hydrothermally altered Central Basalt Sequence pillow lavas that contain an abundance of andradite garnet in hydrothermally-altered interpillow hyaloclastite deposits. Recognition of mineral phases that are not consistent with greenschist-facies metamorphism of original basalt composition protoliths is essential to identifying hydrothermal alteration zones. Detailed mapping of such alteration mineral assemblages provides important data regarding the processes and timing of hydrothermal alteration, which in turn, provide essential clues to economic geologists in their quests to find mineralization. Return to Old Ely Road along the same path used to access Stop 3. 0.45 miles 0.65 miles 1.10 miles At the intersection of the trail leading to Stop 3 and Old Ely Road, turn left (west) and walk approximately 0.2 miles to the intersection of a two-track road that leads to the north. Turn north on the two-track road and proceed north-northeast. Walk approximately 0.27 miles to the first major curve in the road. It is likely to be a bit wet here, so plan to continue walking north-northeast along the edge of the trail. Continue for another approximately 0.27 miles northeast along the two-track road. You will encounter a series of five outcrops along the two-track road that will extend for a distance of approximately 0.1 miles. This sequence of outcrops is Stop 4. Stop 4a: Central Basalt Sequence Pillow Lavas Location: T. 62N, R. 15W, sec. 24, SE, SW, Soudan 7.5-minute quadrangle UTM: 560,440E / 5,297,700N At this location, a series of six small outcrops occurs as one traverses approximately 125 meters up a small hill from southwest to northeast. Here you will observe well-preserved, commonly muffin-shaped, sparsely vesicular variably altered Central Basalt Sequence pillow lavas. In several locations, dark reddish-brown sulfide burn can be observed where sulfide minerals (pyrite, locally minor chalcopyrite) have been oxidized. Near the central part of the outcrop exposure (outcrop number four as one procedes from south to north), locally strong silicification and actinolite alteration may be observed. 1.2 miles 1.3 miles From the northernmost outcrop associated with Stop 4, proceed north approximately 0.1 miles. You will see a series of small outcrops that extend for approximately 0.1 miles along the east side of the two-track road. 62 �Stop 4b: Hydrothermally Altered Central Basalt Sequence Pillow Lava Location: T. 62N, R. 15W, sec. 24, SE, SW, Soudan 7.5-minute quadrangle UTM: 560,525E / 5,297,860N This stop comprises a series of east-northeast striking, north-topping, locally silicified and actinolite-altered Central Basalt Sequence basalt to basaltic-andesite pillow lavas that extend for approximately 0.1 miles as one traverses toward the north. Note the brown stains within both the interpillow hyaloclastite deposits and the cores of the pillow lavas that result from weathering of minor amounts of pyrite in the rock. 1.4 miles 1.75 miles 1.77 miles From the northernmost outcrop exposure, continue walking northwest along the twotrack road. Gather on the two-track road, and follow the field trip leader approximately 0.02 miles (~30 meters) through the bush. Gather on the north slope of the north-south trending outcrop. This is Stop 5. Stop 5: Contact Between Soudan Member Banded Iron Formation and Soudan Basalt Location: T. 62N, R. 15W, sec. 24, NW, SW, Soudan7.5-minute quadrangle UTM: 560,165E / 5,298,240N Mapping by Jirsa et al. (2001), Peterson and Patelke (2003), Radakovich et al. (2010) and Heim et al. (2011) has shown that the Soudan Member of the Ely Greenstone Formation is composed of Algoma-type oxide-facies banded iron formation horizons interbedded with massive and pillowed basalt lava flows, aphyric- to quartz-phyric rhyolite tuffs, and locally, polymict quartz- and plagioclase-phyric dacitic to rhyodacitic lapilli tuff deposits. To the south and east, as well as within the 2700 drift of the Soudan Underground Mine, shearing of the interbedded iron formation and basalt horizons has resulted in a rock comprising chaotically intermingled chlorite schist and banded iron formation which Peterson and Patelke (2003) most appropriately termed “Schist ‘n’ BIF”. The Algoma-type iron formations of the Soudan Member comprise laminated- to medium-bedded iron formation containing dark gray to black magnetite-rich bands, bluish-gray to red hematite-rich bands, red jasper bands, and light gray to black chert bands. Planar bedding is most common, with tight, commonly chaotic folds present that have been, in part, interpreted to be the result of soft sediment deformation. The unit is typically strongly magnetic, but is locally moderately to weakly magnetic where dominated by hematite-rich horizons or chert horizons. The Soudan Basalts comprise medium-green to dark green, aphyric to sparsely plagioclase- ± pyroxenephyric massive to amygdaloidal basalt. Typically, the recrystallized matrix (now chlorite-epidoteactinolite) contains up to 3% <1mm subhedral to euhedral tabular plagioclase phenocrysts and locally, 5- 63 �7% <1mm dark green actinolite pseudomorphs of pyroxene phenocrysts. Locally, amygdaloidal basalt flows contain 5-7% oval to round, light gray to white, quartz- ± epidote- ± chlorite-filled amygdules ranging from <1-4mm in diameter. Locally, brownish-tan colored ankerite alteration and dark green chlorite alteration are present. Figure 10. Detailed (1:5000 scale) map illustrating the complex contact relationships between Soudan Member oxide facies iron formation and basalt units in the vicinity of Stop 5. Figure 10 is a reproduction of a detailed field map (originally mapped at 1:5000 scale) in the general vicinity of Stop 5. The map illustrates the complex contact relationships between Soudan Member banded iron formation and basalt units at this location. Here we will see the nature of the contact between one of the banded iron formation units and an adjacent massive basalt lava flow. 1.77 miles 1.79 Miles 2.14 miles Walk north-northwest through the bush back to the two-track road. Proceed west-southwest along the two-track road. At the fork in the road, proceed to the northwest along the two-track road. Continue walking along the two-track road for approximately 0.35 miles. At this point you will encounter a series of outcrops within, and along the south and north edges, of the two-track road. This will be Field Trip Stop 6a. 64 �Stop 6a: Folded Soudan Iron-Formation Member Banded Iron Formation Location: T. 62N, R. 15W, sec. 23, NE, SE, Soudan 7.5-minute quadrangle UTM: 559,725E / 5,298,155N As is well exemplified at the “classic” outcrop of Soudan Member banded iron formation located in the NE ¼ NE ¼ Sec. 27, T.62N, R. 15W (see stop 7-10 in Peterson et al., 2009), the Soudan Member banded iron formation commonly displays multiple generations of tight folds which can result in complex interference patterns (Figure 11). At this location, and at several other small outcrops along the north side of the twotrack trail, we can observe highly folded, moderately- to strongly magnetic, magnetite-rich Soudan Member oxide facies iron formation. 2.14 miles 2.16 miles 2.17 miles Continue walking southwestward along two-track road for approximately 0.02 miles. You will see several outcrops between 0.01 and 0.02 miles into the bush on the northwest side of the two-track road. Proceed to these outcrops. This will be Stop 6b. Figure 11. Tightly folded Soudan Member Algoma-type banded iron formation 65 �Stop 6b: Contact between Folded Soudan IronFormation Member Banded Iron Formation and Diabase/Gabbro Intrusion Location: T. 62N, R. 15W, sec. 23, SE, SE, Soudan 7.5-minute quadrangle UTM: 559,560E / 5,298,035N Note: Be extremely careful on this outcrop, especially if it is wet. The glacially polished surface combined with wet moss makes for very slippery conditions. This outcrop is once again composed primarily of laminated to thinly-bedded Soudan Member oxidefacies banded iron formation. On the far western side of the outcrop, as well as in several small outcrops to the northeast, we can observe a massive, fine- to medium-grained, dark green to grayish-green rock which has been interpreted to represent a diabase sill. This sill appears to intrude the contact between Soudan Member oxide-facies iron formation (to the south) and Soudan Member massive basalt lava flows (to the north). Based on the outcrop distribution in the park, this unit appears to get coarser grained to the east, where it represents dark green to blackish-green medium-grained gabbro. 2.17 miles 2.19 miles 2.30 miles 2.49 miles 2.60 miles Return to the two-track road. Walk southwest, then west, down the hill along the two-track road. Cross bridge over small creek between unnamed pond (to the south) and creek/swamp (to the north). Be extremely careful crossing this bridge as it may be wet and slippery! Continue west, then southwest along the two-track road and begin to climb a moderately steep hill. Continue walking up the hill. The outcrop ridge on both sides of the two-track road comprise poorly exposed interbedded Soudan Member banded iron formation and Soudan Member basalt lava flows exhibiting both sheet flow facies and associated flow breccia facies. Continue walking another 0.11 miles up to the top of the hill. We are now at the top of the hill. We will reassemble here before moving on to the remainder of the outcrops along our traverse. To the west, the prominent hill and low lying outcrops along the trail comprise magnetite-rich Soudan Member banded iron formation. After reassembling the group, we will proceed north, then northeast, along the two-track trail for 0.09 miles. 2.69 miles 2.80 miles 2.81 miles Take the left fork and proceed to the northwest along the two-track road. We will once again reassemble the group at this location. Once reassembled, we will walk north-northeast approximately 0.01 miles up a hill through the bush to Field Trip Stop 2.7. Field Trip Stop 2.7 66 �Stop 7: “Contact” Between Soudan Member Banded Iron Formation and Rhyodacite Polymict Lapilli Tuff/Tuff Breccia of the Gafvert Lake Sequence Location: T. 62N, R. 15W, sec. 23, NW, SE, Soudan 7.5-minute quadrangle UTM: 558,995E / 5,298,230N Here we will see one of the few places where the nature of the contact between the Soudan IronFormation Member oxide facies iron-formation and the overlying dacitic to rhyodacitic volcaniclastic rocks associated with the informally named Gafvert Lake Sequence (which is part of the Lake Vermilion Formation) can be observed. Based on regional mapping, Sims and Southwick (1980), Southwick (1993), and Southwick et al. (1998) have indicated that the contact between the underlying Soudan Iron-Formation Member of the Ely Greenstone Formation and the overlying Lake Vermilion Formation is locally an unconformity. Geochronological work in the Vermilion District (Peterson et al., 2001; Lodge et al., 2013), combined with detailed field mapping in the limited number of locations where the contact between the Soudan Iron-Formation Member and the Lake Vermilion Formation occurs, bears out this interpretation. Peterson et al. (2001) obtained a U-Pb zircon age of 2722 ± 0.9 Ma from a quartz-phyric rhyolite dome within the Fivemile Lake Sequence at the Fivemile Lake prospect, located approximately 850 meters stratigraphically below the base of the overlying Soudan Member Iron-Formation unit. Regionally extensive detailed mapping in the stratigraphic units that occur between the Fivemile Lake Sequence rhyolite dome and the base of the Soudan Member Iron-Formation has been completed by a number of researchers (Peterson and Jirsa, 1999; Peterson, 2001; Hudak et al., 2002a; Hudak et al., 2002b; Peterson and Patelke, 2003; Hoffman, 2007; Radakovich et al., 2010; Heim et al., 2011). Based on this detailed mapping, there are no indications of any unconformities within the Fivemile Lake Sequence or the Central Basalt Sequence that comprise the footwall to the Soudan Iron-Formation Member. As well, unconformities at the contacts between the Fivemile Lake Sequence and the Central Basalt Sequence, and the Central Basalt Sequence and the overlying Soudan Iron-Formation do not appear to be present. Hudak et al. (2007; 2012) have noted that the contact between the Central Basalt Sequence and the overlying Soudan Iron-Formation Member is transitional over several hundred meters, with the presence of ironformation horizons near the top of the Central Basalt Sequence increasing in abundance, and the abundance of basalt lava flows and associated volcaniclastic rocks decreasing in abundance, as one approaches the base of the Soudan Iron-Formation Member. Therefore, it appears that the Fivemile Lake Sequence is stratigraphically overlain by the Central Basalt Sequence, which in turn is stratigraphically overlain by the Soudan Iron-Formation Member. Furthermore, there appears to be a major period of volcanism and associated hydrothermal activity between 2722 and 2718 Ma in the western part of the Wawa Abitibi Terrane in Ontario that produced both the volcanic rocks and volcanogenic massive sulfide orebodies that occur at the Winston Lake and Geco deposits (Lodge et al., 2013). Based on both the geochronological work and detailed mapping in the Vermilion District, as well as the regional volcanic and hydrothermal events in the western Wawa-Abitibi belt, we currently believe that the Soudan IronFormation Member was deposited between 2722 and 2718 Ma. Further geochronological studies within the stratigraphic package that comprises the Soudan Iron-Formation Member of the Ely Greenstone Formation will need to be completed in order to verify our current interpretation. Based on field relationships recognized by Radakovich et al. (2010), Lodge et al. (2013) collected a sample of the basal part of the Gafvert Lake polymict dacite- to rhyodacite lapilli-tuff / tuff-breccia deposits that occur at this outcrop in order to determine the age of volcanism of the Gafvert Lake 67 �Sequence relative to the ages of the Lower and Soudan Iron-Formation members of the Ely Greenstone Formation. Zircons from the sample of polymict rhyodacite tuff-breccia from this outcrop approximately 2m north of the contact with the Soudan Iron-Formation Member produced a high precision U-Pb age of 2689.7 ± 0.8 Ma using thermal ionization mass spectrometry (Lodge et al., 2013). Given that the basal Gafvert Lake Sequence deposits contain angular intraclasts of chert and banded iron formation, and that there appears to be no intense structural fabric in either the Soudan Iron-Formation Member or the Gafvert Lake volcaniclastic rocks, Lodge et al. (2013) interpreted the contact here to represent a disconformity, a type of unconformity characterized by strata that are essentially parallel on either side of the erosional or non-depositional surface. Several outcrops occur at this location, but at 1:5000 scale mapping they have been combined into a single east-northeast trending outcrop (Figure 12). The majority of the outcrop, which extends east up the hill, is composed of laminated to medium bedded Soudan Iron-Formation Member. Alternating magnetite-rich horizons, chert horizons, and jasper horizons display planar bedding and are locally folded. Moving toward the northwest part of the outcrop, we observe a small break in the outcrop exposure. This break occurs directly above the contact between the Soudan Member Iron-Formation and the Gafvert Lake volcaniclastic rocks. In this area, note the lack of deformation in both lithological units. The lack of structural deformation at this contact, as well as geochronological data obtained from the Gafvert Lake volcaniclastic rocks near this contact (Lodge et al., 2013), supports the interpretation of a disconformity. Figure 12. Detailed (1:5000 scale) map illustrating the disconformable contact between the Soudan Member Algoma-type banded iron-formation (unit S4a) and the Gafvert Lake Sequence quartz- and plagioclase-phyric polymict dacite-rhyodacite tuff-breccia / lapilli tuff deposits (unit US2eh). We will 68 �start our investigation where Stop 7 is indicated, and traverse along the path indicated by the red dashed line over a series of outcrops. We will assemble on the two-track trail where indicated by the star symbol before proceeding to Stop 8. Moving to the northwest, we observe the basal several meters of the Gafvert Lake Succession volcaniclastic rocks. Here, the rock is composed of a very thickly bedded quartz- and plagioclase-phyric polymict dacite-rhyodacite tuff-breccia / lapilli tuff. The rock is characterized by up to 5% 1-3mm diameter subhedral to euhedral gray to blue-gray quartz phenocrysts and locally, 5-10% subhedral to euhedral light gray to tan tabular plagioclase phenocrysts set in a fine-grained quartzo-felspathic matrix that is locally sericite altered. Accidental fragments comprising laplli-sized light gray to grayish black angular to subangular chert, gray to dark gray subangular to angular banded iron formation (Figure 13), and rare angular to subangular reddish brown jasper fragments are present. As well, juvenile fragments comprising lapilli- to locally block-sized pumice are present. Lapilli- to block-sized accessory fragments of quartz- and plagioclase-phyric coherent dacite and rhyodacite are also present, in abundances up to 5%. B A A Figure 13. Gafvert Lake Sequence quartz- and plagioclase-phyric polymict dacite-rhyodacite tuff-breccia / lapilli tuff from the Gafvert Lake Sequence. A. Typical appearance of very thickly bedded quartz- and plagioclase-phyric polymict dacite-rhyodacite lapilli tuff. B. Closeup of unit illustrating tannish-white subhedral to euhedral tabular plagioclase phenocrysts, gray to gray-blue anhedral quartz phenocrysts, and 1cm diameter angular accidental fragment composed of jasper-rich banded iron formation. 2.81 miles 2.90 miles 3.32 miles 3.41 miles 3.48 miles 3.50 miles We will proceed north down the north-sloping hillside for about 0.09 miles over a series of outcrops comprising Gafvert Lake Succession rhyodacite tuffs, lapilli tuffs, and tuff breccias. Observe the subtle changes in crystal content and fragment compositions and abundances while moving down the hill toward the two-track trail. We will reassemble the group on the two-track trail. We will proceed to walk eastnortheast along the two-track road for approximately 0.42 miles. Cross bridge – there will be a large unnamed pond to the east. Continue 0.09 miles up hill to intersection and wait where the two-track road makes a turn from north-trending to east-trending. This is where we parked our truck each day during our 2010 capstone mapping project for the PRC Field Camp. We will now walk northeast along the road for 0.07 miles. Follow the field trip leader approximately 0.01-0.02 miles southeast into the bush to Stop 8. Stop 8. 69 �Stop 8: Gafvert Lake Sequence Tuffs and Lapilli Tuffs Location: T. 62N, R. 15W, sec. 23, SE, NE, Soudan 7.5-minute quadrangle UTM: 559,675E / 5,298,700N We will stop here to observe several small outcrops of the Gafvert Lake Sequence tuffs and lapilli tuffs. These deposits comprise very thickly bedded, light gray, quartzand plagioclase-phyric dacitic to rhyodacitic tuffs and lapilli tuffs. The light gray recrystallized matrix generally contains 10-15% <1-2mm subhedral to euhedral tabular plagioclase phenocrysts which locally appear to be broken, as well as 3-8% <1-2mm pale gray anhedral, locally broken, anhedral to subhedral quartz phenocrysts. Various types of lapilli may be observed, including: 1) 10-20% 1-3cm diameter quartz- and plagioclase-phyric coherent dacite to rhyodacite lapilli; 2) 5-7% <3cm diameter pale gray green, lens-shaped, locally quartzand plagioclase-phyric pumice lapilli; 3) <1mm dark gray to light gray angular chert lapilli ranging from 0.5-3cm in diameter; and 4) 1-3% 0.5-5cm dark gray to black magnetite-rich banded iron formation lapilli. 3.50 miles 3.51 miles 4.00 miles Traverse approximately 0.01 miles north through the bush back on to the two-track road. Proceed northeast, then northwest, then northeast along the two-track trail for approximately 0.48 miles. We will assemble the group at this location before the group follows the leader on a 0.01mile traverse north-northwest into the bush to Stop 9. Stop 9. Stop 9: Quartz ± plagioclase-phyric Rhyodacite Sill (informally named the Gafvert Lake Intrusive Complex) Location: T. 62N, R. 15W, sec. 24, NW, NW, Soudan 7.5-minute quadrangle UTM: 560,015E / 5,299,175N At this location we will observe a spectacular light gray, massive, quartz- ± plagioclase-phyric coherent rhyodacite which, based on regional mapping (Peterson and Jirsa, 1999; Peterson, 2001; Hudak et al., 2002b; Heim et al., 2011) comprises a sill-dike complex that extends from the northern extents of Lake Vermilion State Park over 20km eastward to Mitchell Lake. This intrusion is most prevalent in the vicinity of Gafvert Lake, where it comprises several sills and dikes that intrude into the thickest section of Gafvert Lake Sequence volcaniclastic rocks. Based on the distribution of sills and dikes, coherent-facies Gafvert Lake Sequence deposits, and an abundance of coarse polymict breccias in this region, Peterson (2001) has interpreted this area to be the remnants of a stratovolcano that produced the Gafvert Lake Sequence dacitic to rhyodacitic volcaniclastic rocks. For this reason, this unique quartz-feldspar porphyry intrusion has been informally named the Gafvert Lake Intrusive Complex (GLIC). Lithogeochemical work recently completed at the University of Wisconsin Eau Claire by Geoff Pignotta and Kelly Schwierske indicates that the GLIC and Gafvert Lake volcaniclastic rocks have very similar major, trace and rare earth element characteristics suggesting that they may be genetically related. However, geochronological studies will need to be 70 �performed to determine unambiguously if the GLIC and Gafvert Lake volcaniclastic rocks are genetically related. The GLIC comprises light gray, massive, quartz ± plagioclase-phyric coherent rhyodacite. The light gray aphanitic groundmass contains 3-7% gray to light blue subhedral rounded to euhedral square quartz phenocrysts that range from 3-10mm in diameter, and 10% pale gray to tan, subhedral to euhedral tabular plagioclase phenocrysts ranging from 1-4mm in length. A variety of xenoliths may be found in this intrusion, including: 1) brown mudstone lapilli; 2) green to gray-green massive and/or amygdaloidal basalt lapilli; and 3) light gray aphyric coherent rhyodacite lapilli. In the field, the presence of large 5mm-10mm diameter gray to blue gray quartz phenocrysts distinguishes the GLIC from other quartzfeldspar-porphyry intrusions in the Vermilion District. 4.00 miles 4.01 miles 4.59 miles Traverse south through the bush 0.01 miles back to the two-track trail. Once you return to the two-track trail, walk 0.58 miles east-northeast. At the end of the two-track trail you will intersect a well-maintained dirt road. The bus will be parked at this location. Obtain refreshments and get on the bus. End of traverse log Bus Log (Continued) 56.9 miles Pick up field trip participants after their traverse through the park. Drive east on wellmaintained dirt road 0.48 miles back to the intersection with Old Ely Road. 57.7 miles Turn right (southwest) and follow the Old Ely Road approximately 1.3 miles to the intersection with the dirt road immediately before the gate to Lake Vermilion State Park. 59.0 miles Turn south on dirt road and return to Highway 1/169. 59.1 miles Turn west on Highway 1/169 South. Proceed approximately 5.2 miles past Soudan and through Tower to the intersection between Highway 1/169 and Highway 135. 64.3 miles Pull bus off on to shoulder of Highway 1/169 South just west of the intersection of Highways 1/169 and Highway 135. The outcrop on the south side of the road is Field Trip Stop 10. Stop 10: Recently Exposed Outcrop of Gafvert Lake Sequence Lapilli-tuffs and Tuff-Breccias Location: T. 62N, R. 15W, sec. 32, SW, SW, Tower 7.5-minute quadrangle UTM: 553,500E / 5,294,510N NOTE: Take extreme care when crossing the highway at this location! The final stop on our field trip is to a recently exposed (~2012) outcrop comprising polymict Gafvert Lake Sequence tuff, lapilli-tuff and tuff-breccia deposits. Although never mapped in detail by any of the coauthors in this guidebook, this exposure appears to contain several individual volcaniclastic units that may be distinguished by the size and abundance of the phenocrysts and fragments present. The pale gray to gray aphanitic matrix contains variable percentages and sizes of quartz and plagioclase phenocrysts. Locally, abundant (up to 10%) <1mm euhedral pyrite cubes are disseminated in the matrix. Fragment 71 �composition is also variable, with gray chert, light gray quartz- and plagioclase-phyric coherent rhyodacite, light gray to tan pumice, and rare massive pyrrhotite fragments present. The rock also possesses a moderately- to well-defined schistosity with lineations that plunge moderately to steeply to the northeast (Sims, 1973). 64.3 miles 112.5 miles Field trip participants should return to bus. Take extreme care when crossing Highway 1/169. Follow Highway 1/169 south and retrace route to Vermilion District back to the Hibbing Park Hotel in Hibbing, Minnesota. End of field trip at parking lot in Hibbing Park Hotel. Acknowledgements Characterizing and evaluating the detailed geology of Lake Vermilion State Park involved a team effort between Minnesota Department of Natural Resources (MDNR) personnel, the Minnesota Geological Survey, NRRI geologists, and students and faculty from the Precambrian Research Center Field Camp as well as the Univesity of Wisconsin Eau Claire. The lead author would like to thank Jim Essig (Manager, Soudan Underground Mine State Park and Lake Vermilion State Park) and James Pointer (Interpretive Supervisor, Soudan Underground Mine State Park and Lake Vermilion State Park) from the MDNR for their support, assistance, and guidance while planning and conducting detailed geological mapping by PRC students and faculty in Lake Vermilion State Park in 2010 and 2011. Also, Minnesota Geological Survey geologists Amy Radakovich, Mark Jirsa, and Terry Booerboom are thanked for their assistance (and patience!) during the development of this field trip guide. As well, Dean Peterson, the late Richard Patelke, Mark Severson, John Heine, Peter Jongewaard, Steve Hovis and Adam Hoffman are thanked for their excellent mapping in the southern part of what was to become Lake Vermilion State Park. This work by former and current NRRI colleagues became the foundation upon which new mapping in the park was based. Additionally, Geoff Pignotta, Kelly Schwierske, and the Department of Geology at the University of Wisconsin, Eau Claire are thanked for intellectual efforts and financial support to further evaluate the geology and geochemistry of Lake Vermilion State Park. Finally, PRC students Chris Heim, Rob Kilduff, Chris Mahr, Charlie Parent, Molly Partidge, Rita Pierce, Amy Radakovich, Christine Rahtz, Andrew Ritts, Heather Scott and Andrew Vial are thanked for their outstanding field mapping, compiling, computer map generation, and companionship during the four weeks in 2010 and 2011 that it took to produce the recent detailed geologic maps in Lake Vermilion State Park. Without these exceptional students, our knowledge of the fascinating geology of Lake Vermilion State Park would not be nearly what it is today. REFERENCES Bakst, B., 2013, Minnesota’s Lake Vermilion State Park evolving at a cautious pace: TwinCities.com, Pioneer Press, http://www.twincities.com/localnews/ci_23939683/lake-vermilion-state-parkminnesotas-newest-evolving-at. Batiza, R., and White, J. D. L., 2000, Submarine lavas and hyaloclastite: in Sigurdsson, H., Encyclopedia of Volcanoes: Academic Press, San Diego, CA, p. 361-381. Cas, R. A. F., and Wright, J. V., 1987, Volcanic Successions – Modern and Ancient: George Allen and Unwin, London, 528 p. Dimroth, E., Cousineau, P., Leduc, M., and Sanschagrin, Y., 1978, Structure and organization of Archean subaqueous basalt flows, Rouyn-Noranda area, Quebec: Canadian Journal of Earth Sciences, v. 15, p. 90918. Driese, S. G., Jirsa, M. A, Ren, M., Brantley, S. L., Sheldon, N. D., Parker, D., and Schmitz, M. D., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, v. 189, p. 1-17. 72 �Fisher, R. V., 1961, Proposed classification of volcaniclastic sediments and rocks: Geological Society of America Bulletin, v. 72, p. 1409-1414. Fisher, R. V., 1966, Rocks composed of volcanic fragments and their classification: Earth Science Reviews, v. 1, p. 287-298. Fisher, R. V., 1998, Out of the Crater: Princeton University Press, 180 p. Gibson, H. L., Morton, R. L., and Hudak, G. J., 1999, Submarine volcanic processes, deposits, and environments favorable for the location of volcanic-associated massive sulfide deposits: Reviews in Economic Geology, v. 8, p. 13-51. Hart, T. R., Gibson, H. L., and Lesher, C. M., 2004, Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic Cu-Zn-Pb sulfide deposits: Economic Geology, v. 99, p. 1003-1013. Heim, N., Scott, H., Kilduff, R., Rahtz, C., Vial, A., Young, S., Mahr, C., and Hudak, G., 2011, Preliminary bedrock geology map of the eastern part of Lake Vermilion State Park, St. Louis County, NE Minnesota: Precambrian Research Center Map Series Map PRC/Map – 2010-01, 1:5000 scale. Hoffman, A. T., 2007. Lithostratigraphy, Hydrothermal Alteration, and Lithogeochemistry of Neoarchean Rocks in the Lower and Soudan Members of the Ely Greenstone Formation, Vermilion District, NE Minnesota: Implications for Volcanogenic Massive Sulfide Deposits: unpublished M. S. thesis, University of Minnesota – Duluth, 295 p. Hudak, G. J., Heine, J., Hocker, S. M., and Hauck, S., 2002b, Geological mapping of the Needleboy Lake – Sixmile Lake area, northeastern Minnesota: a summary of volcanogenic massive sulfide potential: Natural Resources Research Institute Report of Investigation NRRI/RI-2002/14, 15 p. Hudak, G.J., Heine, J., Jirsa, M.A., and Peterson, D.M., 2004, Field Trip 1 - Volcanic stratigraphy, hydrothermal alteration, and VMS potential of the Lower Ely Greenstone, Fivemile Lake to Sixmile Lake area: 50th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 50, Part 2 – Field Trip Guidebook, p. 1-45. Hudak, G. J., Heine, J., Lodge, R. W. D., and Jansen, A., 2012, Recent developments understanding the volcanic, magmatic, tectonic, and metallogenic evolution of the Ely Greenstone Formation, Vermilion District, NE Minnesota: Geological Association of Canada – Mineralogical Association of Canada, Abstracts and Program, v. 35, p. 59. Hudak, G. J., Heine, J., Newkirk, T., Odette, J., and Hauck, S., 2002a, Comparative geology, stratigraphy, and lithogeochemistry of the Fivemile Lake, Quartz Hill, and Skeleton Lake VMS occurrences, Vermilion district, NE Minnesota: A report to the Minerals Coordinating Committee, DNR, Minerals Division, State of Minnesota: Natural Resources Research Institute Technical Report NRRI/TR2002/03, 390 pages. Hudak, G. J., Hocker-Finamore, S. M., and Heine, J., 2006, Field distribution, petrography, and lithogeochemistry of epidosites in the vicinities of Fivemile, Needleboy, and Sixmile Lakes, Vermilion District, NE Minnesota: 52nd Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 57, Part 1 – Programs and Abstracts, p. 30-31. Hudak, G. J., Hoffman, A. T., Peterson, D. M., and Heine, J., 2007, Recent developments understanding the volcanic, magmatic, tectonic, and metallogenic evolution of the Ely Greenstone Formation, Vermilion District, NE Minnesota: 53rd Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 53, Part 1 – Program and Abstracts, p. 443. Jansen, A. C., Hudak, G. J., Heine, J. J., and Peterson, D. M., 1999, Lithogeochemical evaluation of Neoarchean mafic volcanic rocks comprising the footwall to the Soudan Member of the Ely Greenstone Formation, northeastern Minnesota: 55th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 55, Part 1 – Program and Abstracts, p. 46-47. Jirsa, M. A., Boerboom, T. J., and Peterson, D. M., 2001, Bedrock geological map of the Eagles Nest Quadrangle, St. Louis County, Minnesota: Minnesota Geological Survey, Miscellaneous Map M-114, scale 1:24,000. Jirsa, M. A., Boerboom, T. J., Green, J. C., Miller, J. D., Morey, G. B., Ojakangas, R. W., and Peterson, D. M., 2004, Field Trip 5 – Classic outcrops of northeastern Minnesota: 50th Annual Meeting, 73 �Institute on Lake Superior Geology, Proceedings Volume 50, Part 2 – Field Trip Guidebook, p. 129169. Jirsa, M., and Hillman, M., 2009, Field Trip 4 – Pioneer Mine (Miners Lake) Canoe Excursion: 55th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 55, Part 2 – Field Trip Guidebook, p. 110-115. Jirsa, M. A., Starns, E. C., and Schmitz, M. D., 2012, Bedrock geologic map of the 2006 Cavity Lake forest fire area, Boundary Waters Canoe Area Wilderness, northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map M-193, 1:24,000 scale. Kawachi, Y., and Pringle, I. J., 1988, Multiple-rind pillow structures in pillow lava as an indicator of shallow water: Bulletin of Volcanology, v. 50, p. 161-168. Larson, P., and Mooers, H., 2009, Field Trip 2 – Glacial geology of the Vermilion Moraine: 55th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 55, Part 2 – Field Trip Guidebook, p. 81-99. Lodge, R. W. D., Gibson, H. L., Stott, G. M., Hudak, G. J., Jirsa, M. A., and Hamilton, M. A., 2013, New U-Pb geochronology from the Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa Subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research, v. 235, p. 264-277. McPhie, J., Doyle, M., and Allen, R., 1993, Volcanic Textures: A Guide to the Interpretation of Textures in Volcanic Rocks: CODES Key Centre, University of Tasmania, Hobart, Tasmania, 198 p. Mercier-Langevin, P., Hannington, M. D., Dubé, B., and Bécu, V., 2010, The gold content of volcanogenic massive sulfide deposits: Mineralium Deposita, v. 46, p. 509-539. Moosavi, S., Johnson, T., Wendland, C., Anderson, A., and Hudak, G., 2007, Bedrock geology map of the footwall of the Soudan Iron Formation south of Twin Lakes, St. Louis County, northeastern Minnesota: Precambrian Research Center Map Series Map PRC/Map – 2007-04, 1:5000 scale. Mueller, W. U., and White, J. D. L., 2004, 4.2 – Terminology of Volcanic and Volcaniclastic Rocks: in Eriksson, P. G., Altermann, W., Nelson, D. R., Mueller, W. U., and Catuneanu, O., (eds.), The Precambrian Earth: Tempos and Events: Developments in Precambrian Geology, v. 12, Elsevier, Amsterdam, p. 273-277. Peterson, D. M., 2001, Development of Archean lode-gold and massive sulfide deposit exploration models using geographic information system applications: targeting mineral exploration in northeastern Minnesota from analysis of analog Canadian mining camps: unpublished Ph. D. dissertation, University of Minnesota, Duluth, Minnesota, 503 p. Peterson, D. M., and Jirsa, M.A., 1999, Bedrock geologic map and mineral exploration data, western Vermilion district, St. Louis and Lake Counties, northeastern Minnesota: MGS Miscellaneous Map M-98, scale 1:48,000. Peterson, D., Jirsa, M., and Hudak, G., 2009a, Field Trip 7 – Architecture of an Archean Greenstone Belt: Stratigraphy, Structure, Mineralization: 55th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 55, Part 2 – Field Trip Guidebook, p. 178-215. Peterson, D. M., and Patelke, R. L., 2003, National Underground Science and Engineering Laboratory (NUSEL): Geological site investigation for the Soudan Mine, northeastern Minnesota: Natural Resources Research Institute Technical Report NRRI/TR-2003/29, 88 p.a Peterson, D. M., and Patelke, R. L., 2004, Field Trip 7 – Economic geology of Archean gold occurrences in the Vermilion District, northeast of Soudan, Minnesota: 50th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 50, Part 2 – Field Trip Guidebook, p. 200-226. Peterson, D. M., Pointer, J., and Marshak, M., 2009b, Field Trip 3 – Soudan Iron Mine and Physics Lab Tour: 55th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 55, Part 2 – Field Trip Guidebook, p. 100-109. Pearce, J. A., 1996, A user’s guide to basalt discrimination diagrams: in Wyman, D. A., ed., Trace Element Geochemistry of Volcanic Rocks: Applications for Massive Sulphide Exploration: Geological Association of Canada, Short Course Notes, v. 12, p. 79-113. 74 �Philips, W. R., and Griffen, D. T., 1981, Optical Mineralogy – The Nonopaque Minerals: W. H. Freeman and Company, San Francisco, 677 p. Radakovich, A. L., Parent, C. T., Partridge, M. E., Ritts, A. D., Pierce, R., and Hudak, G. J., 2010, Reconnaissance bedrock geological map of the northern part of Soudan Underground Mine State Park and the northwestern part of Lake Vermilion State Park, St. Louis County, Minnesota: Precambrian Research Center Map Series Map PRC/Map – 2010-04, 1:5000 scale. Ross, P.-S., and Bédard, J. H., 2009, Magmatic affinity of modern and ancient subalkaline volcanic rocks determined from trace-element discrimination diagrams: Canadian Journal of Earth Sciences, v. 46, no. 11, p. 823-839. Schmid, R., 1981, Descriptive nomenclature and classification of pyroclastic deposits and fragments; recommendations of the IUGS subcommission on the systematics of igneous rocks: Geology, v. 9, p. 41-43. Schwierske, K.L., Pignotta, G. S., and Hudak, G. J., in press, The 2.7 billion year old Mt. St. Helens of northern Minnesota: Petrography, geochemistry, and economic significance of the Neoarchean Gafvert Lake Sequence: 60th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 60, Part 1 – Programs and Abstracts. Sims, P. K., 1973, Geologic map of the western part of the Vermilion District, Northeastern Minnesota: Minnesota Geological Survey, Miscellaneous Map M-13, scale 1:48,000. Sims, P. K., and Southwick, D. L., 1985, Geologic map of Archean rocks, western Vermilion district, northern Minnesota: U. S. Geological Survey, Miscellaneous Investigations Map I-1527, scale 1:48,000. Southwick, D. L., (compiler), 1993, Bedrock geologic map of the Soudan-Bigfork area, northern Minnesota: Minnesota Geological Survey, Miscellaneous Map M-79, scale 1:100,000. Southwick, D. L., Boerboom, T. J., and Jirsa, M. A., 1998, Geologic setting and descriptive geochemistry of Archean supracrustal and hypabyssal rocks, Soudan-Bigfork area, northern Minnesota: implications for metallic mineral exploration: Minnesota Geological Survey, Report of Investigations 51, 69 p. Stott, G., Corkery, T., Leclair, A., Boily, M., and Percival, J., 2007, A revised terrane map for the Superior Province as interpreted from Aeromagnetic Data: 53rd Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 53, Part 1 – Program and Abstracts, p. 74-76. Stott, G. and Mueller, W., 2009, Superior Province: The nature and evolution of the Archean continental lithosphere: Precambrian Research, v. 168, p. 1-3. Vallowe, A. M., Thalhamer, E. J., Rhoades, D. L., and Peterson, D. M., 2010, Surface and subsurface geologic maps of the Soudan Underground Mine State Park, St. Louis County, northeastern, Minnesota: Precambrian Research Center Map Series Map PRC/Map – 2010-01, 1:2500 and 1:5000 scale. White, J. D. L., and Houghton, B. F., 2006, Primary volcaniclastic rocks: Geology, v. 34, no. 8, p. 677680. Winchester, J. A., and Floyd, P. A., 1977, Geochemical discrimination of different magma series and the differentiation products using immobile elements: Chemical Geology, v. 20, p. 325-343. 75 ��FIELD TRIP 3 Wednesday, May 14, 2014 WESTERN MESABI RANGE MINING OPERATIONS LEADERS: Douglas Halverson (Cliffs Natural Resources—Duluth) Daniel Cervin (Cliffs Natural Resources—Hibbing Taconite), William Everett and Kevin Kangas (Essar Steel); and Joseph Nielsen (Magnetation). INTRODUCTION This field trip will visit three distinct iron mining operations along the western part of the Mesabi Iron Range (Fig. 1). The morning will be spent at the Hibbing Taconite operation managed by Cliffs Natural Resources. There, participants will get a pit tour of the mining and reclamation operations, followed by a tour of the processing facility. The afternoon will include a tour through the new construction of Essar Steel’s taconite processing facility near Nashwauk (near the old Butler Taconite site), and may view some core that intersects strata including what is inferred to be the 1850 Ma Sudbury Impact Layer. This will be followed by a tour of Magnetation’s two-year old, 1.2 million ton per-year facility located near Bovey, where iron concentrate is being extracted from the tailings of historic mining of natural (hematite) ores. Figure 1. Bedrock geologic map of the western Mesabi Iron Range showing the 3 operations that will be visited during this trip. Map is clipped from MGS Miscellaneous Map M-163 (Jirsa and others, 2005; published scale=1:100,000). Reddish unit represents subsurface extent of Biwabik Iron Formation. 76 �HIBBING TACONITE PIT AND PLANT Hibbing Taconite Company, managed by Cliffs Natural Resources Doug Halverson and Dan Cervin (Cliffs Natural Resources) Field Trip participants will get a pit tour of the mining and reclamation operations. There may be an opportunity to view a blast, dependent upon blasting schedule and blast location. Hibbing Taconite Company (HTC) is managed by Cliffs Natural Resources, an international mining and natural resources company. Cliffs is the largest producer of iron ore pellets in North America, a major supplier of directshipping lump and fines iron ore out of Australia and a significant producer of metallurgical coal. The annual production capacity of Hibbing Taconite Company is 8.0 million tons of taconite pellets, operating 24 hours per day, year-round, employing 770 people. Through 2013, HTC has produced 260 million gross tons of pellets. HTC is jointly owned by Arcelor Mittal (62.3%), Cliffs Natural Resources (23%), and US Steel Canada (14.7%). The HTC plant is located approximately four miles northwest of Hibbing, Minnesota (Fig.2), just north of the Laurentian Divide. The initial taconite pit was developed in 1975. Since inception, this pit has expanded east, west, and south along the northern crest of the historic Hull-Rust Mahoning natural ore mine. The Hull-Rust Mahoning Mine, actually a combination of 30 separate mines, was developed along an east/west-trending fault structure and operated from 1895 to 1979. Material movement from this "largest open pit iron mine in the world" totaled more than 1.1 billion tons. Figure 2. Airphoto image of Hibbing Taconite’s plant, tailings facility, and mines that extend nearly 7 miles along the strike of iron-formation. The four main subdivisions of the Biwabik Iron Formation are present in the vicinity of HTC. From bottom to top they are Lower Cherty, Lower Slaty, Upper Cherty, Upper Slaty. Erosion has removed the Upper Slaty and most of the Upper Cherty members within the area of the current pit. Where present, the 77 �Lower Slaty member and the upper 30 feet of the Lower Cherty member are stripped as rock waste. Approximately 150 feet of cherty and slaty taconite is mined from the central portion of the Lower Cherty. The formation strikes northeast and dips 6°-8° to the southeast. The Pokegama Quartzite forms the footwall of the iron-formation and outcrops to the north, along the south edge of the divide. The only significant structural features are common, but minor, northwest-trending normal faults. Numerous natural ore mines were located along these oxidized structures. The taconite mined by HTC averages 20 percent magnetic iron, with the general mineralogy consisting of quartz, magnetite, siderite, ankerite, minnesotaite, stilpnomelane and hematite. Ore units in the Lower Cherty (120 feet thick) are predominantly "cherty" taconite, with 6- to 12-inch-thick massive silicate-chert zones separated by 1/8- to 2-inch-thick slaty bands. The lower two units (30 feet thick) are predominantly "slaty" taconite with inter-bedded argillite, magnetite, and minor hematite forming slaty bands from 2-10 inches in thickness separated by 2- to 4-inch massive cherty zones. Stripping materials include glacial overburden, waste rock, lean oxidized taconite, and old stockpiles, all varying widely in thickness from area to area. Standard rotary drilling, blasting, electric shovel loading, and 240-ton truck haulage are the mining methods utilized. The processing flowsheet differs significantly from the standard Mesabi Range taconite plant in the area that involves crushing and grinding. By contrast, the HTC plant utilizes autogenous mills, which do not contain grinding media. A single stage of gyratory crushing in one of two 60-inch crushers reduces the crude to 10 inches. This is followed by autogenous grinding in one of nine 36-foot¬diameter mills. Water is added and as the ore tumbles, it reduces itself to powder fineness. Field Trip Stops at Hibbing Taconite: Stop 1 – The first stop in the field trip will be in Hibbing Taconite’s Group 4 to view the stratigraphy of the Biwabik Iron Formation at the mine site. In this currently inactive portion of the pit, overlying stockpiles from natural ore mines, glacial till, Lower Slaty rock stripping and Lower Cherty ore horizons are exposed in the high wall of the mine pit. Opportunity for the collection of typical cherty taconite ore will be available at this site. Stop 2 – The tour will travel east to Group 2 and view the “footwall” of the mine. Opportunity to collect samples of more slaty, jasper rich ore horizons will be available while discussing the in-pit enrichment processing that is used to improve these horizons prior to plant processing. Stop 3 – The tour will continue east to view active mining activities in Group 1. Typical mine production activities such as production drilling, loading and hauling will be observed from this location. Stop 4 – View a production blast or plant tour. Depending on the production blasting schedule and the visibility of the blast from a safe location the tour may have the opportunity to view a production blast, typically involving 500,000 to 1,000,000 tons of taconite ore or overlying rock. Alternate - If scheduling or location does not allow the viewing of a blast, the tour will view the Hibbing Taconite processing plant. Portions of the process that include crushing, grinding, concentration, balling and induration will be toured to show the flow of material from crude ore to finished product. Stop 5 – Lunch at mine view in old North Hibbing. The scale of the Hibbing Taconite mining operation will be seen from this scenic overview . 78 �ESSAR STEEL’S NEW TACONITE PIT AND PLANT Essar Steel Minnesota, LLC, Nashwauk, MN William Everett (Area Manager of Mining) Essar Steel Minnesota, LLC (ESML) is an iron ore mining company engaged in the development of a fully integrated iron ore mine and pellet plant located on the western end of the Mesabi Range. The Project is adjacent to the City of Nashwauk located in Itasca County, approximately 15 miles (24 km) west of Hibbing and 20 miles (32 km) east of Grand Rapids. This mining project is probably the last major taconite facility to be built on the Mesabi Range. The facility construction is nearly 65% completed and will cost $1.8 billion dollars when finished, with an expected design capacity of 7.0 million tons per year of fluxed, standard and DR-grade pellets. All of the required permits required for construction and operation are in place for the designed capacity. The Crushing & Concentration Facility is separated from the Pelletizing Facility, as shown in the Figure 3. When completed, we believe this mining operation will be one of the lowest cost iron ore pellet producers in North America. Figure 3 – Essar Steel Minnesota site layout. The taconite resource at this project site was originally mined by Butler Taconite Mining Company, a company managed by Hanna Mining Company. Butler Taconite was a jointly owned mining operation and was in production from 1967 to 1985. When one of the owners declared bankruptcy, the other two owners closed the operation, and the facility underwent demolition. In 2007, Essar Steel Holdings acquired Minnesota Steel Industries (MSI), a development company owned by some of the mineral owners on the property. In 2008, Essar renamed the MSI to Essar Steel Minnesota, LLC (ESML) The ESML deposit is a low grade iron-formation with magnetite as the predominate resource. In 2011, ESML conducted a diamond drilling program to bring the deposit into compliance with Canadian 79 �NI43-101 ore reserve standards. A total of 63 diamond drill holes were drilled across the length of the deposit totaling 41,720 feet. The new drilling information combined with the historic drilling, defined the ore zone over the strike length and down dip of the deposit. The drilling program identified 1.7 billion tons of proven and probable ore, having an average stripping ratio of 1.69 and an average weight recovery of 29.1%. Within the project area, the Biwabik Iron Formation is underlain by the Pokegama Quartzite and is overlain by the Virginia Formation. The Biwabik Iron Formation subcrop and the Virginia Formation are overlain by scattered Cretaceous marine deposits, and all these formations are covered by glacial drift. The Biwabik Formation strikes generally E-NE (065°) with a 5° S-SE dip. The stratigraphy in the ESML Project area was characterized in detail by Hanna Mining Company geologists. The Biwabik Iron Formation has four distinct members: Upper Slaty Member – Slaty non-magnetic taconite; Upper Cherty Member – Cherty weakly magnetic taconite; Lower Slaty Member – Slaty non-magnetic taconite; and Lower Cherty Member – Cherty magnetic taconite. The Lower Cherty Member has been divided into ten distinct subunits. The ore zone lies in the LC4A, LC4B, LC4C, and LC5A subunits of the Lower Cherty Member, and averages about 200 feet thick. The La Rue fault system runs along the strike length of the formation, bisecting the historic pits within the project area. The Patrick Shear Zone crosscuts the deposit between Pits 2 and 5 of the original Butler Taconite pits. The taconite has been locally oxidized along these fault zones. Figure 4 is a map depicting the geology and historic drilling across the project site. A basic geologic column at the ESML Project site is shown in Figure 5. Figure 4. Map of bedrock geology and historic drilling (yellow dots) within the project site. Geologic contacts are black; faults are red. 80 �Figure 5. Geologic Column at the Essar property. During the 2011 diamond drilling campaign, down-dip drill holes intersected a well-defined contact between the Virginia Formation and the Upper Slaty Horizon of the Biwabik Iron Formation. This very distinctive contact zone is characterized by a chaotic mix of broken and deformed strata, uncharacteristic for the uniformly bedded strata above and below this zone. The zone of turbulence matches contacts found in the Gunflint Iron Formation which are attributed to the Sudbury Meteorite Impact event. Above the zone, little evidence of magnetic beds would seem to suggest a disruption in the deposition of ferrous iron minerals by this catastrophic event. Figure 6 provides photographic documentation of the contact in two separate diamond drill holes. 81 �Figure 6. Images of core that intersected the Upper Slaty – Virginia Contact. 82 �Taconite mining will start with a development cut adjacent to the new crushing complex and progress down dip into the old Butler Pit 5 taconite pit. Pumping is presently in progress to remove water from the historic mine pit. To date, the water level in Pit 5 has been lowered approximately 60 feet, with another 80 feet needed to reach pit bottom. The mine will be developed using a level bench configuration, hydraulic shovels, and a small fleet of 240-ton haul trucks. The initial mining area has been pre-stripped of glacial overburden north of the Butler Pit 5, and the first blast will be directly in the taconite production zone, as shown in Figure 7 (labeled “Mine Development”). Figure 7. Map showing facilities and planned location for initial mine development adjacent to the historic Butler Taconite pit. 83 �MAGNETATION Joe Nielsen (Magnetation Inc.). This trip will visit a unique new iron ore venture currently operating in the Bovey area. Founded in 2006, the privately-held Magnetation Inc. was created with the intention of utilizing magnetic separation technology to capture iron ore particles left over from previous mining operations that existed on the iron range dating back to the 1890s. Owners Al Fritz and Rod Hunt focused the company’s early efforts on research and development of a beneficiation process centered on the Ferrous Wheel®, a technology Al Fritz invented in the 1970s that uses permanent magnets to separate iron ore from waste materials and produce an upgraded iron ore concentrate. The location near Bovey is the second of three plants operating on the Mesabi Iron Range; the others are near Keewatin and Chisholm. The Bovey plant commenced operations in May 2012, and produces about 1.2 million tons per-year of iron ore concentrate. The operations consist of excavation and transport of iron-bearing tailings to the concentrator facility, where the iron-rich portions are reclaimed. The resulting iron ore concentrate is then trucked to the Jessie LoadOut, where it is shipped by rail to Magnetation customers. A new pelletizing plant under construction in Reynolds, Indiana, will begin processing ore concentrate late 2014. The Ore Magnetation mines the iron ore tailings from mining days long ago. We dig up all the discarded tailings in the Pit and bring them to the Plant one truck load at a time. In the plant the tailings go through various stations to become Iron Concentrate. The trucks then take the Concentrate to our train loading station to ship to our customers. The Pit Unlike a conventional open mine pit our material requires no drilling and blasting. We dig up what was left behind as waste, remove the iron, and return the material back to the basin to be eventually replanted with vegetation. The Plant In the plant, we are using Magnetation’s unique processes and equipment to remove the iron from our feed material. We are constantly adjusting processes to accommodate the variations in feed from the pit and continually improving to further the iron yield from the feed material. www.magnetation.com 84 ��FIELD TRIP 5 Saturday, May 17, 2014 VISIONS OF MATURI: THE GEOLOGY OF THE SOUTH KAWISHIWI INTRUSION LEADER: Dean M. Peterson (Duluth Metals, Ltd.) INTRODUCTION Twin Metals Minnesota (TMM), a private joint venture company owned by Duluth Metals Ltd (60%) and Antofagasta plc (40%), is currently finishing a robust prefeasibility study (due mid 2014) to develop the Maturi Cu-Ni-PGE deposit in the northern part of the South Kawishiwi Intrusion (SKI). The mineralization at Maturi is confined to the basal 500’ of the intrusion and has been the focus of numerous ILSG-based field trips in recent years. In October 2013, UMD’s Precambrian Research Center hosted a workshop on Cu-Ni-PGE deposits in the Lake Superior area and the Duluth Complex fieldtrip guidebook of that workshop (Severson et al., 2013) is perhaps the most up-to-date geologic description of Duluth Complex Cu-Ni-PGE deposits. As such, this field trip is focused mainly on the “rest of the rocks” of the SKI as a way to put the Cu-Ni-PGE ores in better context to the vast majority of the rocks of the intrusion. The importance of understanding these rocks and the overlying glacial deposits will be ever increasing as the TMM Project goes into bankable feasibility, environmental review and permitting, since virtually all of the water in the region (surficial and deep groundwater) interacts mostly with the “other rocks” of the SKI. Duluth Metals’ understanding of the SKI has been facilitated by extensive bedrock geologic mapping in the local region (22,200 outcrops mapped), drilling (2,300,000 feet in 1,556 holes), geochemistry (4,500 tills, 1,800 rocks, and ~110,000 drill core assays), and geophysics (3 recent airborne VTEM surveys covering 178,600 acres, borehole EM surveys, Titan-24 survey, and RIM cross-hole imaging). REGIONAL GEOLOGIC SETTING, DULUTH COMPLEX The Duluth Complex and associated intrusions of Keweenawan age (~1.1 billion years) in northeastern Minnesota constitute one of the largest mafic intrusive complexes in the world, second only to the Bushveld Complex of South Africa (Miller et al., 2002). These rocks cover a 2,200 square mile (5,700 square km) arcuate area associated with the two strongest gravity anomalies (+50 and +70 milligals) in North America, implying intrusive roots over 8 miles (13 km) deep (Allen and others, 1997). The comagmatic flood basalts and intrusive rocks underlying much of northeastern Minnesota were emplaced during development of the Mesoproterozoic Midcontinent rift, which can be traced geophysically from exposures in the Lake Superior region along a 1250 mile (2,000 km) long, segmented, arcuate path to Kansas and Lower Michigan. The Duluth Complex is defined as the more or less continuous mass of mafic to felsic plutonic rocks that extends for >170 miles (275 km) in an arcuate fashion from Duluth nearly to Grand Portage (Fig. 1). It is bounded by a footwall of Paleoproterozoic sedimentary rocks and Archean granite-greenstone terranes (Peterson and Severson, 2002), and a hanging wall largely of comagmatic, rift-related flood basalts and hypabyssal intrusions of the Beaver Bay Complex. In genetic terms, the Duluth Complex is composed of multiple discrete intrusions of mafic to felsic tholeiitic magmas that were episodically emplaced into the base of a volcanic edifice between 1108 and 1098 Ma. The geology of the Duluth Complex and adjacent areas has recently been described in two major publications by the Minnesota Geological Survey (MGS). These include a 1:200,000 scale regional bedrock geological map of northeastern Minnesota (Miller et al., 2001), and a comprehensive written description of the geology depicted on this map (Miller et al., 2002), commonly referred to as the “bible” by geologists working on Duluth Complex geology. Readers’ interested in more detailed descriptions of 86 �the geologic setting of the Duluth Complex should begin their quest for knowledge by downloading these publications from the MGS website (ftp://mgssun6.mngs.umn.edu/pub2/). Within the nearly continuous mass of intrusive igneous rock forming the Duluth Complex, four general rock series are distinguished on the basis of age, dominant lithology, internal structure, and structural position within the complex. Felsic series—Massive granophyric granite and smaller amounts of intermediate rock that occur as a semi-continuous mass of intrusions strung along the eastern and central roof zone of the complex, that were emplaced during early stage magmatism (~1108 Ma). Early gabbro series—Layered sequences of dominantly gabbroic rocks that occur along the northeastern contact of the Duluth Complex, emplaced during early stage magmatism (~1108 Ma). Anorthositic series—Structurally complex suite of foliated, but rarely layered, plagioclase-rich gabbroic anorthosite emplaced throughout the complex during main stage magmatism (~1099 Ma). Layered series—Suite of stratiform troctolitic intrusions that comprises at least 11 variably differentiated mafic layered intrusions that occur mostly along the base of the Duluth Complex. These intrusions were emplaced shortly after the Anorthositic series (~1099 Ma). Figure 1. Generalized geologic map of northeastern Minnesota (modified from Miller et al., 2002). 87 �LOCAL GEOLOGIC SETTING, THE SOUTH KAWISHIWI INTRUSION The SKI consists almost entirely of troctolitic rocks that generally dip gently to the southeast. However, it is not well known that shallow southwesterly dipping troctolite of the upper SKI in the northeastern portion of the intrusion defines an asymmetric funnel-shaped body that emerged from the Nickel Lake macrodike. The basal mineralization of the SKI is exposed in an arc-shaped area that extends from the Serpentine deposit, in the southwest, to the Spruce Road deposit, in the northeast (Fig. 2). Footwall rocks include the Paleoproterozoic Virginia Formation, Biwabik Iron Formation and Archean Giants Range Batholith, the latter is the dominant footwall rock type. Figure 2. Simplified geological and ore deposit map of the northwestern Duluth Complex. GEOLOGIC MAPPING Detailed geological mapping, generally at a scale of 1:5,000 or greater, has been the most important component of Duluth Metals’ understanding of the geology of the SKI. Geologists associated with the company have mapped over 17,000 outcrops (~1,400 total acres of outcrop covering >77,000 acres of ground) within the SKI and adjacent rock units. True understanding of the SKI rocks in their natural environment, In The Field, has led to much improved interpretation of the rocks observed in drill core, the geochemistry of glacial tills, and the interpretation of geophysical studies. Geologists working on the mineral deposits within the Duluth Complex that have not spent a considerable amount of time in the field mapping the rocks will have a difficult time interpreting what they see and log in drill core. Such in the field knowledge is especially important as projects advance and true 3D geological models have to be 88 �constructed for mine planning in feasibility studies. All geological interpretations will be scrutinized and audited once the banks get involved in project financing. A historical account of geological mapping programs within the SKI is presented in Figure 3, and the aerial extent of mapped rock types within the bounds of the SKI are given in Table 1. It is important to note that there are nearly 4,600 acres of sulfide-bearing bedrock exposed on the Earth’s surface within the SKI (1,230 gossanous outcrops mapped). Figure 3. History of geological mapping in the SKI. Table 1. Distribution of rock types exposed on the surface in the SKI. Rock Type Acres % Area 1 6 14 14 206 482 3,112 0.00% 0.01% 0.02% 0.02% 0.31% 0.72% 4.67% Sulfide-bearing Troctolite Augite Troctolite Anorthositic Troctolite 4,594 5,554 20,364 6.89% 8.34% 30.56% Troctolite Grand 32,286 66,633 48.45% 100.00% Diabase dike Iron Formation xenoliths Sandstone xenoliths Ultramafic rocks Gabbroic xenoliths Basalt xenoliths Anorthosite xenoliths 89 �IGNEOUS STRATIGRAPHY & LITHOGEOCHEMISTRY Integration of geological, geochemical, and geophysical data over the last few years by Duluth Metals has resulted, via integration, in a new interpretation of the bulk igneous stratigraphy of the SKI. The proposed new stratigraphy of the intrusion is presented in Figure 4 and consists of five regionally extensive units. From the top down these units include: Upper SKI – Medium to coarsegrained, locally layered troctolite and anorthositic troctolite. Well layered as defined by olivine-rich horizons. SKI Break – Chromium oxide-rich, heterogeneous dunite and melatroctolite. Interpreted to be a magmatic unconformity within the SKI. Middle SKI – Medium to coarsegrained, locally layered troctolite and anorthositic troctolite. Layering defined by olivine. Main AGT – Coarse-grained, homogeneous, augite troctolite with highdensity ophitic Augite grains. This unit is never layered and is interpreted to be the solidified basaltic liquid that carried the phenocrysts and immiscible sulfide droplets of the BMZ. BMZ – Heterogeneous, sulfide-bearing troctolitic rocks. Interpreted to have formed from a sulfide-rich, crystalladen magmatic slurry. Figure 4 Interpreted igneous stratigraphy of the SKI. A simplified composite lithogeochemical compilation profile through the SKI is presented in Figure 5. This geochemical compilation clearly displays the strong correlation of economically important base (Cu, Ni) and precious (Pt, Pd, Au) metals into the basal mineralized zone (BMZ) and adjacent footwall granitoids. In addition, the common 3:1 Cu:Ni ratio of the BMZ is clearly completely different than the vast majority of the intrusion, where in fact Cu averages about 100 ppm and Ni averages about 200 ppm. Strong olivine layering in the Middle and Upper SKI can easily be seen in the Mg % profile and the break between the Middle and Upper SKI is clearly displayed in the Cr (ppm) profile. A geologic cross section roughly along Minnesota State Highway #1 is presented in Figure 6, and integrates hundreds of thousands of feet of drilling into this new regional context. Note the extremely large xenolith of Anorthositic Series rocks and North Shore Volcanic Group lavas in the center of the intrusion. This xenolith of older rocks played an important role in the development of higher grades of Cu, Ni, and PGEs in the Maturi deposit compared to other deposits in the district (Peterson and Boerst, 2013). 90 �Figure 5. Generalized lithogeochemical compilation profile through the SKI. Data from continuous sampling of Duluth Metals drill holes within the western SKI in the Maturi Deposit in 2007 – 2009 (drill holes MEX-072, Mex-109, and MEX-155), and along the northeastern margin of the intrusion in 20122013 (drill holes 12-DM-14, 12-DM-15, and 13-DM-45). Figure 6. Geologic cross section through the northern South Kawishiwi Intrusion. 91 �DRILLING Exploration for Cu-Ni deposits at the base of the Duluth Complex began in 1948, about 12 miles southeast of Ely, MN, when strongly mineralized rocks were uncovered in an excavation used to source road material for Spruce Road. Local prospector Fred S. Childers of Ely noted copper stains in the material and he, along with Roger V. Whiteside of Duluth, began searching along the basal contact in the vicinity of the Kawishiwi River. In 1951, they diamond drilled a 188 foot deep hole and intersected mineralized troctolite that averaged 0.36% Cu and 0.13% Ni. In 1952, the International Nickel Company (INCO) began intensive exploration efforts along the zone that coincided with the basal contact and eventually picked up the Childers-Whiteside properties (Spruce Road and Maturi deposits). Since this initial hole, an additional 1,555 holes (Fig. 7) have been drilled in the intrusion (2,300,000 feet of total drilling) in eight prospective areas (Maturi, Spruce Road, Dunka Pit, Maturi SW, Serpentine, Filson Creek, Birch Lake, and East Shore). Figure 7. Historic perspective of the amount of drilling completed in the SKI since 1951. 92 �ORE DEPOSITS, TWIN METALS MINNESOTA A detailed description of Maturi deposit has recently been published (Peterson and Boerst, 2013) and field trip participants that wish to delve deeper into the geology and geochemistry of the basal mineralized zone (BMZ) should acquire that guidebook. However, since publication of that deposit description, Duluth Metals has received an updated independent NI 43-101 Technical Report completed by AMEC E&C Services Inc. (AMEC) on the Maturi and Maturi SW deposits. The extent of the resource categories for the Maturi, Maturi SW, Birch Lake, and Spruce Road deposits are presented in Figure 8. The updated study utilizes 922 drill holes and 312 wedge offsets, and reports a significant portion of the Maturi deposit upgraded to the Measured Resource category. The mineral resources have been estimated using CIM Definition Standards for Mineral Resources and Reserves dated November 2010. Figure 8. Map of the NI 43-101 qualified resources of Twin Metals Minnesota within the SKI. The majority of the increase in total contained metals in the 2014 resource estimates reflects the addition of the Maturi Southwest Deposit. The updated mineral resources estimate has 295 million tons in the Measured category at a 0.3% copper cut-off in the Maturi Deposit, which may potentially provide an early start-up area for future mining. The change in category for a significant portion of the Indicated Resource to Measured Resource reflects the excellent continuity of the resource demonstrated by the close-spaced fence drilling completed at Maturi. Base case qualified resources for these deposits are given in Table 2, and the combined metal contents for the measured, indicated, and inferred mineral resources are provided in Table 3. The enormity of the metal resource in these deposits is certainly quite staggering and clearly shows why so much time, effort, and money has been spent on these deposits in recent years. 93 �Table 2. Twin Metals Minnesota’s NI 43-101 base case Qualified Mineral Resources Resource Cut-off Tons Cu Deposit Ni Pt Pd (Mst) (%) Name Class Cu (%) (%) (ppm) (ppm) Au (ppm) Maturi Maturi Maturi Measured Indicated Inferred 0.3 0.3 0.3 295 774 562 0.63 0.58 0.51 0.20 0.19 0.17 0.148 0.160 0.138 0.345 0.360 0.317 0.084 0.085 0.071 Maturi Southwest Maturi Southwest Indicated Inferred 0.3 0.3 103 32 0.48 0.43 0.17 0.15 0.080 0.065 0.185 0.157 0.048 0.041 Birch Lake Birch Lake Indicated Inferred 0.3 0.3 100 239 0.52 0.46 0.16 0.15 0.233 0.180 0.511 0.370 0.114 0.087 Spruce Road Inferred 0.3 480 0.43 0.16 Table 3. Contained Metals in the TMM Resource (effective date October 8, 2013) * Metal Copper Nickel Platinum Palladium Gold Measured Resource 3.7 billion lbs. 1.2 billion lbs. 1.3 million ozs. 3.0 million ozs. 0.7 million ozs. Indicated Resource 11.0 billion lbs. 3.5 billion lbs. 4.5 million ozs. 10.2 million ozs. 2.5 million ozs. Measured + Indicated 14.7 billion lbs. 4.7 billion lbs. 5.8 million ozs. 13.2 million ozs. 3.2 million ozs. Inferred Resource 12.3 billion lbs. 4.2 billion lbs. 3.6 million ozs.** 8.0 million ozs.** 1.8 million ozs.** * Based on mineral resources estimated at base case 0.3% Cu cut-off grade; for tons and grade see Tables 2 further below. ** Contained ounces of platinum, palladium, and gold in the Inferred category do not include the Spruce Road deposit. Additional exploration potential highlighted by AMEC outside of the four mineral resources (Maturi, Maturi Southwest, Birch Lake and Spruce Road deposits) and in addition to the TMM defined mineral resource are considered targets for further exploration. An estimate of the exploration potential is between 1.3 to 2.1 billion tons contiguous to the boundaries of the four deposits (Fig. 8). For Maturi and Maturi Southwest, AMEC Assumed that mining, processing and G+A costs would be approximately $15/t, $8/t and $2.50/t respectively for a total of $25.50/t. At Birch Lake and Spruce Road, AMEC assumed that mining, processing and G+A costs would be approximately $16/t, $12/t and $2/t respectively for a total of $30/t. This indicates a breakeven NSR of approximately $30 per ton. Resources meeting an NSR cutoff of $30/t approximately equate to a copper cutoff of 0.3%. The geological subunit within basal mineralized zone (BMZ) in the Maturi Deposit, known as the Stage 3 (Peterson and Boerst, 2013), hosts higher grades ore (172 million short tons at 0.72% Cu, 0.23% Ni, 0.188 ppm Pt, 0.438 ppm Pd and 0.104 ppm Au in the Measured category) that is a subset of the base case mineral resource estimate that may have potential as an early start-up area. The AMEC 2014 Technical Report update on the Measured, Indicated and Inferred categories is presented below for Maturi (Table 4), Maturi SW (Table 5), Birch Lake (Table 6), and Spruce Road (Table 7) deposits. 94 �Table 4. Maturi Deposit Mineral Resources by Copper Cutoff Grade (base case is highlighted) Measured Mineral Resources Cut-off Cu (%) Tons (Mst) Cu (%) Ni (%) Pt (ppm) Pd (ppm) Au (ppm) 0.2 0.3 0.4 0.5 0.6 312.5 295.3 262.2 224.9 174 0.61 0.63 0.66 0.70 0.74 0.20 0.20 0.21 0.22 0.24 0.143 0.148 0.157 0.168 0.179 0.334 0.345 0.366 0.392 0.419 0.081 0.084 0.089 0.094 0.101 Indicated Mineral Resources 0.2 829.4 0.3 774.2 0.4 678.2 0.5 518.2 0.6 366.8 0.56 0.58 0.61 0.66 0.71 0.18 0.19 0.20 0.21 0.22 0.154 0.160 0.171 0.192 0.209 0.345 0.360 0.384 0.431 0.470 0.082 0.085 0.091 0.101 0.109 Inferred Mineral Resource 0.2 804 0.3 562 0.4 399 0.5 266 0.6 147 0.43 0.51 0.57 0.63 0.70 0.14 0.17 0.19 0.20 0.22 0.117 0.138 0.162 0.194 0.233 0.265 0.317 0.370 0.437 0.523 0.060 0.071 0.083 0.097 0.114 *Effective Date is 8 October 2013. *Dr. Harry Parker, RM SME, is the QP for the estimate and is a Professional Geologist licensed in Minnesota. *The resources are based on a US$30/t NSR that assumes a mining cost of $15.00/t, a process cost of $8.00/t and G&A charges of $2.50/t; global metallurgical recoveries of 94.3% (Cu), 60.8% (Ni), 82.3% (Au), 36.1% (Pd), and 42.5% (Pt); and long-term consensus metal prices of $3.00/lb Cu, $9.50/lb Ni, $1,200/troy oz Au, $700/troy oz Pd and $1,650/troy oz Pt. *The NSR equates to a 0.3% Cu cut-off grade. *Figures have been rounded and may not sum. * Mst = million short tons. Table 5. Maturi Southwest Deposit Mineral Resources by Copper Cutoff Grade (base case is highlighted) Indicated Mineral Resources Cut-off Cu (%) Tons (Mst) Cu (%) Ni (%) Pt (ppm) Pd (ppm) Au (ppm) 0.2 0.3 0.4 0.5 0.6 131 103 71 40 16 0.43 0.48 0.53 0.59 0.67 0.15 0.17 0.18 0.20 0.22 0.071 0.080 0.093 0.108 0.124 0.164 0.185 0.217 0.256 0.294 0.042 0.048 0.055 0.064 0.071 Inferred Mineral Resource 0.2 57 0.3 32 0.4 16 0.5 7.2 0.6 3.2 0.35 0.43 0.51 0.60 0.66 0.13 0.15 0.17 0.20 0.22 0.052 0.065 0.082 0.102 0.115 0.126 0.157 0.197 0.251 0.279 0.033 0.041 0.050 0.063 0.069 *Footnotes the same as in Table 4. 95 �Table 6. Birch Lake Deposit Mineral Resources by Copper Cutoff Grade (base case is highlighted) Indicated Mineral Resources Cut-off Cu (%) Tons (Mst) Cu (%) Ni (%) Pt (ppm) Pd (ppm) Au (ppm) 0.2 0.3 0.4 0.5 111.9 99.7 85.4 54.9 0.49 0.52 0.55 0.60 0.15 0.16 0.17 0.18 0.217 0.233 0.247 0.269 0.474 0.511 0.543 0.591 0.106 0.114 0.120 0.130 0.41 0.46 0.51 0.58 0.13 0.15 0.16 0.18 0.156 0.180 0.203 0.228 0.320 0.370 0.423 0.480 0.076 0.087 0.098 0.111 Inferred Mineral Resource 0.2 313.1 0.3 239.2 0.4 158.4 0.5 76.8 * Effective Date is 15 September 2012. * Dr. Harry Parker, RM SME, is the QP for the estimate and is a Professional Geologist licensed in Minnesota. * The resources are based on a US$30/t NSR that assumes a mining cost of $16.00/t, a process cost of $12.00/t and G&A charges of $2.00/t; global metallurgical recoveries of 90.8% (Cu), 57.4% (Ni), 863.3% (Au), 63.6% (Pd), and 55.2% (Pt); and long-term consensus metal prices of $3.00/lb Cu, $9.38/lb Ni, $1,050/troy oz Au, $850/troy oz Pd and $1,840/troy oz Pt. * The NSR equates to a 0.3% Cu cut-off grade. * Figures have been rounded and may not sum. * Mst = million short tons. Table 7. Spruce Road Deposit Mineral Resources by Copper Cutoff Grade (base case is highlighted) Inferred Mineral Resource Cut-off Cu (%) Tons (Mst) Cu (%) Ni (%) 0.2 0.3 0.4 0.5 674 480 254 101 0.38 0.43 0.50 0.57 0.14 0.16 0.18 0.21 * Effective Date is 15 September 2012. * Dr. Harry Parker, RM SME, is the QP for the estimate and is a Professional Geologist licensed in Minnesota. * The resources are based on a US$30/t NSR that assumes a mining cost of $16.00/t, a process cost of $12.00/t and G&A charges of $2.00/t; global metallurgical recoveries of 90.8% (Cu), 57.4% (Ni); and long-term consensus metal prices of $3.00/lb Cu, $9.38/lb Ni. * The NSR equates to a 0.3% Cu cut-off grade. * Figures have been rounded and may not sum. * Mst = million short tons. 96 �DESCRIPTION OF FIELD TRIP STOPS The location of the South Kawishiwi Intrusion field trip stops is presented in Figure 9, and short descriptions of the geology and important take-away knowledge (bedrock geology, glacial geology and dynamics, exploration geochemistry, environmental review, etc.) from each stop is given in the following descriptions. It is important to note that the whole northern SKI is located in the scoured bedrock terrain of the Wisconsinan cycle of the Pleistocene Laurentide ice sheet (herein the Rainy Lobe about 12,000 years ago). The local end moraine of the Rainy Lobe is located immediately south of Stop #1 along the new Tomahawk Road. Work by Duluth Metals has positively shown that the mean transport length (the distance where ½ of material in the glacial deposits is sourced from) in the field trip area is <0.5 miles. Figure 9. Location map of the field trip stops in the South Kawishiwi Intrusion. 97 �STOP 1: Upper SKI Troctolite UTM NAD83 Coordinates: 598664E, 5287937N. PLS: T61N, R11W, S26 Glacially scoured outcrops of weakly layered anorthositic troctolite perfectly exposed at the bottom of a State of Minnesota gravel pit. As is typical in much of the Upper and Middle SKI, the rocks dip shallowly to the southeast with olivine layers striking 35° and dipping 8° to the southeast. Exposures such as this are extremely important in environmental reviews of proposed mining operations as they display the natural occurrence of bedrock surfaces beneath the thin veneer of glacial sediments. As we at Duluth Metals have become aware during the course of prefeasibility studies of TMM’s Maturi deposit, many hydrogeologists (consulting, governmental, academic) believe that there has to be several hundred feet of fractured bedrock immediately below the Quaternary-Mesoproterozoic contact. If these envisioned fractured bedrock zones exist, they would clearly be permeable (even to the Precambrian geologist author) to near-surface groundwater flow, thus requiring identification and study any environmental review of a mining operation. Please note the glacial scours and examine the glacial till bank on the north edge of the gravel pit. PGEenriched sulfide-bearing boulders (with >0.5% Cu) in these glacial deposits have been obtained from this till bank. The nearest identified and exposed up-ice Cu-Ni-PGE occurrence is the Spruce Road deposit, approximately 7 miles to the NNE, which is over 15 times the mean glacial transport distance. STOP 2: Mesabi Black Quarry, Coldspring Granite Company UTM NAD83 Coordinates: 599131E, 5289159N. PLS: T61N, R11W, S24 Coldspring’s Mesabi Black® quarry opened in 2000 and furnishes the dimension stone industry with poikilitic gabbroic anorthosite. The company utilizes a mix of mining techniques at the quarry to harvest the blocks but the technique that will most interest field trip participants is the diamond wire cuts. Get ready, this is perhaps the best locality in existence where one can examine Anorthositic Series rocks of the Duluth Complex in 3D and wrap your mind around the viscosity of a plagioclase crystal mush. The quarry is in the heart of the USGS mapping of the Harris Lake area (Foose and Cooper, 1978). Troctolitic rocks in the area dip (as defined by olivine layering and plagioclase crystal foliation) moderately to the southeast at approximately 25°. The operation ships blocks of rocks from a large Anorthositic Series xenolith from within the Upper SKI. For the ILSG regulars, an anecdote herein is required… Duluth Metals geologists invited Dr. Paul Weiblen to visit the quarry in the spring of 2013. After about an hour looking at Anorthositic Series rocks in perfect 3D cuts, Paul stated to the author (and I quote), “I learned by far more today about the Anorthosite Series than I have over the last 50 years mapping and studying these rocks”. STOP 3: The SKI Break - Middle to Upper SKI Troctolite Contact UTM NAD83 Coordinates: 600417E, 5292338N. PLS: T61N, R10W, S18 Duluth Metals geologists first took interest in the rocks of this field trip stop in 2011 during follow up field work around several highly anomalous till geochemistry (Cu, Ni, Pt and Pd) samples. Numerous angular boulders of melatroctolite and peridotite were discovered containing highly anomalous Cu-NiPGE geochemistry. A detailed mapping and sampling program ensued and the locally sulfide-bearing, chromium oxide-rich ultramafic SKI Break was discovered in outcrops along the valley of Keeley Creek. This stop will involve a moderately long walk (½ mile) walk through large exposures of anorthosite xenolith-rich, shallow dipping (here to the west-southwest), foliated troctolite of the basal portion of the Upper SKI into the recessive weathering SKI Break. As we walk along U.S. Forest Service road 1468 please note the rugged topography of the glacially scoured outcrops of the Upper SKI. The rugged nature of the topography in the Upper SKI becomes more subdued in glacially polished outcrops of the Middle SKI. Subtle differences in geochemistry/mineralogy/texture of these SKI units must have resulted in differential weathering and saprolite development of these similar troctolite units of the SKI. 98 �STOP 4: Middle SKI Troctolite UTM NAD83 Coordinates: 597134E, 5293919N. PLS: T61N, R11W, S11 In the summer of 2013, geologists from Duluth Metals and Twin Metals Minnesota completed detailed structural mapping of nearly 100 sites around the Maturi deposit in the initial field phase of an upcoming geohydrology program for the project. Nearly 800 structural elements were measured and 1,321 new outcrops were mapped in detail. The goal of this work was to attempt to define the dip direction and angle of numerous topographic lineaments that cross the SKI (refer back to Figures 6, 8, and 9). Many workers and NGO activists believe that each and every one of these topographic features represent deepseated faults that directly transport groundwater throughout the entire mass of the intrusion. We at Duluth Metals would beg to differ and it is hoped that a conversation on this topic originates on the outcrop. A simplified single example of this type of geological mapping is presented in Figure 10, and the original field sheet will be available to examine during the field trip. At this stop, massive, moderately layered and foliated troctolite (strike 45°, dip 22°) of the Middle SKI is exposed in numerous outcrops. This stop epitomizes the “Sea of Troctolite” that occurs throughout the vast majority of the SKI. We will take a short walk to the south onto large massive exposures of glacially scoured troctolite to investigate weak jointing developed along the eastern margin of the bedrock exposures and discuss the interpretation of the NNE trending topographic lineaments. Figure 10. Example of detailed structural mapping. STOP 5: Main Augite-Troctolite, the “Main AGT” UTM NAD83 Coordinates: 594743E, 5296267N. PLS: T62N, R11W, S33 Recent road cut along the south side of Minnesota Highway #1 of massive, extremely homogeneous augite troctolite of the Main AGT unit of Severson (1994). Troctolite of the Main AGT unit differs from the Middle and Upper SKI troctolite in two distinctive ways: 1) ophitic augite crystals are black, distinctly associated with Fe-Ti oxides + apatite, and occur as high-density ophitic crystals from 1 to 3 inches in diameter. In the Middle and Upper SKI, ophitic augite crystals are brown, not associated with Fe-Ti oxides, and occur as large (up to 15 inches) low-density grains; and 2) The Main AGT is never layered. Geologists at Duluth Metals interpret the units’ homogeneity and lack of layering as evidence that the Main AGT magma lacked phenocrysts of olivine and plagioclase and represents the end product of topdown and bottom-up solidification of a basaltic liquid. We currently interpret the Main AGT as the solidification of much of the “carrier liquid” of the underlying sulfide-bearing BMZ magmatic slurry. STOP 6: Basalt Xenolith in the BMZ, the Spruce Road Deposit UTM NAD83 Coordinates: 599404E, 5298990N. PLS: T62N, R11W, S24 A short field trip stop up onto a small knob of basalt hornfels within the center of the Spruce Road Cu-Ni deposit. Sulfide-bearing troctolitic rocks of the Spruce Road Cu-Ni deposit are distinctly different in several ways to similar rocks within the Maturi Cu-Ni-PGE deposit. First and foremost is the fact that the precious metal content (Pt-Pd-Au) of Spruce Road ores are much less than within Maturi. The second 99 �fact is that the Spruce Road deposit contains a large amount of sulfide-barren xenoliths. In fact, the mapped proportion of barren xenoliths at Spruce Road approaches 15% of the total rocks within the heart of the deposit (Table 8). At Maturi, such accessory and exotic xenoliths account for <<1% of the Cu-NiPGE mineralized zone. Table 5-8. Extent of mapped rock types in the Spruce Road deposit. SKI Rock Type Acres Extent 187.3 82.5% 6.8 3.0% 14.7 6.5% Barren troctolite (early chill margin?) 9.9 4.4% Biwabik Iron Formation 5.1 2.3% Anorthosite 3.0 1.3% Virginia Formation 0.2 0.1% Sulfide-bearing, heterogeneous troctolite Sulfide-bearing melatroctolite to dunite Xenoliths Basalt The difference between these deposits is interpreted to be the result of the timing of magma injection. The Spruce Road deposit is believed to have formed prior to Maturi and the lithology and amount of xenoliths are the end product of the system cleaning out the pathways that the magma traveled upwards from depth (Peterson and Boerst, 2013). STOP 7: U.S. Forest Service Borrow Pit, BMZ in the Spruce Road Deposit UTM NAD83 Coordinates: 598826E, 5298384N. PLS: T62N, R11W, S25 Beginning in the late 1940s, the U.S. Forest Service utilized locally derived glacial tills and weathered bedrock gossans as road building materials during the construction of the Spruce Road. As we take a short hike into one of these borrow pits, we will walk by the 1973 INCO bulk sample site in the Spruce Road deposit. This short stop will examine the bottom of an old borrow pit where participants can walk on and sample sulfide-bearing troctolite gossans. Please note the friable nature of the rocks in the weakly saprolitic exposure and look for rounded core-stones where weathering over the eons was less intense. STOP 8: Sulfide-bearing Troctolite & layered Melatroctolite, Maturi SW Deposit UTM NAD83 Coordinates: 590572E, 5293036N. PLS: T61N, R11W, S7 Classic roadside exposures of heterogeneous sulfide-bearing troctolite and layered melatroctolite of Severson’s (1994) Basal Heterogeneous (BH) and Ultramafic 3 (U3) units of the SKI. A large core-stone is well exposed in the weakly saprolitic heterogeneous troctolite outcrop. Several small xenoliths of finegrained troctolite can be observed on top of the outcrop and are interpreted as Stage 1 chilled margin autoliths (Peterson and Boerst, 2013). Within the exposure of the overlying U3 layered melatroctolite, olivine layers strike 17° and dip steeply 51° to the ESE. The steep dip is apparently associated with two defined north-south trending faults east of these exposures. Recent drilling by Twin Metals Minnesota in this area has led to the definition of the Maturi SW deposit (see Fig. 8 and Tables 2 and 5). STOP 9: Basal Heterogeneous, Sulfide-poor Troctolite UTM NAD83 Coordinates: 590313E, 5292211N. PLS: T61N, R11W, S18 At this stop, we’ll examine perhaps the best exposure of Severson’s (1994) BH unit in the whole SKI. The heterogeneous troctolitic rocks at this stop are generally poorly mineralized and thus lack a gossanous saprolitic weathering profile which lets one see the true nature of the heterogeneity within the troctolite. I believe that all geologists who ever will log drill core within the Cu-Ni deposits of the Duluth Complex (or who attempt to model such deposits for mine planning purposes) should be required to spend several days examining the rocks within the area around both Stops 8 and 9. All participants should imagine a drill core cutting this exposure and how they would interpret the geology of that core without first examining this outcrop. Such thoughts are why the Precambrian Research Center’s field camp has for many years had its students complete a 1:5,000 scale bedrock geology map of this area. 100 �STOP 10: Giants Range Batholith, the Footwall UTM NAD83 Coordinates: 590518E, 5296544N. PLS: T62N, R11W, S31 The footwall rocks of the whole northern SKI consist of the Neoarchean Giants Range batholith (GRB), such as is exposed along Highway 1 at this field trip stop. The rocks here consist of porphyritic hornblende quartz monzonite with distinctive 1-2 cm potassium feldspar phenocrysts. The massive nature of this unit creates an excellent footwall for the intrusions Cu-Ni-PGE deposits as it lacks bedding and thus rarely (if ever) gets incorporated into the mineralized zone as barren xenoliths. In addition, the melting of the GRB beneath long-lived magma channels (Peterson and Boerst, 2013) at the base of the Maturi deposit has contaminated the SKI and induced additional sulfide immiscibility and the formation of Ni- and Co-rich massive sulfide bodies (see the bottom of Fig. 5). REFERENCES Allen, D.J., Hinze, W.J., Dickas, A.B., and Mudrey, M.G., Jr., 1997, Integrated geophysical modeling of the North American Midcontinent rift system: New interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota, in Ojakangas, R.W., Dickas, A.B., and Green, J.C., eds.: Middle Proterozoic to Cambrian rifting, central North America: Geological Society of America, Special Paper 312, p. 47-72. Foose, M.P., and Cooper, R.W., 1978, Preliminary geologic report on the Harris Lake area, northeastern Minnesota: U.S. Geological Survey, Open-File Report 78-385, 24 p., 1 pl., scale 1:12,000. Miller, J.D., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.M., 2001, Geological map of the Duluth Complex and related rocks, Northeastern Minnesota; Minnesota Geological Survey, Miscellaneous Map M119, scale 1:200,000. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., 2002, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations RI-58, 207 p. Peterson, D.M. and Severson, M.J., 2002, Chapter 4, Archean and Paleoproterozoic rocks forming the footwall of the Duluth Complex, in Geology and mineral potential of the Duluth Complex and related intrusions of northeastern Minnesota, Minnesota Geological Survey, Report of Investigations 58, pp. 76-93. Peterson, D.M. and Boerst, K., 2013, Twin Metals Minnesota’s Maturi Deposit, in Severson, M.J., Peterson, D.M., Ware, A., and Boerst, K., 2013, Cu-Ni-PGE Deposits of the Duluth Complex, Geology and Development: Precambrian Research Center, Workshop on the Copper, Nickel, Platinum Group Element Deposits of the Lake Superior Region, October 6-13, 2013, Field Trip Guidebook, pp. 45- 57. Severson, M.J., 1994, Igneous stratigraphy of the South Kawishiwi intrusion, Duluth Complex, northeastern Minnesota: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report NRRI/TR 93/34, 210 p. (with plates). Severson, M.J., Peterson, D.M., Ware, A., and Boerst, K., 2013, Cu-Ni-PGE Deposits of the Duluth Complex, Geology and Development: Precambrian Research Center, Workshop on the Copper, Nickel, Platinum Group Element Deposits of the Lake Superior Region, October 6-13, 2013, Field Trip Guidebook, 60 p. 101 �FIELD TRIP 6 Saturday, May 17, 2014 THE ST. LOUIS SUBLOBE AND GLACIAL LAKE UPHAM LEADERS: Phil Larson (Duluth Metals Limited) Alan Knaeble (Minnesota Geological Survey) Howard Mooers (University of Minnesota Duluth) with contributions by Lisa Marlo (Halcon Resources Corporation) INTRODUCTION The flat plain lying to the south of the Mesabi Iron Range at first glance stands in stark contrast to the varied topography and geology of northeastern Minnesota. Largely covered by peatland, early travelers on the St. Louis and Savanna Rivers no doubt appreciated their quick passage through the mosquitoinfested terrain north to Lake Vermilion and west to Sandy Lake. After iron ore was discovered on the Mesabi, mine developers at first appreciated the gentle grades and straight lines of the railroads vital to transporting iron ore to market, but learned to respect the difficulties inherent in maintaining lines across water-logged ground prone to sink beneath the heavy traffic. Agricultural settlement was almost an afterthought, memorable more for its heroic efforts than lasting success. Even so, on a long drive south from the Range, a traveler cannot help to wonder, “Why so flat, and why here?” As it turns out, the Glacial Lake Upham Plain (as it is called) is the location of one of the more interesting episodes in the long history of glaciation in Minnesota and the Upper Midwest (Fig. 1). As the last Laurentide Ice Sheet retreated to the north, the long-standing pattern of ice flowing from Hudson Bay in the north to central Minnesota and beyond to the south was dramatically disrupted by a surge of ice from the Red River Valley to the west. As the ice entered the flat Upham Plain, flow took right-turns to either flank, advancing to the southwest as far as Aitkin, and more remarkably, advancing from south to north as far as the Giant’s Range. Despite this dramatic entry, almost as rapidly, it melted away and was gone. The glacial advance by the St. Louis sublobe was rapid, short-lived, and barely left a mark on the landscape. Nevetheless, the signs of this light touch are there for the careful observer to see. This field trip explores sites illustrating the landforms and sediments associated with the St. Louis sublobe and its associated glacial lakes. HISTORICAL BACKGROUND Leverett (1932) first noted the presence of a distinct clayey, calcareous grey till in the region south of the Mesabi Iron Range. He correlated the till with similar fine-textured, calcareous till associated with the Des Moines lobe to the west, and first applied the name St. Louis sublobe to the corresponding glacial advance. He also noted the presence of similar reddish clayey till in the immediate vicinity of the Mesabi, attributing the red color to local incorporation of red hematite iron-ore. He also mapped the extent of the lobe, as well as the associated overlying glacial lake basins, named Glacial Lakes Aitkin and Upham. Although Wright (1955) initially attributed the red color in the drift to reddish drift from a Superior lobe advance from the southeast, later work by Baker (1964) and Wright and Watts (1969) returned to Leverett’s conclusion that the St. Louis sublobe advanced from the northwest. Baker (1964) recognized that two tills were associated with the St. Louis sublobe advance (named the Alborn Phase by him): a reddish clayey till, and an overlying grey (brown when oxidized) silty till, named the Alborn till and Prairie Lake till, respectively. Wright and Watts (1969) attributed the red color of the Alborn till to 102 �incorporation and mixing of reddish lake clay from Glacial Lake Upham I into a glacial debris load composed of grey (-brown) silty till. The red lake clay was believed due to reddish sediment ultimately sourced from the northern margin of the Superior lobe. Farnham and others (1964) published two radiocarbon dates attributing relatively young dates to the Alborn phase. Wright and Watts (1969) recognized that meltwater from the St. Louis sublobe and its successor lakes flowed around the northern margin of the Thomson-Nickerson moraine during the Nickerson phase of the Superior lobe. Winter (1971) and Winter and others (1973) investigated St. Louis sublobe deposits along its northern extent in detail. They questioned that the grey(-brown) Prairie Lake till and reddish Alborn till could be sourced from the same provenance, noting the difference in texture and clast composition. Hobbs (1983) described the history of meltwater flow associated with the St. Louis sublobe and its successor lakes, including the flow of meltwater from Lake Koochiching (potentially a Glacial Lake Agassiz precursor) through the Upham and Aitkin basins. Ballantine (1991) conducted additional stratigraphic investigations in the southern Upham basin. Meyer (1993) defined the St. Louis sublobe as referring only to ice flowing east of the Giant’s Range, distinguishing it from the much more extensive northwestern provenance ice advance to the west. More recently, Knaeble and others (2004) and Knaeble and Hobbs (2009) conducted the first detailed modern investigations of the southern margin of the St. Louis sublobe in Crow Wing and Carlton Counties. Marlow (2004) investigated the widespread occurrence of eolian deposits on the beds of Glacial Lakes Aitkin and Upham II, and Jennings and Reynolds (2005) re-mapped glacial stratigraphy along the axis of the Mesabi Iron Range. GEOLOGY Regional Background Bedrock in the area glaciated by the St. Louis sublobe is underlain by shales, mudstone, and greywackes of the Paleoproterozoic Virginia Formation. This formation has proven relatively less resistant to weathering and erosion than the iron-formation and Archean granite-greenstone terrane to the north, the Mesoproterozoic Duluth Complex to the east, and the fold-and-thrust belt of the Paleoproterozoic Penokean orogen to the south. Consequently, a natural basin existed in the area prior to glaciation. The area was repeatedly glaciated over the course of the Pleistocene. Most recently, during the Wisconsinan glaciation, ice of the Rainy lobe advanced from the north-northeast from the Labradoran sector of the Laurentide ice sheet (bearing approximately 215°), carrying relatively coarse-grained, sandy textured sediment dominated by crystalline igneous and metamorphic rocks of the Canadian Shield. At about the same time, to the south, the Labradoran ice funneled into the Lake Superior basin as the Superior lobe advanced roughly parallel to the Rainy lobe. The Superior lobe carried an abundance of reddish sediment eroded from rift-filling sedimentary rocks of the Mesoproterozoic Midcontinent Rift. During the waning stages of the last glaciation, the Rainy lobe margin retreated back to the northnortheast across the area. Retreat was characterized by a relatively steady rate of margin retreat, punctuated by the occasional minor re-advance and moraine building event. When the line of retreat reached the region underlain by the Virginia Formation (Animikie Basin), proglacial lakes developed in the natural topographic low, Glacial Lake Aitkin to the southwest and Glacial Lake Upham to the northeast (Fig. 1). These lakes received meltwater-borne sediment from both the Rainy lobe to the northeast and the Superior lobe to the south. Glacial Lake Upham I Lake Upham I formed as a proglacial lake as the Rainy lobe ice margin retreated to the northeast. Little is known about the extent or duration of the lake, as its existence is largely inferred from the red clayey lacustrine sediment incorporated into the debris load of the later St. Louis sublobe advance. In situ Upham I sediment is rarely observed. 103 �The distinctive red color of Upham I sediment is due to the presence of hematite in the clay size fraction. Leverett (1932) believed this red color reflected erosion and incorporation of soft hematite iron ore from the Mesabi Iron Range to the north. Leith (1903) recognized that significant glacial erosion of soft iron ores had taken place, and suggested that much of this eroded material would be resident in finegrained glacial sediments. Wright (1955) correlated Upham I sediments with the Superior lobe, believing the red color came from incorporation of red shale from the Mesoproterozoic Midcontinent Rift System. Both are viable hypotheses, and it is not unlikely that both sources contributed to the red color of Upham I sediments. The lake basin was bounded by stagnant ice-cored topography to the north and west, and glacial ice of the Rainy lobe to the northeast and the Superior lobe to the south. Large patches of ice-cored topography likely persisted in the lake basin, particularly along northwest-southeast oriented moraines formed by the retreating Rainy lobe. The outlet to the lake was likely to the southwest through Lake Aitkin I and ultimately the Mississippi River. The age of Upham I is constrained by two bracketing events: retreat of the Rainy lobe from the St. Croix moraine in central Minnesota, tentatively dated at no later than ~15.1 14C kyr BP (Birks, 1976; Mooers and Lehr, 1997), and the Alborn phase advance of the St. Louis sublobe. The lake was likely contemporaneous with the Automba phase of the Superior lobe. St. Louis Sublobe Gradual retreat of the Rainy lobe and deglaciation of the Upham basin was abruptly interrupted by advance of the St. Louis sublobe from the northwest (Fig. 1). In contrast to the Rainy lobe, formed by ice flowing from the northeast and the Labradoran accumulation center of the Laurentide Ice Sheet, the St. Louis sublobe originated from ice flow southward from the Lake Winnipeg basin into the Red River Valley. For most of the glacial cycle, a relatively high flux of ice over the Canadian Shield of northwestern Ontario and northeastern Minnesota blocked eastward expansion of ice streaming south from Lake Winnipeg, funneling this ice into the Minnesota River Valley. Retreat of the Rainy lobe ice margin from the Itasca-St. Croix moraines opened a low elevation, ice-free corridor to the north of the Itasca moraine. Continued high ice flux in the Red River Valley rapidly surged into this gap, expanding south of the Mesabi Range to the northeast and southwest across the fine-grained lacustrine sediments of Lakes Aitkin I and Upham I. This advance and its associated glacial deposits are grouped as the Alborn phase. Figure 1. Maximum extent of St. Louis sublobe advances (white) and inferred location of Lakes Aitkin I and Upham I (stippled pattern). Arrows indicate inferred ice flow directions. Western limit of Alborn member/’red clayey’ till along the Mesabi Range (red line) from Winter and others 104 �(1973). Proglacial Lake North of Nashwauk is the high-level (elev. >1500’) proglacial lake dammed by the initial phase of the St. Louis sublobe. Chronology The absolute timing of the Alborn phase is poorly constrained. As mentioned above, the advance could not have occurred prior to ~15.1 14C kyr BP. Two radiocarbon dates have been attributed directly to St. Louis sublobe associated deposits (Farnham and others, 1964). A buried soil within Lake Aitkin II sediment near Aitkin, MN was dated at 11.6 14C kyr BP. Wood recovered from a red clayey till at the Mariska Mine in Gilbert, MN, near the northeastern limit of the Alborn phase, was dated at 11.2 14C kyr BP. However, the Lake Aitkin II soil was developed on lacustrine sediment, and therefore significantly post-dates the Alborn phase and formation of Lake Aitkin II. The Mariska Mine wood is unlikely to have been incorporated during advance of the St. Louis sublobe into a forested environment. Rather, it perhaps represents wood incorporated into a flow till developed on ice-cored topography, and also significantly post-dates the Alborn phase. These young dates therefore are minimum bracketing dates for the Alborn phase. Relative age relationships provide additional insight. During the Split Rock phase, meltwater from the Superior lobe apparently flowed northwest from the Cloquet moraine into the ice-free Upham I basin (Wright and Watts, 1969). In the later Nickerson phase, meltwater from the eastern margin of the St. Louis sublobe and from post-glacial Lake Upham II flowed down the proto-St. Louis River and around the Superior lobe margin (Thomson moraine) into the St. Croix River drainage. Significant meltwater flow down the St. Louis River persisted after retreat of the Superior lobe from the Thomson moraine and formation of Lake Duluth. However, meltwater inflow from Lake Upham II into Lake Duluth ceased prior to readvance of the Superior lobe to the western end of the Lake Duluth basin during the Marquette phase (Mooers and others, 2005), dated at 10.0 14C kyr BP (Lowell and others, 1999). Glacial Dynamics The St. Louis sublobe was a thin, temperate glacier. It was on the order of perhaps 100-200 m thick at its maximum extent (Knaeble and others, 2005). The apparent confinement of the lobe by topography during its advance is a direct consequence of this relative thinness; flow was apparent confined by stagnant ice in the Itasca and Outing moraines to the south and west, and by stagnant Rainy lobe ice to the north. Notwithstanding the relatively thin ice, expansion into the Lakes Aitkin I and Upham I basins was facilitated by the low shear strength lacustrine sediment substrate. The base of the glacier was apparently at the pressure melting point throughout the Alborn phase. Preexisting proglacial lakes precluded development of permafrost in the Aitkin I and Upham I basins, so much of the glacial advance was over a ‘warm’ bed. Little evidence for freeze-on and entrainment of sediment into glacial ice exists, either in the form of thick Alborn phase glacial sediment accumulations or stagnation topography. Alborn phase tills tend to be compositionally homogeneous, and relatively uniform in thickness. Local incorporation of basal sediment (lacustrine clays and silts) is evident in the basal Alborn member tills, however erosion and entrainment occurred at very low rates relative to the polythermal Rainy lobe just to the north. Smeared pods of Alborn till are often incorporated into the overlying Prairie Lake member till at the contact. However, the overall composition of Prairie Lake till is homogeneous and distinct from Alborn till, suggesting erosion, entrainment, and mixing of substrate into the debris load was occurred at very low rates. These features suggest sediment transport by the St. Louis sublobe is best explained as having occurred as a subglacial deforming layer. Other than Lake Upham I lacustrine sediments, Alborn phase deposits show little evidence of erosion, entrainment, or other modification of the older Rainy lobe glacial deposits over which the St. Louis sublobe advanced. Rainy lobe landforms such as eskers, drumlins, and moraines are clearly visible beneath a veneer of Alborn phase drift, and the hummocky ‘moraines’ found at the margins of the St. Louis sublobe are composed primarily of older, relatively coarse-grained Rainy and Superior lobe drift. 105 �The inability of the St. Louis sublobe to form its own landforms, or modify older landforms, is a consequence of a general lack of englacial sediment, and an inability to transmit shear stress into the substrate. The relative thinness of the ice and the low shear strength substrate suggest that the St. Louis sublobe advance was as one, or possibly two, surge(s) from the main Red River lobe ice mass. By analogy with modern ice streams in Antarctica and Greenland, ice may have streamed into the Upham and Aitkin basins at flow rates up to 10 km/yr, and may have taken only a few decades for the St. Louis sublobe to reach its maximum extent. It is unclear how many individual surges, or advances, from the main Red River lobe ice mass may have occurred over the history of the St. Louis sublobe. The two distinct tills associated with the Alborn phase likely record two distinct phases of ice flow through the St. Louis sublobe. The Goodland esker (Knaeble and others, 2005) formed in a large englacial meltwater conduit developed in response to the initial west-to-east St. Louis sublobe advance blocking southward meltwater flow from the Rainy lobe to the north. A reconstruction indicates around 75 m of ice was necessary to block this flow. However, the Goodland esker and the surrounding ice-cored Rainy lobe deposits are mantled by upper (later) Prairie Lake member tills, suggesting this area was later overrun by thicker ice in a later phase of the St. Louis sublobe advance. Similar to other surging glaciers, the lobe may have stagnated en masse, with an abrupt retreat of the ‘active’ ice margin by as much as 150 km. Ice down-flow of the ‘active’ margin ceased flowing and melted without forming recessional moraines. The lack of significant englacial debris transport precluded formation of significant supraglacial sediment accumulation, even in marginal areas. Consequently, during stagnation and wastage of the glacier an insulating debris layer did not form to retard melting. Perhaps 5 m of surface melting of the clean ice may have occurred each season, meaning even a 200 m thick glacier would have persisted only 40 years after cessation of ice flow. Alborn Phase Deposits Glacial deposits associated with the Alborn phase are placed in the Aitkin formation lithostratigraphic unit (Johnson and others, in press). Two members have been formally defined: the Alborn member, and the Prairie Lake member (Baker, 1964). The Alborn member is the lower member, and is interpreted to have been derived to a great degree from erosion of underlying Lake Upham I lacustrine sediment. The overlying Prairie Lake member contains a significant component of Paleozoic carbonate and Cretaceous shale lithologies, derived from the Red River Valley and Winnipeg basin. The distribution of the two members is poorly understood within and adjacent to the Upham basin, however the Prairie Lake member is less extensive than the underlying Alborn member. Both members are recognized primarily as till lithofacies, however minor glaciofluvial and glaciolacustrine elements locally occur. Alborn Member The Alborn member consists predominantly of reddish-brown to dark reddish-grey, clay to clay loam till (Fig. 2a). The pebble, cobble, and boulder clast content is distinctly lower than underlying Rainy lobe deposits (Knaeble and Hobbs, 2009) (Fig. 2b). Alborn member tills along the northern margin of the Upham basin, adjacent to the Mesabi Range, are distinctly more clay rich than tills along the southern margin of the Upham basin (Winter and others, 1973) (Fig. 2a). This may reflect incorporation of a greater amount of clayey lacustrine sediment by the glacier along a longer flow path across the bed of Upham I. Alborn till is patchily distributed near the margins of the St. Louis sublobe. The once continuous till sheet deposited on ice-cored topography has been disrupted by subsequent meltout and collapse of the underlying substrate. 106 �Figure 2. Matrix texture (a. left triangle) and lithologic composition (b. right triangle) of the 1-2 mm sand fraction of Alborn member tills. Circles are from QDI (Quaternary Data Index) database (Minnesota Geological Survey), triangles are ‘red clayey till’ of Winter and others (1973), and large circle is mean in Carlton County (Knaeble and Hobbs, 2009). The ‘red clayey till’ from the northern end of the Upham basin is distinctly more clay-rich than Alborn member tills from the southern end of the basin. The tills contain only a minor amount of carbonate, and no gray shale, reflecting the composition of recycled underlying northeastern provenance (Rainy lobe) drift. Figure 3. Matrix texture (a. left triangle)and lithologic composition (b. right triangle) of the 1-2 mm sand fraction of Prairie Lake member tills. Circles are from QDI (Quaternary Data Index) database (Minnesota Geological Survey), triangles are ‘brown silty till’ of Winter and others (1973), large circle is mean in Carlton County (Knaeble and Hobbs, 2009). The tills are variably enriched in carbonate and gray shale, reflecting a northwest (Red River Valley) provenance. 107 �Prairie Lake Member The Prairie Lake member consists predominantly of yellow-brown to brown to dark grey, loam to clay loam till (Fig. 3a). The pebble, cobble, and boulder clast content is similar to Alborn member till, and likewise distinctly lower than underlying Rainy lobe deposits. Unleached (grey) Prairie Lake tills average ~10% carbonate and ~17% grey shale clasts in the very coarse sand fraction, indicative of a northwestern provenance (Fig. 3b). Carbonate leaching averages about 1 m depth (Knaeble and Hobbs, 2009). In contrast to Alborn till, Prairie Lake till shows no apparent systematic textural variation across the Upham basin. Glacial Lake Upham II Advance and stagnation of the St. Louis sublobe was followed by a rapid melting of the thin, warm ice, as discussed above. Melting of the ice was accompanied by formation of a series of proglacial lakes, which ultimately coalesced to form Lakes Aitkin II and Upham II (Fig. 4). Initially, Aitkin II and Upham II were contiguous. Opening of successively lower elevation outlets led to drawdown and drainage of both lakes into the proto-St. Louis River. Continued isostatic rebound raised the sill between Aitkin II and Upham II, leading to re-inundation of Aitkin II. Aitkin II ultimately drained with the opening of an outlet into the Mississippi River on the southwestern end of the basin. Figure 4. Location of major meltwater inflow and discharge points for Lakes Aitkin II and Upham II, St. Louis sublobe areal footprint (olive), and active Rainy and Superior lobe ice (white) during main phase of Upham II (>ca. 11.6 kyr BP). As much as 67 m (220’) of differential isostatic rebound occurred over the 160 km extending from the southwestern extent of Aitkin II to the northeastern extent of Upham II, based on strandline correlations (Marlow, 2004). Adjusting for isostatic rebound, the absolute elevation difference between the uppermost Upham II outlet and the final outlet is about 38 m (125’), a number also corresponding to the maximum depth of Upham II. 108 �Meltwater Drainage In addition to their immediate catchments, Aitkin II and Upham II received meltwater from up to 500 km of the margin of the Laurentide Ice Sheet (Fig. 5). Advance of the St. Louis sublobe into the area recently vacated by the retreating Rainy lobe blocked the flow of meltwater south from the Rainy lobe ice margin. Meltwater apparently pooled in the interlobate area forming proglacial lakes, however discharge necessitated flow either through subaerial channels around the northeastern margin of the glacier, or through sub-, en-, or supra-glacial meltwater conduits in the glacier itself. Morphology of the Goodland esker suggests that a significant amount of meltwater flowed in a supraglacial channel southward across the surface of the glacier, at least during the earlier phase of the advance. Later meltwater flow may have been accommodated at least in part by subaerial channels along the northeastern margin of the glacier. Following stagnation and melting of the glacier, and formation of Lakes Aitkin II and Upham II, meltwater entered the Aitkin II and Upham II basins at three major inlets. Retreat of the Rainy lobe north of the Laurentian divide led to formation of proglacial Lake Norwood (Winchell, 1901). Discharge from Norwood entered northeastern portion of the Upham II basin through the Embarrass Gap (Fig. 5). Norwood’s initial outlet had an elevation of 1490’, with successive outlets developed at 451 m (1480’) and 450 m (1476’) elevation (Lehr and Hobbs, 1992). A lower outlet at 443 m (1454’) elevation is correlated with the Mizpah phase of Lake Koochiching (Hobbs, 1983; Lehr and Hobbs, 1992). The final outlet through the Gap has an elevation of around 428 m (1405’), however meltwater drainage ceased when the sill between the Pike and Embarrass Rivers (435 m (1427’) elevation) was exposed. Meltwater also flowed down the proto-Prairie River southward into Lake Aitkin II. This outlet has been nominated as a potential discharge point for the lower elevation Gemmell phase of Lake Koochiching, and even as a potential outlet for Lake Climax, an early, high level of Lake Agassiz (Hobbs, 1983). However, the sill over the Laurentian divide into the Prairie River drainage stands at 418 m (1373’), approximately the same as the lower Koochiching outlets in the Embarrass Gap (corrected for rebound), and channel morphology in the vicinity of the sill shows no evidence of high meltwater discharge. The lower reach of the Prairie River does show evidence for significant meltwater discharge. This was likely limited to a short interval after stagnation and melting of the St. Louis sublobe, and before retreat of the Rainy lobe ice margin from the Laurentian divide and expansion of Lake Norwood. West of Grand Rapids, MN, Lake Sucre formed in the area covered by the ‘upstream’ expanse of the stagnant St. Louis sublobe (Larson and others, 2004). A considerable amount of meltwater from the Red River lobe flowed into the western end of this lake, eventually reaching the western end of Aitkin II. At the time of maximum meltwater discharge through Upham II, flow was ultimately channeled around the northern margin of the Nickerson phase Superior lobe into the Kettle River meltwater channel (Wright and Watts, 1969). Assuming an average melt rate of 2.5 m/yr over an ablation zone extending 200 km from the ice margin, average annual discharge through this channel was roughly 8000 m3/s. Discharge was seasonably variable, so maximum discharge rates were considerably higher. Carney (1996) evaluated peak discharge through this channel examining a variety of parameters, concluding that maximum peak discharges of 12000 to 17000 m3/s were reasonable. This number suggests that meltwater discharge through Upham II and the Kettle River channel solely reflected ablation from the ice margins immediately adjacent to the lakes, and does not contain a Lake Climax/Agassiz discharge component. Significant meltwater discharge through Lakes Aitkin II and Upham II was ultimately restricted to the period between stagnation and melting of the St. Louis sublobe and opening of the Macintosh channel into Lake Climax in the Red River Valley, at the end of the Cass phase (c. 11.6 kyr BP) (Fenton and others, 1983). High Level Outlets A series of outlets to high level proto-Upham II proglacial lakes are present along the eastern margin of the Upham II basin, including include the Chicken Creek (three outlets; 1500’ to 1465’) and Us-Kab- 109 �Wan-Ka River (428 m to 422 m; 1404’ to 1386’) outlets. Initially, these channels drained the eastern margin of the St. Louis sublobe, and a portion of the southern margin of the Rainy lobe. Formation of Lake Norwood may have occurred about the time the Us-Kab-Wan-Ka River outlets formed. As Norwood expanded to the west, an increasing discharge of meltwater was channeled along the eastern margin of the St. Louis sublobe. Meltwater flowed over an unstable, ice-cored landscape, resulting in frequent shifts in outlet location, and drops in outlet elevation. The Us-Kab-Wan-Ka outlets were succeeded by a series of outlets formed in the vicinity of Hellwig Creek ranging from 422 m (1386’) to 418 m (1370’) elevation. Main Lake Stage The highest well-developed strandlines in the Upham II basin correlate to a series of outlets between 413 m (1354’) and 407 m (1336’) elevation at Hellwig Creek. The Hellwig Creek outlets were abandoned as ice-cored topography in the Culver moraine continued to melt and collapse, in favor of lower level outlets in the Artichoke River 404 m (1324’) and Spider Creek 396 m (1300’) channels. The highest strandlines in Upham II are commonly obscured by collapsed topography, indicating these shorelines formed on topography still underlain by stagnant Rainy lobe ice. The final outlet to Upham II was established with opening of an outlet down the proto-St. Louis River channel. Terraces correlating with the uppermost outlet had an elevation of about 392 m (1285’). The broad, wide floodplain and terraces corresponding to the upper St. Louis River outlet indicates a significant amount of meltwater continued to flow through Aitkin II and Upham II at the time this channel was established. At least four subsequent stable lower outlets are present, at 390 m (1278’), 386 m (1266’), 381 m (1250’), and 377 m (1238’) elevation. In contrast to the higher elevation outlets abruptly abandoned with opening of lower elevation outlets, the St. Louis River outlet records stepwise drops within a single channel to an ultimate elevation of 375 m (1230’). Figure 5. Location of outlets for early ice marginal drainage channels (Chicken, Us-Kab-Wan-Ka, Hellwig) and main stage Lake Upham II outlet channels (Hellwig, Artichoke, Spider, and St. Louis). 110 �Eolian Activity in the Upham II basin Upham II was apparently stable at the higher Hellwig Creek and Artichoke River outlet levels for a considerable length of time, long enough to allow deposition of well-sorted lacustrine sediments in the lake, ranging from very fine to fine sand in the nearshore areas to clay in the deepest portion of the basin. Later lowering of lake level and exposure of these nearshore sands triggered eolian activity, leading to extensive dune field development (Marlow, 2004). Final Drainage of Lake Upham II Separation of Upham II from Aitkin II occurred when the lake level in Upham II dropped below the elevation of the Swan River sill (382 m; 1252’). Outflow occurred through a broad, shallow channel, suggesting meltwater inflow into the Aitkin II basin had ceased. Downcutting of the St. Louis River outlet may have been triggered by cessation of meltwater influx into the lakes. Discharge over the Swan River sill eroded dune fields developed on Upham II’s bed, indicating significant eolian activity does not post-date final cessation of drainage from Aitkin II into Upham II. This relationship further indicates that significant eolian activity in the Upham II basin was largely confined to the interval between opening of the lower Hellwig Creek outlets and the 381 m (1250’) elevation St. Louis River outlet. Final drainage of Upham II occurred as the St. Louis River down-cut to its modern level of about 375 m (1230’) at the final outlet to Upham II. Final Drainage of Lake Aitkin II Lake Aitkin II substantially drained as the outlet to Upham II dropped to below 382 m (1252’); this drawdown in lake level may have triggered development of the peat dated by Farnham and others (1964) (11.6 14C kyr BP). Continued differential rebound between the Swan River sill on the northeastern side, and the southwestern side of the basin led to re-inundation of the lake (and deposition of the upper lacustrine sequence reported by Farnham and others (1964)). Hobbs (1983) reported a radiocarbon date of 9.1 14C kyr BP from a snail shell recovered from a marl deposited in Aitkin II; this date indicates Aitkin II persisted for at least 2500 years after separation from Upham II. Aitkin II ultimately drained with opening of a new outlet into the Mississippi River in the southwest corner of the basin at about 366 m (1200’) elevation. 111 �DESCRIPTION OF FIELD TRIP STOPS Figure 6. Location of field trip stops relative to major features of the Upham basin. 1. Glacial Lake Upham II Beach 499660E/5236650N (UTM Zone 15, NAD83 datum) Silica 7.5’ USGS Quadrangle NENE, Section 4, T55N, R21W This site is located on the uppermost relatively well-developed beach associated with Lake Upham II. It likely formed in response to establishment of a relatively stable outlet in the vicinity of Hellwig Creek, on the opposite time of the basin. Subsequent to the time of upper beach formation, Lake Upham II experienced a series of relatively gradual drops in water level at this site. Downward stepping clinoforms visible in a ground penetrating radar profile are interpreted to represent a forced regressive shoreface (Fig. 6). Progradation of the bedforms occurs as a result of shoreline regression, and indicates a constructional shoreline. At this site, regression was triggered by a combination of gradual relative lake level drop due to differential isostatic rebound relative to the more southerly outlets, and relatively abrupt lake level drops triggered by development of new, lower elevation outlets. 112 �Figure 6. West to east ground penetrating radar profile of Site 1. Profile length 120 m, vertical scale ~9 m. From Knaeble and others (2005). 2. Toivola Esker 514200E/5226890N (UTM Zone 15, NAD83 datum) Toivola 7.5’ USGS Quadrangle NENE, Section 1, T54N, R20W Most of the landforms within the Glacial Lake Upham basin predate development of the lakes. This exposure is an example of an esker deposited during retreat of the Rainy lobe that was later overrun by the St. Louis sublobe and subsequently modified by wave action. This esker and others like it became wave-washed "islands" once Glacial Lakes Aitkin and Upham I and II formed. Exposed at the base of the sequence are coarse-grained gravels containing northeast provenance clasts, including granites, ironformation and locally derived shale and greywacke of the Paleoproterozoic Virginia Formation. This esker segment, and numerous similar examples in the Upham basin, was deposited by a beaded esker system during retreat of the Rainy lobe. Overlying the gravels is a yellow-brown fine-grained till (Prairie Lake Member). Overlying the till is a sequence of nearshore sands and gravels. These presumably eroded from that portion of the esker rising above the level of glacial Lake Upham II; the strandline formed at about 397 meters (1,300 feet) elevation. The uppermost portion of the sequence is a blanket of eolian sand and silt derived from the surrounding lake plain to the upland after final drainage of Lake Upham II. 3. Prairie Lake Member Till 507820E/5182550N (UTM Zone 15, NAD83 datum) Prairie Lake 7.5’ USGS Quadrangle SWSE, Section 20, T50N, R20W There are two distinct tills at this road cut on the east side of State Highway 73 on the northeast side of Prairie Lake (Figure 1). Both units are deposits of the St. Louis sublobe. The elevation at the top of the exposure is about 1330 and the upper till, the Prairie Lake Member (Baker, 1964) of the Aitkin Formation (Johnson and others, in press), 113 �is approximately 15 feet thick, yellow-brown (2.5Y5/4) to brown (10YR5/4), calcareous, and loam textured. Three samples have textures averaging 37-35-28 (sand-silt-clay percentages, respectively) and lithologic percentages of the 1-2 mm coarse-grained sand fraction, averaging 46-12-42 (crystallinecarbonate-gray shale, respectively). There are trace amounts (<1%) of red Superior-source sand grains. Five feet of the lower till, the Alborn Member (Baker, 1964) of the Aitkin Formation, is exposed at the base of the ditch. Above the contact between the two tills there are, in places, streaks of red-brown till incorporated into the base of the yellow-brown till. An auger boring in the ditch at the base of the outcrop penetrated another 30 feet. The upper 29 feet detected calcareous red-brown clay loam till. Six samples of this till have textures averaging 23-37-40 (sand-silt-clay percentages, respectively) with coarse-grained sand fraction amounts averaging 3% carbonate, no gray shale, and 11% red Superior-source. There are some, but not many pebbles in the till, which tends to become finer with depth, possibly due to incorporation of underlying lake sediment. The last foot was gray clayey silty lake sediment (possibly Glacial Lake Aitkin I). Previous interpretations suggest that the tills of these two members were deposited by one ice advance (Baker, 1964; Wright, 1972). The Prairie Lake till represents the original yellow-brown and brown characteristics of the sediment in the ice as it advanced into glacial Lakes Aitkin I and Upham I, and the Alborn till depicts the incorporation and mixing of red lake sediments of glacial Lakes Aitkin I and Upham I into the basal portion of the ice as the glacier advanced across the basin. This produced ice deposits with brownish sediment overlying and/or intermixed with red sediments. In contrast, subsequent interpretations suggest that each member was a separate ice advance (Knaeble and Hobbs, 2009). 4. Alborn Member Till 521450E/5190450N (UTM Zone 15, NAD83 datum) Martin Lake 7.5’ USGS Quadrangle NWNW, Section 35, T51N, R19W At this private pit located just south of St. Louis CR 856 there are three tills exposed along the west wall (Figure 2). The elevation at the top of the exposure is about 1350. The upper 8 feet is composed of redbrown (5YR4/3 to 7.5YR4/3), non-calcareous, clay loam Alborn Member till with some pebbles. A sample at 6 foot has a textural analysis result of 22-44-34 (sand-silt-clay percentage, respectively), with no carbonate (leached) or gray shale, and ~10% red Superiorsource in the coarse-grained sand fraction. Below a sharp contact there is 3 feet of red-brown (5YR4/3) sandy loam Cromwell Formation (Wright, 1972; Johnson and others, in press) till of the Automba phase of the Superior lobe. Textural and lithologic analysis results for two samples of this till average 39-47-14 (sand-silt-clay percentages, respectively) with 1% carbonate (one sample was leached), no gray shale, and 21% red Superior-source in the 1-2 mm coarse-grained sand fraction. Below another sharp contact is 2 feet of brown (10YR6/3) cobbly, sandy Independence Formation (Johnson and others, in press) till, a deposit of the Rainy lobe. Two samples of this till averaged 53-40-7 (sand-silt-clay percentage, respectively) with trace amounts of carbonate, no gray shale, and 17% red Superior source in the coarse-grained sand fraction. The pebble concentration basically doubled in each underlying till unit. Underlying the 3 till units at the base of the exposure there is pebbly, cobbly sand and gravel. 114 �This site is about a mile or two north of the southern extent of St. Louis sublobe ice deposits. Here Alborn Member till thinly covers older Superior and Rainy lobe deposits. This same stratigraphic sequence is evident in other pit exposures as far east as Brookston. The Toimi drumlins (Wright and Ruhe, 1965) east of Brookston and the St. Louis River are surface exposures of the Rainy lobe deposits that at this site were covered, first by Automba phase deposits of the Superior lobe and later by the Alborn Member deposits of the St. Louis sublobe. 5. St. Louis River Outlet Channel 530800E/5191070N (UTM Zone 15, NAD83 datum) Brookston 7.5’ USGS Quadrangle SWSW, Section 26, T51N, R18W This stop is located at the intersection of the Artichoke and St. Louis Rivers, where there is a prominent terrace at 375 meters (1,230’) elevation that extends 1 kilometer (0.6 mile) across. The terrace is predominantly composed of a thick deposit of relatively coarse-grained gravel, ultimately derived from erosion of Rainy lobe drift in the Culver moraine. The terrace gravels are overlain and partially infilled by loess. The loess likely originated from lacustrine sediment from the bed of the Lake Upham II, eroded by wind as the littoral zone was episodically exposed by rapid drawdown associated with establishment of new, lower elevation outlets. 6. Spider Creek Outlet Channel 531600E/5201600N (UTM Zone 15, NAD83 datum) Alborn 7.5’ USGS Quadrangle NWNE, Section 26, T52N, R18W Spider Creek occupies one of a series of around 10 successive outlets that drained Lakes Upham II, and by extension Aitkin II. The broad (800 m wide), flat-bottomed channel formed when collapse of underlying ice-cored Rainy lobe drift opened an outlet some 24’ in elevation below the Artichoke River outlet. The channel morphology indicates a significant amount of meltwater was still discharging out of Upham II. Baker (1965) reported a bulk radiocarbon date of 13,000 ± 400 14C yr bp from a sequence of lacustrine marl (sample W-1234) within the Spider Creek outlet. The marl must post-date the cessation of drainage through the channel because marl formation requires shallow, still water. Baker (1965) expressed concern that this date was too old due to possible contamination by lignite. However, this date is consistent with the other evidence presented in this field trip description. The Spider Creek date places the minimum age of Glacial Lakes Aitkin and Upham II, and therefore the maximum limit of the St. Louis sublobe, prior to 13.0 14C kyr B.P. 115 �7. Birch Esker 530580E/5208810N (UTM Zone 15, NAD83 datum) Payne 7.5’ USGS Quadrangle NESE, Section 34, T53N, R18W There are multiple exposures in this large county pit (Figure 3). The eastern most exposure is a 25 foot cut adjacent to the railroad crossing revealing 3 separate tills. At a surface elevation of approximately 1350 the soil has been stripped from a 1 foot thick layer of leached mixed till and sand lenses. Underlying the leached layer is 5 feet of yellow-brown (2.5Y6/3 to 10YR6/4), calcareous, silt loam textured Prairie Lake Member till. Near the base of the unit there are shear bands (streaks, pods, and lenses) of incorporated material from the underlying red-brown till. A sample at a depth of 5 feet had a texture of 28-56-16 (sand-silt-clay percentages, respectively), and 8% carbonate, 2 % gray shale, and 1% red Superior-source in the 1-2 mm coarse-grained sand fraction. The underlying red-brown (7.5YR5/4 to 5YR4/4), slightly calcareous, loam textured Alborn Member till is about 3 feet thick with more pebbles and cobbles than the overlying till. There is a cobble-boulder stone line or lag at the base of the unit. A sample at the depth of 8 feet had a texture of 40-40-20 (sand-silt-clay percentages, respectively), and 5% carbonate, no shale, and 9% red Superior-source in the 1-2 mm coarse-grained sand fraction. The lowest unit is about 9 feet thick and exposes brown (10YR6/3) to gray-brown (10YR6/2) non-calcareous, sandy Independence Formation till with abundant pebbles and cobbles. A sample at the depth of 15 feet had a texture of 55-37-8 (sand-siltclay percentages, respectively), and 1% carbonate, no gray shale, and 15% red Superior-source in the 1-2 mm coarse-grained sand fraction. There is about 10 feet of slump to the pit floor below these units. At this site deposits of both members of the St. Louis sublobe are present in typical stratigraphic position. At Stop 2 the Alborn Member was above both the Cromwell Formation Automba phase till and the Independence Formation till. Here at Stop 3 we are further northeast beyond the depositional extent of the Automba phase deposits and therefore have only the Independence Formation till at the base. 8. Alborn Member Till on Collapsed Rainy Lobe Ice-cored Topography 520160E/5251950N (UTM Zone 15, NAD83 datum) Kirk 7.5’ USGS Quadrangle NESE, Section 15, T57N, R19W The Aitkin and Upham basins were occupied by glacial lakes on two separate occasions during the Late Wisconsin glaciation. Retreat of the Rainy lobe from the Mille Lacs and Outing moraines (Mooers 1988) led to the formation of Lakes Aitkin I and Upham I. The extent and timing of these lakes is poorly understood, as their presence is largely inferred from incorporation of lacustrine sediment into the overlying St. Louis sublobe till. Exposed at the base of the sequence are steeply south-dipping foreset beds of a subaqueously deposited fan. These sediments are Rainy lobe provenance, deposited along the southern margin of stagnant ice lying on the southern slope of the Giant’s Range, an area now characterized by collapsed ice-cored topography. The upper portion of the sequence is a St. Louis sublobe till. Between the fan sediments and till are a number of elongate slabs of fine-grained lacustrine sediment derived from Glacial Lake Upham I. This lacustrine sediment was eroded from deeper water and thrust onto the fan during the advance of the St. Louis sublobe (Figs. 11 and 12). The relationships visible in this exposure indicate that the St. Louis sublobe advance occurred while a substantial amount of debrismantled, stagnant Rainy lobe ice was still present south of the Giant’s Range. 116 �REFERENCES Baker, R.G., 1964, Late-Wisconsin glacial geology and vegetation history of the Alborn area, St. Louis County, Minnesota: University of Minnesota Master’s Thesis, 44 pp. Baker, R.G., 1965, Late-glacial pollen and plant macrofossils from Spider Creek, So. St. Louis County, MN: Geological Society of America Bulletin v. 45, p. 645-665. Ballantine, J.W., 1991, Late-Wisconsin Stratigraphy and Glacial History of Southwestern St. Louis County, Minnesota: Unpublished Master’s Thesis, University of Minnesota Duluth, 154 pp. Birks, H.J.B., 1976, Late Wisconsinan vegetational history at Wolf Creek, central Minnesota: Ecological Monographs v. 46, p. 395-429. Carney, S.J., 1996, Paleohydrology of the Western Outlets of Glacial Lake Duluth: Unpublished Master’s Thesis, University of Minnesota Duluth, 129 pp. Farnham, R.S., McAndrews, J.H., and Wright, H.E., 1964, A Late-Wisconsin Buried Soil Near Aitkin, Minnesota, and its Paleobotanical Setting: American Journal of Science, v. 262, p. 393–412. Fenton, M.M., Moran, S.R., Teller, J.T., Clayton, L., 1983. Quaternary stratigraphy and history in the southern part of the Lake Agassiz Basin. In Teller, J.T., Clayton, L., eds., Glacial Lake Agassiz, Geological Association of Canada Special Paper v. 26, p. 49-74. Hobbs, H.C., 1983, Drainage relationships of Glacial Lakes Aitkin and Upham and early Lake Agassiz in northeastern Minnesota. In Teller, J.T. and Clayton, L., eds., Glacial Lake Agassiz. Geological Association of Canada Special Paper v. 26, 245-259. Jennings, C.E., and Reynolds, W.K., 2005, Surficial Geology of the Mesabi Iron Range, Minnesota: Minnesota Geological Survey Miscellaneous Map Series, v. 164. Johnson, M.D., Adams, R.S., Gowan, A.S., Harris, K.L., Hobbs, H.C., Jennings, C.E., Knaeble, A.R., Lusardi, B.A., and Meyer, G.N., in press, Quaternary lithostratigraphic units of Minnesota: Minnesota Geological Survey Report of Investigations RI-68. Knaeble, A.R., and Hobbs, H.C., 2009, Surficial geology, pl. 3 of Boerboom, T.J., project manager, Geologic atlas of Carlton County, Minnesota: Minnesota Geological Survey County Atlas, C-19, pt. A, 6 pls., scale 1:100,000. Knaeble, A.R., Meyer, G.N., and Hobbs, H.C., 2004, Surficial geology, pl. 3 of Setterholm, D.R., project manager, Geologic atlas of Crow Wing County, Minnesota: Minnesota Geological Survey County Atlas C-16, pt. A, 6 pls., scale 1:100,000. Knaeble, A.R., Meyer, G.N., Marlow, L.M., Larson, P.C., and Mooers, H.D., 2005, Deposits and Landforms in the Region Glaciated by the St. Louis Sublobe, in Robinson, L. ed., Field Trip Guidebook for Selected Geology in Minnesota and Wisconsin, Minnesota Geological Survey, Minneapolis, p. 40–79. Larson, P.C., Mooers, H.D., and Marlow, L.M., 2004, Early advance of the St. Louis sublobe: A revised chronology of the deglaciation of northeastern Minnesota [abstract]; Institute on Lake Superior Geology Proceedings, 50th Annual Meeting, Duluth, MN, v. 50, part 1, p.100-101. Lehr, J.D., and Hobbs, H., 1992, Field Trip Guidebook for the Glacial Geology of the Laurentian Divide Area, St. Louis and Lake Counties, Minnesota: Minnesota Geological Survey Guidebook Series 18. Leith, C.K., 1903, The Mesabi Iron-Bearing District of Minnesota: USGS Monograph, v. 43, pp. 345. Leverett, F., 1932, Quaternary Geology of Minnesota and Parts of Adjacent States: USGS Professional Paper, v. 161, 149 pp. Lowell, T. V., Larson, G.J., Hughes, J.D., and Denton, G.H., 1999, Age verification of the Lake Gribben forest bed and the Younger Dryas Advance of the Laurentide Ice Sheet: Canadian Journal of Earth Sciences, v. 36, p. 383– 393. Marlow, L.M., 2004, Late Glacial and Early Holocene history of the Glacial Lakes Aitkin and Upham basin, NorthCentral Minnesota: Implications for the timing of post-glacial eolian activity. Unpublished Master’s Thesis, University of Minnesota Duluth, 82 pp. Meyer, G.N., 1993, Surficial geologic map of parts of Koochinging, Itasca, and Beltrami Counties, north-central Minnesota: Minnesota Geological Survey Miscellaneous Map M-76, scale 1:250,000. Mooers, H.D., and Lehr, J.D., 1997, Terrestrial record of Laurentide ice sheet reorganization during Heinrich events: Geology v. 25, p. 987-990. Mooers, H.D., 1988, Quaternary history and ice dynamics of the late Wisconsin Rainy and Superior lobes, central Minnesota. Unpublished Doctoral Thesis, University of Minnesota, 200 pp. Winchell, N.W., 1901, Glacial Lakes of Minnesota. Geological Society of America Bulletin v. 12, p. 109-128. Winter, T.C., 1971, Sequence of Glaciation in the Mesabi-Vermilion Iron Range Area, Northeastern Minnesota: USGS Professional Paper, v. 750-C, p. C82–C88. 117 �Winter, T.C., Cotter, R.D., and Young, H.L., 1973, Petrography and Stratigraphy of Glacial Drift, Iron Range Area, Northeastern Minnesota: USGS Bulletin, v. 1331-C, p. 50. Wright, H.E., Jr., 1955, Valders drift in Minnesota: Journal of Geology, v. 63, p. 403-411. Wright, H.E., and Watts, W.A., 1969, Glacial and Vegetational History of Northeastern Minnesota: Minnesota Geological Survey Special Publication, v. 11, 59 pp. 118 ��FIELD TRIP 7 Saturday, May 17, 2014 GEOLOGY AND GOLD MINERALIZATION OF THE VIRGINIA HORN AREA LEADERS: Mark Jirsa (Minnesota Geological Survey), William Rowell and Richard Sandri (Vermillion Gold LLC), and Jason Richter (Minnesota Department of Transportation) INTRODUCTION The local term “Virginia horn” applies to an area near the town of Virginia, where the generally easttrending, Paleoproterozoic, Biwabik Iron Formation makes an abrupt bend to the southwest, creating a marked anomaly in the map pattern (Fig. 1). The iron-formation unconformably overlies well-exposed Neoarchean bedrock within an uplifted, wedge-shaped block. This trip visits exposures that provide an overview of the Archean metavolcanic, meta-igneous, and metasedimentary rocks—including a Timiskaming-type successor-basin sequence, and Paleoproterozoic iron-formation and associated strata. Archean quartz-feldspar porphyry intrusions were the locus of deformation, alteration, and associated gold mineralization. The area has a long history of intermittent gold prospecting, and the potential for mineable quantities is currently under investigation by Vermillion Gold, LLC. Channel sampling and drill core from that exploration work will be displayed. Some of the most recently acquired core in the area is a product of highway relocation work underway to accommodate proposed new iron mining. The following field guide is modified from Jirsa and Green (2011). The stops are ordered for expeditious travel, rather than stratigraphy or geochronology. All UTM coordinates are given using NAD 83, Zone 15N. It is likely that more stops are described here than can reasonably be covered in a single day. Stops missed during this excursion can be visited by individuals using the guide, with the caveat that some stops may require permission from land owners. Figure 1. Generalized geologic map of northeastern Minnesota showing location of the Virginia Horn area (black outline). 119 �Geology of Archean Rocks (STOPS 1-5, 9, 10) The Archean rocks in the Virginia horn area (Fig. 2) are part of the Wawa subprovince of Superior Province, and are similar in most respects to other greenstone-granite terranes of the subprovince. The supracrustal rocks in the horn are separated from the well-known Vermilion district to the north by the Giants Range Batholith—a large, composite body consisting of several intrusive generations and compositions. The mafic volcanic and hypabyssal intrusive components in the Virginia horn may be equivalent to the Ely Greenstone in the Vermilion district, which has a 207Pb/206Pb zircon age of associated felsic strata of 2722.6±0.9 Ma (Peterson and others, 2001). The supracrustal rocks in the horn are subdivided into northern and southern panels on the basis of metamorphic grade and deformation style. The northern panel, immediately south of the Giants Range batholith, contains intensely lineated, amphibolite-grade schist having volcanic, intrusive, and clastic protoliths (Minntac sequence). The southern panel contains lithologically and stratigraphically similar rocks that were metamorphosed to much lower grades, ranging from prehnite-pumpellyite to low greenschist (Mud Lake sequence). The two panels are separated by the east-trending, post-metamorphic, Laurentian fault. The two sequences are stratigraphically, lithologically, and geochemically identical; suggesting that they may represent different crustal exposure levels of the same stratigraphic package. The metamorphic cleavage-forming event in both panels was the second (D2) of three regional scale deformation events—no metamorphic effects are recognized from the other two deformations. The first (D1) involved upright folding, soft-sediment deformation, and complex faulting. Strata of the southern panel form the broad, southwest-plunging, Mud Lake syncline (Fig. 2.B) and many smaller sympathetic folds—all inferred to be D1 structures. The syncline is cored by graywacke, slate, and minor felsic tuff, and has outer limbs of calc-alkalic and tholeiitic volcanic strata. The Mud Lake strata and D1 structures (F1 fold axes) developed in it were cut by felsic quartz- and feldspar-phyric (QFP) intrusions. However, the QFP intrusions are lithologically similar to some layers within graywacke, suggesting the possibility of temporal overlap between dacitic magmatism and sedimentation. Strata in the Mud Lake syncline, and the quartzofeldspathic dikes that intrude them, are bisected by a fault- and unconformitybounded, alluvial fan-fluvial-volcanic succession known as the Midway sequence (see discussion below). All three sequences described above were metamorphosed and deformed during D 2, which has been bracketed locally between about 2674 Ma and 2682 Ma (Boerboom and Zartman, 1993). The third deformation event (D3) produced localized semi-brittle crenulation of D1 and D2 structures, and selective reactivation of earlier-formed faults. Based on seismic and geochronologic work in adjacent Ontario (e.g., Percival and Helmstaedt, 2006), D1 may be equated with the Shebandowanian orogeny at about 2695 Ma. It may represent collision of the Wawa subprovince with the composite Superior superterrane to the north. The second deformation (D 2) occurred at about 2680 Ma during the Minnesotan orogeny. It can be attributed to oblique collision of the Minnesota River Valley subprovince with the Superior superterrane along a north-dipping suture known as the Great Lakes Tectonic Zone in south-central MN (Percival and others, 2006). The three major deformation events are more or less coaxial, reflecting a continuum of roughly NW-SE-directed compression. The D3 event appears to have been a late manifestation of this, perhaps at a time when rocks were at sufficiently high crustal levels to elicit mainly partitioned brittle responses. The Midway sequence—a Timiskaming-type assemblage (STOPS 2, 5) The Midway sequence forms a wedge of strata <500 meters thick that youngs consistently southward. The basal (NW) contact is not exposed, but intersections in 10 drill holes indicate the contact is a fault in some localities, and an unconformity in others. The latter is indicated in drill core by the presence of sand-filled fractures in the subjacent Viking porphyry intrusion, and the abundance of QFP clasts in the “overlying” Midways sequence. The Midway sequence contains attributes of Timiskaming-type successions, including temporal and geographic association with large fault systems, alkalic and calcalkalic volcanic and intrusive rocks (hornblende trachyandesite), and conglomerate containing clasts of trachyandesite, together with those derived from older plutonic and strongly foliated substrate bedrock 120 �(Jirsa and Boerboom, 2003). The origin of such Timiskaming-type assemblages has been variously ascribed to localized extension in regional transpressional regimes (Jirsa, 2000; Corcoran and Mueller, 2007), or regional extension in response to imbrication and crustal loading during terrane accretion (Bleeker, 2012). Isolated occurrences of sequences similar to the Midway are exposed in other parts of Minnesota—near International Falls (Seine Group), in the Vermilion district (Gafvert Lake sequence), and in the Knife Lake area (Ogishkemuncie conglomerate). The ages for deposition of the Seine Group are 2693±1 to 2692±1 Ma (Fralick and Davis, 1999); the Gafvert Lake sequence is 2689.7± 0.8 Ma (Lodge and others, 2013); and the Ogishkemuncie conglomerate contains clasts of 2690.83 Ma Saganaga Tonalite (Driese and others, 2011). The broad equivalence of these sequences, and the fact that most young southward, indicates more or less synchronous development, which is consistent with an origin involving a single event of regional extension. Filler—taken from an old GAC GEOLOG publication 121 �Fig ure 2. Geologic map (A.) and schematic cross-section (B.) of the Virginia Horn area (modified from Jirsa and others, 1998) showing details of field trip STOPS 1 to 10. Block arrows on B indicate directions of stratigraphic facing. 122 �Gold Mineralization (STOP 3) In the 1930s visible gold was discovered by Minnesota Geological Survey geologist J.W. Gruner (Grout, 1937) in Archean rocks adjacent to a railroad grade cutting through the Virginia Horn (Fig. 2A and STOP 3). Perhaps because of the regional emphasis on iron ore mining, there was no systematic exploration for gold in the Virginia Horn until the 1980s. During the 1980s Newmont Mining, Rhude and Fryberger, Resources Limited, and American Shield conducted exploration programs that included 43 drill holes totaling 20,000 feet, geologic mapping, soil and outcrop geochemical surveys, and ground geophysical surveys. Most of the 1980s exploration focused on well exposed knobs of what is known locally as the Viking quartz-feldspar porphyry (QFP), which strikes west-southwest from the western side of the Pike River fault (Fig. 2B). Within the QFP, gold is concentrated in zones of variable brittle-ductile deformation (likely both D2 and D3), with associated carbonate-sericite alteration and abundant quartz veins. Proximal to the Pike River fault, higher grade gold mineralization occurs in one to three cm-thick quartz veins with pyrite, arsenopyrite and free gold, and in greyish quartz flooded zones with acicular arsenopyrite but no visible gold. One kilometer to the west of the Pike River Fault, most of the known gold mineralization is associated with the Viking QFP and variably sericitized porphyritic dacite that rarely outcrops. Gold is predominantly concentrated in one to three cm anastomosing quartz veins with pyrite and arsenopyrite concentrated along vein margins. Arsenopyrite occurs in irregular masses and not in the acicular habit associated with gold mineralization one km to the east. Recent studies of the geology of the Virginia Horn area by the Minnesota Geological Survey have shown that the Virginia Horn prospect is located within a Timiskaming-type geologic setting with good potential for gold mineralization beyond the porphyry and into adjacent metasedimentary and metavolcanic rocks (Jirsa and Boerboom, 2003; Bleeker, 2012). Subsequent analyses of samples from outcrop and drill core have confirmed that these rock types are also gold-enriched. Geologic Setting of Paleoproterozoic rocks (STOPS 6-8) The Paleoproterozoic strata exposed in the Virginia Horn are part of the Animikie Group, a sequence of sedimentary rocks, including basal quartzite and siltstone (Pokegama Quartzite), medial iron-bearing strata (Biwabik Iron Formation), and upper graywacke and shale of turbidite origin (Virginia Formation). Current models indicate deposition in a back-arc basin that evolved into a northward-migrating fore-deep during the compressional phase of the Penokean orogen—largely complete by about 1850 Ma (Schulz and Cannon, 2007; Pufahl and others, 2010). A depositional age for iron-formation can be inferred from interbedded volcanic tuff in the equivalent Gunflint Iron Formation to the northeast, which produced a UPb zircon date of approximately 1878 Ma (Fralick and others, 2002). The contact between ironformation and overlying slate of the Virginia Formation is marked by the presence of breccia and ejecta formed during the 1850 Ma Sudbury meteorite impact event. The ejecta contain abundant petrographic evidence of impact origin, including the presence of zoned spherules and quartz fragments displaying multiple planar deformation features. The impact-related horizon, known as the Sudbury Impact Layer, is well exposed in the Gunflint Lake area of northeast Minnesota (Jirsa and others, 2011), in the Thunder Bay area of adjacent Ontario (Addison and others, 2005), and in Michigan (Cannon and others, 2010; Pufahl and other, 2007); however, it can be seen only in drill core on the Mesabi range. Tuffaceous layers in basal strata of Virginia Formation were sampled a few meters above the Sudbury Impact Layer and produced an age of 1832±3 Ma (Addison and others, 2005). The belt of exposure forming the Mesabi Iron Range defines a regional monocline striking ENE and dipping shallowly (0-12 degrees) southward. The exception to this trend is in the Virginia Horn, where strike varies from N-S to NE, and dips as great as 25 degrees are recorded. This paired syncline-anticline is inferred to be related to uplift of a horst, now manifest in the Archean core of the structure, which formed by a combination of folding and faulting (Morey, 2003). Offset along the Laurentian fault, which was south-side down after the D2 Archean metamorphic and deformation event, was subsequently reactivated during the Paleoproterozoic to produce north-side down movement during and after deposition of the Animikie Group. 123 �DESCRIPTION OF FIELD TRIP STOPS STOP 1 Archean pillowed and massive greenstone—Old Gilbert school Location: UTM (NAD 83, Zone 15): 539,820E/5,259,750N; north edge of athletic fields for former Gilbert Junior High School off Wisconsin Avenue, 4 blocks northwest of State Highway 37. Description: This outcrop of pillowed and massive basalt is part of the Archean Mud Lake sequence, metamorphosed to low greenschist-grade. Pillow shapes indicate stratigraphic facing is to the northwest, consistent with the location of this outcrop on the south side of a major D 1 structure known as the Mud Lake syncline. Note also the presence locally of fractures and shallow depressions on the outcrop surface that are filled with reddish jasper, presumably deposited by overstepping of Paleoproterozoic seas onto the eroded surface of Archean bedrock during deposition of the Biwabik Iron Formation. Discussion: Detailed structural study by Jirsa and others (1998) and Jirsa and Boerboom (2003b) demonstrate that the tholeiitic and calc-alkalic volcanic rocks and tholeiitic intrusions are conformably overlain by graywacke and slate of the Mud Lake sequence (STOPS 3 and 4). In detail, the Mud Lake sequence forms a broad, twice-deformed, west-plunging syncline that has been segmented by faults of several generations. STOP 2 Archean volcanogenic conglomerate of the Midway sequence Location: UTM: 539120E/5261040N; Old Railroad trail Description: This former railroad cut exposes parts of the volcanic and lower conglomerate facies of the Timiskamingtype Midway sequence. Figure 3 shows the position of this exposure within a composite stratigraphic section. The rocks here are characterized by disorganized beds of poorly sorted conglomerate and volcanic breccia. Dark red, green, and purple clasts of hornblende- and plagioclase-phyric trachyandesitic composition are most abundant. Both normal and reversed grading are preserved locally. Clasts are as large as 25cm, and diamictites containing outsized clasts are common in this unit. Flattening of clasts in the plane of D2 is apparent, and a matrix of varied grain size locally displays anastomosing S 2 cleavage. Overall, the conglomerate contains clasts representing all lithologic components of the Mud Lake sequence and the porphyry dikes that intruded it. However, significant variations in clast content and internal organization of bedding characterize these units, and these attributes vary gradationally both laterally and with stratigraphic height. Figure 4 shows the map and cross-section distribution of the Midway sequence based on both drill core and surface exposures. The Upper conglomerate facies will be visited at STOP 5. Discussion: Remarkably, the polychromatic trachyandesite clasts are identical with those in parts of the Shebandowan assemblage exposed some 240 km, more or less along strike to the NE in Ontario (e.g., Aubut and Campbell, 2012), despite the intervening Giants Range batholith and other terranes. Figure 3. Composite cross section of the Midway sequence. Dark polygons represent clasts of trachyandesitic to trachybasaltic composition; white polygons represent clasts of quartzofeldspathic porphyry identical with the Viking QFP (STOP 3); which is represented diagrammatically by the dot pattern. (From Jirsa and Boerboom, 2003, Fig. 2.4). 124 �STOP 3 STOP 2 Figure 4. Surface and drill core-based geology of the “Viking Prospect area (STOPS 2 and 3). A) shows geologic map view; B) shows drill holes that intersected the northeastern (basal) contact of Midway sequence with adjacent rocks. From Jirsa and Boerboom, 2003, Fig. 2.6. 125 �STOP 3 Archean Graywacke, argillite, and quartzofeldspathic porphyry with Au mineralization Location: UTM: 537715E/5261240; Old Railroad trail Description: This former railroad cut and outcrops nearby expose interlayered graywacke and argillite of the Mud Lake sequence, intruded locally by quartzofeldspathic porphyry. One of the earliest reports of gold in Minnesota (1930) was made at this locality, and visible gold can still be found associated with small quartz veins. The porphyry intrusions here are identical with and presumably apophosial offshoots of the larger Viking QFP exposed to the east along and south of the trail (Fig. 4A). Regionally, the porphyry contains ornately embayed phenocrysts of quartz as large as 2cm, smaller albite phenocrysts, and minor mica in an aphanitic quartzofeldspathic groundmass. Groundmass is commonly crossed by anastomosing shear planes and altered to combinations of quartz, sericite, dolomite, iron-carbonate (ankerite and ferroan dolomite), pyrite, and locally arsenopyrite. Channel samples on this outcrop (Fig. 5) and core from several nearby exploration drill holes will be displayed and discussed. Discussion: During 2009 and 2010, Vermillion Gold, LLC (http://vermilliongold.com) completed nine drill holes designed to reevaluate areas where gold mineralization was intersected by the 1980s drill holes, and test new targets outside of the Viking QFP. Five of the nine drill holes have focused on gold mineralization in altered porphyritic dacite that outcrops adjacent to a railroad grade on the western side of the property (STOP 3). A channel sample of intercalated porphyritic dacite and metasediments taken from the outcrop by Newmont in the 1980s averaged 1.2 gpt over a sample length of 77.5 ft. Vermillion Gold’s 2009 drill hole (Fig. 6) beneath the railroad grade outcrop intersected 1.1 gpt gold over an interval of 195.7 feet and includes intersections of 16.1 gpt/3.6 ft and 11.4 gpt/4.2 ft. A 2010 drill hole, located approximately 150 m south of the railroad outcrop, intersected a thick unit of porphyritic dacite locally cut by thin fingers of Viking QFP. A 224.9 ft section of the drill hole averages 1.0 gpt gold and includes a 77.9 gpt/2.3 ft sample with visible gold in a quartz vein. Within the thick, gold-enriched intersections, values of 100 to 300 ppb are associated with disseminated pyrite and arsenopyrite. Samples with multi-gram gold values include quartz veins with free gold. There is a strong correlation between gold and arsenic values, however, samples with the highest grade gold values reflect free gold and not increased arsenopyrite content. 126 �Figure 5. Location (upper image) and analytical results (lower image) of channel samples on outcrops at STOP 3 (imagery from Vermillion Gold, LLC) 127 �Figure 6. Partial log of core from drill hole VH-09-4, showing lithologic and analytical results (from records of Vermillion Gold, LLC). 128 �STOP 4 Archean graywacke and slate, intruded by quartzofeldspathic porphyry—Bourgin Road Location: UTM: 536,311E/5,260,659N; road cut on east side of Bourgin Road. Description: Outcrops along this side of the road expose quartzofeldspathic porphyry (QFP) intruded into variably deformed graywacke, siltstone, and slate of the Mud Lake sequence. The sedimentary rocks here are moderately deformed, but much of that deformation is inferred to predate the main cleavage-forming event D2, and some may have occurred prior to lithification. The QFP is large and continuous to the east, but at this locality, appears to be segmented into a zone of multiple anastomosing dikes. Both graywacke and QFP are intensely altered to some combination of iron-carbonate minerals (ankerite, ferroan dolomite) and sericite. Regionally, this style of alteration is commonly, though not always associated with the QFP intrusions—presumably because they remained more structurally rigid than the enclosing sedimentary rocks during the shear-related deformation event that accompanied alteration late in D2. Most gold mineralization in the area is closely allied to this alteration, yet this outcrop is apparently barren. STOP 5 Archean conglomerate—Midway sequence Location: UTM: 535,713E/5,259,459N; driveway at No. 7 Mesabi Lane, village of Midway. NOTE: This is private property! Permission must be obtained before entering. Description: The Archean conglomerate and lithic sandstone that form this driveway surface represents the Upper conglomerate facies of the Timiskaming-type Midway sequence. It differs from the Lower conglomerate facies at STOP 2 in containing more diverse clast content, more rounded clasts, graded bedding, and more abundant sandy beds. These attributes imply submarine deposition. Taken in the context of the sequence stratigraphy (Fig. 3), this may indicate basin deepening over time, which is consistent with observations of other Timiskaming-type successions (Bleeker, 2012). The conglomerate contains clasts of basalt, graywacke, quartzofeldspathic porphyry (QFP), and porphyritic trachyandesite. This provenance indicates that the older Archean rocks of the Mud Lake sequence were intruded by QFP, deformed, uplifted and eroded to provide detritus to what was probably a “pull-apart” or extensional basin developed along a major structure now occupied by the Pike River fault zone. The southward-younging basin is bounded by both faults and an inferred unconformity (Fig. 7). Discussion: Note also the presence of red jasper in depressions and joints as at STOP 1, indicating that this outcrop surface represents paleo-seafloor during deposition of the Paleoproterozoic Biwabik Iron Formation. Figure 7. Schematic illustration of Midway basin geometry prior to steepening by D2 compressional deformation (from Jirsa and Boerboom, 2003). STOP 6 Paleoproterozoic Biwabik Iron Formation-Highway 53 Location: UTM: 536,263E/5,256,200N; Outcrops along north-bound exit ramp from Hwy 53 to Hwy 37, Eveleth. Description: This exposure of gently south-dipping strata is part of the Lower Cherty member of the Biwabik Iron Formation. It lies nearly at the crest-line of the anticline that forms half of the Horn structure. The ironformation forms a transitional contact with the underlying Pokegama Quartzite exposed at Stop 7. Both formations have fine- to coarse-grained sandy textures and cross-bedding, consistent with a high-energy, 129 �near-shore depositional environment. Bimodal-bipolar cross-strata in the iron-formation indicate that tidal currents may have played an active role in deposition (Ojakangas, 1993), though tidal bundles have not been documented. The most significant difference between these two formations is the abrupt change in sediment source from the extrabasinal quartz grains in the Pokegama, to recycled, chemically precipitated chert nearly devoid of detrital grains in the Biwabik. Discussions: 1) One possible explanation for the abrupt change in sediment source in the transition from Pokegama Quartzite to basal Biwabik Iron Formation may be related to topography of the watershed. If the terrane was a relatively flat peneplain, the continued rise of sea level may have essentially drowned the detrital source region. 2) The Biwabik Iron Formation is generally divided into 4 members; termed lower Cherty, Lower Slaty, Upper Cherty, and Upper Slaty (Fig. 8). These are convenient field names based on readily apparent bedding attributes; however, they are somewhat misleading. The cherty units are beds of recycled granular chemical precipitates including chert, iron oxides, iron carbonates, and iron silicates. They are interbedded on all scales with “slaty” units of fine-grained, laminated iron silicates and iron carbonates. In most of the Mesabi range, the iron-formation and associated strata were not significantly metamorphosed, and much of the textural and mineralogic attributes are products of diagenesis and subsequent fluid movement. As a result, no slaty cleavage exists, and thus the term “slate” is applied only as a field identifier. Most, though not all of the iron ore mined on the Mesabi range is extracted from cherty members (Fig. 8B). Figure 8. Simplified geologic map (A) of the Mesabi Iron Range showing locations of taconite mines (black) and drill holes (numbered), and cross-section (B) showing subdivision of the Biwabik Iron Formation based on mined sections and drill core, and approximate intervals mined for taconite at each locality (modified from Jirsa and others, 2008). 130 �STOP 7 Paleoproterozoic Pokegama Quartzite-Highway 53 Location: UTM: 535,956E/5,256,913N; Outcrops along north-bound entrance ramp onto Hwy 53 from Hwy 37, Eveleth. Description: In the area of the Virginia horn, the Pokegama consists largely of siltstone and shale. This exposure represents the sandy, upper member of the Pokegama Quartzite, which is only of fraction of the units’ total thickness of 26-51 m. It is quartz arenite characterized by coarse grain size, intraclasts of shale and siltstone, and massive beds as thick as 1.5 m, separated by thin beds of shale and siltstone. Ojakangas (1993) interpreted that the deposition of this facies occurred within a high-energy, lower tidal or subtidal environment. Because stratigraphic dip is southward, the outcrop at this location is inferred to be several meters stratigraphically beneath the iron-formation at STOP 6. Discussions: 1) The basal strata of the Pokegama Quartzite is marked locally by conglomerate composed of a poorly sorted array of clasts derived from underlying Archean and Paleoproterozoic (diabase dikes) bedrock. The patchy distribution of conglomerate, and the presence of red jasper in fractures on some Archean exposures (as at STOPS 1 and 5), implies that chemical sedimentation abruptly overstepped clastic deposition during early evolution of the Animikie basin. 2) The contact between the Pokegama Quartzite and overlying Biwabik Iron Formation is conformable and gradational. In the transition zone, both units contain similar sedimentary structures and grain size, implying continuity of depositional process. The primary difference between them is grain composition—the Pokegama grains are epiclastic vs. those in the Biwabik are reworked from poorly lithified or unlithified chemical precipitates. The absence of epiclastic grains in the nearly 200 m-thick stratigraphic section of the Biwabik begs the question: How was this detritus abruptly shut off from the watershed? STOP 8 Abandoned and flooded Rouchleau “natural ore” mine Location: UTM: 535,710E/5,261,650N; Mineview in the Sky overlook near Virginia. Description: North from this overlook is a 3-mile-long complex of abandoned mining properties, known collectively as the Rouchleau mine, all developed within the Paleoproterozoic Biwabik Iron Formation. Actually, within this view there were some 15 separately named mines that collectively shipped ore during the period 1893-1986. All of them, and nearly 400 more along the 150-mile long Mesabi Iron Range, extracted oxidized (hematite- or goethite-rich) and leached (silica-depleted) iron-formation referred to locally as “natural ore.” Iron-formation at this point lies on the north-trending limb separating the syncline to our west and the anticline to the east. The natural ore deposits here are localized along a set of faults (Fig. 2A) that presumably provided the plumbing system for fluids that first oxidized the formation and produced permeability, then leached silica from the porous zones, thus increasing ore tenor. Natural ores typically contained as much as 50 percent iron and less than 10 percent silica. Since about the 1950’s, the principal “ore of choice” has shifted from hematite- to magnetite-rich deposits. The mammoth open-pit mine in the distance to the northwest, and another just southwest of the highway, are developed in unoxidized magnetite ore containing about 30 percent iron, and 50 percent silica. The ore mined at these locations, and several others along the range is the source of the iron concentrate known as taconite. The name taconite has also been applied generally to magnetite-bearing iron-formation where it contains sufficient iron content to be mined for a profit using today’s technology. 131 �Discussions: 1) Origin of “natural ore” Nearly 70 percent of the 3.6 billion metric tons of iron ore produced on the Mesabi range between years 1892 and about 2000 was extracted as natural ores. Although it is generally accepted that these ores formed by oxidation and leaching along folds, faults, and bedding planes, the source and composition of altering solutions and the timing of alteration have been subjects of considerable debate among economic geologists for nearly 80 years. Much of the literature and geologic observations on the issue are reviewed in Morey (2003). Many writers support the concept of descending meteoric waters to account for the dissolution of silica and oxidation of iron minerals. Others, including Gruner (1930) believed the geologic features were better explained by ascending hydrothermal solutions. Gruner’s theory failed to gain common acceptance, in part because no driving mechanism for such a hydrothermal system could be envisioned. The integration of Animikie Group strata into the tectonic context of the Penokean orogen in east-central Minnesota revived the theory of hydrothermal fluid flow within the Pokegama Quartzite and ultimately the iron-formation, as part of a continent-scale, gravity-driven ground-water system (Morey, 1999). The debate continues—fueled in part by the observation that most of the alteration occurs near the present land surface. Field trip #1 in this guide explores these issues more fully. 2) Highway relocation This overlook will soon be gone, as taconite mining to the southwest is slated to expand into the Rouchleau pit area. As a result, U.S. Highway 53 will be rerouted to skirt the new mining. Currently the Minnesota Department of Transportation (MnDOT), the Department of Natural Resources, and the mining company are engaged in geotechnical work and discussions to evaluate the various potential new routes for the highway. Exploratory drilling (for both ferrous and non-ferrous metallic mineral potential), geotechnical drilling, and geophysical surveying have been conducted by MnDOT to assess the financial and engineering risks of two relocation alternatives proposed through the Rouchleau Pit (Fig. 9), as well as a third alternative through the mine site to the southwest. This type of apparent conflict between surface infrastructure and mining has characterized development along the Mesabi Iron Range for more than 120 years. It was particularly acute during the shift in the “ore of choice” in the 1950’s and 1960’s from natural ores that typically occur in narrow and steep, structurally-controlled deposits; to taconite that occurs in more widely distributed layers. Entire cities have been moved and removed, as at Hibbing (See Field Trip B in this guide book for example). 3) Wind turbines on northern horizon Minnesota Power’s Taconite Ridge Energy Center is visible to the north, located on U.S. Steel property in Mt. Iron. It consists of 10 wind turbines that generate 25-megawatts, capable of powering the equivalent of 8,000 homes annually. (Reference=http://www.hometownfocus.us/news/2013-09-06/Mining_Features; accessed 2/2014) 132 �Figure 9—Map showing potential new routes for a portion of Highway 53 (blue and green lines through Rouchleau mine complex), and associated test drilling (magenta dots=holes to iron-formation; green dots=holes to Archean) to evaluate both the engineering and resource implications of proposed realignments. Colored overlay on air photo imagery shows geologic units and faults from Jirsa and others (2012). Image created using data from MnDOT. STOP 9 Archean Giants Range batholith at “Confusion Hill” Location: UTM: 534,337E/5,269,458N; outcrops at Laurentian Wayside, near Highway 169 south of its split with Highway 53. Description: Exposures at Confusion Hill are a small part of the Giants Range batholith, which forms the core bedrock of the Laurentian (drainage) divide. The batholith is a belt of intrusions that can be traced on geophysical maps and outcrop east to the Mesoproterozoic Duluth Complex, and west beyond the western border of Minnesota. It separates Archean supracrustal sequences in the Virginia Horn from those of the Vermilion District to the north—making stratigraphic correlation between the two districts speculative in the near absence of high-precision geochronologic data. Exposed near this wayside and in road cuts on both sides of the highway is an array of variably layered intrusions having both tonalitic (white) and dioritic (black) compositions. A cursory look shows intrusive relationships that conclusively demonstrate that diorite was emplaced into tonalite at one locality; and at another, tonalite was emplaced into diorite. In detail, all compositions intermediate between the two end members are also present locally. Although the dioritic component is abundant here, the bulk of the mapped unit is tonalitic. Emplacement of this unit, now known as the Lookout Mountain tonalite, probably involved some degree of magma mingling. It may be equivalent to tonalitic gneiss exposed along strike to the east and having a somewhat imprecise U-Pb zircon age of 2718±67 Ma (Southwick, 1994). Dikes of tonalite that cut the adjacent high-grade supracrustal rocks of the highgrade Minntac sequence contain metamorphic fabrics, yet little evidence of metamorphic origin can be 133 �seen in the interior of the body, implying it is syntectonic to pre-tectonic with respect to D2 deformation. U-Pb zircon dates (Boerboom and Zartman, 1993) of two components of the batholith exposed to the north bracket the age of D2 deformation between about 2674 and 2682 Ma. STOP 10 Archean diorite in Giants Range batholith Location: UTM: 534,337E/5,269,458N; outcrops at Laurentian Wayside, near Highway 169 south of its split with Highway 53. Description: Rock type exposed in this now partially reclaimed quarry is a massive hornblende-pyroxene-biotite diorite. Currently, little is known about the intrusion, as no petrologic study, geochronologic analysis, or mapping of its contacts has been conducted by the authors. Nevertheless, it is similar to other small alkalic plutons in and adjacent to the Giants Range Batholith. These vary in composition—in some cases within a single intrusion—from syenite to monzodiorite to lamprophyre and pyroxenite (Boerboom, 1994). It is interesting to speculate that this intrusion may fall into the category of the late alkalic to calcalkalic intrusions that are temporally, and likely geochemically, related to Timiskaming-type assemblages such as the Midway sequence. REFERENCES Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005, Discovery of distal ejecta from the 1850 Ma Sudbury impact event: Geology, v. 33, p. 193-196. Aubut, A., and Campbell, D., 2012, Field Trip 4—Shebandowan Mine area: in Hollings, P., MacTavish, A., and Addison, W., Institute on Lake Superior Geology Proceedings, Part 2, p. 67-73. Bleeker, W., 2012, Targeted Geoscience Initiative 4. Lode Gold deposits in ancient deformed and metamorphosed terranes: The role of extension in the formation of Timiskaming basins and large gold deposits, Abitibi Greenstone Belt—A discussion: in Summary of Field Work and other activities 2012, Ontario Geological Survey Open File Report 6280, p. 47-1 to 47-12. Boerboom, T.J., 1994, Alkalic plutons of northeastern Minnesota: in Southwick, D.L., ed., Short contributions to the geology of Minnesota, Minnesota Geological Survey Report of Investigations 43, p. 20-38. Boerboom, T.J., and Zartman, R.E., 1993, Geology, geochemistry, and geochronology of the central Giants Range batholith, northeastern Minnesota: Canadian Journal of Earth Sciences, v. 30, p. 2510-2522. Cannon, W.F., Schultz, K.J., Horton, W. Jr., and Kring, D.A., 2010, The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA: Geological Society of America Bulletin, v. 122, p. 50-75. Corcoran, P.L., and Mueller, W.U., 2007. Time-transgressive Archean unconformities underlying molasse basin-fill successions of dissected oceanic arcs, Superior Province,Canada. Journal of Geology 115, 655–674. Driese, S.G., Jirsa, M.A., Ren, M., Sheldon, N.D., Brantley, S.L., Parker, D., and Schmitz, M., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, v. 189, p. 1-17. Fralick, P., and Davis, D.W., 1999, The Seine-Coutchiching problem revisited: Sedimentology, geochronology and geochemistry of sedimentary units in the Rainy Lake and Sioux Lookout areas: in Harrap, R.M., and Helmstaedt, H., eds., 1999 Western Superior Transect Fifth Annual Workshop 70, Lithoprobe Secretariat, University of British Columbia, p. 66-75. Fralick, P., Davis, D.W., and Kissin, S.A., 2002, The age of the Gunflint Formation, Ontario, Canada: single zircon U-Pb age determinations from reworked volcanic ash: Canadian Journal of Earth Sciences, v. 39, p. 1085-1091. Grout, F.F., 1937, Petrographic study of gold prospects of Minnesota: Economic Geology, v. 37, p. 56-68. Gruner, J.W., 1930, Hydrothermal oxidation and leaching experiments; their bearing on the origin of Lake Superior hematite-limonite ores: Economic Geology, v. 21 pp. 697-719, 837-867. Jirsa, M.A., 2000, The Midway sequence: A Timiskaming-type, pull-apart basin deposit in the western Wawa subprovince, Minnesota: Canadian Journal of Earth Sciences, v. 37, p. 1-15. 134 �Jirsa, M.A., and Boerboom, T.J., 2003, Geology and mineralization of Archean bedrock in the Virginia Horn: in Jirsa, M.A., and Morey, G.B., eds., Contributions to the geology of the Virginia Horn area, St. Louis County, Minnesota: Minnesota Geological Survey Report of Investigations 53, p. 10-73. Jirsa, M.A., Boerboom, T.J., Chandler, V.W., 2012, Geologic map of Minnesota—Precambrian bedrock geology: Minnesota Geological Survey State Map Series S-22, scale 1:500,000. Jirsa, M.A., Boerboom, T.J., and Morey, G.B., 1998, Bedrock geologic map of the Virginia Horn, Mesabi Iron Range, St. Louis County, Minnesota: Minnesota Geological Survey, Miscellaneous Map Series M-85, scale 1:48 000. Jirsa, M.A., Fralick, P.W., Weiblen, P.W, and Anderson, J.L.B., 2011, The Sudbury impact layer in the western Lake Superior region: in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 147-169. Jirsa, M.A., and Green, J.C., 2011, Classic Precambrian geology of northeast Minnesota: in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 25-45. Jirsa, M.A., Miller, J.D. Jr., and Morey, G.B., 2008, Geology of the Biwabik Iron Formation and Duluth Complex: Regulatory Toxicology and Pharmacology, v. 52, p. S5-S10. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., Jirsa, M.A., and Hamilton, M.A., 2013, New U-Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research v. 235, p. 264-277. Morey, G.B., 1999, High-grade iron ore deposits of the Mesabi range, Minnesota—Product of a continental-scale Proterozoic ground-water flow system: Economic Geology, v. 94, p. 133-142. Morey, G.B., 2003, Paleoproterozoic Animikie Group, related rocks, and associated iron-ore deposits in the Virginia Horn: in Jirsa, M.A., and Morey, G.B., eds., Contributions to the geology of the Virginia Horn area, St. Louis County, Minnesota: Minnesota Geological Survey Report of Investigations 53, p. 74-102. Ojakangas, R.W., 1993, Pokegama Quartzite: in Institute on Lake Superior Geology Proceedings, 39th Annual Meeting, Eveleth Minnesota, v. 39, Part 2, p.19-21 and 46-48. Percival, J.A., and Helmstaedt, H., 2006, The Western Superior Lithoprobe and NATMAP transects: Introduction and summary: Canadian Journal of Earth Science, v. 43, p. 743-747 and articles therein. Percival, J.A., Sandborn-Barrie, M., Skulski, T., Stott, G.M., Helmstaedt, H., and White, D.J., 2006, Tectonic evolution of the western Superior Province from NATMAP and Lithoprobe studies: Canadian Journal of Earth Science, v. 43, p. 1085-1117. Peterson, D.M., Gallup, C., Jirsa, M.A., and Davis, D.W., 2001, Correlation of Archean assemblages across the U.S.-Canadian border: Phase I geochronology (abs): Institute on Lake Superior Geology, 47 th Annual Meeting, Madison, Wisconsin, Proceedings v. 47, Part 1, p. 77-78. Pufahl, P.K., Hiatt, E.E., and Kyser, T.K., Does the Paleoproterozoic Animikie Basin record the sulfidic ocean transition?: Geology, v.38, p. 659-662. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, G.J., and Edwards, C.T., 2007, Physical and chemical evidence of the 1850 Ma Sudbury impact event in the Baraga Group, Michigan: Geology, v. 35, p. 827-830. Southwick, D.L., 1994, Assorted geochronologic studies of Precambrian terranes in Minnesota: A potpourri of timely information: in Southwick, D.L., ed., Shorter Contributions to the Geology of Minnesota, Minnesota Geological Survey Report of Investigations 43, p. 1-19. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4-25. 135 �FRIDAY AFTERNOON FIELD TRIPS MAY 16, 2014 136 �FIELD TRIP A Friday, May 16, 2014 STATE DRILL CORE LIBRARY—HIBBING MINNESOTA Minnesota Department of Natural Resources—Division of Lands and Minerals LEADERS: Dave Dahl, (MnDNR), and Dean Rossell (Kennecott) INTRODUCTION The Minnesota Department of Natural Resources maintains a Drill Core Library in Hibbing, Minnesota. It serves as the single State of Minnesota repository for archiving bedrock and earthen material core samples collected during minerals exploration, engineering, and geoscience research programs across the state. The library attracts a worldwide audience of scientists who use core samples to develop new ideas about the capacity of the state’s bedrock and glacial materials to host mineral resources, and to model the geologic forces and features that have shaped the state’s foundation. This trip will provide opportunities for visitors to tour the facility and view some very old exploration data collections, century-old historical cores, and recent scientific and exploratory cores. The latter include cores taken from iron-formation, Midcontinent rift peridotite and gabbro, greenstone belt prospects, and sedimentary and glacial settings. Figure 1. Interior of Building #3 showing stacked core boxes. 137 �Figure 2. Map of Minnesota showing locations of drill holes from which cores stored at DNR were extracted. 138 �The three buildings that comprise the core library facility house more than 3 million lineal feet of drilled core samples archived from approximately 9,000 exploratory and scientific borings. Some archived samples are well cuttings (depending on the drilling method employed during sampling). The archive collection contains approximately 7,000 mineral exploration cores, 1,500 roadway and bridge foundation cores, and 500 cores collected during scientific, governmental and academic research investigations (Fig. 2). Building #1, built in 1972 has a storage capacity of 400,000 lineal feet of core. Building #2, constructed in 1979 has a 600,000 lineal foot storage capacity. Building #3 (Fig. 1), originally constructed in 1989 with a capacity of 800,000 lineal feet, has been expanded twice. In 1995 the building was doubled in size through an addition, and in 2009 the building was nearly doubled in size again through addition of a wing. In present configuration, the three buildings have capacity to store approximately 4 million lineal feet of NQ-sized core samples. Core samples are normally transferred to the facility in two-foot long boxes, 5 core segments per box, or 10 feet of core per box. Box storage capacities range from 7 segments for small diameter (A- and E-size) core to 2 segments for large diameter (PQ—size) core. Boxes are generally designed to hold 50 lbs weight or less. Today, most exploratory boring samples are delivered to the library in fulfilment of statutory requirements (M.S. 103 I.601 and 103 I.605) which have been in effect since 1980. The library archives are augmented by substantial collections of historical (pre-1980) core samples that have been received via donation from mineral exploration companies or through consolidation of agency core collections. The earliest known collar date in the core collection is 1905, for exploratory cores taken along the Gunflint trail. Mineral exploration archive documents housed in Hibbing indicate that core samples were obtained in the Vermilion district as early as the 1870’s. Core samples are expensive to obtain and maintain, but that money is well spent. Samples acquired to meet one investigative objective commonly are “recycled and reused” several times in subsequent programs. They can be used to test new working models of geologic processes, and to provide new insights on the disposition and location of mineral resources. Analytical techniques (geochemical, geochronologic, and geophysical) are constantly evolving, and these cores provide ready materials for testing within well constrained geologic contexts. Archived cores have provided the basis for advancement in the evaluation of several copper-nickel, gold, titanium, and iron deposits and prospects (Tamarack, Birch Lake, Maturi, Spruce Road, Serpentine, Mud Creek, Lost Lake, Virginia Horn, Longnose, TiTac, Buckeye, Emily and others). New private investments to advance these properties, some of which include School Trust or other state-owned mineral lands, are on the order of $200 million over the past decade. The DNR recently convened a working group to increase efficient delivery of core library services and to attract research and investment in the evaluation and understanding of Minnesota resources. 139 ��FIELD TRIP B Friday, May 16, 2014 HIBBING’S IRON MINING AND CULTURAL HISTORY LEADERS: Henry Djerlev, Bob Kearney, Erica Larson, and Hibbing Historical Society Staff INTRODUCTION The nearly 120 years of iron ore mining in the Hibbing area has certainly played a large part in shaping the history and culture of the towns and residents and the rest of the Mesabi Iron Range. Within the outline of what is now called the Hull-Rust-Mahoning open pit, more than 30 separate mines operated from 1895 to the present. The early miners emigrated from dozens of countries, and all of their various languages were spoken in the mines. The mix of cultures made for awkward communications in the mines, and often created ethnic neighborhoods in the mining locations and larger towns. Initial underground mining quickly changed to open pit due to the nature of the ore body and the introduction of large steam operated drills, shovels and trains. Very quickly, with eastern monies invested, some of our larger corporations emerged, such as United States Steel. Very small "mining locations" grew into formal towns like Hibbing founded by Frank Hibbing and A. J. Trimble. Open pit mining quickly progressed from small individual pits, into one large open pit that was nicknamed "The Grand Canyon of the North". This expansion made it necessary to physically move, from 1919 to 1921, what was called North Hibbing to the south in order to make room for increased mining. This gave Hibbing another nickname: "The Town that Moved". During this short tour of Hibbing, the National Historic Landmark of the Hull-RustMahoning Mine will be visited and several of the historic buildings that were made possible by iron mining dollars. These include the spectacular Hibbing High School and the Hibbing Historical Society Museum. Stop locations given in latitude/longitude. *Departure Point- Hibbing Park Hotel (47o 25' 38.71'' N/92o 35' 26.22''W) On the short trip from the Hibbing Park Hotel to Stop 1 the bus will pass by several historic points in Hibbing such as the one time home of Andrew "Bus Andy" Anderson, built in 1920. Andy with his partner Carl Wickman initiated what was to become The Greyhound Bus Company. Carl, after losing his job at the Alice mine started as a salesman for the Hupmobile company. In 1914, after seeing the failing sales of the seven passenger Hupmobile, he tried to show Figure 1: Early photo of one of Mesabi Transportations Hupmobile Buses. his clients what a great product the 140 �Hupmobile was by taking people for short rides. Wickman was soon giving miners rides to and from work for a cheap fifteen cents a ride. When they found out that giving the miners transportation was more profitable, Wickman and Anderson created the Mesaba Transportation Company. Three years after starting the company, they were running 18 buses and were making $40,000 a year. In 1922, he sold the company for $60,000. In 1933, the company was formally named The Greyhound Corporation and was running nationally. Other historic points along this leg of the tour will be the Sons of Italy Hall (1923), Mesabi Railway Company(1921)—now Zimmy’s Restaurant, the Androy Hotel (1923), Bennett Park, and the Greyhound Bus Museum. The small city got its start by a German immigrant named Frank Hibbing who founded the town in 1893. Originally named Frans Dietrich Von Ahlen, Frank took his mother's last name of "Hibbing" which comes from English descent. Frank decided to do this out of honor for his mother who passed away in his infancy and thought that this would be a good move for his exploration of the "New World." Frank Hibbing originally settled in Beaver Dam, Wisconsin where he worked at a farm and shingle mill. Originally he had hopes of becoming a lawyer but after finding out the extreme differences between the German and English language, he decided to forego that dream and then became interested in the Figure 2 - Frank Hibbing area's most abundant resource: timber. In 1887, Frank Hibbing moved to Duluth and became a real estate salesman which eventually lead him further north into the Vermillion Range. It was not until 1892 when he and thirty men set out to cut a road from Mountain Iron to what was then called Section 22. While cutting this road, Frank Hibbing found iron ore on the ground and realized what that meant to the area's economy. Little did Frank Hibbing know at the time, this ore deposit would be one of the largest in the world! In 1893, the city of Hibbing was laid out and named in honor of Frank Hibbing. The city even has a statue to the German who had the sense to notice the reddish soil and the value that it had. Artist Robert Mitchell, born in Alice location, created the statute which was dedicated on October 21, 1941. Robert was the son of a Hanna Company mining captain for whom Mitchell location was named. Frank took so much pride in his new town that he used personal means to finance the first water plant, electrical plant, hotel, saw mill, and bank building. Frank Hibbing made Duluth his home for the last ten years of his life until his death from appendicitis on July 30, 1897. He did retain close communication with "his" town during that time. Frank Hibbing was only 40 years old. Figure 3- Frank Hibbing Park 141 �Stop 1- HULL-RUST MAHONING MINE (47o 26' 50.53'' N/92o 56' 45.86''W) "The Largest Open Pit Iron Ore Mine in the World" is more than three miles long, two miles wide and 600 feet deep. This man-made "Grand Canyon of the North" was the one of the first open pit mines on the Mesabi Iron Range. This amazing view continues to grow as the Hibbing Taconite Company mine expands its mining operations. Rotary drills, 33-cubic-yard shovels and 240-ton production trucks can been seen in action at this National Historic Site. Occasionally, you may witness a production mining blast of nearly 1 million tons used to clear bedrock away and break the taconite ore for processing in the plant. Figure 4: Recent photo of the Hull-Rust-Mahoning Pit Since 1895 more than 1.4 billion tons of earth have been removed on its 2,000 acres of land, and more than 800 million gross tons of iron ore have been shipped from the mine. At peak production in the 1940's, as much as one quarter of the ore mined in the United States came from the Hull Rust Mine. Currently Hibbing Taconite Company (Cliffs) produces approximately 8 Million tons of taconite pellets annually. Over 520 million tons of waste material and 690 million tons of iron ore have been removed. A slide presentation in the observation building explains the colorful history of the mine and early mining. An observation building, mine exhibits, mine shovel bucket, mining truck, interpretive graphics and a walking trail complete the trip to the Hull Rust Mine View. A miner poses near the edge of the pit. This area of the Mesabi Range was first explored in 1893, shortly after the Mountain Iron Mine was established in 1892. The early development was as an underground mine, but open pit mining soon proved to be a better choice because of the shallow nature of the ore deposits. The many smaller open pit mines developed in the area soon merged into one large mine. The growth of the mine even resulted in the town of Hibbing being relocated to accommodate expansion. The move started in 1919 and took two years to complete at a cost of $16,000,000. A total of 185 houses and 20 businesses were moved, and some of the larger buildings had to be cut in half for the move. Only a portion of the network of city streets and foundations from old North Hibbing remain in the vicinity of the Mine Observation Figure 5- Early Miner along the crest Overlook. of the Hull-Rust Pit Figure 6- Street sign in Other historic points along the second leg of the tour will be Frank Hibbing North Hibbing Park (1941), the Godfrey House (1920's) and the Mitchell-Tappan House (1897). 142 �Stop 2- The Hibbing High School - 1922 (47o 25' 33.30'' N/92o 55' 57.08''W) One of the buildings that was built during the Oliver Mining Company's relocation of North Hibbing was the Hibbing High School. Built in 1921 by the Oliver Mining Company, the school building originally cost almost $4 million dollars. In today's dollars that would be close to $45 million dollars! Figure 7- Hibbing High School under construction. Figure 8- Present Day Hibbing High School Why did it cost so much? Well it is simple, in order to lure prospective workers and miners to work in dangerous situations such as tunnels and around explosions, they had to provide a first class environment for their families and especially their children. So they went all out for the education. The auditorium of the high school was modeled after the famous Capitol Theatre in New York City. The auditorium has cut glass chandeliers which were imported from Belgium which light the 1800 velvet seated venue. The chandeliers were originally priced at $15,000 when built and today are insured for over $250,000 each. The auditorium has a rarity in it as well, which is a Barton pipe organ. Only two are in existence in North America. With 1800 pipes, it can synthesize any instrument excluding the violin. Figure 9- Hibbing High School Auditorium A few celebrities have attended this high school including basketball star and musician Robert Zimmerman aka "Bob Dylan". Robert Zimmerman may have been born in Duluth, Minnesota on May 24, 1941 but it is Hibbing, Minnesota where he grew up. Robert grew up in Duluth, MN until he was six years old. It was when his father was sick with polio that his family moved back to his mother's hometown of Hibbing, MN. This is where he fell in love with music and formed many bands throughout his high school career including such Figure 10- Dylan Family Home names as "The Shadow Blaster" and "The Golden Chords." 143 �During the high school talent show, Danny and the Juniors, played so loud that the principal cut off the microphone. Robert Zimmerman left Hibbing in 1959 to move to Minneapolis so he could enroll at the University of Minnesota. This is where Robert Zimmerman did two things; he fell in love with folk music and changed his name to Bob Dylan. The reason for the change was that Bob was very familiar with the poetry of Dylan Thomas. In a 2004 interview, Bob stated, "You're born, you know, the wrong names, wrong parents. I mean, that happens. You call yourself what you want to call yourself. This is the Figure 11- Bob Dylan, 1963 land of the free." Stop 3- Hibbing Historical Society Museum (47o 25' 25.92'' N/92o 56' 13.36''W) The current focus of the Hibbing Historical Society is the documentation and presentation of the early "Mining Locations” that grew up in the vicinity of the surrounding mines. There were three types of mining locations, but all had the common factor that they placed on company-owned land. There were over 175 locations between 1892 and the 1920's on the Mesabi Range. Initially miners and sometimes their families just found an open piece of land near the mine they worked at where they built a very small home from whatever building materials they could locate. These were "squatter's locations" and often called 'chicken towns' as there were often farm animals in the yards or even in the living quarters themselves, during the winter months. Figure 12- Mahoning Location and Rail Shops As time progressed the companies decided that if they organized a mining location and built modest homes for its miners, they had better success in keeping a stable and highly trained workforce. These were called "company locations" and were actually townsites that were surveyed in and could included paved streets with water and sewer utilities. Mahoning location is a fine local example of a company location in the vicinity of Hibbing. The homes in these Locations were built by the mining firms and then leased or rented to the employees. In some cases the residents were permitted to purchase their house, but not the land. Common monthly rental fees might be $1.00 per $100 invested by the mining company. Very rarely a mining company, such as United States Steel, would design and build “Model Locations". These communities would consist of more elaborate homes with full services plus community buildings such as recreation halls, hospitals and fire departments. Stellar examples of a Model Location would be Morgan Park in Duluth and the community of Coleraine on the western end of the Mesabi Range. The company plan here was for a more attractive community with higher quality constructed homes that would show them in a better light. 144 �More than 30 squatter's locations and company locations were located just to the north, east and west of Hibbing to service a like number of small early mines. In 1915, the town of Hibbing had 20,000 people who all had to uproot their homes and families and move them south to the small village of Alice. Many of the buildings were actually lifted and rolled down to Alice. The Oliver Mining Company (later to become US Steel) was the brainchild behind this move and agreed that if the town relocated 2 miles south to Alice, they would develop the downtown buildings with low interest loans for the retailers. At the Hibbing Historical Society Museum there is a large scale model of the Hibbing with excellent exhibits that describe the physical moving of that "North Hibbing" portion of the town. The move started in 1919 after four years of careful planning and was completed in 1921. The buildings were all moved down what was called at the time, "the First Avenue Highway" which is still in existence today. In total, about 200 structures were moved to the new town while new structures were also built including the Hibbing High School, the Androy Hotel, the Rood Hospital, and the Village Hall. These buildings were created with mining company money to Figure 13- North Hibbing at crest of Mining. help ease the settlers' mood about having to move the entire town. Only one structure did not make it to the new town during the move. A hotel tumbled off the rollers and crashed into a million pieces. One eyewitness referred to it as "an enormous pile of kindling." The city of Alice was then renamed to Hibbing and annexed. The land size of the city of Hibbing is the largest in Minnesota, even surpassing both Minneapolis and St. Paul's city limits! A children's book chronicles the amazing story of the "Town that Moved". On the very short final leg of the tour we'll pass-by the family home of Bob Dylan (Robert Zimmerman). End of tour and return to the Hibbing Park Hotel: Sources:  City of Hibbing website- (http://www.hibbing.mn.us/index.asp?Type=B_LIST&SEC={0BA3C1786F53-4831-B739-0315936323C6})  Hibbing Historical Society  Hibbing High School website- (http://www.hibbing.k12.mn.us/)  "Hibbing Historical Walking Tour" pamphlet - Hibbing Daily Tribune & Roger Saccoman Architecture  Hibbing Chamber of Commerce website- (http://www.hibbing.org/pages/History/)  Hibbing: The Town That Moved website(http://minnesotaghosts.com/index.php/library/mnhistory/81-hibbing-the-town-that-moved)  "The "Locations" - Company Communities on the Minnesota's Iron Ranges",1982, by Arnold R. Alanen (Minnesota Historical Society). 145 �FIELD TRIP C Friday, May 16, 2014 MINNESOTA DISCOVERY CENTER LEADERS: Discovery Center Staff Figure 1. Miners statue near MDC entrance The Minnesota Discovery Center museum and research library in Chisholm (a few miles north of Hibbing) houses artifacts, examines mining methods, explores regional geology, and hosts traveling exhibits that highlight the story of the predominantly European immigrants who migrated to this region at the turn of the 20th century to find work in the burgeoning iron ore industry. Their stories document the development of the Mesabi Iron Range, a region that became the nation’s largest producer of iron ore. The museum, formerly known as “Ironworld,” is perched at the edge of a lakefilled gorge that represents the collective footprint of many open-pit and underground mine properties. This field trip includes a guided tour of the museum, and a trolly ride across a portion of the mine. 146 �FIELD TRIP D Friday, May 16, 2014 COLERAINE MINERALS RESEARCH LABORATORY Natural Resources Research Institute University of Minnesota-Duluth LEADERS: Dick Kiesel (Director CMRL) Dave Hendrickson (Director Strategic Planning) Matt Mlinar (Program Coordinator Mineral Processing) Basak Anameric (Program Coordinator High Temperature Process) This trip will tour the Coleraine Minerals Research Laboratory (CMRL) in Coleraine, about 25 miles SW of Hibbing. The CMRL conducts applied research that supports technology-based economic development for iron ore mining, non-ferrous minerals, industrial minerals, environmental remediation, alternative iron making, and the use of taconite mining products for various value-added aggregate applications. The facility consists of an analytical laboratory, mineral processing and pyrometallurgical processing capabilities from bench to pilot scale for applied research and development projects. Geographic proximity to the nation’s largest iron mining district (the Mesabi Iron Range) has meant that the CMRL has historically conducted minerals development research, and contributed to the training and development of a substantial number of iron mining and minerals industry professionals. Demand varies from solving short term problems, identifying unique market niches, to providing medium to long range technical innovation and developing products and processes for the future. 147 �FIELD TRIP E Friday, May 16, 2014 MINEVIEW FROM A CANOE LEADER: Mark Jirsa (Minnesota Geological Survey); with assistance from Daniel Jordan (Iron Range Resources and Rehabilitation Board), and Dale Cartwright (Minnesota Department of Natural Resources, Division of Lands and Minerals) Figure 1. Airphoto image of the collection of inactive natural (hematite-goethite) ore mines that form what is referred to here as “Ironworld Pit Lake,” just south of Chisholm. Image shows general locations of 2 main geologic features (ovals) that will be viewed from the gunnels, and other local landmarks. Width of photo ~ 2.5 miles. This trip offers a unique duck’s-eye view of the geology at “Ironworld Pit Lake” in Chisholm, a few miles north of Hibbing. The lake occupies the abandoned footprints of as many as 15 separate natural (hematite-goethite) ore mine properties. These include the Pillsbury (operating years1898-1969), Glen (1902-1957), Leonard-Burt (1909-1974), Leonard (1903-1974), Clark (1900-1925), Monroe-Tener (19051981), South Tener (1928-1981), Bruce Annex (1929-1937), Dunwoody (1917-1977), Douglas (19421977), Neville R (1947-78), Duncan (1914-1970), Pillsbury-Brown (1951-1978), Chisholm (1901-1967), and Godfrey (1926-1963) mines. One wall of the pit exposes a 30-foot thick slab of what is inferred to be Cretaceous iron-rich conglomerate that was glaciotectonically dislodged and thrust over till (Fig. 1). In the early days of mining (1892-1950’s), these hematite-pebble conglomerates were prized as extremely high-grade ore. 148 �Another wall portrays fold and fault structures that likely were genetically related to the formation of natural ores by oxidation and leaching of various layers of Biwabik Iron Formation. The central part of mine pit appears to follow major NW-trending fault/fold structures and subparallel joints (Fig.2), and thus obliquely crosses the strike of iron-formation. As a result, a comparatively thick section of strata is exposed along pit walls—perhaps including parts of the Lower slaty, Upper cherty, and Upper slaty members, depending on water level (SEE Field Trip 1, this guidebook for vernacular). Intuitively, most of the natural ore was exhausted from this site; however, exposures of oxidized (near-ore) can be seen locally on pit walls, depending on water level. The various types of natural ore can generally be color-correlated with the inferred protore (via Gruner, 1946). For example, the precursor of “blue ore,” composed of semi-massive martite (magnetite pseudomorphed by hematite), may have been cherty magnetite-rich layers. Yellow-colored ores that contain primarily goethite and limonite (a generic term for undifferentiated, hydrated iron oxides; typically hydrated goethite) formed from layers of thinly bedded to laminated (slaty) iron-silicates. Brown-colored ores consist of mixtures of goethite, limonite, hematite, and martite, and likely were derived from intimately interbedded cherty and slaty layers. The processes of oxidation and localized leaching of silica and iron-carbonate results in considerable volume loss (as much as 50%), and some of the slump structures visible on pit walls are a product of collapse related to this alteration. Some are also undoubtedly related to collapse into historic underground workings. In general, it is difficult, and in some cases impossible, to assign specific episodes of deformation to individual structures (SEE discussion of iron-formation structures in Field Trip 1, this guidebook). Figure 2. Bedrock geologic map of the Chisholm area showing mine pit lakes (pale blue), faults (solid and dashed thick black lines with letters denoting inferred fault movement; Up, Down), fold axes (thick blue lines), and subsurface extent of shallowly southeast-dipping Biwabik Iron Formation (reddish). Brown line in iron-formation approximates the surface trajectory of the Intermediate slate unit, which marks the stratigraphic top of the Lower cherty member. Dashed green lines represent known extent of Cretaceous strata. Width of photo ~ 6 miles (gray section lines). Map is clipped from 1:100,000-scale (Jirsa and others, 2005). Historically, mining geologists and engineers on the Mesabi Iron Range classified natural ore bodies into 3 main types: trough, fissure, and flat-lying (Wolff, 1917). Trough ore bodies are as large 3000 feet long, 1000 feet wide, and 200-400 feet deep. They formed typically along permeable faults or joint sets 149 �by selective leaching and oxidation. Consequent collapse into linear zones of reduced volume produced the synclinal or trough shapes. Fissure ore bodies are similar, but smaller (≤ 200’X2-5’X50’)—having formed along lesser joint structures and typically involving only minor collapse. Flat-lying ore bodies represent oxidation and leaching along select stratigraphic intervals. They are irregular in shape, commonly follow bedding planes outboard of vertical faults or joints, and therefore persist over considerable distances. The Godfrey Mine that lies just south of Ironworld Pit Lake was developed in such an ore body. Ore was mined there from a 20-30 foot-thick silicate horizon that lies at the stratigraphic top of the Lower cherty member, just beneath what’s known as the Intermediate Slate or Paint Rock horizon. The Godfrey Mine’s underground workings extend for nearly a mile, and produced more than 12 million tons of ore. The ores were extracted from this area utilizing first underground mining, followed by open-pit methods. The underground workings were recently digitized by the Lands and Minerals Division of the Department of Natural Resources (DNR) using historic paper records. The resulting 3D imagery reveals an extensive network of underground workings at various depths (Fig. 3). The deepest of the mines shown here was the Monroe-Tener, at ~ 220 feet below surface. Much of the current lake basin was created by subsequent open-pit mining. Figure 3. Airphoto image of the Chisholm area showing extensive underground workings digitized in 3D. Color coding for drifts and shafts differs for each mine, but represents various depths of workings. From Cartwright and others, 2011. Width of photo ~ 1.5 miles. See website http://www.dnr.state.mn.us/lands_minerals/underground/index.html for more details. The work by DNR and associated geologic mapping by the Minnesota Geological Survey (Jirsa and Meyer, 2007; Jirsa and others, 2002, 2005; Jennings and Reynolds, 2005) was undertaken in large part to evaluate connectivity of ground and surface waters between historic and active mines on the Mesabi Iron Range. Obviously, the underground workings have significant influence on water movement—at least in 150 �the upper several hundred feet—and they present engineering challenges for potential mining of taconite in the future. On-going surface subsidence into these historic underground workings (via sink holes) continues to damage local infrastructure. Despite extensive underground operations on parts of the Mesabi Iron Range between 1892 and 1961, the only remaining head frame is that for the Bruce Mine just north of “Ironworld Pit Lake” (Fig. 4). Figure 4. Head-frame from underground mining at the Bruce Mine; the last of its kind on the Mesabi Iron Range. REFERENCES Cartwright, D.F., Oreskovich, J.A., and Oberhelman, M.W. 2011. Central Iron Range Underground Mine Mapping: Minnesota Department of Natural Resources, Division of Lands and Minerals, DVD Disk #1. Gruner, J.W., 1946, Mineralogy and geology of the Mesabi Range: publications of the Office of the Commissioner of the Iron Range Resources and Rehabilitation, 127 p. Jennings, C.E., and Reynolds, W.K., 2005, Surficial geology of the Mesabi Iron Range, Minnesota: Minnesota Geological Survey Miscellaneous Map M-164, scale 1:100,000. Jirsa, M.A., Chandler, V.W., and Lively, R.S., 2005, Bedrock geology of the Mesabi Iron Range, Minnesota: Minnesota Geological Survey Miscellaneous Map M-163, scale 1:100,000 Jirsa, M.A. and Meyer, G.N., 2007, Bedrock and Quaternary geology of the Central Mesabi Iron Range, northeastern Minnesota: Minnesota Geological Survey Open-File Report OFR-07-03. Jirsa, M.A., Setterholm, D.R., Bloomgren, B.A., and Lively, R.S., 2002 , Bedrock topographic and depth to bedrock maps of the western half of the Mesabi Iron Range, northern Minnesota: Minnesota Geological Survey Miscellaneous Map M126, scale 1:100,000. Wolff, J.F., 1915, Ore bodies of the Mesabi range: Engineering and Mining Journal, v. 100, p. 219-224. 151 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2014 Description An account of the resource Institute on Lake Superior Geology. Hibbing, Minnesota. May 14-17, 2014. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2014 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/15972bb6d69bb343e7dfe2bf006624d3.pdf 80e63708d307845e80958b75b276229d PDF Text Text 61st ANNUAL MEETING InstItute on Lake superIor GeoLoGy Dryden, Ontario - May 20-24, 2015 Part 1 – Proceedings and Abstracts �Sponsors The following organizations made generous contributions to the 61st Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. All of the funds contributed this year go toward travel awards for student registrants. For the past 60 years this organization has thrived as a result of the interest of individuals, corporations, universities and government agencies. The dedication to an exchange of scientific ideas and a passion for field trips has enabled the Institute to provide one of its primary objectives – to promote better understanding of the geology of the Lake Superior Region. Mary Arthur Steve Baumann Leonard Espinosa Gordon Medaris Jr. Allan MacTavish Jim Miller Paul Weiblen Canadian Institute of Mining and Metallurgy Thunder Bay Branch �61st annuaL MeetInG InstItute on Lake superIor GeoLoGy Supported by ONTARIO MINISTRY OF NORTHERN DEVELOPMENT AND MINES May 20-24, 2015 Dryden, Ontario HOSTED BY: Rob Cundari & Peter Hinz Co-Chairs Ontario Geological Survey Proceedings - Volume 61 Part 1 – Proceedings and Abstracts Edited by Mark Smyk Cover photos: Top - Max the Moose (courtesy of Peter Hinz), Middle - Pickle Crow, No. 3 Headframe, Pickle Lake, ca.1989 (courtesy John Scott), Bottom - Field Trip, Thierry Mine, Pickle Lake (courtesy Mark Smyk) �61st InstItute on Lake superIor GeoLoGy VoLuMe 61 consIsts of: part 1: proGraM and abstracts part 2: fIeLd trIp GuIdebook trIp 1: The CenTral red lake Gold BelT trIp 2: WesTern WaBiGoon suBprovinCe TranseCT, dryden To MeGGisi lake trIp 3: CanCelled trIp 4: Thunder lake (GoliaTh) projeCT trIp 5: ClassiC ouTCrops of The dryden area trIp 6: GOLD OCCURRENCES OF VAN HORNE TOWNSHIP, VAN HORNE GOLD PROPERTY FLAMBEAU EXPLOSURES trIp 7: unique MineralizinG evenT aT The pidGeon MolyBdenuM deposiT sTripped surfaCe exposure trIp 8: GeoloGy and Mineral deposiTs of The piCkle lake GreensTone BelT trIp 9: The GhosT lake BaTholiTh and relaTed peGMaTiTes trIp 10: MaTTaBi/sTurGeon lake hisToriC vMs CaMp Reference to material in Part 1 should follow the example below: Arts, A., and Fralick, P., 2015.Iron-rich siliceous stromatolites from the upper algal unit of the Gunflint and Biwabik iron formations. In; Smyk, M., (Ed.), Institute on Lake Superior Geology Proceedings, 61st Annual Meeting, Dryden, Ontario, Part 1 - Program and Abstract, v.61, part 1, 7-8. Published by the 61st Institute on Lake Superior Geology and distributed by the ILSG Secretary: Pete Hollings - ILSG Secretary Department of Geology Lakehead University 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Email: peter.hollings@lakeheadu.ca ILSG website: www.lakesuperiorgeology.org ISSN 1042-9964 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Table of Contents Institutes on Lake Superior Geology, 1955-2015 ii Sam Goldich and the Goldich Medal v Goldich Medal Guidelines vii Goldich Medalists and Goldich Medal Committee ix Citation for Goldich Medal Award to Rodney Ikola x Eisenbrey Student Travel Awards xiii Joe Mancuso Student Research Awards xiv Doug Duskin Student Paper Awards and Award Committee xv Board of Directors, Local Committee, and Banquet Speaker xvi Session Chairs and Field Trip Leaders xvii Corporate and Individual Sponsors of Student Travel Scholarships xviii Report of the Chairs of the 60th Annual Meeting xix Program xxii Poster Presentations xxviii Abstracts 1-90 Some figures in this volume were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the volume when it is available on-line at http://www.lakesuperiorgeology.org. i �Proceedings of the 61st ILSG Annual Meeting - Part 1 Institutes on Lake Superior Geology, 1955-2015 # Date Place Chairs 1 1955 Minneapolis, Minnesota C.E. Dutton 2 1956 Houghton, Michigan A.K. Snelgrove 3 1957 East Lansing, Michigan B.T. Sandefur 4 1958 Duluth, Minnesota R.W. Marsden 5 1959 Minneapolis, Minnesota G.M. Schwartz & C. Craddock 6 1960 Madison, Wisconsin E.N. Cameron 7 1961 Port Arthur, Ontario E.G. Pye 8 1962 Houghton, Michigan A.K. Snelgrove 9 1963 Duluth, Minnesota H. Lepp 10 1964 Ishpeming, Michigan A.T. Broderick 11 1965 St. Paul, Minnesota P.K. Sims & R.K. Hogberg 12 1966 Sault Ste. Marie, Michigan R.W. White 13 1967 East Lansing, Michigan W.J. Hinze ii �Proceedings of the 61st ILSG Annual Meeting - Part 1 14 1968 Superior, Wisconsin A.B. Dickas 15 1969 Oshkosh, Wisconsin G.L. LaBerge 16 1970 Thunder Bay, Ontario M.W. Bartley & E. Mercy 17 1971 Duluth, Minnesota D.M. Davidson 18 1972 Houghton, Michigan J. Kalliokoski 19 1973 Madison, Wisconsin M.E. Ostrom 20 1974 Sault Ste. Marie, Ontario P.E. Giblin 21 1975 Marquette, Michigan J.D. Hughes 22 1976 St. Paul, Minnesota M. Walton 23 1977 Thunder Bay, Ontario M.M. Kehlenbeck 24 1978 Milwaukee, Wisconsin G. Mursky 25 1979 Duluth, Minnesota D.M. Davidson 26 1980 Eau Claire, Wisconsin P.E. Myers 27 1981 East Lansing, Michigan W.C. Cambray 28 1982 International Falls, Minnesota D.L. Southwick 29 1983 Houghton, Michigan T.J. Bornhorst 30 1984 Wausau, Wisconsin G.L. LaBerge 31 1985 Kenora, Ontario C.E. Blackburn 32 1986 Wisconsin Rapids, Wisconsin J.K. Greenberg 33 1987 Wawa, Ontario E.D. Frey & R.P. Sage 34 1988 Marquette, Michigan J. S. Klasner 35 1989 Duluth, Minnesota J.C. Green 36 1990 Thunder Bay, Ontario M.M. Kehlenbeck 37 1991 Eau Claire, Wisconsin P.E. Myers 38 1992 Hurley, Wisconsin A.B. Dickas 39 1993 Eveleth, Minnesota D.L. Southwick 40 1994 Houghton, Michigan T.J. Bornhorst 41 1995 Marathon, Ontario M.C. Smyk 42 1996 Cable, Wisconsin L.G. Woodruff 43 1997 Sudbury, Ontario R.P. Sage & W. Meyer 44 1998 Minneapolis, Minnesota J.D. Miller & M.A. Jirsa 45 1999 Marquette, Michigan T.J. Bornhorst & R.S. Regis 46 2000 Thunder Bay, Ontario S.A. Kissin & P. Fralick 47 2001 Madison, Wisconsin M.G. Mudrey & Jr., B.A. Brown iii �Proceedings of the 61st ILSG Annual Meeting - Part 1 48 2002 Kenora, Ontario P. Hinz & R.C. Beard 49 2003 Iron Mountain, Michigan L. Woodruff & W.F. Cannon 50 2004 Duluth, Minnesota S. Hauck & M. Severson 51 2005 Nipigon, Ontario M. Smyk & P. Hollings 52 2006 Sault Ste. Marie, Ontario A. Wilson & R. Sage 53 2007 Lutsen, Minnesota L. Woodruff & J. Miller 54 2008 Marquette, Michigan T. Bornhorst & J. Klasner 55 2009 Ely, Minnesota J. Miller, G. Hudak, & D. Peterson 56 2010 International Falls, Minnesota M. Jirsa, P. Hollings, & T. Boerboom, P. Hinz & M.Smyk 57 2011 Ashland, Wisconsin T. Fitz 58 2012 Thunder Bay, Ontario P. Hollings 59 2013 Houghton, Michigan T. Bornhorst & A. Blaske 60 2014 Hibbing, Minnesota J. Miller & M. Jirsa 61 2015 Dryden, Ontario P. Hinz & R. Cundari iv �Proceedings of the 61st ILSG Annual Meeting - Part 1 Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970’s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 v �Proceedings of the 61st ILSG Annual Meeting - Part 1 INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL vi �Proceedings of the 61st ILSG Annual Meeting - Part 1 Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. vii �Proceedings of the 61st ILSG Annual Meeting - Part 1 Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. viii �Proceedings of the 61st ILSG Annual Meeting - Part 1 Goldich Medalists 1979 Samuel S. Goldich 1997 Ronald P. Sage 1980 not awarded 1998 Zell Peterman 1981 Carl E. Dutton, Jr. 1999 Tsu-Ming Han 1982 Ralph W. Marsden 2000 John C. Green 1983 Burton Boyum 2001 John S. Klasner 1984 Richard W. Ojakangas 2002 Ernest K. Lehmann 1985 Paul K. Sims 2003 Klaus J. Schulz 1986 G.B. Morey 2004 Paul Weiblen 1987 Henry H. Halls 2005 Mark C. Smyk 1988 Walter S. White 2006 Michael G. Mudrey 1989 Jorma Kalliokoski 2007 Joseph Mancuso 1990 Kenneth C. Card 2008 Theodore J. Bornhorst 1991 William Hinze 2009 L. Gordon Medaris, Jr. 1992 William F. Cannon 2010 William D. Addison & Gregory R. Brumpton 1993 Donald W. Davis 2011 Dean M. Rossell 1994 Cedric Iverson 2012 James D. Miller 1995 Gene La Berge 2013 Tom Waggoner 1996 David L. Southwick 2014 Laurel Woodruff 2015 GOLDICH MEDAL RECIPIENT RODNEY J. IKOLA Goldich Medal Committee Serving through the meeting year shown in parentheses. Bernhardt Saini-Eidukat (2015) North Dakota State University Mark Smyk (2016) Ontario Geological Survey Hélène Lukey (2017) Cliffs Natural Resources ix �Proceedings of the 61st ILSG Annual Meeting - Part 1 Citation for the Goldich Medal Award to Rodney J. Ikola Rodney J. (Rod) Ikola was born and raised in the Finnish community of Esko, Minnesota. As he puts it, he had learned his first “foreign language” (ENGLISH) by the end of his first year in school. He has continued to be active in Finnish organizations throughout his life, such as Festival Finlandia at Ironworld in Chisholm, MN, and FinnfestUSA in Duluth in 2008. He has dual citizenship in the U.S. and Finland and served as a member of the Expatriate Parliament of the Republic of Finland for a number of years. Rod has a list of accomplishments in the field of geophysics that would make any Finnish Mother proud, and we suggest that they are worthy of adding Rodney Ikola’s name to the list of Goldich Medal recipients. He did his undergraduate studies at the University of Minnesota Duluth, with a major in geology and a minor in mathematics. He continued his education at the University of Utah on a scholarship in geophysics from the Continental Oil Company, during which time he obtained sufficient math credits to fulfill the requirements for a degree in mathematics. After one year he transferred to the University of Minnesota to complete his Masters degree. During his time at the University of Minnesota, Professor Hal Mooney knew that Rod had been doing gravity work around the southern end of the Duluth Complex and west into Carlton County, simply out of personal curiosity, using a gravimeter made available to him compliments of U. S. Steel. Mooney suggested that Rod should write up the gravity work he had already done and he would accept that as a Master’s thesis. This gravity survey showed several gravity anomalies at the western edge of Carlton County, which became known as the Tamarack Intrusion; this is currently being drilled for copper, nickel and PGE by Rio Tinto-Kennecott. During his days at the U of M, Rod worked on several geophysical projects. In 1959 he conducted his first geophysical survey; a magnetic study of the Barden’s Peak Intrusive of the Duluth Complex. During the summers of 1960 and 1962 he worked on field geophysical exploration for U.S. Steel under the supervision of their geophysicist, George Durfee. The 1960 project consisted of running magnetometer lines across every dip needle anomaly in northwestern Wisconsin (most of which were located in swamps). In 1962 he conducted an extensive gravity survey of Jackson County in central Wisconsin, for the Jackson County Iron Company, a subsidiary of Inland Steel; a taconite mine was developed a decade later in the Archean rocks in Jackson County. In 1965 Director Paul Sims obtained funding to significantly expand the activities of the Minnesota Geological Survey and asked Rod if he would join the Survey to start a systematic gravity survey of the State. Rod accepted the offer and he and G. B. Morey started working for the Survey on the same day. Most of Rod’s time with the Survey was consumed with the gravity survey of the state. This resulted in the publication of several Bouguer gravity maps at a scale of x �Proceedings of the 61st ILSG Annual Meeting - Part 1 1:250,000. During this time he produced the first gravity map of the entire Duluth Complex. He was also periodically on loan to the U.S. Army Topographic Command to establish geodetic control over the central U.S. At the AIME meeting in Duluth in 1970, Fred Chase, chief geologist of the Hanna Mining Co., asked Rod if he would be interested in joining Hanna. They had started a large exploration program in the greenstone belts of Minnesota and needed a geophysicist for the project. Rod accepted the offer, eventually was appointed chief geophysicist for Hanna in 1974, and became responsible for world- wide exploration, a position he held until 1982. Initially his work with Hanna was mainly involved with geophysical exploration in support of Hanna’s geological activities in Minnesota’s greenstone belts and the Duluth Complex. However, he soon began working on all of Hanna’s projects around the world. The list of projects is too long to list here, but demonstrates a wide range of geophysical techniques that Rod mastered. With the downturn of the iron ore market in 1982, Hanna eliminated their entire exploration program as the first step in the eventual demise of the entire company. With this event, Rod decided to go on his own as a consulting geophysicist, which he has been for the last thirty years. Almost immediately he became heavily involved in geophysical consulting in the Lake Superior region for many of the world’s major mining companies. Some of these include Newmont, INCO, American Copper and Nickel, Cominco American, Noranda, Phelps Dodge, DeBeers, (through their regional affiliate), plus numerous junior companies. Most of this work remains proprietary but one project in particular can be mentioned. He did all the geophysical work for Noranda that led to the discovery of the large Lynne (Cu-Zn) Deposit in northern Wisconsin. Only environmental issues prevented the development of the project into a commercial venture. He also became extensively involved in Freeport-McMoRan for many years and acted as de facto geophysicist on many of their worldwide exploration efforts. He worked extensively in the Iberian Pyrite Belt of Spain and Portugal. One of these efforts led to the discovery of the Agua Blanca nickel deposit, which is Europe’s largest nickel producer. He also spent considerable time doing geophysics in the Grasberg area of Indonesia. And he worked on porphyry exploration for Freeport in Baja California. Consulting work has taken Rod to numerous mining camps around the world. He has spent time at Noril’sk in Russia studying their geophysical exploration techniques. He also did some work for a consortium of companies exploring for gold in the “reefs” south of Lake Victoria in Africa, diamond exploration in Brazil, nickel exploration in Western Australia, uranium exploration in the Athabasca area of Canada, and deeply buried porphyry deposits in the southwest U.S.. In recent years, with the upsurge of mineral exploration in the Lake Superior region, he has spent more time close to home. He has been involved with Polymet Mining and Duluth Metals in the Duluth Complex, and Keweenaw Copper Co. and Bitteroot Exploration in Michigan. However, the results of this work remain proprietary. In recent years, Rod has been involved in numerous geophysical projects pertaining to environmental and groundwater problems. On a project with Barr Engineering, he helped develop a groundwater resource at the Sherco Power Plant in Becker, MN. For this work the Minnesota Society of Professional Engineers awarded the group a Distinguished Achievement Award. On another project the group used a unique application of the SP geophysical method to delineate karst features in the Keweenawan sandstones near Askov, MN, to help prevent pollution from their sewage treatment plant. xi �Proceedings of the 61st ILSG Annual Meeting - Part 1 Rod has always been interested in the use of geophysics in archaeological investigations, and he has participated in many studies. He spent two summers in Greece (one on top of Mt. Olympus with the Gods!!) looking at sites from the Homeric Age. On another occasion he used geophysics to look for buried Mayan tombs in Belize. Rod has been affiliated with many professional organizations during his career. He is an emeritus member of the Society of Exploration Geophysicists, belongs to the Australian Society of Exploration Geophysicists, a member of the American Institute of Professional Geologists and has belonged to the Society of Mining Engineers for many years (and served on seven national committees for them). He is also a founding member, and was on the board of directors of the Minnesota Exploration Association (now Mining Minnesota). He is a Registered Earth Scientist in Minnesota and a Professional Geophysicist in California. He has been involved with the Institute on Lake Superior Geology for over fifty years. The first Institute meeting he attended was the fourth one, in Duluth in 1958, and has subsequently attended approximately 45 meetings. During his career as a geophysicist, Rod Ikola has made many contributions to our understanding of the geology of the Lake Superior Region, particularly in the areas of government surveys and industry. These accomplishments, as well as his geophysical studies at so many other places around the world, make him highly qualified for the Goldich Medal. Respectfully submitted, Gene L. LaBerge Richard W. Ojakangas xii �Proceedings of the 61st ILSG Annual Meeting - Part 1 Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xiii �Proceedings of the 61st ILSG Annual Meeting - Part 1 Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive onehalf of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. In 2014, the ILSG Board of Governors awarded a $1000 award from the Student Research Fund to Justin Beermaert. It should be noted that an especially generous donation was once again provided by Ron Seavoy. xiv �Proceedings of the 61st ILSG Annual Meeting - Part 1 Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and long-time friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. Student Paper Awards Committee Amy Radakovich– Minnesota Geological Survey Mark Severson – Teck American Dave Good – Western University xv �Proceedings of the 61st ILSG Annual Meeting - Part 1 Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Robert Cundari (2015-2018) – Ontario Geological Survey Jim Miller (2014-2017) – University of Minnesota Duluth Allan Blaske (2013-2016) – AECOM Peter Hinz (2012-2015) – Ontario Geological Survey Pete Hollings – Secretary (2013-2016) – Lakehead University Mark Jirsa – Treasurer (2014-2017) – Minnesota Geological Survey Local Committee Chairs Peter Hinz – Co-Chair Ring of Fire Secretariat, Ministry of Northern Development and Mines Robert Cundari – Co-Chair Resident Geologist Program, Ontario Geological Survey Volume Editors Mark Smyk – Proceedings Volume Resident Geologist Program, Ontario Geological Survey Allan MacTavish – Field Trip Guidebook Panoramic PGMs (Canada) Limited Banquet Speaker Steve Beneteau (Senior Diamond Advisor / Chief Gemmologist for the Province of Ontario and the Manager of the Diamond Sector Unit, Ministry of Northern Development and Mines) “Ontario’s Diamonds: A Journey from Mine to Market” xvi �Proceedings of the 61st ILSG Annual Meeting - Part 1 Session Chairs Jim Miller – University of Minnesota Duluth Anthony Pace - Ontario Geological Survey Dean Peterson – Peterson Geoscience LLC Mark Smyk - Ontario Geological Survey Ann Wilson – Ontario Geological Survey Laurel Woodruff – United States Geological Survey Field Trip Leaders Field trips have been the mainstay of the ILSG since its inception 60 years ago. We want to give a special thanks to the field trip leaders who volunteered their time and talent in carrying that tradition forward. Pre-Meeting: 1. Red Lake Geology (2-day) Tuesday May 19th and Wednesday May 20th, 2015 Leaders: Andreas Lichtblau (OGS) and Carmen Storey (OGS) 2. Western Wabigoon Subprovince Transect (Dryden to Meggisi Lake) Wednesday May 20th, 2015 Leaders: Mark Puumala (OGS) and Dorothy Campbell (OGS) 3. Mine Reclamation and Legacy Issues at the South Bay Mine Wednesday May 20th, 2015 Leader: Rob Purdon (MNDM) [CANCELLED DUE TO UNFORESEEN CIRCUMSTANCES] 4. Geological Setting of the Thunder Lake Gold Deposit Wednesday May 20th, 2015 Leaders: Treasury Metals Inc. Half-Day: 5. Classic outcrops of the Dryden Area Friday May 22nd, 2015 Leader: Peter Hinz (MNDM) 6. Gold Occurrences of Van Horne Township Friday May 22nd, 2015 Leader: Steve Meade (OGS) 7. Unique mineralizing event at the Pidgeon Molybdenum Occurrence Friday May 22nd, 2015 Leaders: Craig Ravnaas (OGS) Post-Meeting: 8. Historic Pickle Lake Camp (1.5-day) Friday May 22nd and Saturday May 23rd, 2015 xvii �Proceedings of the 61st ILSG Annual Meeting - Part 1 Leaders: Mark Smyk (OGS), Pete Hollings (Lakehead University) and Neil Pettigrew (PC Gold Inc.) 9. Ghost Lake Batholith and Related Pegmatites Saturday May 23rd, 2015 Leader: Shannon Zurevinski (Lakehead University) 10. Mattabi/Sturgeon Lake Historic VMS Camp (2-day) Friday May 22nd and Saturday May 23rd, 2015 Leader: George Hudak (Natural Resources Research Institute – University of Minnesota Duluth) Sponsors The following organizations made generous contributions to the 61st Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. All of the funds contributed this year go toward travel awards for student registrants. For the past 60 years this organization has thrived as a result of the interest of individuals, corporations, universities and government agencies. The dedication to an exchange of scientific ideas and a passion for field trips has enabled the Institute to provide one of its primary objectives – to promote better understanding of the geology of the Lake Superior Region. Mary Arthur Steve Baumann Leonard Espinosa Gordon Medaris Jr. Allan MacTavish Jim Miller Paul Weiblen Canadian Institute of Mining and Metallurgy Thunder Bay Branch xviii �Proceedings of the 61st ILSG Annual Meeting - Part 1 REPORT OF THE CHAIRS OF THE 60TH ANNUAL MEETING INSTITUTE ON LAKE SUPERIOR GEOLOGY HIBBING, MINNESOTA The Precambrian Research Center (PRC) of the University of Minnesota Duluth (UMD) and the Minnesota Geological Survey (MGS) of the University of Minnesota – Twin Cities hosted the 60th Annual Institute on Lake Superior Geology on May 14– 17, 2014 at the Hibbing Park Hotel in Hibbing, Minnesota in the heart of the Mesabi Range. This was the first time that the ILSG has been held in Hibbing. We are pleased to report a very vigorous attendance of 225 registrants, including 57 students. The organizing committee for the meeting was comprised of Jim Miller of UMD-PRC (technical program, meeting and field trip logistics, and registration) and Mark Jirsa of the MGS (field trip coordinator and guidebook editor, student travel awards, and sponsorships). Amy Radakovich of the MGS helped with the procurement and design of T-shirts and beer glasses. Terry Boerboom, also of the MGS, assisted with the organization of the field guidebook. Julie Ann Heinz, an executive office administrator at the UMD Natural Resources Research Institute, provided assistance with meeting registration and creating name tags. During the meeting, a number of UMD students assisted with on-site registration and other institute business. The two-day technical session held at the Hibbing Park Hotel on Thursday and Friday (5/15 and 5/16) included 30 talks and 35 posters presentations, including 17 oral and 20 poster presentations by students. The number of student presenters and attendees are both ILSG records. The meeting opened with a remembrance of Jack Everett and Ernie Lehmann who passed away in August, 2013 and December, 2013, respectively. Both Jack and Ernie were exceptional exploration geologists who devoted much of their careers to minerals exploration in the Lake Superior region. Both were long-time supporters of the ILSG, with Ernie receiving the Goldich medal in 2002. This year’s Goldich medal recipient was Laurel Woodruff of the US Geological Survey. Laurel was recognized for her long and productive 30 year career with the USGS’s mineral resources research program, mostly in the Lake Superior region and for her scientific contributions and service to the ILSG, especially her serving as meeting chair in 1996 (Cable, WI), 2003 (Iron Mtn, MI), and 2007 (Lutsen, MN). Laurel was presented the medal at the annual banquet by Bill Cannon, her colleague and mentor at the USGS who received the Goldich medal in 1992. The evening banquet talk was presented by Dr. Francis M. Carroll of the University of Manitoba - Winnipeg and St. Johns University. The title of his talk was: "A Line in the Trees: History of the US-Canadian Boundary from Lake Superior to Lake of the Woods". The meeting offered six full-day and three Friday afternoon field trips that highlighted various aspects of the geology, ore deposits, and culture of the central Mesabi Range. Most trips were filled to capacity with a cumulative total of 255 field trip attendees. Three pre-meeting field trips run on Wednesday, May 14 included: 1) Stratigraphy, Sedimentology, Structure, and Mineralization of the Biwabik Iron Formation, Central Mesabi Iron Range led by Phil Larson (Duluth Metals), Marsha Patelke (UMD-NRRI), Jakob Wartman (Cliffs NR), Michael Totenhagen (Arcelor xix �Proceedings of the 61st ILSG Annual Meeting - Part 1 Mittal), Mark Jirsa (MGS), Steven Losh (Minnesota State University-Mankato), and Peter Jongewaard (Cliffs NR); 2) A Walk in the Park – Neoarchean Geology of Lake Vermilion State Park led by George Hudak (UMD-NRRI), Amy Radakovich (MGS), Geoff Pignotta (UW - Eau Claire), and Kelly Schwierske (UW - Eau Claire); and 3) Western Mesabi Range Mining Operations led by Doug Halverson (Cliffs NR), Dan Cervin (Cliffs NR), William Everett (Essar Steel), Kevin Kangas (Essar Steel), and Joey Nielsen (Magnetation). Three Friday afternoon trips included: 1) State Drill Core Library – Hibbing, Minnesota led by Dave Dahl, Barry Frey, and other MNDNR staff and Dean Rossell (Kennecott Exploration, Rio Tinto); 2) Hibbing’s Iron Mining and Cultural History led by Henry Djerlev, Bob Kearney, Erica Larson and other Hibbing Historical Society staff; and 3) Mineview from a Canoe led by Mark Jirsa (MGS), Dan Jordan (IRRRB), and Dale Cartwright (MNDNR). Three post-meeting trips were run on Saturday, May 17 and included: 1) Visions of Maturi: The Geology of the South Kawishiwi Intrusion led by Dean Peterson (Duluth Metals Ltd.); 2) The St. Louis Sublobe and Glacial Lake Upham led by Phil Larson (Duluth Metals Ltd.), Alan Knaeble (MGS), Howard Mooers (UMD), and Lisa Marlo (Halcon Resources Corp.); and 3) Geology and Gold Mineralization of the Virginia Horn Area led by Mark Jirsa (MGS), Bill Rowell (Vermillion Gold), Rick Sandri - (Vermillion Gold), and Jason Richter (MN DOT). The student paper committee comprised of Andrew Ware (PolyMet Mining), Prajukti Bhattacharyya (University of Wisconsin-Whitewater), and Rob Cundari (Ontario Geological Survey had the onerous task of judging 17 oral and 20 poster presentations by students. The committee awarded 4 Doug Duskin Student Paper Awards with a cash prize of $500 each to Amanda Van Lankfelt (U Mass.), Adrian Arts (Lakehead), Monica Karman (Lakehead), and Darcy Jacobson (Michigan Tech.). To defray student’s expenses for travel and registration, a total of $6400 was distributed to 32 students representing 8 different schools. This generous aid was provided by the Eisenbrey Student Travel Awards. Additional student travel support was provided by funds contributed by 11 meeting registrants (Mary Arthur, Jack Berkley, Karl Everett, John Green, George Hudak, Peter Jongewaard, Steven Losh, Al MacTavish, Gordon Medaris, Michael Mudrey, and Jill Peterman),and several corporations and organizations (Eagle Mine, Teck American, Midwest Institute of Geosciences and Engineering, and GEI Consultants). Ron Seavoy provided a particularly generous donation to establish and maintain the Joe Mancuso Student Research Grant program. Three $500 Mancuso research grants were awarded in 2014 to Michael Doyle (UMD), Michael Fedorchuk (UW-Milwaukee), and Sarah Sauer (UMD). The Institute’s Board of Directors met on May 15, 2014. The meeting was attended by meeting co-chair Jim Miller, Treasurer Mark Jirsa, Secretary Peter Hollings, and board members Allan Blaske (2013 chair) and Al MacTavish (2012 chair). Incoming chair for the 2015 ILSG, Pete Hinz, also attended. Secretary Hollings took the minutes of the Board meeting that are as follows: 1. Accepted report of the Chairs for the 59th ILSG, Houghton, Michigan; as printed in the Proceeding Volume (Blaske), and minutes of last Board meeting, May 9, 2013 (Hollings) xx �Proceedings of the 61st ILSG Annual Meeting - Part 1 2. 3. 4. 5. Received, discussed, and accepted 2013-2014 ILSG Financial Summary (Jirsa). Received, discussed, and accepted 2013-2014 report of the Secretary (Hollings). Approved Jim Miller as on-going ILSG Board member Approved Dryden as the site for the 61st annual ILSG meeting. The meeting will be hosted by the Ontario Geological Survey with Peter Hinz and Rob Cundari serving as co-chairs. 6. Discussed and approved renewal of Mark Jirsa as Institute Secretary (end of term 2017). This was later approved by a vote of the membership 7. Discussed and approved replacing Graham Wilson as the “member from industry” on Goldich Committee (end of term 2017) with Helene Lukey 8. Discussed student attendance and presentations at future meetings. We would like to thank the participants, especially the students, for supporting the Institute by their attendance and enthusiasm, the field trip leaders for their hard work, the presenters for their high quality and informative talks and posters, the session chairs and subcommittee members for their important contributions, and the meeting sponsors for their generosity in helping students participate in the Institute. Respectfully submitted, Jim Miller and Mark Jirsa Co-Chairs, 60th Institute on Lake Superior Geology xxi �Proceedings of the 61st ILSG Annual Meeting - Part 1 PROGRAM SCHEDULED ACTIVITIES AND FIELD TRIPS Tuesday, May 19th, 2015 8:00 a.m. – 6:00 p.m. Field Trip 1 Red Lake Geology (Day One) Wednesday, May 20th, 2015 8:00 a.m. – 6:00 p.m. Field Trip 1 Field Trip 2 Field Trip 3 Field Trip 4 7:00 p.m. – 10:00 p.m. Red Lake Geology (Day Two) Western Wabigoon Subprovince Transect Mine Reclamation and Legacy Issues at the South Bay Mine [cancelled] Geological Setting of the Thunder Lake Gold Deposit / Gold occurrences of Van Horne Township Registration at Best Western Poster Session (Gemini Room) / Ice Breaker Social (Centennial Room) Thursday, May 21st, 2015 8:00 a.m. – 12:00 p.m. Registration continues 9:00 a.m. – 12:00 p.m. Technical Session I 12:00 p.m. – 1:30 p.m. Lunch (provided) 1:30 p.m. – 4:00 p.m. Technical Session II 6:00 p.m. – 7:00 p.m. Mixer / Cash Bar 7:00 p.m. – 10:00 p.m. Annual Banquet, Keynote Speaker and Awards Presentation (Sunset Ballroom) Friday, May 22nd, 2015 9:00 a.m. – 11:10 a.m. 11:10 a.m. – 12:00 p.m. 12:00 p.m. – 1:30 p.m. 1:30 p.m. – 6:00 p.m. 6:00 p.m. – 7:00 p.m. Saturday, May 23rd, 2015 8:00 a.m. – 6:00 p.m. Sunday, May 24th, 2015 8:00 a.m. – 6:00 p.m. Technical Session III Student Awards presentations Lunch (provided) Field Trip 5 Classic outcrops of the Dryden Area Field Trip 6 Gold Occurrences of Van Horne Township Field Trip 7 Unique mineralizing event at the Pidgeon Molybdenum Occurrence Field Trip 8 (departs for Pickle Lake) Field Trip 10 (departs for Ignace) Field Trip 8 Field Trip 9 Historic Pickle Lake Camp (return to Ignace / Dryden) Ghost Lake Batholith and Related Pegmatites (return to Dryden) Field Trip 10 Mattabi / Sturgeon Lake Historic VMS Camp (Day One; return to Ignace) Field Trip 10 Mattabi / Sturgeon Lake Historic VMS Camp (Day Two; return to Dryden) TUESDAY, MAY 19TH 8:00 am – 6:00 pm Pre-Meeting Field Trip: 1. Red Lake Geology (Day One; overnight in Red Lake) Leaders: Andreas Lichtblau (OGS) and Carmen Storey (OGS) xxii �Proceedings of the 61st ILSG Annual Meeting - Part 1 WEDNESDAY, MAY 20TH 8:00 am – 6:00 pm Pre-Meeting Field Trips: 1. Red Lake Geology (Day Two; return to Dryden) Leaders: Andreas Lichtblau (OGS) and Carmen Storey (OGS) 2. Western Wabigoon Subprovince Transect Leaders: Mark Puumala (OGS) and Dorothy Campbell (OGS) 3. Mine Reclamation and Legacy Issues at the South Bay Mine Leader: Rob Purdon (MNDM)[cancelled due to unforeseen circumstances] 4. Geological Setting of the Thunder Lake Gold Deposit Leaders: Treasury Metals Inc. 7:00 p.m. – 11:00 p.m. Registration at Best Western Poster Session (Gemini Room) / Ice Breaker Social (Centennial Room) THURSDAY, MAY 21ST 8:00 – 12:00 pm Registration continues 8:50 - 9:00 am OPENING REMARKS, UPDATES Peter Hinz and Robert Cundari, Co-Chairs, 2015 ILSG 9:00 – 9:10 am Welcoming Remarks TECHNICAL SESSION I (*denotes a student eligible for Best Student Paper Award) Session Chairs: Laurel Woodruff – United States Geological Survey Anthony Pace - Ontario Geological Survey 9:10 – 9:50 am Bill Cannon, Bill Addison, Greg Brumpton and Mark Jirsa The Sudbury Impact Event in the Lake Superior region: Ten years of research on ten minutes of geologic time 9:50 – 10:10 am Daniel Lafontaine* and Mary Louise Hill Structural control on the Borden Gold deposit, Chapleau, ON 10:10 – 10:40 am COFFEE BREAK AND POSTER VIEWING xxiii �Proceedings of the 61st ILSG Annual Meeting - Part 1 10:40 – 11:00 am Jim Miller Role of Felsic and Feldspathic Rocks in Triggering Subvolcanic Emplacement of Mafic Intrusions: Evidence from the Midcontinent Rift in Northeastern Minnesota 11:00 – 11:20 am David Good, Peter Hollings, Robert Cundari and Doreen Ames Significance of LREE-enriched mantle source to genesis of basalt in the Coldwell Alkaline Complex, Midcontinent Rift, Ontario 11:20 – 11:40 am Michael Doyle* and Jim Miller Geologic and geochemical attributes of the Beaver River Diabase and Greenstone Flow: Testing a possible intrusive-volcanic correlation in the 1.1 Ga Midcontinent Rift 11:40– 12:00 pm Sarah Sauer* and Jim Miller Petrologic study of the "Chill" zone of the Layered Series at Duluth: Testing a possible plutonic-volcanic correlation within the Midcontinent Rift 12:00 – 1:30 pm LUNCH (provided) / MEETING OF THE BOARD OF DIRECTORS TECHNICAL SESSION II (*denotes a student eligible for Best Student Paper Award) Session Chairs: Dean Peterson – Peterson Geoscience LLC Ann Wilson – Ontario Geological Survey 1:30 – 1:50 pm Robert Mahin The Eagle Mine in Production: U.S.A.’s Only Primary Nickel Producer 1:50 – 2:10 pm Kristofer Asp*, Christian Schardt, and Lev Spivak-Birndorf Evidence of high temperature Ni isotopic fractionation during the formation of Cu-Ni-PGE sulfide deposits in the Duluth Complex xxiv �Proceedings of the 61st ILSG Annual Meeting - Part 1 2:10 – 2:30 pm David Good, Louis Cabri and Doreen Ames Comparison of PGM assemblages for the Marathon, Geordie Lake and Area 41 deposits, Coldwell Alkaline Complex, Ontario 2:30 – 2:50 pm Jeffrey Mauk, Poul Emsbo and Peter Theodorakos Evaporated seawater formed sediment-hosted stratiform copper orebodies and second-stage copper mineralization in the Mesoproterozoic Nonesuch Formation of the Midcontinent Rift 2:50 – 3:20 pm COFFEE BREAK AND POSTER VIEWING 3:20 – 3:40 pm Amos Albert*, Jessica Eagle-Bluestone* and Bernie Saini-Eidukat Chemistry and Mineralogy of Nopeming metasiltstone at the Grandview Site, Duluth, Minnesota 3:40 – 4:00 pm Adrian Arts* and Phil Fralick Iron-rich siliceous stromatolites from the upper algal unit of the Gunflint and Biwabik Iron Formations 4:00 – 4:20 pm Christopher Yip* and Phil Fralick Exposure Surfaces of the Gunflint Iron Formation, Northwestern Ontario 4:20 – 4:40 pm Riku Metsaranta and Phil Fralick Sedimentology and Geochemistry of a 1.4 Ga Continental Playa System, the Lower Sibley Group, Northwestern Ontario: Implications for the Mesoproterozoic Hydrosphere and Atmosphere 4:40 –5:00 pm Paul Fix* and Tamara Diedrich Characterization of secondary minerals formed on weathered Duluth Complex Cu-Ni-PGE deposit rock and implications for controls on metal mobility 6:00 pm RECEPTION – CASH BAR 7:00 pm ANNUAL BANQUET (Sunset Ballroom) − Announcement of 62nd Annual Meeting Location xxv �Proceedings of the 61st ILSG Annual Meeting - Part 1 − 2015 Goldich Award Presentation to Rodney Ikola − Banquet Presentation by Steve Beneteau (MNDM) “Ontario’s Diamonds: A Journey from Mine to Market” FRIDAY, MAY 22ND 8:50 – 9:00 am OPENING REMARKS, UPDATES Peter Hinz and Rob Cundari, Co-Chairs, 2015 ILSG TECHNICAL SESSION III Session Chairs: Mark Smyk - Ontario Geological Survey Jim Miller – University of Minnesota Duluth 9:00 – 9:20 am Brent Trevisan, Pete Hollings, Doreen Ames and Nicole Rayner The petrology, mineralization and regional context of the Thunder mafic to ultramafic intrusion, Midcontinent Rift, Thunder Bay, Ontario 9:20 – 9:40 am Seamus Magnus Geology and geochemistry of the Lang Lake greenstone belt, Uchi Domain, Superior Province 9:40 – 10:00 am Phil Fralick Lateral Geochemical Gradients and Physical Processes Associated with the Genesis of Iron Formations: Examples from the Paleoproterozoic to Mesoarchean of Superior Province 10:00 – 10:30 am COFFEE BREAK AND POSTER VIEWING 10:30 – 10:50 am Steve Kissin Rainy River, northwestern Ontario's first meteorite 10:50 – 11:10 am Dennis Smyk, William Ross and Mark Smyk Images on stone: Pictographs of the Ignace area, northwestern Ontario xxvi �Proceedings of the 61st ILSG Annual Meeting - Part 1 11:10 – 11:30 am Dean Peterson So, an Environmental Impact Statement is required: Some Lake Superior area Geologic Parameters for Geologists, Consultants, Companies, and Regulators 11:30 – 12:00 pm BEST STUDENT PAPER AWARDS AND STUDENT TRAVEL AWARDS 12:00 – 1:30 pm LUNCH (provided) 1:30 – 6:00 pm FRIDAY AFTERNOON FIELD TRIPS 5. Classic outcrops of the Dryden Area Leader: Peter Hinz (MNDM) 6. Gold Occurrences of Van Horne Township Leader: Steve Meade (OGS) 7. Unique mineralizing event at the Pidgeon Molybdenum Occurrence Leader: Craig Ravnaas (OGS) 8. Historic Pickle Lake Camp (departs Dryden for Pickle Lake) Leaders: Mark Smyk (OGS), Peter Hollings (Lakehead University) and Neil Pettigrew (PC Gold Inc.) 6:00 – 7:00 pm 10. Mattabi / Sturgeon Lake Historic VMS Camp (departs Dryden for Ignace) Leader: George Hudak (NRRI – University of Minnesota Duluth) SATURDAY, MAY 23RD 8:00 am – 6:00 pm POST-MEETING FIELD TRIPS 8. Historic Pickle Lake Camp (returns to Ignace / Dryden in evening) 9. Ghost Lake Batholith and Related Pegmatites Leader: Shannon Zurevinski (Lakehead University) 10. Mattabi/Sturgeon Lake Historic VMS Camp (Day One) SUNDAY, MAY 24TH 8:00 am – 6:00 pm POST-MEETING FIELD TRIPS 10. Mattabi/Sturgeon Lake Historic VMS Camp (Day Two; returns to Ignace / Dryden in evening) xxvii �Proceedings of the 61st ILSG Annual Meeting - Part 1 POSTER PRESENTATIONS (*denotes a student eligible for Best Student Paper Award) Eric Anderson, Ashley Quigley, Patrick Quigley and Thomas Monecke Geophysical imaging of the bedrock geology of the Pembine-Wausau terrane, Wisconsin: Constraints on the setting of volcanogenic massive sulfide deposits Eric Anderson, V. Grauch, Michael Powers and Bill Cannon Seismic, gravimetric, and magnetic modeling over the Bayfield Peninsula, Wisconsin: Testing hypotheses on the source of a gravity low Jordan Baird* and Mary Louise Hill Fold analyses in the Gunflint Formation: working towards a characterization of regional deformation in the Animikie Group near Thunder Bay, Ontario Steven Baumann, Alexandra Cory and Sandra Dylka Interpretation of the St. Amour Deep Stratigraphic Test Well, Alger County, Michigan Greg Brumpton and Steve Kissin Large hypervelocity impacts on Earth: Empirical observations and validation of computational model predictions for Sudbury and Chicxulub Tom Buchholz, Alexander Falster and W. B. Simmons Tainiolite from the Stettin Intrusion, Wausau Complex, Marathon County, Wisconsin Benjamin Drenth, Chad Ailes and Eric Anderson Re-digitized public aeromagnetic data for the Baraga basin and surrounding region, Upper Peninsula, Michigan Espree Essig*, George Hudak, Geoff Pignotta and Robert Lodge Petrographic Analysis of Felsic Tuffs within the Neoarchean Soudan Member of the Ely Greenstone Formation, Northeastern Minnesota V.J.S. Grauch, Michael Powers, Eric Anderson and Bill Cannon Preliminary 3D model of the Midcontinent Rift System in western Lake Superior region Steve Hauck, John Heine, Mark Severson, Sara Post, Sarah Chlebecek, Stephen Monson Geerts, Julie Oreskovich, Sarah Gordee and George Hudak Geological and Geochemical Reconnaissance for Rare Earth Element (REE) Mineralization in Minnesota Jonathan Haynes*, Joyashish Thakurta and Tom Quigley Petrological and geochemical evaluation of the Sturgeon Falls Igneous Body and its relationship with the Penokean Orogenic Belt Benjamin Hinks*, Joyashish Thakurta and Bob Mahin xxviii �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geochemical and petrological studies on the origin of Ni-Cu sulfide mineralization at the Eagle Intrusion in Marquette County, Michigan George Hudak, Stephen Monson Geerts, Larry Zanko, Sara Post and Bryan Bandli The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter - 2015 Update Mark Jirsa, Terry Boerboom, Val Chandler and Mark Schmitz Geology and geochronology of Archean rocks in the International Falls and Littlefork 30X60’ quadrangles, north-central Minnesota Steve Kissin and Greg Brumpton Studies on PDFs in shocked quartz from distal Sudbury ejecta in the Thunder Bay area compared with Chicxulub Benjamin Krogmeier, Dylan McKevitt, Elizabeth Roepke, Michael Sara, Paul Szkilnyk and Mark Jirsa Geologic mapping of Neoarchean and Proterozoic rocks near Knife Lake, Northeastern Minnesota, by students of the Precambrian Research Center’s 2014 field camp Nathan Lentsch* and Jim Miller Incorporation of Duluth Complex maps into GIS platform Jim Miller, Christopher Beaver, Timothy Hahn, Nikolas Miller, Joseph Puliese and Erick Wright Geology of the North and South Temperance Lakes Area of the Boundary Waters Canoe Area, Cook County, Minnesota - 2014 Precambrian Field Camp Capstone Mapping Doug Nikkila* and Shannon Zurevinski The mineralogy and petrology of a newly discovered REE occurrence within the Coldwell Complex near Marathon, Ontario Sean O’Brien*, Pete Hollings and Jim Miller Petrology, geochemistry and mineral chemistry of the Crystal Lake and Mount Mollie mafic intrusions, Northwestern Ontario Mark Puumala, Rob Cundari, Dorothy Campbell, Desmond Rainsford and Riku Metsaranta New airborne geophysical data for the Lake Superior Region of northwestern Ontario: A new tool for the identification of Neoarchean to Mesoproterozoic structures and associated maficultramafic intrusions Patrick Quigley* and Thomas Monecke Spectrum of Volcanogenic Massive Sulfide Deposits in the Penokean Volcanic Belt, Great Lakes Region, USA Andrew Sasso* and Joyashish Thakurta xxix �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geochemical and Petrologic Characterizations of Peridotite, Marquette County, Michigan Laurel Woodruff and Carrie Jennings Bedrock and Soil Chemistry in Paired Watersheds in Northeastern Minnesota xxx �Proceedings of the 61st ILSG Annual Meeting - Part 1 Chemistry and Mineralogy of Nopeming metasiltstone at the Grandview Site, Duluth, Minnesota ALBERT, Amos, EAGLE-BLUESTONE, Jessica, and SAINI-EIDUKAT, Bernhardt Department of Geosciences, North Dakota State University, Fargo, ND 58102 USA The series of outcrops in the Grandview area of Duluth, Minnesota, is considered a classic geologic exposure (Mattis, 1972; Jirsa and Morey, 1987). Here, we examined the contact between the Middle Proterozoic Nopeming Sandstone and the lower magnetically reversed Ely’s Peak Basalts of the North Shore Volcanic Group. We investigated whether prominent light/dark banding in the uppermost siltstone portion of the Nopeming formation resulted from sedimentary deposition or metamorphism, and whether any chemical and mineralogical differences exist. The sample was taken from the metasiltstone layer directly beneath the basalt (Fig. 1). After carrying out petrographic analysis, polished samples were examined by SEM-EDS (NDSU). In preparation for XRD (NDSU) and ICP analysis (Activation Laboratories), we crushed the sample and hand-sorted the light and dark grains. Grain size varies from fine silt in the light bands to clay size in the dark colored bands. Although there are areas of disrupted banding where light and dark materials are intimately mixed, the bands counted (n= 30 total) have average widths of 2.8 mm and 2.25 mm respectively (Fig. 2). Light bands contain coarser grains compared to dark bands. Fining from light to dark bands represents a graded bedded sequence of sedimentation. A few thin bands, less than ~40 µm in width, of opaque grains with reaction rims surrounding each grain were also observed. XRD and petrography show the light bands are richer in quartz and albite, plus some augite. Dark bands consist mainly of quartz, albite, hornblende, and actinolite. SEM-EDS indicates a reaction rim assemblage of ~ 25 µm grains consisting of ilmenite cores with titanite rims (Fig. 3), in a matrix of actinolite and albite. Zircon, diopside, potassium feldspar and quartz were also observed. Whole rock chemistry shows both light and dark bands are silica rich, but the dark bands contain less silica (61.59 wt. %) than the light bands (69.22 wt.%) (Table 1). Higher amounts of Al2O3, Fe2O3, MnO, MgO, CaO, K2O and TiO2 were found in the dark bands. Dark bands also contain more total REE (147 ppm) vs. light (131 ppm), consistent with the concept of higher original clay content in the dark bands. Both light and dark bands show marked CN LREE enrichment with a small negative Eu anomaly and flat HREEs. The overall pattern is similar to average sedimentary and crustal REE patterns (McLennan, 1989) (Fig. 4). On a PAAS normalized diagram, both light and dark bands plot near unity, although the LREES have ratios slightly <1 while the HREE's are slightly >1 (Fig. 4). It is unclear why a disproportional amount of barium (1467 ppm) was found in the dark bands, vs. 273 ppm in the light ones. Ilmenite-titanite reaction rims appear as sharp boundaries between the two minerals. Ilmenite has high FeTiO3 content suggesting that it crystallized in conditions of higher T and/or lower fO2. Titanite often occurs as product of late stage oxidation. Textural evidence suggests that titanite + ferroactinolite assemblage is due to a hydration reaction, following such as below from Harlov et al. (2006) (Fig. 5, Reaction 4). 6 Hedenbergite + 3 Ilmenite + 5 Quartz + 2 H2O = 2 Fe-actinolite + 3 Titanite We conclude the banding resulted from a process producing fining upward graded bedding, perhaps by small turbidities in a shallow aqueous environment. The presence of hornblende and actinolite indicates the emplacement of the Ely’s Peak Basalt altered the original mineralogy to a 1 �Proceedings of the 61st ILSG Annual Meeting - Part 1 metamorphic grade between hornblende-hornfels and amphibolite facies. This event also produced the observed ilmenite-titanite assemblage. Fig. 1. Metasiltstone in outcrop. UTM 15T 555,616E/5,174,699N Fig. 4. Chondrite and PAAS normalized REE spiderplots for the light and dark bands. Fig. 2. Cut sample showing banding Fig. 5. Schematic phase relationships in the system CaOFe2O3-TiO2-SiO2-H2O-O2 involving titanite (Ttn), ilmenite (Ilm), magnetite (Mt), hedenbergite (Hed), Fe-actinolite (Act), and quartz (Qtz) as a function of logfO2 and logfH2O at constant temperature and pressure (Harlov et al., 2006) Fig. 3. SEM image of ilmenite w titanite rim. 1: Ilmenite; 2: Titanite; 3,4: Actinolite; 5: Albite wt. % Light Dark SiO2 69.22 61.59 Al2O3 7.98 8.94 Fe2O3t 3.34 5.3 MnO 0.179 0.203 MgO 6.02 9.28 CaO 9.37 9.74 Na2O 1.95 2.03 K 2O 0.98 1.3 TiO2 0.678 0.834 P 2O 5 0.09 0.1 LOI n.a. 0.89 Total 99.81 100.21 Table 1. Whole rock major element chemistry for the light and dark bands of the metasiltstone. n.a.: not analyzed Funding from the Three Affiliated Tribes of North Dakota to J.E.-B.is gratefully acknowledged. REFERENCES Harlov, D., Tropper, P., Seifert, W., Nijland, T., Förster, H.-J., 2006. Formation of Al-rich titanite (CaTiSiO4OCaAlSiO4OH) reaction rims on ilmenite in metamorphic rocks as a function of fH2O and fO2, Lithos 88, 72– 84. Jirsa, M.A. and Morey, 1987, Jay Cooke State Park and Grandview areas: evidence for a major early Proterozoic - middle Proterozoic unconformity in Minnesota, in Biggs, D.L., ed., Centennial Field Guide 3: Boulder, CO, Geological Society of America, p. 67-72. Mattis, A.F., 1972, The petrology and sedimentation of the basal Keweenawan sandstones of the north and south shores of Lake Superior. Unpubl. M.S. Thesis, Univ. Minn. Duluth, 123 p. McLennan, S.M., 1989. Rare earth elements in sedimentary rocks; influence of provenance and sedimentary processes. Reviews in Mineralogy and Geochemistry, v. 21, p. 169-200. 2 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Seismic, gravimetric, and magnetic modeling over the Bayfield Peninsula, Wisconsin: Testing hypotheses on the source of a gravity low ANDERSON, Eric D.1, GRAUCH, V.J.S.1, POWERS, Michael H.1, and CANNON, William F.2 1 US Geological Survey, MS 964, PO Box 25046, Denver, CO 80225 USA 2 US Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA 20192 USA A prominent gravity low lies over the Bayfield Peninsula in northern Wisconsin (Figure 1). The mapped bedrock geology includes sedimentary rocks of the Oronto and Bayfield Groups that overlie Midcontinent rift-related volcanic and intrusive rocks. The nearly 100 mGal amplitude anomaly has been interpreted to reflect low density Archean granite that is surrounded by higher density basalt (White, 1966; Allen and others, 1997). Two-dimensional (2D) gravity and magnetic models, constrained by limited seismic reflection data, are being developed to test possible sources for the gravity anomaly. Seismic reflection data acquired in 1984 were licensed from Seismic Exchange International for portions of several lines on the Bayfield Peninsula. Plots of two-way travel times across the western gradient of the gravity low (Figure 1) show relatively flat and continuous reflections within the Bayfield and Oronto Group rocks. Calculated depths from two-way travels times for the base of the sedimentary rocks vary slightly from 2.7 km in the west to 3.6 km in the east, indicating that the rocks have an apparent dip to the east. On the western side of the seismic profile are reflections that dip moderately to the west which can be observed to estimated depths of 7.6 km. These westerly dipping reflections pinch-out at around 3.5 km depth near the eastcentral part of the profile. The contrasting dip direction indicates that the source rocks are in unconformable contact with the overlying, gently dipping reflections attributed to the sedimentary rocks. This angular unconformity has been interpreted to represent the contact between the Oronto Group and the Midcontinent rift-related volcanic rocks (Allen and others, 1997). Seismic reflections are not apparent beneath the volcanic rocks where the geology is inferred to be Archean granite. Publically available gravity and magnetic data map contrasting physical properties that are related to changes in subsurface geology. The gravity data with stations spaced approximately 2 km on-shore and 5 km off-shore indicate that low-density crustal material underlies the Bayfield Peninsula. Magnetic anomaly data show that a moderate amplitude, long wavelength anomaly high occurs over much of the gravity low which likely reflects a relatively deep magnetic source. Forward models of gravity and magnetic data along a 160 km east-west profile that spans the Bouguer gravity low (Figure 1; line A-A`) were constructed using reported physical property values (Chandler and Lively, 2011). The results confirm that the gravity anomaly can be explained by a ridge of low density material, possibly Archean granite, flanked by west-dipping high density rocks, both of which are overlain by low density sedimentary rocks. Magnetic depth estimates and model sensitivity to changes in source magnetization at depths ranging 3 to 12 km suggest that the west-dipping rocks, constrained by the seismic data, are magnetic. However, reported basalt magnetizations did not produce acceptable model response, indicating that true magnetizations are much lower and possibly indistinguishable from the adjacent Archean rocks. These results suggest basalt is present, but may have reversed-polarity or reduced magnetization compared to elsewhere. 3 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: Generalized geology map of western Lake Superior showing location of prominent gravity low over the Bayfield Peninsula. Seismic lines provide detailed subsurface imaging that helps constrain 2D gravity and magnetic models to test possible sources for the gravity low. Forward models of gravity and magnetic data were constructed along line A-A`. REFERENCES Allen, D.J., Hinze, W.J., Dickas, A.B., and Mudrey, M.G., 1997. Integrated geophysical modeling of the North American Midcontinent rift system: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota. In Ojakangas, R.W., Dickas, A.B., and Green, J.C. (Eds.), Middle Proterozoic to Cambrian Rifting, central North America: Geological Society of America Special Paper 312: 47-72. Chandler, V.W., and Lively, R.S., 2011. Compilation of Minnesota and western Wisconsin geoscience for the USGS National Geologic Carbon Dioxide Sequestration Assessment: Enhanced geophysical model for extent and thickness of deep sedimentary rocks. Minnesota Geological Survey Open-File Report 2011-03: 37 pages. White, W.S., 1966. Tectonics of the Keweenawan basin, western Lake Superior region. U.S. Geological Survey Professional Paper 524-E: E1-E23. 4 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geophysical imaging of the bedrock geology of the Pembine-Wausau terrane, Wisconsin: Constraints on the setting of volcanogenic massive sulfide deposits ANDERSON, Eric1, QUIGLEY, Ashley2, QUIGLEY, Patrick2, and MONECKE, Thomas2 1 US Geological Survey, MS 964, PO Box 25046, Denver, CO 80225 USA 2 Department of Geology and Geological Engineering, Colorado School of Mines, 1516 Illinois St., Golden, CO 80401 The Pembine-Wausau terrane in Wisconsin (Figure 1) represents a major Paleoproterozoic belt of metavolcanic rocks that have formed in an island-arc setting at the southern limit of the Superior Craton (Schulz and Cannon, 2007). The terrane is known to host a number of significant volcanogenic massive sulfide deposits, including the world-class Crandon deposit which comprises a total resource of 65.8 million metric tons of massive sulfides. Despite its potential future economic significance, little is known about the bedrock geology of this terrane. A thick cover of glacial deposits makes field observations difficult. Existing geologic reconnaissance map compilations indicate that many of the known VMS prospects occur within a succession dominated by bimodal metavolcanic rocks (Nicholson and others, 2004). Existing regional-scale potential field data provide a continuous set of observations across the entire terrane. This study reinterprets these data and applies filters to highlight changes in rock properties that may in part reflect magmatic controls on the location of the VMS deposits. The interpretations are being integrated with on-going geochemical and geochronological studies to better understand the observed geophysical anomalies over an accreted island-arc setting. The gravity compilation contains stations spaced approximately 1.6 km where access was not limited (Snyder and others, 2004). The data were gridded to a 400 m cell size from which filtered data sets were generated. The complete Bouguer anomaly map highlights dense mafic volcano-plutonic rocks that were intruded by less dense, syn- and post-tectonic granite-tonalite rocks. Large northeast-southwest and east-west trending gradients coincide with mapped and inferred buried faults that indicate offset crustal blocks of varying densities within, or beneath, the bimodal metavolcanic rocks. Such structures may reflect bounding faults enclosing grabens within which the VMS deposits may have formed. Aeromagnetic data were collected along north-south flight lines spaced 800 m at a nominal height of 150 m (Karl, 1986). These data were contoured using a cell size of 250 m from which filtered data sets were generated. The reduced-to-pole (RTP) transformation shows that the mafic volcano-plutonic rocks produce strong magnetic anomaly highs. Moderate amplitude RTP anomaly highs are observed over the younger granite-tonalite plutons. Bimodal volcanic rocks produce magnetic lows; however, within these lows are circular and linear magnetic highs that trend northeast-southwest and east-west, some of which are associated with gabbro rocks. The analytic signal (AS) transformation shows high gradients over the mafic volcano-plutonic rocks. The AS highlights isolated magnetic anomalies within the bimodal volcanic rocks, some of which may be imaging synvolcanic plutons that may have acted as heat sources for VMS hydrothermal systems. Several circular and linear AS anomalies occur along the gravity gradients and mapped faults. The tilt derivative (TDR) transform highlights a northeast-trending magnetic fabric within both the mafic volcano-plutonic rocks and the bimodal volcanic rocks. TDR lineaments occur in higher concentrations proximal to mapped faults. These lineaments are parallel to bedding orientations and major structures and, are therefore, interpreted to reflect the strike of mafic flows within the volcanic rock package. 5 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Together, the gravity and magnetic data are able to identify the structural framework and the buried extents of the bimodal metavolcanic rocks that are the targets for VMS exploration. In addition, these data help map the location of plutons that are critical to understanding the thermal evolution of the region. Figure 1: Generalized geology map showing major rock types within the Pembine-Wausau terrane between the Niagara fault and Eau Pleine shear zone (Nicholson and others, 2004). Red triangles and circles depict VMS deposits and prospects, respectively. REFERENCES Karl, J.H., 1986. Total magnetic intensity map of northern Wisconsin: Wisconsin Geological and Natural History Survey Map 86-7: scale 1:250,000. Nicholson, S.W., Dicken, C.L., Foose, M.P., and Mueller, J.A.L., 2004. Preliminary integrated geologic map databases of the United States: Minnesota, Wisconsin, Michigan, Illinois, and Indiana. U.S. Geological Survey Open-File Report 2004-1355. Schulz, K.J. and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research, 157: 4-25. Snyder, S.L., Geister, D.W., Daniels, D.L., and Ervin, C.P., 2004. Principal facts for gravity data collected in Wisconsin: A website and CD-ROM for distribution of data. U.S. Geological Survey Open-File Report 03157. 6 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Iron-rich siliceous stromatolites from the upper algal unit of the Gunflint and Biwabik iron formations ARTS, Adrian and FRALICK, Philip Department of Geology, Lakehead University, Oliver Rd. Thunder Bay, ON, P7B 5E1, Canada The Gunflint and Biwabik formations comprise the middle units of the Proterozoic Animikie Group that crop out along the north shore of Lake Superior. Two stromatolitic units are present within these formations; a basal unit that the stromatolites grow directly on the peneplained Archean basement or on the conglomerate which forms the base of the Animikie, and a second upper unit roughly 45 meters above the base. The Animikie stromatolites are unique in that they are mainly composed of finely laminated, fine grained, bands of iron-rich silica. This is unusual as most stromatolites, modern and ancient are composed primarily of carbonate. Since their initial discovery, speculation as to their original mineralogy have been raised (Barghoorn & Tyler, 1965; Cloud 1965; Lougheed 1983; Sommers et al, 2000). This study was conducted on the two stromatolitic horizons to determine whether these iron-silica-rich stromatolites represent a primary mineralogy or if there is evidence for silicification of an earlier carbonate phase. The utilization of high resolution, field emission scanning electron microscopy (SEM), x-ray diffraction (XRD), whole rock geochemistry, and transmitted light microscopy revealed several pieces of compelling evidence within the upper algal unit. Hand samples cut horizontally, (Fig. 1A) show the fine grained, hematite-rich laminae located within the columns (white arrows). In thin section, erosive scouring of the siliceous stromatolite column tops is common, with new siliceous bacterial mat truncating the old (Fig. 1B). The sharp contact between the scour-truncation suggest lithification prior to the development of the younger laminae. The microquartz which the stromatolitic laminae are composed of, is in sharp contrast to the mega quartz cement found within the interspace between the columns (Fig. 1C). Note the thin (≤10µm) microquartz wisps overlaying the coated grains bridging the columns, suggesting a fossilized bacterial mat (white arrow). Intraformational clasts containing pieces of lithified stromatolite are common, especially as nucleation sites of ooids (Fig. 1D, 1E). Finally, energy dispersive x-ray (EDX) mapping show distinct alternation of silica-iron-manganese in ooid coatings and stromatolite laminae (Figs. 1F-1I). The above strongly indicates the Gunflint and Biwabik stromatolites were originally siliceous and formed by a different precipitation mechanism than Proterozoic carbonate stromatolites or modern agglutinated forms did. REFERENCES Barghoorn, E.S. and Tyler, S.A. 1965. Microorganisms from the Gunflint chert, Science, 147, 563-577. Cloud, P, 1965. Significance of the Gunflint (Precambrian) microflora. Science, 148(3666), 27-35. Lougheed, M.S. 1983. Origin of Precambrian iron-formations in the Lake Superior region. Geological Society of America Bulletin, 94, 325-340. Sommers, M.G., Awramik, S.M., Woo, K.S., 2000. Evidence for initial calcite-aragonite composition of lower algal chert member ooids and stromatolites, Paleoproterozoic Gunflint Formation, Ontario, Canada. Canadian Journal of Earth Sciences, 37(9), 1229-1243. 7 �Proceedings of the 61st ILSG Annual Meeting - Part 1 (A ) (B) (C ) (D ) (E) (F) (G ) (H ) (I) Figure. 1. Features of the silica-iron-rich stromatolites within the Gunflint and Biwabik Formations. (A) Horizontal section through stromatolite columns showing fine grained hematite laminae within the column (white arrows). (B) Photomicrograph illustrating the truncation and subsequent overgrowth of stromatolite columns by fine grained, thin laminae composed of iron-rich jasper (red/brown) and quartz (clear). (C) XPL photomicrograph highlighting difference between microquartz stromatolitic laminae, and blocky quartz cement in interspace. The bridging by a thin siliceous algal mat between two columns (white arrow), suggests a rapid growth of mat over the grainstone. (D) Clast containing coated grains and piece of stromatolite column. (E) SEM-BSE image of an ooid containing a broken piece of stromatolite as its nucleation point. This suggests the stromatolite was lithified prior to the development of the ooid lamination. (F-I) EDX false colour images of alternating silica (G), iron (H), manganese (I) stromatolitic laminae. 8 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Evidence of high-temperature Ni isotopic fractionation during the formation of Cu-Ni-PGE sulfide deposits in the Duluth Complex ASP, Kristofer1, SCHARDT, Christian1, and SPIVAK-BIRNDORF, Lev2 1 Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 1114 Kirby Dr., Duluth, MN 55812 USA 2 Department of Geological Sciences, University of Indiana, 1001 East 10th St., Bloomington, IN 47405 USA The Duluth Complex in northeastern Minnesota is an extensive, largely gabbroic body that formed during the 1.1 Ga Midcontinent Rift event. The basal zones of the complex, the South Kawishiwi (SKI) and Partridge River intrusions (PRI), host extensive Cu-Ni-PGE mineralization in a number of recognized deposits that are actively being explored. Previous research, using sulfur isotopes, indicates the underlying Virginia formation as a source of sulfur during the formation of these deposits [1]. Recent studies have shown significant Ni isotopic variations of up to 1.1‰ in high-temperature magmatic rocks associated with magmatic sulfide mineralization [24]. The heterogeneous mineralogy and mafic lithologies in the basal Duluth Complex indicate a range of different processes active during crystallization. The primary goal of this study is to examine mineralized and unmineralized Duluth Complex material to assess the potential of Ni isotopic fractionation between early cumulates, subsequent Cu-Ni-PGE mineralization, and weathering products. Of particular interest is the exploration potential of Ni isotopes for magmatic sulfides recorded in weathered products at the surface. Sample material collected includes till, outcrop material, and in-situ mineralization as well as control samples unconnected to the SKI or PRI. Massive sulfides, disseminated sulfides, and non-mineralized gabbro were selected to obtain Ni isotopic signatures from both a sulfide and silicate source. Till and surface samples were collected in the vicinity of known Cu-Ni-PGE deposits, including Spruce Road, Maturi, Mesaba, Serpentine, and NorthMet. In-situ mineralized material from drill core was provided by local mining companies (Duluth Metals/Twin Metals, PolyMet, Teck, Encampment Minerals) that hold mineral rights to individual deposits. Material from other deposits, including Birch Lake, Wetlegs, and Wyman Creek was sampled from drill core available from the Minnesota DNR. Samples were processed to produce thin sections and polished thick to observe representative textures and mineral compositions for the Cu-Ni-PGE mineralization and host rock gabbro. Material from each deposit was analyzed using XRD, whole rock, and trace element geochemistry. Electric pulse disaggregation (EPD) was used to separate 1 cm3 samples into individual mineral grains and olivine was separated for isotopic analysis. EPD olivine crystals, along with till, massive sulfide material, and weathered surface samples were ground to < 70 µm using a shatterbox and send for nickel isotope analysis to the University of Indiana using the double-spiking method outlined in [2]. Ni isotope ratios are reported relative to the NIST SRM 968 standard with conventional delta notation and a general 2σ error of 0.06‰. Isotopic results show a spread of δ60/58Ni values from -0.97‰ to +0.21‰, within the range of Ni isotopic values reported previously [2,3]. The least fractionated values come from unmineralized mafic intrusives (-0.07‰), while Ni isotopic ratios become progressively lighter with increasing sulfide content, ranging from -0.16 ‰ to -0.97‰ (Fig. 1). Till samples record intermediate values (-0.02 ‰ to -0.77‰) and weathered surface samples can span the entire range. 9 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: δ60/58 Ni values for sulfide, olivine, and till compared to background values and data from previous studies. The isotopic data indicate an isotopic fractionation trend in the Duluth Complex from unfractionated values around zero, assumed to be bulk silicate earth, for early crystallizing phases to notably fractionated sulfide mineralization, accumulated at some later stage. This fractionation of up to 0.81‰ suggests that Ni was fractionated during the Ni sulfide formation by incorporating preferentially lighter Ni into the accumulating sulfide melt and resulting Ni sulfides. Ni isotopic values for till and mineralized surface samples, and their correlation with known deposits, may be useful in distinguishing regions overlying Cu-Ni-PGE mineralization from barren areas. Data may also help to identify the entry point of the mineralizing magma based on the location and isotopic signature of individual sulfide deposits. This will require a more detailed sampling of selected locations and materials. REFERENCES Gueguen B., Rouxel O., Ponzevera E., Bekker A., Fouquet Y. (2013) Ni isotope variations in terrestrial silicate rocks and geological reference materials measured by MC-ICP-MS. Geostandards and Geoanalytical Research 3: 297-317 Hiebert RS., Rouxel, O., Houlé, MG., Bekker, A. (2014) Ni isotope fractionation between komatiite and sulfide mineralization at the Neoarchean Hart deposit, Abitibi greenstone belt, Canada. Geological Society of America Abstracts 46: 467 Ripley, E. (2006) Sulfur isotopic studies of the Dunka Road Cu-Ni deposit, Duluth Complex, Minnesota. Economic Geology 76: 610-620 Wasylenki, L.E, Howe, Haleigh D., Spivak-Birndorf, L.J., Bish, DL. (2015) Ni isotope fractionation during sorption to ferrihydrite: implications for Ni in banded iron formations. Chemical Geology 400: 56-64 10 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Fold analyses in the Gunflint Formation: working towards a characterization of regional deformation in the Animikie Group near Thunder Bay, Ontario BAIRD, Jordan, and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Deformation in the Animikie Group near Thunder Bay is characterized by intense, localized fold-and-thrust belt deformation within units that are otherwise relatively undeformed and flat-lying. This deformation is interpreted to be the result regional stress during Proterozoic tectonism. Stereographic analyses of structural measurements and outcrop observations are used to determine beta-axes (mean fold axes) for fold populations observed, as well as to determine structural relationships within the region. Two main fold populations are observable in the data. The primary population exhibits a north-south trending beta-axis, indicating east-west compressive stress. The secondary population exhibits an east-west trending beta-axis, indicating either north-south compressive stress or the presence of lateral ramps. Slickenlines are also observed to trend north-south and east-west, depending on outcrop location. East-west trending slickenlines tend to be in areas of more intense folding, indicating that they may be older than the north-south trending slickenlines. Older slickenlines may have been destroyed when fault surfaces were reactivated in areas of less intense folding, replaced by the slickenlines associated with the most recent deformation. Additional observations include the presence of fold-hinge breccia associated with non-cylindrical folding, which may indicate lateral ramp formation, as well as the presence of a possible cleavage duplex structure, which may indicate repetitive east-verging thrusting. A specific tectonic history for the Animikie Group has been suggested based on these observations. It has been proposed here that primary east-verging thrusting was associated with the Trans-Hudson orogeny in the Paleoproterozoic. Following this, there may or may not have been a secondary compressional phase due to the Yavapai-Mazatzal orogenies during the early to middle Proterozoic; the effects of these orogenies remain unclear. Deformation likely culminated in the late Proterozoic with north-south extension related to the Midcontinent Rift.   11 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Interpretation of the St. Amour Deep Stratigraphic Test Well, Alger County, Michigan BAUMANN, Steven D.J.1, CORY, Alexandra B.1, and DYLKA, Sandra K.1 1 Geology Section, Midwest Institute of Geosciences and Engineering, 1321 W. Touhy Ave. 2S, Chicago, IL 60626 The Amoco Production St. Amour (1-29R) was a petroleum deep test well, drilled in 1988 to a depth of 7238 feet below ground surface (bgs). Surface elevation of the well was approximately 910 feet above mean sea level. The GPS coordinates for the well are 46.354591o 86.584370o. The well passed through 110 feet of glacial material at the surface. Below the glacial material is 411 feet of Paleozoic sediments. Under the Paleozoic is 2132 feet of the Precambrian Jacobsville Formation (this differs from Ojakangas’ 2002 interpretation). The Jacobsville can be subdivided into three distinct units that are very similar to the formations of the Bayfield Group in Wisconsin. Chequamegon type lithology is encountered at 521-1470 feet bgs, Devils Island type lithology is encountered at 1470-1939 feet bgs, and Orienta type lithology is encountered at 19392653 feet bgs. Beneath the Jacobsville lies 3197 feet of the Freda Formation, which can also be divided into three units. There is an upper red to gray and light brown, fine to medium grained arkose (2653-2830 feet bgs). The middle unit is composed of mostly red, silty, very fine to medium grained arkose, with beds of red siltstone and shale (2830-3260 feet below the surface). The basal Freda is a fining upwards sequence of red to gray mottled pale brown (occasionally green), fine to coarse grained arkose with beds of deep red shale, and pebbles are common. Below the Freda is a unit not seen elsewhere in the Oronto Group. There is 675 feet of a red, hematitic quartz arenite with thick beds of quartz conglomerate (unit Y(x), Figure 1). This unit is very mature for its stratigraphic position and appears unique from above and below units. From a depth of 625-6933 was described as a “heterogeneous unit” (Ojakangas 2002). It consists of 203 feet of faulted basalt interbedded with sandstone, siltstone, and conglomerate (unit Y(xx), Figure 1). This unit appears to be unconformable at its base, although this is difficult to determine for sure in the core. Below this unit (6728-6783 feet bgs) is what has been interpreted as 145 feet of the Nonesuch Formation. We agree with this interpretation since the stratigraphy matches up well with the lower three units in the Big Iron River at Bonanza Falls (Susek 1997). Below the Nonesuch lies 50 feet of a basalt flow (unit B1, Figure 1) over a gabbro diabase (unit B2, Figure 1). The gabbro diabase appears to be a later intrusion into units Y(bb) and Y(xxx) and may have been emplaced during the Grenville Orogeny. Under the gabbro is 10 feet of red and gray siltstone and shale underlain by conglomerate, unit Y(xxx). The deepest unit penetrated is 350 feet of rhyolite and ignimbrite, which has been dated at 1.083 + 0.003 Gya old. The only known igneous rocks younger than this within the Mid-continental Rift belong to the Bear Lake Rhyolite Stock in the Freda Formation (1.054 + 0.034 Gya). There are two faults present within the St. Amour core. There is a 32 foot long stretch highly sheered red and green sandstone from 6573-6605 feet bgs. A smaller second fault, which contains 12 feet of gray and red sandstone and shale is present at 6691-6703 feet bgs. Figure 1 shows the interpretive relationship between the geologic units from 6400-7000 feet bgs. In our interpretation both faults are high angle reverse faults that dip about 73o from the horizontal (the strike of the core is not known so the dip direction could not be obtained). The smaller of the two faults has been modeled parallel to the larger fault. However, in reality, it likely branches off from it and represents a faulted shatter zone. Total fault displacement is about 121 feet. There are two basalt units above the Nonesuch within the core. The upper one is 48 feet thick and lies 12 �Proceedings of the 61st ILSG Annual Meeting - Part 1 above the faults. The second is about 43 feet thick and lies between the two faults. Both basalts have mostly sandstone and siltstone with some conglomerate below them. Due to their similar thickness and lithology, we interpret them to be the same unit offset by the faults (units Y(ba), Figure 1). The basalt has not been dated but since it overlies the rhyolite, it has to be younger. If it were to be dated we could be able to bracket the time of deposition for the Nonesuch Formation. Based on our analysis of the core we partially agree with Ojakangas and Dickas’ 2002 interpretation that late Mid-continent Rift volcanism occurred later than in other areas because of the core’s proximity to the mantle plume. We also postulate that the area around the core may represent a buried stratovolcano similar to the one at Porcupine Mountains (located about 150 miles west-northwest of the St. Amour core location). We propose that a high altitude composite volcano existed in the area until about 1.080 Gya, until the complete deposition of the Nonesuch Formation. At that time main volcanism ceased and the area began to rapidly subside allowing for the deposition of thick quartz sandstone (unit Y(x), Figure 1) over the youngest basalt (unit Y(ba), Figure 1), thus burying the volcano. The main fault in the core may have originated as a normal growth fault that was later reactivated as a reverse fault during the Grenville Orogeny, creating smaller offshoot faults. Until more deep cores are obtained from the area of the St. Amour test well, the presence of a subsurface volcano cannot be verified. REFERENCES Bornhorst, T.J., Rose, W.I., 1994. Self-Guided Geologic Field Trip to the Keweenaw Peninsula, Michigan. Institute on Lake Superior Geology, volume 40, pp. 161-164 Dickas, A.B., Mudrey Jr., M.G., 1992. Keweenaw Sedimentary Rock of the South Shore, Lake Superior. Institute on Lake Superior Geology, volume 38, pp. 43-102 Friedhoff-Miller, Diana, 1988. Record of Well Drilling or Deepening, St. Amour 1-29R. State of Michigan Department of Natural Resources, Geological Survey Division Ojakangas, R.W., Dickas, A.B., 2002. The 1.1-Ga Midcontinent Rift System, central North America: sedimentology of two deep boreholes, Lake Superior Region. Journal of Sedimentary Geology 147 (2002) pp. 13-36 Suszek, T. J., 1997. Petrography and sedimentation of the Middle Proterozoic (Keweenaw) Nonesuch Formation, western Lake Superior region, Midcontinent rift system, Geological Society of America, Special Paper 312 13 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: Structural Interpretation of the St. Amour Core from 6400-7000 feet Below the Surface 14 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Linking the Ordovician L-Chondrite Event to the Terrestrial Cratering Record: a NorthAmerican Perspective BLEEKER, Wouter Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, Canada, K1A 0E8 Based on abundant evidence of shock metamorphism and partially or wholly reset Ar ages of a class of meteorites (low metal or “L-chondrites”), meteorite researchers have hypothesized that a catastrophic impact and breakup event took place in the asteroid belt about 500 million years ago (Keil et al., 1994), spawning a large population of meteoritic fragments some of which were perturbed into Earth crossing orbits. This hypothesis received a major boost with the discovery of numerous L-chondrite meteorite fragments in limestone quarries in southern Sweden (Thorslund et al., 1984; Nyström et al., 1988; Schmitz et al., 1996), preserved in-situ across a section of Middle Ordovician stratigraphy (ca. 470 Ma). The Lchondrite-bearing, condensed, shallow marine limestones are also characterized by a sharply increased heavy mineral count of extraterrestrial Ni spinels, also of L-chondrite affinity (Schmitz et al., 2003). This Ni spinel “rain out” has now been documented not only in Sweden but also in China (Heck, et al., 2010), and by all expectations should constitute a global signal. Since then, more detailed Ar-Ar dating of the partially degassed meteorites has refined the likely age of the breakup event to 470±6 Ma (Korochantseva et al., 2007). A number of small meteorite impact craters in Scandinavia has been linked to this event, e.g. the 458 Ma Lockne crater in central Sweden (Grahn et al., 1996; Alwmark and Schmitz, 2007). With Earth moving through a dynamically evolving swarm of asteroid debris (e.g., Nesvorný et al., 2009), the effects of this event should have been global. Numerous, possibly large, impact craters should be linked to this event, particularly in North America with its large cratonic target area and a robust population of ~60 confirmed impact structures. We have previously linked the well-known Brent impact crater to this event (Bleeker, 2011), with a stratigraphically constrained age of ca. 460-450 Ma (Lozej and Beales, 1975) and with a melt/breccia sheet that is known to contain geochemical traces of an L-chondrite bolide (Palme et al., 1981). Among the ~32 confirmed and possible impact craters in Canada alone (e.g., see Grieve, 2006), there could be as many as 5-10 structures that are linked to the same broad event: Brent, Holleford, Skeleton Lake, Nicholson, Pilot Lake, Presqu’ile, Couture, La Moinerie, and the large Slate Island structure in Lake Superior. Scattered and non-definitive K-Ar and Ar-Ar ages could extend this list to the large (~50-55 km-diameter) Carswell structure, and perhaps the buried but unconfirmed Can-Am crater. We have recently redated this structure, using Ar-Ar stepwise heating on pristine adularia crystals, to 481±1 Ma (Bleeker et al., 2015), confirming it as part of the Ordovician impact spike. Even if only the most likely subgroup of this list is indeed related to a ca. 470-440 Ma impact spike of L-chondrites, the proportion of impact craters linked to this event is very large (1 in 5?), as similarly suggested by the more limited sample of just the Swedish crater record alone. The same conclusion is also reached for craters in the remainder of North America (mainly the USA), where among ~30 confirmed craters the following could be linked to the same event: Ames, Calvin, Glasford, Glover Bluff, Newporte, Rock Elm and Versailles (again, a proportionally similar and very large subpopulation, as in Canada). Several of these structures have stratigraphically constrained ages in the 470440 Ma interval (e.g., see Koeberl et al., 2001, for Ames; Millstein, 1994, for Calvin; McHone et al., 1986 for Glasford). It is concluded that the Ordovician L-chondrite event left a major imprint in the North American and by inference global cratering record and, as recognized by Birger Schmitz and coworkers (Schmitz et al., 2008), must have jarred the Earth system throughout much of the Middle and Late Ordovician. During the 470-440 Ma interval, the flux of large impactors appears to have been an order of magnitude higher than during the remainder of the Phanerozoic. To gauge the full scope of this event, an integrated effort to produce better and more precise ages for all major impact structures is needed, with equal emphasis on stratigraphic and isotopic constraints. Important rewards could be an improved understanding of the 15 �Proceedings of the 61st ILSG Annual Meeting - Part 1 dynamical evolution of an asteroidal breakup swarm, a quantification of the overall impact flux during this interval, and a better appreciation of how the Earth system and biota responded to this major event. REFERENCES Awlmark, C., and Schmitz, B., 2007. Extraterrestrial chromite in the resurge deposit of the early Late Ordovician Lockne crater, central Sweden. Earth and Planetary Science Letters, vol. 253, p. 291-303. Bleeker, W., 2011. Linking the Ordovician L-chondrite event to the terrestrial cratering record: A North American perspective. Ottawa 2011, GAC-MAC Joint Annual Meeting, University of Ottawa, May 25-27, Abstracts, vol. 34, p. 19. Bleeker W., LeCheminant, A.N., Alwmark, C., Page, L., Scherstén, A., and Söderlund, U., 2015. The age of the Carswell impact structure. AGU-GAC-MAC-CGU Joint Assemblee, 3-7 May 2015, Montreal. Grahn, Y., Nolvak, J., and Paris, F., 1996. Precise chitinozoan dating of Ordovician impact events in Baltoscandia. Journal of Micropaleontology, vol. 15, p. 21-35. Grieve, R.A.F., 2006. Impact structures in Canada. Geological Association of Canada, Geotext 5. Heck, P.R., Ushikubo, T., Schmitz, B., Kita, N.T., Spicuzza, M.J., and Valley, J.W., 2010. A single asteroidal source for extraterrestrial Ordovician chromite grains from Sweden and China: High Precision oxygen three-isotope SIMS analysis. Geochimica et Cosmochimica Acta, vol. 74, p. 497-509. Keil, K., Haack, H., and Scott, E.R.D., 1994. Catastrophic fragmentation of asteroids: evidence from meteorites. Planetary and Space Science, vol. 42 (12), p. 1109-1122. Korochantseva, E.V., Trieloff, M., Lorenz, C.A., Buykin, A.I., Ivanova, M.A., Schwarz, W.H., Hopp, J., and Jessberger, E.K., 2007. L-chondrite asteroid breakup tied to Ordovician meteorite shower by multiple isochron 40Ar-39Ar dating. Meteroritics & Planetary Science, vol. 42 (1), p. 113-130. Lozej, G.P., and Beales, F.W., 1975. The unmetamorphosed sedimentary fill of the Brent meteorite crater, southeastern Ontario. Canadian Journal of Earth Sciences, vol. 12, p. 606-628. Koeberl, C., Reimold, W.U., and Kelly, S.P., 2001. Petrography, geochemistry, and argon-40/argon-39 ages of impact-melt rocks and breccias from the Ames impact structure, Oklahoma: The Nicor Chestnut 18-4 drill core. Meteoritics & Planetary Science, vol. 36, p. 651-669. McHone, J.F., Sargent, M.L., and Nelson, W.J., 1986. Shatter cones in Illinois: evidence for meteoritic impacts at Glasford and Des Plaines. Meteoritics, vol. 21, p. 446. Millstein, R.L., 1994. The Calvin impact crater, Cass County, Michigan: identification and analysis of a subsurface Ordovician astrobleme. Ph.D. thesis, unpublished, Oregon State University, 114 p. Nesvorný, D., Vokroulicky, D., Morbidelli, and Bottke, W., 2009. Asteroidal source of L chondrite meteorites. Icarus, vol. 2009, p. 698-701. Nyström, J.O., Lindström, M., and Wickman, F.E., 1988. Discovery of a second Ordovician meteorite using chromite as a tracer. Nature, vol., 336, p. 572-574. Palme, H., Grieve, A.F., and Wolf, R., 1981. Identification of the projectile at the Brent crater, and further considerations of projectile types at terrestrial craters. Geochimica et Cosmochimica Acta, vol. 45, p. 24172424. Schmitz, B., Lindström, Asaro, F., and Tassinari, M., 1996. Geochemistry of meteorite-rich marine limestone strata and fossil meteorites from the Lower Ordovician and Kinnekulle, Sweden. Earth and Planetary Science Letters, vol. 145, p. 31-48. Schmitz, B., Häggström, T., and Tassinari, M., 2003. Sediment-dispersed extraterrestrial chromite traces a major asteroid disruption event. Science, vol. 300, p. 961-964. Schmitz, B., Harper, D.A.T., Puecker-Ehrenbrink, B., Stouge, S., Alwmark, C., Cronholm, A., Bergström, Tassinari, M., and Xiaofeng, W., 2008. Asteroid breakup linked to the Great Ordovician Biodiversification event. Nature Geoscience, vol. 1, p. 49-53. Thorslund, P., Wickman, F.E., Nyström, J.O., 1984. The Ordovician chondrite from Brunflo, central Sweden, I. General description and primary minerals. Lithos, vol. 17, p. 87-100. 16 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Large hypervelocity impacts on Earth: Empirical observations and validation of computational model predictions for Sudbury and Chicxulub BRUMPTON, Gregory R. and KISSIN, Stephen A. Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Johnson and Melosh (2014) reported on a high-resolution, two-dimensional computational model of effects during a hypervelocity 10km impactor collision with Earth at 20km/s. This early version of the model focused on three impact products, melt droplets, melt fragments and accretionary impact lapilli. They estimated the size of the ejecta products using simple analytical expressions and information determined from their hydrocode models. Prediction of the size of the ejecta products depends on the impactor size, impact velocity and ejection velocity from the forming crater. Johnson and Melosh sought to find a consensus between model predictions describing the formation of the ejecta products and actual geological observations. Modeled estimates of the sizes of melt droplets and accretionary impact lapilli are generally within one order of magnitude of limited empirical measurements at Chicxulub and Sudbury. This agreement acts as a validation of their model and illustrates a process whereby geologic observations can be applied so as to improve the model. Our studies on Sudbury ejecta from the Thunder Bay area (Addison et al. 2005; 2010) have considered the predictions of the Johnson and Melosh model with the results indicated below: Prediction: The model predicts order of magnitude estimates of the size [mm-scale] of melt droplets and melt fragments. - Our studies of Sudbury ejecta and comparison with literature information on Chicxulub (Yancey and Guillemette 2008;Yancey and Liu 2013) verify the predictions of the model. Prediction: Millimeter-sized melt droplets should be found together with accretionary impact lapilli and rarer melt fragments (tektites). Figure 1. Accretionary lapillus (left), tektite (horizontal arrow) and melt droplet (vertical arrow). MC18A SEM, XPL. Prediction: A wide range of sizes of melt droplets should be found at any given site. Figure 2. Melt droplets illustrating a range of sizes. JN23, XPL. Prediction Accretionary impact lapilli form during the ejection process in the turbulent ejecta curtain from fine-grained, solid fragments and molten silicate acts as a binding agent. Lapilli range from larger than 1cm to less than 1mm diameter. 17 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 3. Accretionary lapillus with multiple rings of opaque silicate, cored with crystal fragment. JN29A-2 SEM, XPL Prediction: There is a tendency for the small solid fragments to follow vapor streamlines so that the small fragments may be swept around the growing lapilli and as a result accrete in rows. Figure 4. Detail of an accretionary lapillus showing streamlined, rows of fragments apparently deposited by turbulent vapors during growth of the lapillus. JN29A-2 SEM, XPL. Prediction: The largest melt fragments (tektites) will come from more lightly shocked, near surface, target material. The composition of melt fragments in Sudbury ejecta is consistent with that of surficial sedimentary rocks and granitic gneiss of the target area. They show little if any isotropic silicate melt on the exterior. Prediction: "A more detailed comparison of our models to known ejecta layers will allow us to test the predicted dependence of ejecta product size on impactor size and may even allow us to empirically constrain some additional products ..." (Johnson and Melosh 2014). REFERENCES Addison, W.D., Brumpton G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L. 2005. Discovery of distal ejecta from the 1850 Ma Sudbury impact event, Geology,33: 193-196. Addison, W.D., Brumpton, G.R., Davis, D.W., Fralick, P.W., and Kissin, S.A. 2010. Debrisites from the Sudbury impact event in Ontario, north of Lake Superior, and a new age constraint: Are they base-surge deposits or tsunami deposits? in Gibson, R.L. and Reihold, U.W., eds. Large Meteorite Impacts and Planetary Evoution IV: Geological Society of America Special Paper, 465: 245-268. Johnson, B.C. and Melosh, H.J. 2014. Formation of melt droplets, melt fragments, and accretionary impact lapilli during a hypervelocity impact. Icarus, 228: 347-363. Yancey, T.E. and Guillemette, R.N. 2008. Carbonate accretionary lapilli in distal deposits of the Chicxulub impact event. Geological Society of America Bulletin, 120: 1105-1118. Yancey, T.E. and Liu, C. 2013. Impact-induced sediment deposition on an offshore, mud-substrate continental shelf, Cretaceous-Paleogene boundary, Brazos River, Texas, U.S.A. Journal of Sedimentary Research, 83: 354-367. 18 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Tainiolite from the Stettin intrusion, Wausau Complex, Marathon County, WI. BUCHHOLZ, THOMAS W.1, FALSTER, Alexander U.2, and SIMMONS2, W. B. 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494 2 Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217. The Wausau Syenite Complex (WSC) is composed of four plutons, of which the Stettin Pluton (1565 Ma +3-5) (Van Wyck, 1994) is the oldest and most alkalic. Tainiolite is a relatively uncommon Li-Mg mica; KLiMg2(Si4O10)F2 and occurs in alkalic igneous rocks such as syenites and associated pegmatites. Recently tainiolite has been identified from two sites located in the Stettin Pluton, and probable tainiolite has been found at one other site within the pluton. The first occurrence is in the Ravine Pegmatite, a small irregular evolved pegmatite located on the west side of the long-dormant Dehnel quarry, near the western edge of the Stettin Pluton. Here, tainiolite occurs as colorless crystals of composition (K0.921Na0.080)Σ1.001Li1.000(Mg1.201Fe0.705Mn0.049Al0.033Ti0.012Ca0.009)Σ2.009(Si3.910Al0.090)Σ4.000O22 [(F1.811OH0.189)]Σ2.000 in the miarolitic core zone of the pegmatite and in adjacent intermediate zones, associated with abundant zircon, pyrochlore, bastnaesite-(Ce), bastnaesite-(La), bastnaesite-(Nd), columbite-(Fe), fersmite, aegerine, riebeckite, microcline, albite, and rare baddeleyite. The given analytical results, determined using EMP, XRD and DCPS, are in good agreement with recent published data for tainiolite (e.g. Armbruster et al, 2007). An additional occurrence was noted in summer 2014 in a cobble of syenite pegmatite recovered from a rock pile located in the western portion of the Stettin intrusion on the north side of Evergreen Drive (southern portion of S. 10, T 29N, R 6 E), near the contact between amphibole syenite and the discontinuous nepheline syenite outer rim of the pluton. Here the tainiolite, of composition (K0.969Na0.032)Σ1.001Li1.000(Mg1.109 Fe0.762Al0.041Mn0.030Ti0.010Ca0.011)Σ1.936(Si3.9100Al0.090)Σ4.000O22[(F1.851OH0.149)]Σ2.000, determined using EMP and DCPS, is found as colorless crystals associated with riebeckite, aegerine and Kspar. Interestingly, despite the likely simultaneous crystallization of riebeckite and probable tainiolite, the riebeckite is virtually Mg-free, suggesting that, given a limited supply of Mg vs Na, Mg may be preferentially taken up by tainiolite as opposed to riebeckite. The final (probable) occurrence was identified in material from the old Summit prospect (a failed attempt to mine U from a Th-rich pegmatite) obtained from a mineral collection assembled in the early 1960’s. The prospect worked a pegmatite located in the south-central portion of the Stettin Intrusion. Here the probable tainiolite occurs as sparse light yellow-brown to colorless flakes in intergrown aegerine, fluorite, zircon, bastnaesite-(Ce) and minor quartz. REFERENCES Armbruster, T., Richards, R. P., Gnos,E., Pettke, T., Herwegh, M. (2007): Unusual fibrous sodian tainiolite epitactic on phlogopite from marble xenoliths of Mont Saint-Hilaire, Quebec, Canada. The Canadian Mineralogist, Vol. 45, pp. 541-549. Van Wyck, N. (1994) The Wolf River A-type magmatic event in Wisconsin: U/Pb and Sm/Nd constraints on timing and petrogenesis. Institute on Lake Superior Geology, 40th Annual Meeting, Part 1, Program and Abstracts, p. 81-82. 19 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The Sudbury impact event in the Lake Superior region: Ten years of research on ten minutes of geologic time CANNON, W.F.1, ADDISON, W.D.2, BRUMPTON, Gregory R.3, and JIRSA, M.J.4 1 U.S. Geological Survey, MS 954, Reston, VA 20192 2 371 Crossbow Court, Thunder Bay, ON, P7G 1H5, Canada 3 Lakehead University, Department of Geology, Thunder Bay, ON, P7B 5E2, Canada 4 Minnesota Geological Survey, 2609 W. Territorial Rd., St. Paul, MN 55114 The ancient meteor impact near Sudbury, Ontario is the second largest impact event preserved in the geologic record. The impactor was most likely an ordinary or enstatite chondrite based on siderophile element concentrations in melt rocks (Huber et al., 2014; Petrus et al., 2015). It struck Earth at 1850 Ma, the age of the Sudbury Igneous Complex (SIC) (Davis, 2008), a remnant of the melt sheet generated by the impact. The impact formed a crater, now deeply eroded, with a diameter variously estimated from 130 to 250 km. The impactor, if chondritic, must have been 10 to 15 km in diameter to provide adequate energy to account for the crater size and melt volume. An impact of that magnitude undoubtedly spread a layer of debris hundreds of kilometers beyond the crater. Sudbury impact debris (ejecta) was first reported in the Lake Superior region in 2005 (Addison et al., 2005). Many other sites have been identified since in the western Lake Superior region (Pufahl et al., 2007; Cannon et al., 2010; Jirsa et al., 2011). The impact produced a variety of sedimentary and seismic features; the rocks in which these have been found are known collectively as the Sudbury Impact Layer (SIL). Compelling evidence of meteor impact is relict planar deformation features in quartz grains (Fig. 1A), because they are uniquely produced by impact-generated shock. Other common features are millimeter-scale spherules of devitrified impact glass (Fig. 1B), angular glass particles (Fig 1C), and accretionary lapilli (Fig. 1D). During deposition and reworking, ejecta became mixed with local material to form a hybrid rock. Seismic effects include fracturing and dislocation of pre-impact rocks and seismically triggered submarine debris flows. The effects of the impact across the Lake Superior region can be simulated using a computational model (Collins et al., 2005). Although there is no unique solution, input parameters of: 1) a chondritic impactor with a density of 3.4 g/cm3, 2) an impact velocity of 25 km/sec, 3) an impactor diameter of 15 km, 4) an impact angle of 45o, and 5) crystalline target rocks, predict: 1) a crater diameter of 193 km, well within the range of estimates, 2) a melt volume of about 12,000 km3, similar to previous estimates (Deutsch et al., 1995), and 3) a melt sheet thickness of about 1.2 km, less than the observed thickness of the preserved SIC. Using those same parameters, the model predicts effects of the impact at various distances from Sudbury, which provides a framework in which to judge observed features (Figure 2). The currently known extent of the SIL stretches from the Dead River Basin in Michigan, 500 km west of the impact site, to Coleraine in the western Mesabi Iron Range in Minnesota, about 1000 km west of the impact site. It spans the transition from proximal to distal deposits and its character changes markedly over that distance. The SIL also spans a north-south distance of about 150 kilometers, across which it was deposited in conditions ranging from dry land in the Thunder Bay area, across a continental shelf to the south, to deep water in the Iron River and Crystal Falls area. This combination of factors makes the Lake Superior region a unique natural laboratory in which to study the range of effects of the Sudbury impact beyond the crater margin. 20 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Seismicity: The Sudbury impact generated an earthquake of Richter magnitude 10 or greater, larger than any possible tectonic-related earthquake. The severity of shaking was as great as Mercalli intensity VIII at the more proximal sites and gradually waned to the west. Because the velocity of seismic waves is greater than the velocity of ejecta, severe shaking began minutes before the arrival of ejecta. This shaking caused liquefaction of pre-impact sediments and brecciation of basement rocks, probably aiding in their incorporation into ground surges of ejecta that followed within minutes. At Gunflint Lake the upper 5-7 m of the Gunflint Iron Formation were liquefied to formed coarse breccias of chert in an iron silicate matrix. This is the most distal locality at which strong seismicity has been documented. In the eastern Gogebic Range, upper parts of the Ironwood Iron Formation were mobilized into submarine slump deposits. In the Iron River-Crystal Falls area, in deep-water, massive debris flows were triggered from both the shelf edge to the north and an island arc to the south, resulting in slump deposits as much as 150 m thick. They remained active long enough for ejecta to reach the deep seabed and be incorporated into the slump deposits. In the Dead River Basin, the most proximal SIL localities, Archean basement rocks were intensely fractured and overlying sediments were injected downward into these fractures, probably a few tens of meters below the sea floor. Ejecta: The Sudbury impact blasted material (ejecta) away from the crater as an “ejecta curtain”, which swept across the entire western Lake Superior region over a span of about 3 minutes with likely velocities of 1.5-2 km/sec. The ejecta included fragments of the pre-impact target rocks of varying sizes, with a predicted average fragment size of 2 cm at the more proximal sites, masses of melted rock, which eventually solidified into spherules and glass fragments (Figure 1B, C), and part of the impacting body itself, largely as high temperature vapor. On landing, the huge mass of material was propelled across the landscape by its forward momentum as a turbulent groundhugging density current (base surge) for hundreds of kilometers. The energy within the base surge allowed entrainment of underlying rock and unconsolidated sediments resulting in hybrid deposits of true ejecta (material from the crater itself) mixed with more local rock ranging from fine particles to meter-scale boulders. As the energy and velocity of the base surges waned, a discontinuous layer of debris was deposited as a variable thickness of bedded material whose internal structure records the chaotic nature of the transport and deposition of the impact debris 21 �Proceedings of the 61st ILSG Annual Meeting - Part 1 (Fralick et al., 2012). The most distal base surge deposits are at Gunflint Lake, about 700 km from the impact site, where a layer of breccia and lapillistone about a meter thick overlies the seismically disturbed upper beds of the Gunflint Iron Formation (Jirsa et al., 2011). At more proximal sites near Thunder Bay, base surge deposits are as thick as 4+ m and contain meter-scale and finer clasts of the underlying Gunflint Formation, ejecta rock fragments, devitrified glass, and accretionary lapilli (Addison et al., 2010). In Michigan, deposits are discontinuous. The SIL is absent in many exposures and drill core, but is as thick as 26 m in the Baraga Basin and 40 m in the Dead River Basin (Cannon et al., 2010). All of these deposits contain variable amounts of relatively local material, so the thickness of ejecta is significantly less than the total thickness of the surge deposits. At sites along the Mesabi Iron Range in Minnesota, beyond the outer fringes of ground surges, the SIL is a discontinuous layer, only tens of centimeters thick, composed mostly of glass spherules and minor clasts of quartz and feldspar, all of which are probably ejecta. These were deposited from a cloud of suspended ejecta material that spread over the region in the minutes to hours after the impact and settled onto the sea floor. The most distal known occurrence of the SIL, near Coleraine, Minnesota, is nearly 1000 km from the impact site, the greatest distance at which rocks of suitable age to record the Sudbury event are exposed in the Lake Superior region. Figure 3. Schematic cross section showing the stratigraphic position and depositional setting of the Sudbury Impact Layer across the Penokean foreland. The SIL was deposited within an active tectonic belt of the Penokean orogeny that varied from a low lying land area on the north, across a marine continental shelf southward, into a foredeep at the southernmost localities in the Iron River-Crystal Falls area (Fig. 3). The nature of its deposition varied accordingly. Ground surge deposition on the northern land area is the most straight-forward to comprehend, lacking the complexities of interaction of ejecta with seawater. To the south, in the Gunflint Lake, Mesabi, Gogebic, Baraga Basin, and Dead River Basin areas, ejecta appears to have been deposited in a relatively shallow sea in which iron-formations and ferruginous cherts were being deposited. The manner in which the extremely energetic ejecta and ground surges interacted with the ambient shallow ocean remains largely speculative. Tsunamis: The last impact-generated event that is recorded in the SIL is the inferred reworking of ejecta and underlying breccia by tsunami waves that were generated by the impact. Because these 22 �Proceedings of the 61st ILSG Annual Meeting - Part 1 waves move slowly relative to seismic waves and ejecta, they swept the area some hours after the ejecta was deposited. The magnitude of the tsunamis is poorly understood in theory and depends in part on whether or not the crater was developed in an eastward extension of the ocean that existed at that time in the Lake Superior region. In any case, tsunami waves must have been substantial and possibly enormous. Features formed by those waves have not been reliably differentiated from waning phases of ground surges in many deposits. Perhaps the best indication of wave reworking of ejecta is at Gunflint Lake, where upper beds of the SIL consist of ejecta intermixed with large clasts that are more rounded than typical of lower parts of the layer, indicating energetic reworking. The day after: An impact of the magnitude of Sudbury surely had global consequences, just as the Chicxulub impact did at the end of the Cretaceous period. A global layer of fine impact material was probably deposited, but has not yet been identified outside of the Lake Superior region. More locally, within the Lake Superior region, land areas were mantled with ejecta that probably dominated both the chemistry of water and physical nature of sediment being carried to the adjacent shallow ocean for a considerable time after the impact. Model studies suggest that the impact also resulted in a global-scale “nuclear winter”, a period of cold and dark conditions, as fine particles in the upper atmosphere blocked sunlight for months or years. This may have severely affected photosynthesizing microorganisms, whose short life cycles coupled with the prolonged lack of sunlight may have led to major population declines if not extinctions. Further study of the immediate post-SIL strata may yield critical information on such effects. REFERENCES Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W.,Kissin, S.A., Fralick, P.W., and Hammon, A.L., 2005, Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology, 33: 193196. Addison, W.D., Brumpton, G.R., Davis, Don W., Fralick, P.W., and Kissin, S.A., 2010, Debrisites from the Sudbury impact event in Ontario, north of Lake superior, and a new age constraint: are they base surge deposits or tsunami deposits?, Geological Society of America Special Paper 465: 245-268. Cannon, W.F., Schulz, K.J., Horton, J.W. Jr., and Kring, D.A., 2010, The Sudbury impact layer in the Proterozoic iron ranges of northern Michigan, USA: Geological society of America Bulletin, 122, : 50-75. Collins, G.S., Melosh, H.J., Marcus, R.A., 2005, Earth Impact Effects Program: A web-based computer program for calculating the regional environmental consequences of a meteoroid impact on Earth. Meteoriteics and Planetary Science, 40:817-840. http://www.purdue.edu/impactearth/ Davis, D.E., 2008, Sub-million year age resolution of Precambrian igneous events by thermal extraction-thermal ionization mass spectrometer Pb dating of zircon: application to crystallization of the Sudbury impact melt sheet. Geology, 36: 383-386 Deutsch, A., Grieve, R.A.F., Avermann, M., Bischoff, L., Brockmeyer, P., Buhl, D., Lakomy, R., Muller-Mohr, V., Ostermann, M., and Stoffler, D., 1995, The Sudbury structure (Ontario, Canada: a tectonically deformed multi-ring impact basin. Geoligische Rundschau, 84: 697-709. Fralick, P., Grotzinger, J., and Edgar. L., 2012, Potential recognition of accretionary lapilli in distal impact deposits on mars: a facies analog provided by the 1.85 Ga Sudbury impact deposit, in Sedimentary Geology of Mars. SEPM Special Publication, 102: 211-227. Jirsa, M.A., Fralick, P.W, Weiblen, P.W., and Anderson, J.L.B., 2011, Sudbury impact layer in the western Lake Superior region. Geological Society of America Field Guides, 24: 147-169. Huber, M.S., McDonald, and Koeberl, C., 2014, Petrography and geochemistry of ejecta from the Sudbury impact event. Meteoritics and Planetary Science, 49: 1749-1768. Petrus, J.S., Ames, D.E., and Kamber, B.S., 2015, On the track of the elusive Sudbury impact: geochemical evidence for a chondritie or comet bolide. Terra Nova, 27: 9-20. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, H.J., and Edwards, C.T., 2007, Physical and chemical evidence of the 1850 Ma Sudbury impact event in the Baraga Group, Michigan. Geology, 35: 827830. 23 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geologic and geochemical attributes of the Beaver River Diabase and Greenstone Flow: Testing a possible intrusive-volcanic correlation in the 1.1 Ga Midcontinent Rift DOYLE, Michael S.1 and MILLER, James D. Jr.1 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 Over the last century, numerous geological studies have fairly well constrained the overall tectonomagmatic evolution of the Midcontinent Rift (MCR). However, until now, the correlation of the numerous flood basalts with their intrusive feeder systems has not been attempted. This study proposes one such correlation between two of the largest igneous bodies in the MCR: the Beaver River Diabase (BRD) intrusive complex in northeastern Minnesota and the Greenstone Flow (GSF) lava sheet in northern Michigan. The objective of this research is to test the validity of this link through a detailed analysis of the field relationships, petrographic characteristics, and geochemical attributes of these two units. The GSF is an enormous (at least 1,650 km3) lava sheet exposed over a ~5,000 km2 area on Isle Royale and the Keweenaw Peninsula in northern Michigan (White, 1960; Longo, 1984). The GSF forms the prominent ridge that runs the length of the Keweenaw Peninsula (~90 km) where it reaches a maximum thickness of 460 m (Cornwall, 1951). The GSF has been correlated across Isle Royale (Lane, 1893; Longo, 1984) where it reaches a maximum thickness of 260 m (Huber, 1973). The BRD is an extensive, composite dike and sill complex exposed over a ~600 km2 area in northeastern Minnesota. Perhaps the most intriguing feature of the BRD is the occurrence of numerous large (≤ 500m in diameter), lower crustal anorthosite xenoliths in the BRD (Miller and Green, 2002). That these diabase feeder dikes were at one time wide enough to accommodate such large blocks within several kilometers of the Earth’s surface implies that such conduits would most certainly have reached the surface and resulted in enormous outpourings of lava such as those that would have created the GSF. Field mapping and previous studies have shown that both the BRD and GSF are composite systems formed by multiple pulses of successively fractionated magma. BRD dikes and sills occur as ophitic olivine diabase, with 0.5 – 10 cm augite oikocrysts, which grades into coarser and more subophitic to intergranular olivine oxide gabbro in the medial portions of larger dikes and sills. A distinctive textural attribute of the ophitic diabase if the occurrence of clustered, often radiating, plagioclase laths. Within these dikes and sills are numerous, smaller composite intrusions of more highly fractionated lithologies (ferrodiorite to quartz ferromonzonite) that locally display modal layering and strong foliation. Contacts between dioritic rocks and the enclosing gabbros are sharp, but unchilled. In addition, ophitic olivine diabase locally occurs as altered xenoliths in the intermediate composite intrusions. These composite bodies are especially prevalent in the southern extent of the BRD where they are termed the Silver Bay Intrusions (SBI) (Miller and Green, 2002). Paces and Miller (1993) reported a U-Pb age of 1095.8 ± 1.2 Ma which is within analytical error of the date reported by Davis and Paces (1990) for the GSF (1094.0 ± 1.5 Ma). Similarly, the GSF can be divided into distinct lithological zones (from bottom to top): the lower ophite, heterolithic, upper ophite, and entablature. The upper and lower ophite zones are composed of ophitic olivine basalt with 0.1-4 cm augite oikocrysts and displaying clustered, often radiating, plagioclase laths. Occupying the central 1/4 to 1/2 of the lava sheet is the heterolithic zone, which is composed of coarser, subophitic to intergranular olivine oxide gabbro to subprismatic, locally foliated, ferrodiorite. Within the gabbroic/dioritic rocks of the heterolithic 24 �Proceedings of the 61st ILSG Annual Meeting - Part 1 zone occur numerous en echelon bodies of granophyre-rich lithologies (ferromonzodiorite to quartz ferromonzonite). As in the BRD-SBI intrusions, contacts between these rocks and the host gabbros/diorites are sharp and display no discernable chilled margins. And, as in the BRD intrusions, ophitic olivine basalt commonly occurs as inclusions in the heterolithic zone gabbros and diorites. Whole-rock geochemical analysis shows significant similarities in major and trace element compositions between BRD and GSF lithologies. REE patterns on primitive mantle-normalized diagrams are nearly identical for comparable rocks types of each unit. Gabbroic rocks are characterized by moderate LREE enrichment (La/Smn 1.54-2.49) and weak HREE fractionation (Gd/Ybn = 1.45-1.79). The more highly fractionated rocks of the SBI and GSF heterolithic zone show higher LREE enrichment (La/Smn = 1.86-2.54) and HREE fractionation (Gd/Ybn = 1.552.74), as well as strongly negative Sr anomalies, moderate Zr-Hf anomalies, and weak negative Eu anomalies. SEM-EDS analysis showed a similar range in pyroxene compositions between comparable rocks of the BRD and GSF. Pyroxenes within the gabbroic rocks of each unit were predominantly augite with lesser pigeonite and minor enstatite while the more highly differentiated lithologies of the SBI intrusions and GSF heterolithic zone were generally more Fe-rich (ferroaugite to ferrosilite). Olivine compositions tended to be more Fe-rich in BRD samples than those in the GSF ophites Fo28-69 and Fo50-66, respectively. Within the GSF, fresh olivine was only found in the upper and lower ophites so rocks of the heterolithic zone could not be compared with comparable rocks in the BRD-SBI. SEM-EDS analysis was also used to measure the An content in plagioclase phenocrysts in GSF samples. Three phenocrysts from GSF ophites of Isle Royale were found to have anomalously high anorthite contents (An71-81) with respect to the groundmass plagioclase (An30-60). The An content of these calcic phenocrysts (xenocrysts?) is consistent with those reported by Morrison (1983) for the anorthosites xenoliths in the BRD (An54-78) and could be indicative of a similar source. Based on the evidence obtained during this research, we propose the first ever intrusivevolcanic link between an MCR flood basalt and its intrusive feeder system. Based on the 1) overlap in U-Pb ages; 2) similar composite lithologies and contact relationships; 3) similar mineralogical and textural attributes, especially the occurrence of clustered plagioclase laths; 4) similar major and trace element compositions; 5) similar primary mineral chemistries; 6) similar An contents between anorthosite xenoliths in the BRD and plagioclase megacrysts in the GSF; and, 7) enormous volumes represented by each unit. Collectively, this data points to the conclusion that the BRD acted as the feeder conduit for the GSF. If this is the case, it more than doubles the total volume of the GSF making it perhaps the largest single lava flow on Earth. REFERENCES Cornwall, H. R. (1951). Differentiation in the lavas of the Keweenawan series and the origin of the copper districts of Michigan. Geological Society of America Bulletin, 62, 159–202. Davis, D. W. and Paces, J. B. (1990). Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system. Earth and Planetary Science Letters, 97(1-2), 54–64. Huber, N. K. (1973a). The Portage Lake Volcanics (Middle Keweenawan) on Isle Royale, Michigan. United State Geological Survey Professional Paper 754-C, C1–C32 Lane, A. C. (1893). Geological report on Isle Royale, Michigan. Geological Survey of Michigan, 6, 1–265. Longo, A. (1984). A correlation for a Middle Keweenawan flood basalt: The Greenstone Flow, Isle Royale and the Keweenaw Peninsula, Michigan. Michigan Technological University. 25 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Miller, J.D., and Green, J.C., 2002a, Geology of the Beaver Bay Complex and related hypabyssal intrusions. In Miller, J.D. Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.E., and Wahl, T.E., Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota. Minnesota Geological Survey Report of Investigations 58, p. 144-163. Morrison, D. A., Ashwal, L. D., Phinney, W. C., Shih, C., and Wooden, J. L. (1983). Pre-Keweenawan anorthosite inclusions in the Keweenawan Beaver Bay and Duluth Complexes, northeastern Minnesota. Geological Society of America Bulletin, 94, 206–221. Paces, J. B., & Miller, J. D. Jr. (1993). Precise U-Pb ages of Duluth Complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagmatic processes associated with the 1.1 Ga Midcontinent Rift System. Journal of Geophysical Research, 98, 13997–14013. White, W. (1960). The Keweenawan Lavas of Lake Superior, an example of flood basalts. American Journal of Science, 258-A, 367–374. 26 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Re-digitized public aeromagnetic data for the Baraga basin and surrounding region, Upper Peninsula, Michigan DRENTH, Benjamin, AILES, Chad, and ANDERSON, Eric 1 Crustal Geophysics and Geochemistry Science Center, U.S. Geological Survey, PO Box 25046 MS 964, Denver, CO, 80225 USA The public aeromagnetic database (Daniels et al., 2009) for Michigan’s western Upper Peninsula (UP) is widely regarded as unsuitable for intermediate- and detailed-scale geologic mapping and mineral exploration applications. There are several limitations of the data, including being available only in a native analog format, being acquired with too wide of a line spacing and too high of a terrain clearance, and being digitized at a lower level of detail than shown on original contour maps. This abstract describes a recent experimental effort to re-digitize sample aeromagnetic data from original contour maps in the greatest detail possible. The area chosen is the Proterozoic Baraga basin, containing the Eagle Ni-Cu deposit, and surrounding Archean rocks (Sheet 2 of Case and Gair, 1965). A fixed-wing total-field aeromagnetic survey was flown in the region in 1950, along north-south lines spaced 400 metres at a nominal terrain clearance of 150 metres (Case and Gair, 1965). After removal of an unspecified base level, the acquired data were interpolated onto contour maps with a minimum contour interval of 50 nT (see Case and Gair, 1965). A subsequent digitization effort from the contour maps followed at an unknown time, and the resulting digitized data are those publically available today from the USGS (e.g., Daniels et al., 2009). However, that digitization effort sampled the contour maps along only every other flightline, effectively simulating a survey with 800 metre line spacing. This resulted in very poor geologic resolution. The same problem plagues several other vintage aeromagnetic datasets acquired in the western UP. As an experiment we re-digitized a portion of this aeromagnetic dataset, sampling each contour from the original contour map and effectively capturing all of the available detail. The experiment is considered a success, as the resulting map is a far more effective representation of the region’s geology. As originally interpreted by Case and Gair (1965), many Keweenawan diabase dikes are imaged against a background of generally weakly magnetized Archean rocks and older Proterozoic metasedimentary rocks. The magnetic anomaly over the Eagle-hosting intrusion is difficult to pick out in the published data due to the wide data spacing, yet is readily apparent in the re-digitized data. In spite of this successful experiment, the recovered data still have several major and minor limitations that must be considered by interpreters. First, the survey was flown at too wide a line spacing (400 meters) and too far above the ground (~150 meters) for detailed geologic mapping and mineral exploration. Second, the minimum contour interval of 50 nT shown on the original contour maps means that more subtle anomalies and geologic details undoubtedly present in the flightline data will never be recoverable. Third, the exact terrain clearance of the magnetometer was not recorded, and in several localities may have varied significantly from the nominal 150 meter clearance. Finally, the base level removed from the magnetic data wasn’t recorded, meaning that the total field intensity and formal total field anomalies cannot be calculated. 27 �Proceedings of the 61st ILSG Annual Meeting - Part 1 REFERENCES Case, J.E., and Gair, J.E., 1965. Aeromagnetic map of parts of Marquette, Dickinson, Baraga, Alger and Schoolcraft Counties, Michigan, and its geologic interpretation: U.S. Geological Survey Geophysical Investigations Map GP-467. Daniels, D.L., Kucks, R.P., Hill, P.L., and Snyder, S.L., 2009, Michigan magnetic and gravity maps and data: a website for the distribution of data: U.S. Geological Survey Data Series 411 available only online at http://pubs.usgs.gov.ds/ds411. 28 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Petrographic Analysis of Felsic Tuffs within the Neoarchean Soudan Member of the Ely Greenstone Formation, NE Minnesota ESSIG, Espree1, HUDAK, George 2, PIGNOTTA, Geoff3, and LODGE, Robert3 1 Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 2 Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN 3 Department of Geology, University of Wisconsin Eau Claire, Eau Claire, WI The Vermilion District of northeastern Minnesota contains one of the classic greenstone belts in the United States, and is composed of a wide variety of greenschist facies metamorphosed Neoarchean volcanic, sedimentary, and intrusive rocks that comprise the southwestern part of the Wawa Abitibi Terrane (Stott et al., 2007). The Ely Greenstone Formation occurs within the Western Vermilion District and is composed of calc-alkaline to tholeiitic massive to pillowed basalt, andesite, dacite, and rhyolite lava flows and volcaniclastic rocks (Lower Member); Algoma-type banded iron formations with interbedded tholeiitic massive to pillowed basalt lava flows, rhyodacitic to dacitic tuffs, and polymict volcaniclastic rocks (Soudan Member); and pillowed to massive tholeiitic basalt lava flows and interbedded Algoma-type banded formation horizons (Upper Member; Peterson and Jirsa, 1999). The purpose of this study is to identify and characterize potential felsic tuff horizons that are interbedded with Algoma-type banded iron formations within the Soudan Member of the Ely Greenstone Formation. This has been accomplished by detailed field mapping, petrographic studies, scanning electron microscopy (SEM) studies, and lithogeochemical studies. In addition to texturally, mineralogically, and chemically characterizing potential felsic tuff units, our research seeks to determine if the felsic tuff units could potentially yield age dates by means of future U/Pb geochronological studies in a manner similar to that which has been done in the Abitibi Belt in northeastern Ontario (Thurston et al., 2008). Despite the long history of geological studies in the Western Vermilion District, relatively few absolute age dates are present (Fig. 1A; Peterson et al., 2001; Lodge et al., 2013), and no absolute age dates exist for the Soudan Member of the Ely Greenstone Formation. Detailed mapping has identified a light gray, 30-50cm thick, laminated to thinly-bedded, possibly resedimented felsic tuff horizon that is interlayered with laminated to very thinly-bedded magnetite-chert Algoma-type banded iron formation within the uppermost 25 meters of the Soudan Member in the central part of Lake Vermilion State Park. In thin section, the tuff is sparsely quartz-phyric, and comprises a matrix of fine-grained, recrystallized polygonal quartz with up to 1%, up to 1mm in diameter subhedral, recrystallized quartz phenocrysts. Hydrothermal alteration of the tuffs varies from moderate to intense (up to 75% alteration minerals), with the greenschist-facies metamorphosed synvolcanic hydrothermal alteration assemblage now composed of variable amounts of iron carbonate (siderite, ankerite), actinolite, chlorite, and various epidote-group minerals (pistacite, clinozoisite/zoisite). Due to the extremely fine-grained texture of the tuff, and the locally pervasive hydrothermal alteration within the tuff, searching for zircons using standard petrographic analysis has proven to be difficult. On-going SEM analysis will be used to constrain the mineralogy of the existing alteration and to seek out zircons in a more systematic manner. Lithogeochemical analysis and lithogeochemical classification by means of immobile trace elements (e.g., Pearce, 1996) is ongoing. 29 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: A: Generalized section of the Ely Greenstone Formation in Lake Vermilion State Park (modified from Hudak and Peterson, 2014); B: Field appearance of felsic tuff unit within the uppermost 25 meters of the Soudan Member; C and D: Cross-polarized and plane polarized appearance of hydrothermally altered felsic tuff. REFERENCES Hudak, G.J., and Peterson, D. ., 2014, Non-Ferrous Mineralization Associated with the Wawa-Abitibi Terrane and Duluth Complex Cu-Ni-PGM Deposits, NE Minnesota: Society of Economic Geologists, SEG Guidebook Series Guidebook 47, 150 p. Lodge, R. W. D., Gibson, H. L., Stott, G. M., Hudak, G. J., Jirsa, M. A., and Hamilton, M. A., 2013, New U-Pb geochronology from the Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa Subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research, v. 235, p. 264-277. Pearce, J. A., 1996, A user’s guide to basalt discrimination diagrams: in Wyman, D. A., ed., Trace Element Geochemistry of Volcanic Rocks: Applications for Massive Sulphide Exploration: Geological Association of Canada, Short Course Notes, v. 12, p. 79-113. Peterson, D. M., Gallup, C., Jirsa, M. A., and Davis, D. W., 2001, Correlation of Archean assemblages across the U.S.- Canadian border: Phase I geochronology: 47th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 47, Part 1 – Programs and Abstracts, p. 77-78. Peterson, D. M., and Jirsa, M.A., 1999, Bedrock geologic map and mineral exploration data, western Vermilion district, St. Louis and Lake Counties, northeastern Minnesota: MGS Miscellaneous Map M-98, scale 1:48,000. Stott, G., Corkery, T., Leclair, A., Boily, M., and Percival, J., 2007, A revised terrane map for the Superior Province as interpreted from Aeromagnetic Data: 53rd Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 53, Part 1 – Program and Abstracts, p. 74-76. Thurston, P. C., Ayer, J. A., Goutier, J., and Hamilton, M. A., 2008, Depositional Gaps in Abitibi Greenstone Belt Stratigraphy: A Key to Exploration for Syngenetic Mineralization: Economic Geology, v. 103, p. 1097-1134. 30 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Characterization of secondary minerals formed on weathered Duluth Complex Cu-Ni-PGE deposit rock and implications for controls on metal mobility FIX, Paul M.1, and DIEDRICH, Tamara R.2 1 Department of Earth and Environmental Sciences, University of Minnesota, Duluth, MN, USA 2 Barr Engineering Company, Duluth, MN, USA Secondary minerals and other weathering products can serve as an important control on the concentration of trace metals in mine-impacted waters (for example, Jambor, 2003). Ultimately, aqueous metal concentrations will reflect combined effect of constituent release from primary minerals (e.g. sulfides), as well as, attenuation during weathering from mechanisms such as precipitation of secondary phases, sorption onto solid phase surfaces, and co-precipitation. We characterized weathering products formed on weathered exposures of mineralized Duluth Complex to specifically investigate possible solubility controls for Cu and Ni. Combining these solid phase characterization results with both standard-method mine waste studies (see Lapakko, 2012) and numerical modeling should yield an improved understanding of the mobility of trace metals that would be released during weathering of potential future waste rock. Opportunistic sampling was conducted at five Duluth Complex exposures at the Mesaba deposit (currently held by Teck American and formerly known as the Babbitt deposit). Exposures included both natural outcrops and a railroad cut. Six samples were collected with the intent of capturing the variety of visual alteration. Powder X-ray diffraction (XRD) was conducted on weathering products isolated by scraping the outer surfaces of hand sample specimens. XRD of iron-rich material (rusty coatings) generally did not produce diffraction patterns that allowed phase identification, suggesting the material structure is very poorly crystalline to amorphous. However, in one case, poorly crystalline goethite (FeOOH) was identified and appeared to be a product of sulfide mineral replacement. In addition, the secondary minerals malachite (Cu2(CO3)(OH)2), rozenite (FeSO4·4H2O), alunogen (Al2(SO4)3·17H2O), and epsomite (MgSO4·7H2O) were identified. The morphology and semi-quantitative chemistry of weathering products was characterized using a JEOL JSM-6590LV scanning electron microscope, combined with an INCA X-ACT energy dispersive spectroscopy system at University of Minnesota- Duluth. SEM observations of rusty coatings reveal micro-scale banded (alternating Fe and Si rich) features (Figure 1). The coatings commonly contain Fe and Si as major components and variable amounts of Al, S, and Cu as minor elements. It should be noted that Ni was not found to be commonly associated with these features. This could be due to sub-detection limit (< 0.1 wt. %) quantities or absence of Ni. Sorption of trace metals by hydrous ferric oxides in mine-water systems is common. SEM-EDS analyses indicate Ni and Cu can be associated with sheet silicate minerals, in concentrations that can exceed several percent by weight. Research on similar materials found Nirich metallic particles along laths of serpentine and chlorite (Suárez, 2011) and sorption of Ni(OH)2 onto primary mineral grains (Plante, 2010). It may be possible that Cu and Ni were incorporated during late stage deuteric alteration and not surface weathering. Planned analyses include SEM-EDS of non-weathered rock from the same lithology to determine if trace metal enrichment of sheet silicates is unique to weathered samples as well as TEM analyses to determine nano-scale mineral properties. Collectively, the mineral characterization techniques employed provide evidence for attenuation of constituents released during weathering of Duluth Complex rock by means of: (1) 31 �Proceedings of the 61st ILSG Annual Meeting - Part 1 secondary mineral precipitation (malachite and sulfate salts); (2) incorporation of Cu and Ni by sheet silicates (likely, but not necessarily occurring during sub-aerial weathering); and (3) formation of iron-oxide rich coatings which may retain trace metals (especially Cu) though adsorption. The degree to which trace metals are attenuated will be function of drainage pH, total iron content, trace metal content, and reactive surface area among other variables. +@A&)!"#$#%& '&()* +, -.(/ 4" 2(5 0# 2(6 :; /(< =, 3(9 78 /(2 7? 9(1 7" 9(3 > .9(3 +@A&)6 !"#$#%& '&()* 0# 12(3 +, 1(1 78 3(9 + -(2 4" -(1 > 19(2 Figure 1: SEM-Backscatter electron image of a weathered Duluth Complex sample. Lower left region shows banding typical of iron rich surface rinds. Upper right phase appears to be a weathered sheet silicate with significant Cu enrichment. Similar phases in our samples have been observed to contain up to 4 wt. % Ni. Iron rich phases (light strips) can also be seen where sheets have parted. Spots indicate locations of semi-quantitative SEM-EDS compositional data (right) and are sized to the approximate analytical volume at these working conditions (15 kV). REFERENCES Jambor, J.L. 2003. Mine-waste mineralogy and mineralogical perspectives of acid-base accounting, In: Jambor, J., Blowes, D., Ritchie, A. (Eds) Environmental Aspects of Mine Wastes. Mineralogical Association of Canada. Short Course Series 31, pp. 117-145. Lapakko, K., Antonson, D. A. 2012 Duluth Complex Rock Dissolution and Mitigation Techniques: A summary of 35 years of DNR research, Minnesota Department of Natural Resources, (p. 56). Plante, B., Benzaazoua, M., Bussière, B. 2010. Study of Ni sorption onto Tio mine waste rock surfaces. Applied Geochemistry 25, 1830–1844. Suárez, S., Nieto, F., Velasco, F., Martín, F.J. 2011. Serpentine and chlorite as effective Ni-Cu sinks during weathering of the Aguablanca sulphide deposit (SW Spain). TEM evidence for metal-retention mechanisms in sheet silicates. European Journal of Mineralogy 23, 179–196. 32 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Lateral geochemical gradients and physical processes associated with the genesis of iron formations: Examples from the Paleoproterozoic to Mesoarchean of Superior Province FRALICK, Philip Department of Geology, Lakehead University, Thunder Bay, ON, Canada, P7B 5E1, The presence of Fe, Si, Mn and P in mineral phases with precursors that were quasi-stable precipitates from the overlying water column necessitates that disequilibrium conditions existed during their formation. Commonly a two box model is used to explain iron precipitation in the Precambrian with an oxic to sub-oxic phytoplankton-rich or CO2-rich surface layer driving iron precipitation in the underlying anoxic, Fe+2-rich ocean. But what is the evidence for this model? 1) On the wide, Paleoproterozoic Gunflint-Mesabi shelf carbonate iron formation (IF) dominated the shallow areas, whereas oxide IF accumulated in more offshore locations. Storm induced geostrophic flows delivered oxygenated inner shelf waters to the offshore instigating formation of the discrete iron to iron+manganese+silica+mud laminae that compose the fine-grained IF. Here frequent storm mixing would have destroyed vertical stratification and Fe+2 transport across the 150km wide shallow shelf indicates anoxic conditions across the shelf during fair-weather times. 2) The Neoarchean delta-top IFs in the Lake St Joseph area record extremely rapid accumulation of iron hydroxides during limited sediment flux to portions of distributary mouth bar complexes. IF accumulated on short-lived reactivation surfaces of gravel bars to ripples in water depths of a few meters. No IF accumulated further offshore. Its chemistry is virtually identical to deep-water IF deposits. The IF was probably formed by high nutrient flux to the shallows promoting cyanobacteria. 3) Carbonate was deposited on the Mesoarchean, stromatolitic, oxygenated Steep Rock platform, while the generation of free oxygen caused Fe and Mn precipitation offshore. Onshore water movement, probably driven by storm events, deposited iron-rich layers in the limestone and shifted the offshore from Fe to chert or mud accumulation. In these examples IF deposition was dependent on lateral, not vertical, geochemical differences augmented by storm induced mixing. 33 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Comparison of PGM assemblages for the Marathon, Geordie Lake and Area 41 deposits, Coldwell Alkaline Complex, Ontario GOOD, David1, CABRI, Louis 2, and AMES, Doreen 3 1 Department of Earth Sciences, University of Western Ontario, London, ON N6A 5B7 2 Cabri Consulting Inc., 700-702 Bank Street, PO Box 14087, Ottawa, ON, K1S 5P5 3 Geological Survey of Canada, 750-601 Booth St., Ottawa, ON, K1A 0E8 Numerous Cu-PGE deposits in the Coldwell Alkaline Complex exhibit widely varying mineralization styles and Cu/Pd and Pd/Pt ratios (Ruthart, 2012; Meghji et al., 2013; Good et al., 2015). However, the host gabbro and ultramafic bodies are believed to be co-genetic and differences between the deposits may be explained by variations in mineralizing processes and the respective magma conduit setting. This study presents results for heavy mineral separates from 5 mineralized zones within three deposits: W horizon and Main zone in the Marathon deposit, Main zone at the Geordie Lake deposit, and the Main and PGE-enriched zones at the Area 41 occurrence. A total of over 9000 PGM comprising 46 PGM species, numerous unknown PGM, and 7 Au/Ag minerals were identified. The mineral separates were prepared by two different methods on two different groups of samples. The first group was prepared by hydroseparation (HS) of screened size-by-size composites and the second group by electric-pulse disaggregation (EPD) of drill core (e.g., Cabri et al. 2008). All mineral separates were mounted on polished sections and analysed by SEM-EDS techniques to characterize the platinum group minerals (PGM). The sum of measured surface areas for each mineral are collated to provide estimates of mineral abundances. Large representative sample sets were prepared from each location. In group one, mineral separates for three composite samples with similar Cu and Pd abundances (Fig. 1a), one from each intrusion, a total of 513 precious metal grains were found and characterized in 15 sized monolayer polished sections. In group two, mineral separates for samples with relatively high PGE abundances were prepared from 12 pieces of drill core, 4 from each of the Main zone, W Horizon, and Area 41 intrusion (Fig. 1b). The results from group 1 show that three mineralized zones (Main zone, W horizon and Geordie Lake) have distinct precious mineral signatures (Fig. 2), as expected based on the variation of Cu/Pd and Pd/Pt values, and differences between local intrusive settings and 34 �Proceedings of the 61st ILSG Annual Meeting - Part 1 mineralization styles. The Pd PGM in the Main zone sample are dominated (mass%) by arsenian PGM (80%) with less bismuthides (12%); in the Geordie Lake sample, arsenides dominate (51%), followed by arsenian-antimonian (33%), and arsenian-nickeloan PGM (16%); and in the Area 41 sample, plumbian PGM dominate (47%), followed by arsenian PGM (17%), arsenian-antimonian PGM (14%), bismuthides (13%), stannides (6%), and tellurides (3%). The Pt PGM for the Main zone and Geordie Lake are the same, where sperrylite predominates, but are very different for the Area 41 sample where Pt alloys are abundant: sperrylite (~53%), isoferroplatinum (~31%), and tetraferroplatinum (~16%). Results are summarized in Figure 2. The PGM assemblage at Area 41 resembles that for the W Horizon and is consistent with the very wide range and notably low Cu/Pd values present at both locations. Further, the host gabbros exhibit similar petrographic features and trace element abundances that suggest they formed by similar processes. REFERENCES Cabri, L.J., Rudashevsky, N.S., Rudashevsky, V.N., and Oberthür, T., 2008, Electric-Pulse Disaggregation (EPD), Hydroseparation (HS) and their use in combination for mineral processing and advanced characterization of ores. Canadian Mineral Processors 40th Annual Meeting, Proceedings, Paper 14, 211-235. Good, D.J., Epstein, R., McLean, K., Linnen, R.L. & Samson, I.M., 2015, Evolution of the Main Zone at the Marathon Cu-PGE sulfide deposit, Midcontinent Rift, Canada: spatial relationships in a magma conduit setting. Economic Geology (in press). Meghji I., Linnen R.L., Samson I.M., Ames D.E., Good D.J., 2013, The character and distribution of Cu-PGE mineralization at the Geordie Lake Deposit within the Coldwell Complex, Ontario, GAC-MAC, Poster presentation. Ruthart R., 2012, Characterization of High-PGE, Low-Sulphur Mineralization at the Marathon PGE-Cu Deposit, Ontario, M.Sc. thesis, University of Waterloo, 145 p. 35 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Significance of LREE-enriched mantle source to genesis of basalt in the Coldwell Alkaline Complex, Midcontinent Rift, Ontario GOOD, David1, HOLLINGS, Peter 2, CUNDARI, Robert 2, 3, and AMES, Doreen 4 1 Department of Earth Sciences, University of Western Ontario, London, ON N6A 5B7 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada 3 Ontario Geological Survey, Ministry of Northern Development, Mines and Forestry, Suite B002, 435 James St. South Thunder Bay, ON P7E 6S7 Canada 4 Geological Survey of Canada, 750-601 Booth St., Ottawa, Ontario, K1A 0E8 At least three distinct basaltic packages occur within the Coldwell Alkaline Complex: Lower and Upper basalt units located within the Eastern Gabbro Suite, and the Coubran basalt. Geochemical evidence suggests the Coubran basalt is co-genetic with the Two Duck Lake gabbro, a late phase of the Eastern Gabbro with an age of 1108±1 Ma (Heaman et al., 2007) and thus formed as part of the Early Magmatic stage of the rifting event. The Upper and Lower basalt were re-crystallized during pyroxene hornfels grade metamorphism during intrusion of the Eastern Gabbro suite, and are older than 1108 Ma and possibly represent magma formed during the Initiation stage of the rifting event. Major element abundances for the Coldwell basalt units are comparable to other Early Stage basaltic units of the Midcontinent rift. Based on Mg number and Ni abundance, the three groups, listed in order from most primitive to most evolved, are Lower basalt, Upper basalt and Coubran basalt. Spider diagrams were prepared following the method of Pearce (2008) whereby data are normalized in two steps: first by N-MORB and then by Tb (Fig. 1). Note that we normalize to Tb and not Ti. All Coldwell units show depleted HFSE relative to LREE, similar to that demonstrated by Groups 1 and 2 at Mamainse Point. The units listed, from most to least depleted, are Coubran basalt, Upper basalt and Lower basalt. Although LREE abundances in the Coldwell basalts are consistent with OIB source, the regular pattern of LREE enrichment relative to HFSE indicates a more complicated origin and cannot be explained by crustal contamination. Crustal contamination, if present, would have resulted in elevated SiO2 and Zr abundances as well as higher Th/La, but these affects were not observed. Further, La enrichment relative to Zr in the Coubran basalt (Fig. 2) cannot easily be explained by crustal contamination. Therefore, it seems more likely that the mantle source was LREE enriched and HFSE abundances are a better indicator of either MORB or OIB source. The HFSE suggest the Coubran basalt magma was intermediate between E-MORB and OIB. Similarly, the Lower and Upper basalt groups show HFSE signatures that are intermediate between E-MORB and N-MORB. Coldwell units exhibit Gd/Yb that are intermediate between E-MORB and OIB. These units, in order of decreasing Gd/Yb, are Lower basalt, Coubran basalt and Upper basalt, whereby Lower basalt resembles that of OIB and Upper basalt resembles MORB. These results suggest a progressive change in the mantle source for the Coldwell basaltic magmas. The mantle source areas have varying degrees of LREE enrichment, possibly introduced by an earlier subduction event as discussed by Shirey et al. (1994) for the MPVG. Our observations for the Coldwell basalts may be explained by partial melting of sources that, from oldest to youngest show, deep depleted mantle having moderate enriched LREE signature (Lower basalt); shallow depleted mantle with significant LREE enrichment (Upper basalt); and finally, 36 �Proceedings of the 61st ILSG Annual Meeting - Part 1 enriched mantle from intermediate depth with significant but heterogeneous LREE enrichment (Coubran basalt). A model to test origin of Eastern gabbros from heterogeneous LREE enriched mantle source similar to that which produced Coubran basalt is proposed in Figure 2. Figure 1: Average data for Coubran basalt and Lower and Upper basalts compared to averages for MPVG 1, 2 and 3a (after Shirey et al. 1994). MORB, OIB and E-MORB after Sun and McDonough (1989).   Figure 2: Model for derivation of Eastern gabbros based on contamination trend of Coubran basalt REFERENCES Cundari R., 2012, Geology and geochemistry of Midcontinent rift-related igneous rocks: M.Sc. thesis, Thunder Bay, ON, Lakehead University, 122 p. Good, D.J., Epstein, R., McLean, K., Linnen, R.L. & Samson, I.M., 2015, Evolution of the Main Zone at the Marathon Cu-PGE sulfide deposit, Midcontinent Rift, Canada: spatial relationships in a magma conduit setting. Economic Geology (in press). Heaman, L.M., Easton, M., Hart, T.R., Hollings, P., Macdonald, C.A., and Smyk, M., 2007, Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario: Canadian Journal of Earth Sciences v. 44, p. 1055–1086. Pearce, J.A., 2008, Geochemical fingerprinting of oceanic basalts with applications to ophiolite classification and the search for Archean oceanic crust. Lithos, v. 100, p. 14-48. Shirey, S.B., Klewin K.W., Berg, J.H. and Carlson, R.W., 1994, Temporal changes in the sources of flood basalts: Isotopic and trace element evidence from the 1100 Ma old Keweenawan Mamainse Point Formation, Ontario, Canada, Geochimica et Cosmochimica Acta, V. 58, P 4475-4490. 37 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Preliminary 3D model of the Midcontinent Rift System in western Lake Superior region GRAUCH, V.J.S., POWERS, Michael H., ANDERSON, Eric D., and CANNON, William F., U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 Over the past several decades, geophysical models have played a large part in developing our understanding of the Midcontinent Rift System (MRS). Technical advances in modeling capabilities, expanded and improved data coverage, and renewed interest in the mineral resources of the MRS provide the motivation for new attempts at 3D digital modeling of the MRS. Efforts are focused on modeling the structure and configuration of sedimentary and volcanic basins in the western Lake Superior region (Fig. 1). Previous workers used geophysical data to identify extremely thick volcanic and sedimentary sections that commonly have unconformable contacts against intervening structural highs. An improved 3D model of the area helps visualize the relations, provides mechanisms to test hypotheses about tectonic history and the spatial distribution of mineralization, and helps identify areas where more detailed analysis is required. The 3D model is regional and intended to show broad variations in geology. Only three generalized, MRS-related geologic packages are represented, following previous 3D models of Allen (1994; Allen et al., 1997). The packages, from oldest to youngest, are 1) undivided Keweenawan plutonic and volcanic rocks, 2) Oronto Group sedimentary rocks, and 3) Bayfield Group and equivalent sedimentary rocks. A fourth package represents undivided pre-rift rocks (basement). In addition, three major fault systems are modeled: the Douglas, Lake Owen, and Keweenaw fault systems (Fig. 1). The modeling strategy involves digitizing the bottoms of the rift-related geologic packages, fault locations, and general orientation data along 2D sections and from a geologic map. The 3D modeling software then connects the digitized points into surfaces and volumes in 3D space, using simple geologic rules for stratigraphic and onlap relations and for the lateral extents of the influences of faults. Steps that have been accomplished for the new 3D model are the following. 1. Images were captured from published 2D geophysical models, interpreted seismic-reflection sections, and geologic cross-sections. They were geo-registered and hung in 3D model space by projecting them onto 20 different sections crossing the model area. Digital and analog geologic maps from several sources provided information in plan view. 2. Contacts between geologic packages were determined for available, industry seismicreflection time sections using surfaces derived from previous 3D models (Allen and others, 1997) as guides. In addition, selected, proprietary industry seismic-reflection data from the Bayfield Peninsula, which only a few workers have previously seen, were licensed to the U.S. Geological Survey. Time was converted to depth for the sections using root-meansquare velocities derived from industry data processing. 3. Selected points representing fault surfaces and the bases of the geologic packages were digitized, then rendered into 3D volumes and surfaces by the modeling software. A 3D perspective view of the modeled volume of the Oronto Group is shown in Figure 2. Although this and the previous model (Allen and others, 1997) are intentionally very similar, there are significant differences. First, the newly acquired seismic data provided corrections for several significant errors in the current digital rendition of the previous 3D model that would not have been known other-wise. Second, due to advances in software capabilities, faults are now properly represented rather than handled by gridded surfaces. Finally, even though all apparent constraints from 2D sections were met, the Oronto Group in the Bayfield Basin does not extend as far south underneath the Bayfield Group in the new model as it does in the old one. This discrepancy 38 �Proceedings of the 61st ILSG Annual Meeting - Part 1 brings up the difficult problem of determining the contact between the Bayfield and Oronto Groups from geophysical data, and is yet to be resolved. Future work involves more extensive analysis of existing seismic data in conjunction with gravity and magnetic modeling; division of intrusive from extrusive Keweenawan igneous rocks; and addition of layers to the volumes of the generalized geologic packages to better represent dips, stratigraphic variations, and unconformities. REFERENCE Allen, D. A., Hinze, W. J., Dickas, A. B., and Mudrey, M. G., Jr., 1997, Integrated geophysical modeling of the North American Midcontinent Rift System: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota, in Ojakangas, R. W., Dickas, A. B., and Green, J.C., ed., Middle Proterozoic to Cambrian Rifting, Central North America: Geological Society of America Special Paper 312, p. 47-72. Figure 1: Generalized rock units of the Midcontinent rift system in the western Lake Superior region and area covered by the 3D model. Geographic boundaries are shown by dashed lines. Figure 2: 3D perspective view of the modeled base of the Oronto Group (green) and surface representing the Douglas fault (dark purple) for the model area located on Fig. 1. View is to the west-northwest. 39 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geological and geochemical reconnaissance for rare earth element (REE) mineralization in Minnesota HAUCK, Steven, HEINE, John, SEVERSON, Mark, POST, Sara, CHLEBECEK, Sarah, MONSON GEERTS, Steven, ORESKOVICH, Julie, GORDEE, Sarah, and HUDAK, George Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Duluth, MN 55811 The purpose of this study was to: 1) collect rock samples from across Minnesota and assay them for Rare Earth Elements (REEs); 2) evaluate the assay results; and 3) via a combination of acquisition and evaluation of both new data and historical geochemical data, identify locations or regions within Minnesota that possess anomalous REE concentrations that may warrant further characterization for potentially economic REE mineral deposits (Hauck et al., 2014). Currently, approximately 90-95% of the world’s REE processing and sales are controlled by China. REEs are vital to the U.S. economy, particularly in components used by the U.S. military, in windmills and other equipment that use REE-containing magnets, and in green economy products such as hybrid vehicles. This study is multi-dimensional, and includes: 1) collection of new REE geochemical data through detailed geological mapping, sampling, and analysis; 2) compilation of previously published REE data from a variety of scientific publications; and 3) re-analysis of one historical sample containing anomalous REE contents to confirm the validity of the historic analysis. Resulting from this study are: 1) a new, detailed 1:5,000 scale geologic map in NE Minnesota showing anomalous REE contents; 2) a new lithogeochemical dataset containing 287 samples compiling both new and historic REE data to provide the exploration industry up-to-date lithogeochemical resource data for designing future REE exploration programs in the State; 3) new maps illustrating both historic and new geochemical sampling location and 4), a new interpretation of areas that may host anomalous REE mineralization. Based upon currently mined REE ore deposits, igneous rocks or ionic clays are most likely to contain anomalous REE contents, and, therefore, sample collection concentrated primarily on silica-rich igneous rocks (rocks containing >67% SiO2). Two hundred twenty-two rock samples were collected. Of these, 147 rock samples were analyzed. Based upon the availability of outcrops and diamond drill core samples, the majority of the samples were from St. Louis, Lake of the Woods, Koochiching, and Lake Counties. The rock samples were analyzed at Acme Labs, Vancouver, B.C., for multi-element chemical analyses, including a complete suite of REEs+Y+Sc. The geochemical data received was combined with lithogeochemistry from a previous study (Klenner et al., 2012) to provide as complete a dataset of REE values of Minnesota rocks for interpretation as possible. This combined data set contains over 280 REE analyses, with sample analyses from diamond drill holes comprising ~33% of the combined geochemical assays and ~67% of the data were from outcrop samples from all regions of the state. This compiled data set indicated that 25 samples from around the State had TREEs > 425 ppm. The highest seven REE analyses occurred in samples from NE MN in Koochiching and St. Louis Counties. Lac Qui Parle County, in SW MN, contained the only anomalous sample outside of the NE section of the State, and this sample contained 760.36 ppm TREE. The most promising sample in this study was identified from a review of publications on Minnesota geology. Morey and McDonald (1989) reported a highly anomalous sample, GSP-47, 40 �Proceedings of the 61st ILSG Annual Meeting - Part 1 from near Ray, MN that had a TREE analysis of 11,313.50 ppm and 1,100 ppm Thorium (a common component of REE mineralization). Although this original TREE analysis was incomplete (some heavy rare earth elements (HREEs) were missing), the location of the sample, as well as a powdered sample split from the original analysis were both available, and afforded the NRRI the opportunity to re-evaluate the original sample, which provided an almost duplicate REE and Th analysis, and added the missing HREEs. The sample location is west of Ray, MN. The Minnesota Geological Survey provided a split of the original sample. Reanalysis confirmed its anomalous nature, i.e., 11,139.46 ppm TREE and 1,162.3 ppm Th. Upon confirmation of the original anomalous REE contents of GSP-47, detailed geological mapping (1:5000 scale) and additional geochemical sampling was done west of Ray, MN. The sample site was mapped and sampled (8 samples), and two samples assayed had > 4,000 ppm TREE and > 425 ppm Th, confirming the anomalous nature of this geologic area. Further work is recommended to more fully characterize the nature of the geological areas associated with anomalous TREE and Th contents identified in this study. Such work recommended for future exploration and characterization includes detailed airborne and ground radiometric surveys of exposed rock outcroppings of favorable REE host rocks, followup/confirmation assaying, detailed geologic mapping, and, if warranted, diamond drilling. As well, future geologic and geophysical studies should be conducted to further characterize, and if possible, confirm the existence of carbonatites in the southwest part of the state as suggested by Southwick (2014). Such carbonatites are currently the major source of REEs in China and the U.S. (Molycorp) and may afford Minnesota an excellent target for future REE exploration and characterization. REFERENCES Hauck, S., Heine, J., Severson, M., Post, S. Chlebecek, S. Monson Geerts, S., Oreskovich, J., Gordee, S., and Hudak, G., 2014, Geological and geochemical reconnaissance for rare earth element mineralization in Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-2014/39, 572 p. Klenner, R., Gosnold, W., Heine, J.J., Severson, M.J., Hauck, S.A., Hudak, G., and Fosnacht, D.R., 2012, New Heat Flow Map of Minnesota Corrected for the Effects of Climate Change and an Assessment of Enhanced Geothermal System Resources: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-2012/01, 109 p. Morey, G.B., and MacDonald, L.L., 1989, Analytical Results of the Public Geologic Sample Program, 1987-1989 Biennium: Minnesota Geological Survey, Information Circular 29, 66 p. Southwick, D.L., 2014, Reexamination of the Minnesota River Valley Subprovince with Emphasis on Neoarchean and Paleoproterozoic Events: Minnesota Geological Survey Report of Investigations 69, 52 p. 41 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Petrological and geochemical evaluation of the Sturgeon Falls Igneous Body and its relationship with the Penokean Orogenic Belt HAYNES, Jonathan1, THAKURTA, Joyashish1, and QUIGLEY, Tom2 1 Department of Geosciences, Western Michigan University, 1903 W. Michigan Ave. Kalamazoo, MI 49048. 2 Aquila Resources Inc., 414 10th Avenue, Suite 1,Menominee MI 49858 The Sturgeon Falls Igneous Body (SFIB) is a mafic to ultramafic intrusion located along the Michigan-Wisconsin border just south of the town of Norway, MI. The SFIB is bounded to the north by the Niagara Shear Zone and the Michigamme Formation (Schulz and Cannon, 2006) and to the south by an unnamed thrust fault zone and the Quinnesec Formation (Sims and Schulz, 1993). Field mapping has shown that the SFIB is composed almost entirely of metagabbro, with isolated outcrops peridotite and serpentinite. The metagabbro reached greenschist facies (Prinz, 1959), and is mostly composed of plagioclase, clinopyroxene, and hornblende. Hornblende is present as both a primary and secondary alteration mineral. The level of alteration within this rock is spatially varied within the SFIB. The portions of the SFIB near the fault zones have been metamorphosed to the point where their original fabric and mineral composition has been lost. This area has been named the Heterogeneous Altered zone. It is rich with secondary amphibole, with weak foliation often present. Prinz (1959) observed that the metagabbro cut the ultramafic rocks indicating that they were formed as part of an earlier magmatic event. Schulz and LeBerge (2003) proposed that the SFIB is part of a larger suprasubduction ophiolite sequence that formed along with the Pembine-Wausau Terrane. Major and trace element geochemistry of approximately 35 samples was compared to other known suprasubduction zone ophiolites, as well as island arcs suites. The results show that the major element geochemistry of the SFIB matches well with both suprasubduction zone ophiolites and island arc terranes. Trace element geochemical signatures such as enrichment in light REE and LILE elements are often used to distinguish suprasubduction zone ophiolites (Shervais, 2001), however in this case, these methods prove problematic for several reasons. First, light REE are highly mobile during metamorphism (Yumal, 1996). Secondly, these characteristics are also common to Island Arcs, so they are not in themselves sufficient to distinguish an ophiolite sequence. In order to identify an ophiolite sequence, a comprehensive study of the lithology, structure, and geochemistry must be used, one method by itself is insufficient. While the geochemistry of the SFIB is roughly in agreement with the ophiolite theory, the evidence available for consideration is not sufficient to distinguish the SFIB as an ophiolite sequence as opposed to an arc related intrusion. 42 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Representative geochemistry from each rock unit within the SFIB Sample   Lithology   SiO2   Al2O3   Fe2O3   MnO   MgO   CaO   Na2O   K2O   P2O5   Cr2O3   TiO2   BaO   LOI   Total                                 SF-­‐149   SF-­‐097   CY-­‐05     MG   HET   UM     46.75   52.49   39.03     17.83   14.61   0.49     11   10.42   11.29     0.166   0.18   0.17     8.14   8.94   36.2     10.802   3.956   0.1     2.41   4.39   0.05     0.27   1.25   <0.01     0.028   0.066   <0.01     0.08   0.02   0.75     0.49   0.67   0.01     <0.004   0.03   <0.01     2.82   3.63   11.85     100.8   100.66   99.94                 Table 1 is a representative geochemistry from each rock unit within the SFIB shown as oxides. MG= metagabbro, HET= Heterogenous Altered Zone, UM= Ultramafic. Concentrations expressed in weight %. REFERENCES Prinz, W.C., Geology of the southern part of the Menominee district, Michigan and Wisconsin: US Geological Survey open-file report. April 17, 1959, 221p. Schulz, K. J., & Cannon, W. F. (2006). The Penokean orogeny in the Lake Superior region. Precambrian Research, 157, 4-25. Schulz, K., and LaBerge, G. (2003). Pembine-Wausau Magmatic Terrane. Institute on Lake Superior Geology 49th Annual Meeting Proceedings, 49, 33-47. Sims, P.K., Schulz, P.K., Geologic Map of Precambrian Rocks of Parts of Iron Mountain 30’ x 60’ Quadrangles, Northeastern Wisconsin and Adjacent Michigan. US geological Survey Miscellaneous Investigation Series Map _2356. Scale 1:100,000 Shervais, J. (2001). Birth, death, and resurrection: The life cycle of suprasubduction zone ophiolites. Geochemistry Geophysics Geosystems, 2. 43 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geochemical and petrological studies on the origin of Ni-Cu sulfide mineralization at the Eagle Intrusion in Marquette County, Michigan HINKS, Benjamin1, THAKURTA, Joyashish1 and MAHIN, Robert2 1 Department of Geosciences, Western Michigan University, 1903 W. Michigan Ave. Kalamazoo, MI 49008 2 Eagle Mine, Lundin Mining Corporation, 4547 County Road, Champion, MI 49814 The ~1.1 Ga Eagle deposit is a small mafic to ultramafic sulfide-bearing intrusion located in the north central portion of Michigan’s Upper Peninsula within Michigamme Township, Marquette County (Figure 1). The Eagle and East Eagle intrusions penetrate Paleoproterozoic rocks of the Marquette Supergroup that were deposited in the 400 km2 Baraga Basin during the ~1.85 Ga Penokean orogeny (Ding et al. 2010). The Eagle and East Eagle intrusions are a part of the east-west trending Marquette-Baraga dike swarm that is associated with ~1.1 Ga Midcontinent Rift System (MRS) magmatism (Ding et al. 2010). The age of the Eagle intrusions were determined using uranium-lead dating to be 1107.3 ± 3.7 Ma, which constrains their formation to the early stages of Midcontinent Rift System formation (Ding et al. 2010). The current proven and probable reserves for Eagle are 5.2 million tonnes with an average grade of 3.11% Ni, 2.55% Cu, 0.08% Co, 0.69 gpt Pt, 0.47 gpt Pd, and 0.28 gpt Au (R. Mahin, pers. comm.). This study will attempt to constrain the mechanisms responsible for the formation of sulfide minerals at the Eagle deposit and the source(s) of external sulfur required to produce the sulfide ores. Principle objectives include identifying the source(s) of external sulfur required to form sulfide minerals and the geochemical/ petrological relationships between the intrusions, country rocks and sulfide deposits. It is well known that the Eagle intrusion hosts high-grade Ni-Cu sulfide ores, while the East Eagle intrusion is weakly mineralized (Ding et al. 2010). As of now a relationship between the two intrusions has not been well established. Further analysis will attempt to constrain the connection between the Eagle and East Eagle intrusions. Geochemical studies done through previous research have suggested multiple sources of external sulfur from Archean and Paleoproterozoic country rocks, even though sulfur isotope signatures are indicative of mantle values (Ding et al. 2012). Further analysis of sulfur isotope data from the Eagle and East Eagle deposits will help to address the question of external sulfur in the magmatic system. Petrographic analysis of rocks from the intrusions and the country rocks will be used to study textural characteristics. From the results generated in this study we hope to identify a set of characteristics that can be used for the identification of other sulfide deposits in the Upper Peninsula of Michigan. Preliminary hand sample analyses of previously collected core samples have yielded three main rock types in decreasing olivine contents: feldspathic peridotite, melatroctolite and olivine melagabbro (Ding et al. 2010). Primary sulfur textures are disseminated, semi-massive and massive sulfide mineralization with major sulfide minerals of pyrrhotite, chalcopyrite, pentlandite and traces of bornite. Disseminated sulfide mineralization tends to occur as scattered blebs with 315% sulfide minerals. Semi-massive sulfide mineralization occurs as a net-textured matrix containing 30-50% sulfide minerals. Massive sulfide mineralization is characterized by a leopardtextured matrix with stringers of sulfide minerals. Massive sulfides contain >50% sulfide minerals. 44 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: Eagle area geology. Mafic dykes determined from magnetics are shown in red. Purple zones are the locations of the Eagle and East Eagle intrusions. Fault zones determined from magnetics are shown as black lines. Blue lines represent gravity lineaments from Resolve electromagnetics. Black dots are representative of borehole locations. From: Rossell et al. 2005 REFERENCES Ding, X., C. Li, E. M. Ripley, D. Rossell, and S. Kamo (2010), The Eagle and East Eagle sulfide ore-bearing maficultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution, Geochem. Geophys. Geosyst., 11, Q03003, doi:10.1029/2009GC002546. Ding, X., E. M. Ripley, C. Li (2010), PGE geochemistry of the Eagle Ni-Cu-(PGE) deposit, Upper Michigan: constraints on ore genesis in a dynamic magma conduit: Miner Deposita, doi: 10.1007/s00126-0350-y. Ding, X., E. M. Ripley, S. B. Shirey, C. Li (2012), Os, Nd, O and S isotope constraints on country rock contamination in the conduit-related Eagle Cu-Ni-(PGE) deposit, Midcontinent Rift System, Upper Michigan: Geochemica et Cosmochimica Acta, 89, pp. 10-30. Owen, M. L. and Meyer L. H. I. (2013), NI 43-101 Technical Report on the Eagle Mine, Upper Peninsula of Michigan, USA. Report for Lundin Mining Corporation, dated July 26, 2013, pp. 1-241. Rossell, D and Coombes, S, (2005), The Geology of the Eagle Nickel-Copper Deposit Michigan, USA. Report for Kennecott Exploration, dated April 29, 2005, pp. 1-35. Schneider, D.A., Bickford, M.E., Cannon, W.F., Schultz, K.J., and Hamilton, M.A., (2002). Age of volcanic rocks and syndepositional iron formations, Marquette Range Supergroup: implications for the tectonic setting of Paleoproterozoic iron formations of the Lake Superior region. Can. J. Earth Sci., vol. 39, pp. 999-1012. 45 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter - 2015 Update HUDAK, George1, MONSON GEERTS, Stephen1, ZANKO, Larry1, POST, Sara1, and BANDLI, Bryan2 1 Natural Resources Research Institute, 5013 Miller Trunk Highway, Duluth, MN, 55811 2 Department of Geological Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 The Natural Resources Research Institute (NRRI) continues to conduct a detailed characterization of mineral dust in northeastern Minnesota. The purpose of this research is to evaluate the effects of present emissions from taconite mining and processing on air quality throughout the Mesabi Iron Range (MIR) (Figure 1) by characterizing airborne mineral particulate matter (PM) within currently operating taconite processing plants, in MIR communities surrounding taconite mining/processing operations, and in population centers in Minnesota not associated with taconite mining. Characterization studies of age-dated lake sediments are also being conducted to determine the composition of past PM deposition. NRRI’s sampling and characterization work represents the community/environmental component of the Minnesota Taconite Workers Health Study, a broad University of Minnesota (UM) research effort involving both the NRRI and the School of Public Health. Figure 1. Locations of taconite processing plants on the Mesabi Iron Range being sampled during this study (after Oreskovich and Patelke, 2006) Air sampling was performed within taconite operations, MIR communities, and non-MIR communities by NRRI scientists during both winter and summer seasons from 2009-2012. Sampling was conducted at four process locations within taconite operations, including: 1) secondary crushers; 2) magnetic separators/concentrators; 3) agglomerators/ball drums; and 4) kiln/pellet discharge areas. Community sampling took place on centrally-located rooftops of public buildings, or in the case of the northern most background site, in a remote sampling location to evaluate the air quality away from the MIR. Airborne particles were collected using: 1) a micro orifice uniform deposit impactor (MOUDI) (Marple et al., 1991, 2014), which enables size46 �Proceedings of the 61st ILSG Annual Meeting - Part 1 fractionated PM collection; and 2) a Total Filter Sampler (TFS). Particulate matter was evaluated via gravimetric analysis and was subsequently subjected to comprehensive characterization that included: 1) scanning electron microscopy (SEM) imaging; 2) energy dispersive x-ray spectroscopy (EDS); 3) electron backscattered diffraction (EBSD); 4) proton induced x-ray emission (PIXE); 5) the Minnesota Department of Health’s 852 Method Transmission Electron Microscopy (TEM) Analysis for Mineral Fibers in Air; and 6) the International Standardization Organization’s Indirect Method 13794 for Ambient air – Determination of Asbestos Fibers. NRRI’s research methods do not produce exposure data, and are not meant to provide data for regulatory purposes. During the past year, the NRRI has been evaluating the physical (gravimetric, morphology, concentration), mineralogical, and chemical characteristics of the PM obtained from sampling at the taconite operations and MIR/non-MIR communities. This includes analysis of 55 taconite plants; 73 northeastern Minnesota community and 6 Minneapolis samples. Lake sediment analysis has been completed, and will provide important historical data regarding potential mineralogical inputs from iron mining and processing from ~1840 (which pre-dates iron mining on the MIR) to the present, which includes the period where the transition from natural ore mining to taconite mining took place. Community results to date are as follows: • measured particulate matter concentrations for PM2.5 in all MIR communities have been below 12 µg/m3, and for total PM have been below 16µg/m3; • particulate matter concentrations on the MIR are similar to those in the two NE Minnesota background sites (Duluth NRRI, Ely Fernberg site), and are lower than those obtained in Minneapolis (UM Mechanical Engineering Building rooftop); • mineral particulate matter in community air samples reflects the mineralogy of the Biwabik Iron Formation and other Minnesota rock types and geological materials; • elongate mineral particles (EMP) are present in MIR community ambient air samples; however, asbestiform amphiboles were rarely observed (1 asbestiform amphibole EMP in ~22,800m3 of air). Taconite plant results to date are as follows: • plant environments can be dusty, with the most dusty environments associated with the agglomerator and kiln discharge areas; • particulate matter levels (PM1, PM2.5, PM10, and total PM) show a slight increase in the five MIR communities during plant/mine activity, but this increase is not statistically significant compared to when the plants were not operating. • significantly higher concentrations of EMPs, including amphiboles, were detected in the eastern most plant compared with the other five plants, but the morphology of these structures more closely resembles cleavage fragments rather than asbestiform morphologies. REFERENCES ISO 13794 (1999), Ambient air — Determination of asbestos fibres — Indirect-transfer transmission electron microscopy method. Marple, V. A., Rubow, K. L., and Behm, S. M., 1991, A micro orifice uniform deposit impactor (MOUDI): description, calibration, and use: Aerosol Science and Technology, v. 14, p. 434-446. Marple, V., Olson, B., Romay, F., Hudak, G., Monson Geerts, S., and Lundgren, D., 2014, Second Generation MicroOrifice Uniform Deposit Impactor, 120 MOUDI-II: Design, Evaluation, and Application to Long-Term Ambient Sampling: Aerosol Science and Technology, v. 48-4, p. 427-433. 47 �Proceedings of the 61st ILSG Annual Meeting - Part 1 MDH. Method 852 (1999) T.E.M. analysis for mineral fibers in air – 852. Minnesota Department of Health, Microparticulate Unit, St. Paul, MN. 42 pp. Oreskovich, J. A., and Patelke, M. M., 2006, Historical use of taconite byproducts as construction aggregate materials in Minnesota: A Progress Report: Natural Resources Research Institute Report of Investigation NRRI-RI-200602, 10 p. 48 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geology and geochronology of Archean rocks in the International Falls and Littlefork 30X60’ quadrangles, north-central Minnesota JIRSA1, Mark A., BOERBOOM1, Terrence J., CHANDLER1, V., and SCHMITZ2, Mark D., 1 Minnesota Geological Survey (MGS); 2Department of Geosciences, Boise State University The International Falls and Littlefork quadrangles provide a transect across parts of 3 subprovinces of the Archean Superior Province—Wabigoon, Quetico, and Wawa (Fig. 1). The interpretation of bedrock geology described here was recently published as MGS Miscellaneous Map M-197. The map incorporates field work and geophysical modeling by the authors, augmented by lidar, air photography, reprocessed aeromagnetic data, and unpublished field notes of former MGS geologists. In addition, 4 samples were submitted for high-precision U-Pb geochronologic analysis of zircons (Fig. 2) to quantify ages for some units and temporally constrain some events. The map depicts a complex history of volcanism, sedimentation, intrusion, multiple episodes of migmatization involving partial melting and melt dispersion, and several periods of deformation and metamorphism. Figure 1. Generalized geologic setting of subject map area showing the approximate locations from which samples were taken for geochronologic analysis (solid circles #1-4). See Fig. 2 for sample details. The structural grain of the subprovinces is largely a product of three major orogenic events— each involving a component of NW-SE-directed compressional and transpressional deformation— referred to here as D1, D2, D3. The Wabigoon and Wawa subprovinces are greenstone-granite terranes inferred to represent oceanic and island arc settings. The intervening Quetico subprovince consists of sedimentary rocks deposited in continental margin and oceanic environs that were subsequently deformed, metamorphosed, partially melted, and multiply intruded. Various estimates bracket volcanism in the Wabigoon subprovince in Ontario between ~ 2728-2725 Ma. Geochronologic analyses conducted for this mapping project indicate that a younger sequence of felsic volcanic strata (2702.9±0.6 Ma) marks the southernmost part of the subprovince (Figs.1 and 2, sample 1.). Similar rock types occur in the northern Wawa subprovince, but ages there are less well constrained, as only a single age of ~2722 Ma exists from a sample acquired 85 miles east of the map area. A new age of 2715.8±0.5 Ma (Figs. 1 and 2, sample 2) acquired from felsic volcanic rocks in this map area establishes broad equivalence across the northern part of the Wawa subprovince. Rocks of the Quetico subprovince consist of biotite-plagioclase schist, granitoid and minor mafic intrusions, and complex migmatite containing multiple paleosomatic and neosomatic components. Based on detrital zircons in Canadian analogs, the schist was derived from graywacke and pelitic sediments deposited ~2698-2692 Ma in an accretionary prism during early 49 �Proceedings of the 61st ILSG Annual Meeting - Part 1 stages of collision between the Wawa subprovince island arc to the south, and the Superior craton (superterrane) to the north. Later stages of this D1 collisional event produced tilting, folding that included broad nappe structures locally, and thrust imbrication of volcanoplutonic rocks in subprovinces north and south of the Quetico, and recumbent folding within the Quetico. This D1 deformation was followed by what may have been a regional extensional event that produced localized calc-alkalic volcanism, and sediments containing clasts derived from all precursor rock types deposited in isolated, unconformity- and fault-bounded basins. The remnants of one such basin in the International Falls area is known as the Seine Group. A new age of 2693.9±0.6 Ma was acquired from a clast of trachyandesite in Seine conglomerate (Figs. 1 and 2, sample 3). A second deformation event (D2) occurred at about 2680 Ma during the Minnesotan orogeny. It produced regional penetrative fabrics, folds, and prograde metamorphism in all 3 subprovinces, and partial melting of Quetico schist to form an early suite of leucogranite, granodiorite, trondhjemite, and tonalite that is interlayered on all scales with biotite schist, forming complex migmatites. A third deformation event (D3) is manifest in shear and fault zones in the volcanoplutonic subprovinces; and broad, east- and west-plunging folds (synforms, antiforms) of D2 fabrics in the Quetico subprovince. At least part of this deformation was synchronous with or just preceded migmatization and emplacement of 2-mica leucogranite and slightly younger, typically red, variably magnetic biotite granite known as the Lac La Croix. An age of 2661.3±0.3 Ma was acquired in this map area from a late granitic intrusion lithologically similar to the Lac La Croix (Figs. 1 and 2, sample 4). The grade of metamorphism is more or less symmetrical along the northeast-trending axis of the Quetico subprovince, having greenschist facies at the margins, and middle to upper amphibolite facies near the axis. Metamorphism was generally syntectonic with D2 and D3 deformations; and a contact metamorphic overprint occurs locally in rocks adjacent to the Lac La Croix Granite and similar late intrusions. The axis of the Quetico is coincident with a large post-metamorphic anticlinorium, increased neosome abundance, and moderately high magnetic and gravity expression, despite the presence of less dense rocks at surface. Collectively, these attributes indicate exposure of more deeply buried crust. The presence of dense, magnetic rocks at depth that may represent the uplifted floor on which sedimentary strata were deposited. Figure 2. Concordia plots showing results of LAICPMS geochronologic analysis of zircons. Ages reported here are taken from calculations of 207 Pb/206Pb. Details of this work are available from published MGS digital files associated with Miscellaneous Map M-197. Sample locations in UTM NAD83 coordinates: 1. 436644E/5382588N 2. 454692E/5316937N 3. 485860E/5383472N 4. 476489E/5354520N Mapping and geochronologic analyses were funded in part by 2013 USGS STATEMAP element of the National Geologic Mapping Program 50 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Rainy River, northwestern Ontario's first meteorite KISSIN, Stephen A. Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The only meteorite previously known from northern Ontario is the Osseo iron meteorite, found in 1931 in the easternmost part of Ontario in the vicinity of the Cobalt silver camp (Buchwald, 1975). In 2000, Robert Weaver found a 3.26 kg iron object in a field on his farm, near Rainy River, Ontario. It is now in the possession of Howard Williams of Winnipeg, who presented the specimen for identification in 2013. A polished slab of 47.26 g has been deposited as the type specimen in the Royal Ontario Museum. The meteorite is oxidized on its exterior and is of a roughly ellipsoidal shape with dimensions of 17.4 cm x 13.2 cm x 9.5 cm (Figure 1). The interior is unoxidized except minor patches near the thin exterior iron oxide crust and along kamacite grain boundaries. The meteorite is an octahedrite, with coarse kamacite lamellae of irregular width and a stubby aspect of L: W= 3:1 to 4:1. The residual taenite lamellae are very narrow where preserved and only one area of net plessite was observed (Figure 2). The other minerals present are troilite, in the form of small, intergranular veinlets, and schreibersite, as small, globular forms and as small, euhedral, crystals known as rhabdites. The rhabdites are very abundant and display strong preferred orientation within kamacite lamellae. The meteorite has experienced moderate cosmic shock as seen in the abundant Neumann lines, some of which have been bent in places. The kamacite lamellae have been polygonalized due to brittle fracture. Hardness of kamacite (VHN=254±13) and taenite (VHN= 492±28) is also evidence of work hardening due to shock. Analysis by neutron activation yielded the following minor and trace element composition: Ni= 7.23 wt%, Co=0.463 wt%, Sb= 410 ppb, all in ppm Cr=19, Cu=114, Ga=91, Ge=170, As 13.9, W=<10, Re=0.64, Ir=3.87, Pt=7.8, Au=1.47. Together with the details of the structure, the meteorite is a member of the group IAB complex, as defined by Wasson and Kallemeyn (2003). However, the Au-content is at low extreme of the group, thus suggesting that the meteorite may belong to the Algarrabo duo of Wasson and Kallemeyn. Moreover, other trace element contents, especially Ge and Ga, are similar to those of one of the duo, Livingston (Tennessee). As well, the structure of Livingston (Tennessee) is similar to that of Rainy River in that Ni-content does not agree with the normally expected kamacite bandwidth, which should be in the range of 2.5-3 mm in well-defined Widmanstaetten structure (Buchwald, 1975). Buchwald further proposed an usual thermal history for Livingston (Tennessee), which may also have applied to Rainy River. 51 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: Main mass of the Rainy River meteorite, less off-cut for type specimen. Scale in cm. Figure 2: Polished surface of the type specimen. Some kamacite lamellae outlined by weathering. Scale in cm. REFERENCES Buchwald, V.F. 1975. Handbook of Iron Meteorites. University of California Press. Wasson, J.T. and Kallemeyn, G.W. 2003. The IAB iron-meteorite complex: A group, five subgroups, numerous grouplets, closely related, mainly formed by crystal segregation in rapidly cooling melts. Geochimica et Cosmochimica Acta 66: 2445-2473. 52 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Studies on PDFs in shocked quartz from distal Sudbury ejecta in the Thunder Bay area compared with Chicxulub KISSIN, Stephen A. and BRUMPTON, Gregory R Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Planar deformation features (PDFs) in Sudbury ejecta in the Marquette, Michigan area have been studied briefly by Cannon et al. (2010) and more extensively by Pufahl et al. (2007) and briefly in the Thunder Bay area by Kissin & Brumpton (2014). We have made an extensive study of PDFs in quartz from ejecta from various localities in the Thunder Bay area. A total of 153 PDFs from 104 quartz grains were measured on the universal stage (Table 1). Of these, 77 PDFs were measured in 55 grains within accretionary lapilli, and 76 PDFs were measured in 49 grains in the matrix of the ejecta. Indexing was carried out by both the use of stereographic projection and use of the ANIE program of Huber et al. (2011). Agreement between the two methods was generally good, although the ANIE program eliminated drawing errors that may be introduced in the stereographic projection method. Comparison of the distribution of indexed PDFs in grains in lapilli vs. those of grains in the matrix revealed no significant difference at 95% confidence using the Wilcoxon matched-pairs signed-ranks test, although the source of the grains may have differed. Differences were noted in that grains within lapilli are smaller (50-100 µm) as opposed to those in the matrix (100-500 µm). As well, grains in the matrix were frequently rounded, whereas those in the lapilli were almost always angular. These observations suggest that the grains in lapilli were predominantly from crystalline basement rocks of the target area, whereas those in the matrix had a considerable contribution from sedimentary surficial rocks. The distribution of the PDFs in this study was compared with the results of Nakano et al. (2008) who studied PDFs in distal ejecta at various distances from the Chicxulub impact crater (Table 2). The PDF sets are grouped in five-degree bins containing most of the major PDFs. Note that Nakano et al. did not include the {1014} set, which is abundant in both their and our data sets. Analysis of all the data sets in Table 2 using the Friedman test yields a Friedman Statistic of 0.9592. Thus, the probability is > 95% that the sum of the ranks (rows) is the same in each column. As well, the PDF sets from each of the Chicxulub sites was compared with the Sudbury ejecta set using the Willcoxon matched-pairs signed ranks test. In all cases except for the DSDP 536 set, the nonparametric Spearman correlation coefficient and the one-tailed P value indicated pairing of the ranks. The DSDP 536 data were slightly less than statistically significantly paired. These tests indicate the similarity of distribution of PDF sets in Sudbury and Chicxulub distal ejecta. The formation of PDFs in quartz has been experimentally calibrated in response to shock pressure, as reviewed by Stöffler and Langenhorst (1994). The PDFs in Sudbury and Chicxulub ejecta correspond to a wide range of shock pressures extending as high as 35 GPa. Even higher shock pressures are indicated by the occurrence of diaplectic quartz glass and incipient melting of quartz grains. This range of shock pressures is indicative of the impact process and the location of the quartz grains to it, as has been modeled by Nakano et al. (2008). 53 �Proceedings of the 61st ILSG Annual Meeting - Part 1 !"""#$% CB,! C-B*! !!"!! !$% *CB.! !!"!! !!! .BE! !!"!! !N! *BC! !!!!! !!! @BA! !!"!! !N! 1B,! !!"!! !N! 1B@! !!!!! !}! 1B,! !!"!! !N! @B-! !!!!! !N! *BC! !!"!! !!! -! !!"!! !!! .BE! !!"!! !}! *BC! !!"!! !}! -B.! !!!!! !}! @BA! !!"!! !N! OF3FP'Q'P! .BE! Table 1. Absolute frequency of PDF sets (153 sets) Table 2 REFERENCES Cannon, W.F., Schulz, K.J., Horton, J.W. Jr. and Kring, D.A. 2010. The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA, Geological Society of America Bulletin 122: 50-75. Huber, M./S., Ferriére, L., Losiak, A. and Koeberl, C. 2011. ANIE: A mathematical algorithm for automated indexing of planar deformation features in quartz grains, Meteoritics & Planetary Science 46: 1418-1424. Kissin S.A. and Brumpton, G.R. 2014. PDFs in Sudbury ejecta in the Gunflint Formation, Ontario: A comparison of methods., Institute on Lake Superior Geology Proceedings 60, Part 1, 69-70. Nakano, Y., Goto, K., Matsui, T., Tada, R. and Tajika, E. 2008. PDF orientations in shocked quartz grains around the Chicxulub crater, Meteoritics & Planetary Science 43: 745-760. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, G.J. and Edwards, C.T. 2007. Physical and chemical evidence of the 1850 Ma Sudbury impact event in the Baraga Group, Michigan, Geology 35: 827-830. Stöffler, D. and Langenhorst, F. 1994. Shock metamorphism of quartz in nature and experiment: I. Basic observation and theory. Meteoritics, 29: 155-181. 54 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geologic mapping of Neoarchean and Proterozoic rocks near Knife Lake, northeastern Minnesota, by students of the Precambrian Research Center’s 2014 field camp KROGMEIER, Benjamin1, McKEVITT, Dylan1, ROEPKE, Elizabeth1, SARA, Michael1, SZKILNYK, Paul1, and JIRSA, Mark2 1 2014 Field Camp Students, Precambrian Research Center, Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Duluth, Minnesota 55811 2 Minnesota Geological Survey, University of Minnesota, 2609 W. Territorial Rd., St. Paul, Minnesota 55114 The University of Minnesota-Duluth’s Precambrian Research Center conducted its eighth annual field camp in 2014, and this presentation is one of a series that show results of “capstone” mapping projects. The projects test student skills by creating new geologic maps in areas of poorly known geology, which benefits both students and mentor organizations. This capstone project involved mapping an area of ~12 mi2 in the Boundary Waters Canoe Area Wilderness (BWCAW), centered on the south arm of Knife Lake (Fig. 1). The resulting map provides details about the complex depositional and tectonic history of a Neoarchean metavolcanic and metasedimentary terrane that is part of the Wawa subprovince of Superior Province, and rare diabasic dikes that intruded it. Figure 1. Generalized bedrock geologic map of NE Minnesota showing the Knife Lake capstone area (solid black polygon). The Neoarchean unit labeled “Supracrustal Rocks” encloses both older volcanic sequences and younger, largely sedimentary ones. Outline of Boundary Waters Canoe Area Wilderness is dashed. The Neoarchean rocks in the central BWCAW comprise a Timiskaming-type extensional basin and its apparent wall- and floor-rocks. The geologic units are parceled into structural lozenges separated by anastomosing shear and fault zones. Although rock types are comparatively pristine within each lozenge, correlation of units from one fault-bounded block to another is challenging. Nevertheless, this project and the several that preceded it in prior years of mapping attempt to “unstrain” the rocks within each parcel to reveal stratigraphic variations that may reflect fluctuations in basin geometry and progressive erosional dissection of basin wall rocks. Understanding the lithologic details and the apparent post-depositional tilt of individual lozenges of rock are essential to this objective. The Knife Lake map area provides a window into this complex terrane. It consists of 2 sequences of broadly folded metasedimentary rocks that are part of the Knife Lake Group, separated by an east-trending block of vertically dipping, southward-facing, variolitic metabasalt as thick as 0.5 km. The sedimentary strata include tuffaceous graywacke and mudstone, locally containing conglomeratic and gritstone layers having clasts of the metabasalt, Saganaga Tonalite (~2.69 Ga), and other calc-alkalic igneous rocks. All these rocks were deformed and metamorphosed to very low greenschist facies during the Minnesotan Orogeny (~2.68 Ga)—thus constraining deposition of the sediments to the approximate interval 2.69-2.68 Ga. The primary objective of this mapping was to delineate and interpret the nature of contacts between the 55 �Proceedings of the 61st ILSG Annual Meeting - Part 1 apparently older metabasalt and superjacent sedimentary strata derived in part from it. Highlights of our mapping include the following: 1) The abundance of graded sequences of sand- to mud-sized detritus implies deposition of much of the strata was submarine (or lacustrine). 2) Local sequences of polymictic and oligomictic conglomerate, and arkose containing matrixsupported, subrounded metabasalt fragments as large as 20 cm implies occasional subaerial deposition of more chaotic debris flow, alluvial fan, and fluvial sediments. This is consistent with episodic uplift of wall rocks adjacent to the developing fault- and unconformity-bounded basin. 3) The abundance of white-weathered tuff and tuffaceous siltstone and mudstone implies calc-alkalic volcanism may have been contemporaneous with deposition, or volcanic strata were not fully lithified at the time of deposition. 4) Although both northern and southern sedimentary sequences are broadly folded; stratigraphic facing near contacts with the medial basalt block is consistently away from the basalt. 5) The northern contact of basalt with sedimentary strata is poorly exposed, but appears to be a fault. It lies along the down-section part of the basalt, and thus is stratigraphically discordant. 6) The southern contact of basalt with sedimentary strata is visible in several areas where it varies from an angular unconformity developed on relatively fresh basalt, to one developed on paleosaprolitic basalt. A spectacular paleosaprolite of basaltic protolith is exposed at one locality. 7) Major fold axes in the northern sedimentary sequence plunge shallowly to the northeast; those in the southern sequence plunge to the southwest. 8) From these observations it appears that the block containing basalt unconformably overlain by the southern sedimentary sequence was uplifted on its north side (Fig. 2). In addition, the divergence of major fold plunges in sedimentary sequences north and south of the fault implies a scissor motion that tilted the southern block down on the west. Judging from inferred structural position and lithologic similarities, it is also likely that some portions of the two sedimentary sequences may have been fault-duplicated. Although the precise age of rare diabase dikes is unknown, most trend northwest and dip nearly vertically, similar to the Paleoproterozoic Kenora-Kabetogama dike swarm. However, a Mesoproterozoic age may be indicated by the trend of one dike that is nearly horizontal, which is quite anomalous for Paleoproterozoic dikes in the Superior Province. We speculate that it may represent emplacement at very high levels in the crust, where comparatively less lithostatic load permitted delamination along horizontal exfoliation structures in host rocks. This and other capstone mapping projects can be viewed at www.d.umn.edu/prc. Figure 2. Schematic cross-sectional model to explain the distribution of map units and partial repetition of stratigraphic sequences by faulting. 56 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Structural control on the Borden Gold deposit, Chapleau, Ontario LAFONTAINE, Daniel and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Rd. Thunder Bay, ON P7B 5E1 Canada The multi-million-ounce Borden Gold deposit is located 20 km east of Chapleau, within the Wawa Subprovince of the Superior Province. Interestingly, it is hosted in upper amphibolite to granulite facies metamorphic rocks at the southern margin of the Kapuskasing Structural Zone. Competent lithons of granulite facies rock appear to be surrounded by more ductile amphibolite facies gneisses and schists, suggesting polymetamorphism with retrograde amphibolite facies metamorphism after granulite facies metamorphism. Competency contrasts between the granulite and retrograde amphibolite facies lithologies created heterogeneous strain, ideal for gold mineralization, during ductile deformation at amphibolite facies metamorphic temperatures. Gold is typically observed in competent rocks with weakly developed foliation and also in competent rocks that are bordered by strongly foliated units. Garnet-biotite geothermometry on unzoned almandine garnets yields temperatures ranging from 579°C to 690°C ±50°C for metamorphism of the garnet-biotite schist. Temperatures increase from the garnet core towards the rim, indicating that garnets equilibrated rapidly during prograde metamorphism from the upper amphibolite to granulite facies. Fieldwork and microstructural analysis have identified a variety of competent lithologies and minerals, which provide low-strain environments for gold mineralization. On the macroscopic scale, the relict granulite facies rock behaves more competently than the retrograde amphibolite facies rock. Competent minerals that provide a low-strain site for fluid transport and gold mineralization include relict orthopyroxene, garnet, pyrite and coarse sillimanite. Preliminary results indicate an important relationship between gold mineralization, metamorphism and deformation, and understanding this relationship will benefit exploration and development of the Borden Gold deposit. 57 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Incorporation of Duluth Complex maps into GIS platform LENTSCH, Nathan1 and MILLER, Jim1 1 Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812 In 2001, Jim Miller and colleagues compiled much of the mapping existing at the time into a digital geologic map and database for the Duluth Complex that was published by the Minnesota Geological Survey (MGS) as Miscellaneous Map M-119 (Miller et al., 2001). Because this compilation was focused on areas of the complex that were open to minerals exploration, a substantial amount of historical field data collected within the Boundary Water Canoe Area Wilderness (BWCAW) was largely excluded from this digital compilation, this includes data collected by Phinney (1972) and Miller (1986). In the 1960’s, William Phinney, an igneous petrology professor at the University of Minnesota-Twin Cities, conducted bedrock mapping for the MGS in the northwestern part of the Duluth Complex. In the summers of 1966 to 1969 Phinney conducted extensive reconnaissance mapping of lakeshore exposures by canoe and floatplane support in areas of the Duluth Complex now contained entirely within the BWCAW (established in 1976) and only accessible by canoe or by foot. Although he did not publish any maps from this field work, Phinney summarized his studies in the MGS’s Centenial Volume on the Geology of Minnesota publication (Phinney, 1972). In 1970, he left the University to take a job with NASA. In 1979, Dr. Phinney passed along all field journals, maps, and thin sections of his Duluth Complex mapping to Dr. Miller, who at the time was a PhD candidate planning his own detailed mapping study in part of an area of the complex that Phinney had previous reconnaissance mapped. In 1981, Dr. Miller conducted four months of detailed mapping in an area focused on the Lake One - Lake Four chain in the Snowbank 7.5’ quadrangle. Miller’s outcrop mapping is preserved only as a blueprint map in his PhD thesis (Miller, 1986), field maps on airphoto bases, and field notebooks. Only his geological linework and structural measurements were digitized for the M-119 map. None of Phinney’s original outcrop data has been digitized. The goal of my research project, which was funded by the University of Minnesota Duluth’s UROP program (Undergraduate Research Opportunity Program), was to compile outcrop-based field data into a digital database using ArcMAP 10 for the area of the Duluth Complex mapped by Phinney (unpublished data) and Miller (1986) in the Snowbank Lake 7.5’ quadrangle. Digitally compiling these field observations, measurements, and sample locations is important for several reasons: 1) it will preserve an important database of geologic information, some of which previously only existed as fragile paper copies; 2) whereas only one copy of Phinney’s and Miller’s field data had existed, digitally archiving their field maps and observations will allow open access to future geologists and researchers; 3) with most of their mapping covering areas deep into the BWCAW, it is likely that in many areas, their mapping will be all that exists for the foreseeable future; and 4) this project gave me the opportunity to learn the powerful geospatial tool that is ArcMap. To efficiently and accurately trace the locations and shapes of the numerous outcrops mapped by Dr. Miller, scans of field maps on aerial photo bases were inserted as a layer under a partially transparent topographic map layer in ArcMAP. Figure 1 shows an example of what this process looked like. As of this writing, over 1,000 outcrops have been digitized, each linked to a table of attributes. Recorded attributes include the outcrop station, the date visited, the major and minor lithologies observed, and a short description for each. An example of this table is shown in Figure 2. 58 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: GIS map showing areas of the BWCAW. Aerial photo as a base layer in bottom left. Light gray outcrops around Lake Two were digitized from Dr. Miller’s 1981 mapping data (Miller, 1986). Figure 2: Table of attributes associated with outcrops in GIS. Each outcrop can be referenced by major/minor lithology and physical description. Given the time constraints of this project, only Dr. Miller’s field data was able to be digitized thus far. This leaves the door open for future work done by another eager undergraduate who would like to become familiar with GIS and the Duluth Complex mapped by Phinney. REFERENCES Miller, J.D., Jr., 1986, The geology and petrology of anorthositic rocks in the Duluth Complex, Snowbank Lake quadrangle, northeastern Minnesota. unpublished Ph.D. dissertation, University of Minnesota, Minneapolis, 280 p. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.E., 2001, Geologic map of the Duluth Complex and related rocks, northeastern Minnesota. Miscellaneous Map Series, M-119, scale 1:200,000, 2 sheets. Phinney, W.C., 1972. Northwestern part of Duluth Complex. In: Sims, P.K. & Morey, G.B. (eds.) Geology of Minnesota -A centennial volume. Minnesota Geological Survey, p. 335-345 59 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geology and geochemistry of the Lang Lake greenstone belt, Uchi Domain, Superior Province MAGNUS, Seamus Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5 Canada The Lang Lake greenstone belt, located 70 km WNW of Pickle Lake, Ontario, lies within the central Uchi domain of the greater North Caribou terrane. The belt is composed of ca. 2750 Ma dominantly greenschist facies volcanic, volcaniclastic and sedimentary rocks, and is intruded by various syn-volcanic and syn-to-post-tectonic plutonic rocks. Following the initial discovery of gold in the Shonia Lake area in 1928 and subsequent mapping of the surrounding greenstone by Laird (1930), the Lang Lake greenstone belt received little attention until mapping by the Ontario Geological Survey in the 1970s (Fenwick 1970, 1971; Fenwick & Srivastava 1972). These 3 maps were compiled and ground-checked as part of “Operation Pickle Lake” by Sage and Breaks (1973), who later released an Ontario Geological Survey Open File Report (Sage and Breaks 1982) which included a brief lithologic description of the greenstone belt. Unlike the neighbouring greenstone belts of the central Uchi domain, which have been subject of modern geological studies (government and academic) throughout the 1990s and 2000s, the Lang Lake greenstone belt represented a significant gap in our regional geoscience knowledge. To fill this gap and to supplement the Cat Lake First Nations land use plan, the Lang Lake greenstone belt was mapped during the 2014 field season at a scale of 1:20,000. 278 hand samples were collected for whole rock major and trace element geochemical analysis, 6 samples were submitted for U-Pb zircon geochronological analysis, and 16 samples were submitted for whole rock Sm-Nd isotopic analysis as part of an HBSc thesis conducted by Matthew Hanewich at Carleton University. Preliminary geochemical analysis of the coherent facies metavolcanic rocks and synvolcanic dikes indicates the presence of primitive (i.e. derived from “Primitive Mantle”) Fetholeiitic basalt, calc-alkaline basaltic andesite and calc-alkaline FII rhyolite. Volcaniclastic rocks contain whole rock major and trace element compositions equivalent to calc-alkalic basalt, andesite, dacite and rhyolite, suggesting that they were proximally sourced, and may represent a mixture of material from the aforementioned tholeiitic and calc-alkalic flows. Interflow volcaniclastic mudstones and wackes are intercalated with bands of magnetitechert iron formation which extend across the entire strike-length of the greenstone belt. Similar bands of iron formation are interbedded with clastic metasedimentary rocks which dominate the eastern half of the greenstone belt. Facies include mudstone, wacke, arenite and several lenses of granitoid-clast-bearing conglomeratic rocks similar to those found in the Billet Assemblage of the nearby Meen-Dempster greenstone belt (Stott and Corfu 1993). Whole rock geochemical trends and concentrations overlap those of the metavolcanic and metavolcaniclastic rocks, thus making them difficult to distinguish geochemically. Whole rock geochemical analysis of the mafic intrusive rocks at McVicar Lake and the tonalitic stock which cuts them suggests that these intrusive rocks represent a high level magma chamber which acted as a feeder conduit for the overlying metavolcanic rocks. Magma mingling, assimilation, hybridization and fractional crystallization textures are visible at an outcrop scale throughout the mineralogically and geochemically diverse pluton, supporting the proposed genetic relationship between the local intrusive and extrusive igneous rocks. 60 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The supracrustal rocks are surrounded and intruded by calc-alkalic, magnesian, peraluminous to metaluminous biotite tonalite to granodiorite. A pluton of alkali-calcic syenodiorite to quartz syenite of unknown age is hosted by volcaniclastic rocks between Lang Lake and McVicar Lake, and contains trace element concentrations indicative of a metasomatised mantle source. In contrast the late alkalic Otoskwin pluton (granodiorite to gabbronorite), just northeast of the greenstone belt, and several intermediate alkalic dikes, appear to have been derived from an unaltered mantle source. Four phases of deformation have been identified within the greenstone belt, including i) one purely compressional event that occurred during emplacement of the surrounding granodioritic intrusions, ii) a sinistral transpressional event which produces the S-asymmetry observed throughout the belt and within the surrounding intrusive rocks, iii) a dextral transpressional event associated with movement along the northwest trending Bear Head Shear Zone at the west end of the belt, and finally another iv) compressional event which formed a series of discrete NNEtrending shear zones that offset all of the previous structural features. This mapping project has provided some broad insights into the petrogenesis and structural history of the Lang Lake greenstone belt. The belt contains abundant prospects for further academic study which would help relate this belt to others within the Uchi province, provide new insights into Archean igneous processes, and aid in the pursuit of economic gold mineralization. REFERENCES Fenwick, K.G. 1970. Lang–Cannon lakes area (west half); Ontario Geological Survey, Preliminary Map P.581, scale 1:31 680. Fenwick, K.G. 1971. Lang–Cannon lakes area (central part); Ontario Geological Survey, Preliminary Map P.665, scale 1:31 680. Fenwick, K.G. and Srivastava, P. 1972. Lang–Cannon lakes area (eastern part); Ontario Geological Survey, Preliminary Map P.738, scale 1:31 680. Laird, H.C. 1930. Shonia Lake area, District of Kenora (Patricia Portion), Ontario; Ontario Geological Survey, Map 39d, scale 1:63 360. Sage, R.P. and Breaks, F.W. 1982. Geology of the Cat Lake–Pickle Lake area, districts of Kenora and Thunder Bay; Ontario Geological Survey, Report 207, 238p. Stott, G.M. and Corfu, F. 1991. Uchi Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.145-238. 61 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The Eagle Mine in Production: U.S.A.’s Only Primary Nickel Producer MAHIN, Robert Eagle Mine LLC, Exploration Department, 200 Echelon Drive, Negaunee, MI 49866 Lundin Mining Corporation’s Eagle Mine is located in Marquette County in the Upper Peninsula of Michigan. The Eagle deposit is an ultramafic-intrusive-hosted high grade Ni-Cu deposit, with associated cobalt, platinum, palladium, silver and gold. Published (2014) P&P reserves for Eagle are 5.2MT @ 3.11% NI and 2.55% Cu. Lundin acquired the partially developed project from Rio Tinto in July, 2013. Capital expenditure, including purchase and construction completion are $770M. As of Q1 2015, full commercial production was achieved significantly ahead of schedule. Annual production over the first three full years (2015 - 2017) is expected to average approximately 23,000 tonnes of nickel and 20,000 tonnes of copper per annum, with additional byproduct credits of precious metals and cobalt. The orebody is accessed from surface by a 1 mile long, 13% grade decline. The mine employs longhole open stoping. The majority of the stopes will be mined as transverse bench and fill stopes, with some thinner zones mined as longitudinal retreat stopes. Stope drilling is accomplished by top hammer vertical drilling from the top sill cut with a breakthrough to the bottom sill cut. Stope dimensions are 10 meters wide by 18 to 29 meters high sill to sill. The stope lengths vary with the thickness of the orebody (15 to 85 meters). The sill will be cut the full width of the stopes at 10 meters wide and 5 meters high. Level spacing varies between 18-29 meters and there are nine mining horizons. Stopes will be mined from the bottom up in an alternating sequence of primary and secondary stopes with cemented rock fill in the primaries and rock fill in the secondary stopes. After mining the uppermost stopes, backfill will be placed tight to the stope backs with a jammer to prevent subsidence. Approximately forty-five truckloads per day deliver ore 65 miles to the Humboldt Mill. The mill is a renovated pellet processing plant with a capacity of 2000 tonne per day. Conventional flotation produces separate nickel and copper concentrates with approximate recoveries of 82% Ni and 93% Cu. Flotation tailings are thickened and deposited subaqueously into the flooded Humboldt open pit. Concentrates are railed directly to either Canadian smelters or to port for overseas shipping. The average life of mine (8 years) production of 17 ktpa Ni and 17ktpa Cu is expected at cash costs (excl. royalties) of approximately $2.50/lb Ni. Since inception the company has striven for open and transparent communication with the community. Eagle has committed to a 75% local hire goal, created programs to support small business development, and hosts semiannual town hall meetings. In addition, the company helped develop a precedent-setting third party community environmental monitoring program (CEMP). The CEMP provides independent verification monitoring for the Eagle Mine, Humboldt Mill, and transportation route. Full-time employment, including long-term primary contractors will be approximately 330. Over the 13 year life of the mine, including construction and closure, the economic impact for Marquette County is predicted to be on the order of $4 billion. Eagle is actively exploring for additional mineralization. Exploration is largely geology driven and based on an open-system magma conduit (chonolith) model. Efforts have focused on identifying and tracing the feeder dikes to Eagle and a sister intrusion, Eagle East (a.k.a. the Yellowdog Peridotite). In 2014, exploration successfully intersected significant mineralization in what is interpreted as the feeder dike to Eagle East. 62 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Evaporated seawater formed sediment-hosted stratiform copper orebodies and second-stage copper mineralization in the Mesoproterozoic Nonesuch Formation of the Midcontinent Rift MAUK, Jeffrey L., EMSBO, Poul, and THEODORAKOS, Peter U.S. Geological Survey, MS-973 Denver Federal Center, P O Box 25046, Denver, CO 80225-0046 The Mesoproterozoic North American Midcontinent Rift System contains sediment hosted stratiform Cu deposits at White Pine and Copperwood, which combined host at least 4 Mt Cu and 75 Moz Ag (Nicholson et al., 1992; Bornhorst and Williams, 2013). Main stage sediment-hosted Cu mineralization formed at both deposits during diagenesis at temperatures near or cooler than 100°C. White Pine also contains structurally controlled second-stage mineralization that was likely synchronous with Keweenaw Peninsula native Cu mineralization (Mauk et al., 1992). Here, we report chemical data from fluid inclusions from main- and second-stage mineralization at White Pine to constrain the origin of these mineralizing fluids. We measured the solute compositions of ore-forming brines from fluid inclusions in mainstage chalcocite, and second-stage calcite and chalcocite from the White Pine deposit. Fluid inclusions were extracted from 100-600 mg of calcite and chalcocite, and analyzed for Na+1, NH4+1, Ca+2, Mg+2, K+1, Rb+1, Sr+2, Ba+2, Cl-1, Br-1, F-1, S2O3-2, SO4-2, and acetate using the ion chromatography methods described by Viets et al. (1996). The Cl-Br-Na data from main- and second-stage minerals from White Pine plot in a relatively small compositional field, with Cl/Br molar ratios that are less than 300, and Na/Cl molar ratios that are less than 0.3. Main- and second-stage fluids may occupy slightly different fields, but this minor difference may only be apparent due to the relatively few analyses of mainstage chalcocite. The Cl-Br-Na data plot close to or along the seawater evaporation curve with no evidence for the dissolution of salt or input from non-marine brines, which would have much greater Cl/Br ratios (Fig. 1). The Na/Cl molar ratios are distinctly depleted compared to those of most basinal fluids worldwide, suggesting that the brines evolved significantly beyond halite precipitation, and approached Mg- and K-salt saturation. Terrestrial fluvial environments played a key role in the sediment deposition that followed the volcanic phase of the Midcontinent Rift, but debate continues on whether the fine-grained clastic sedimentary rocks of the Nonesuch Formation, which host the White Pine and Copperwood deposits, formed in a marine or lacustrine environment (e.g., Cumming et al., 2013, and references therein). Our data require seawater that evaporated to the point of salt deposition, which supports a marine depositional environment for the Nonesuch Formation. Furthermore, the evaporation of seawater beyond halite saturation required by our data, plus the enormous volume of brine required to form the Cu deposits, would require a significant evaporite basin filled with gypsum, halite, and potentially even bittern salts occurred somewhere in the rift basin. If preserved, the most likely location of this thick evaporite sequence was in the thickest and deepest axial portion of the rift, which lies under present-day Lake Superior. The similar composition of main- and second-stage brines raises the intriguing possibility that these two stages of mineralization, despite apparently being separated by nearly 60 m.y. (Ohr, 1993), formed from the same brine. If so, a large basin was required to store the large volume of brine necessary to form second-stage Cu at White Pine. 63 �Proceedings of the 61st ILSG Annual Meeting - Part 1 $#!!" ,+"-& !"#$%& .&/01000& *+,-./0+12" /234-5./0+12" -4-.462"3+674-+02" 82+9+026" !"#$%&'()& $!!!" #!!" 4-" 46+= :;+< !" !%!" !%&" !%'" !%(" *+#!"&'()& !%)" $%!" Figure 1: Cl/Br versus Na/Cl molar ratios for main-stage chalcocite and second-stage calcite and chalcocite from the White Pine deposit. Integrating these results with current understanding of basin architecture and location of deposits may provide new insights into why some areas of the rift produced world-class deposits and other segments are barren. Furthermore, analyses of minerals from the Keweenaw native Cu deposits, which presumably formed synchronously with White Pine second-stage mineralization (Mauk et al., 1992; Bornhorst, 1997), could test whether these large Cu endowments formed from the same fluid, or whether different brines produced mineralization in different portions of the rift. REFERENCES Bornhorst, T. J., 1997, Tectonic context of native copper deposits of the North American Midcontinent Rift System: Geological Society of America Special Papers, v. 312, p. 127-136. Bornhorst, T. J., and Williams, W. C., 2013, The Mesoproterozoic Copperwood sedimentary rock-hosted stratiform copper deposit, Upper Peninsula, Michigan: Economic Geology, v. 108, p. 1325-1346. Cumming, V. M., Poulton, S. W., Rooney, A. D., and Selby, D., 2013, Anoxia in the terrestrial environment during the late Mesoproterozoic: Geology, v. 41, p. 583-586. Mauk, J. L., Kelly, W. C., van der Pluijm, B. A., and Seasor, R. W., 1992, Relationships between deformation and sediment-hosted copper mineralization: Evidence from the White Pine portion of the Midcontinent rift system: Geology, v. 20, p. 427-430. Nicholson, S. W., Cannon, W. F., and Schulz, K. J., 1992, Metallogeny of the Midcontinent rift system of North America: Precambrian Research, v. 58, p. 355-386. Ohr, M., 1993, Geochronology of diagenesis and low-grade metamorphism in pelites: Unpub. PhD thesis, University of Michigan., 147 p. Viets, J., Hofstra, A. H., and Emsbo, P., 1996, Solute composition of fluid inclusions in sphalerite from North America and European Mississippi Valley-type ore deposits: Ore fluid derived from evaporated seawater, in Sangster, D. F., ed., Carbonate-Hosted Lead-Zinc Deposits, Society of Economic Geologists Special Publication 4 p. 465483. 64 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Sedimentology and Geochemistry of a 1.4 Ga Continental Playa System, the Lower Sibley Group, Northwestern Ontario: Implications for the Mesoproterozoic Hydrosphere and Atmosphere METSARANTA, Riku T.1 and FRALICK, Philip2 1 Ontario Ministry of Northern Development and Mines, Sudbury, Ontario, Canada, 2 Department of Geology, Lakehead University, Thunder Bay, Ontario, Canada, P7B 5E1 The 900 m thick Sibley Group consists of playa to deltaic to aeolian deposits outcropping north of Lake Superior and east of Thunder Bay. The lowermost 100 m thick succession of highly oxidized siliciclastic rocks and dolostone was deposited in a north-south trending half graben. The sediments can be divided into 15 lithofacies associations representing distinct depositional environments. The lower siliciclastic unit contains: boulder conglomerate-sandstone-dolocrete (proximal ephemeral braided stream), pebble to cobble conglomerate (ephemeral braided stream), trough cross-stratified sandstone (braided stream), green sandstone-siltstone (delta), massive cobble conglomerate (transgressive shoreline lag), planar cross-stratified sandstones (nearshore lacustrine sandwaves), and thinning-upward sandstones (lacustrine storm sand sheets). The overlying mixed siliciclastic-carbonate unit contains: red siltstone (non-saline lake), red siltstonedolostone or dolomitic sandstone (saline lake), and halite-mudstone (ephemeral salt pans). Next is the upper siliciclastic unit with: sheet sandstones (lake infilling) and stromatolitic dolostone-chert (shoreline). After final desiccation of the lake terra rosa soils, collapse breccias and intraformational conglomerates developed. Paleocurrents, detrital zircon geochronology and sulfur isotopes indicate a change in drainage directions resulting in sand sheets infilling the saline lake. Sr isotopes reflect shallow groundwater circulation and lacustrine dolostone containing significant radiogenic Sr. Carbon and O isotopes are heavier upward in the saline lake deposits, probably due to evaporation and residence time effects. Most interestingly, REE patterns for dolomite in the dolocrete, stromatolitic shoreline deposits and overlying intraformational conglomerates have patterns similar to modern oxygenated groundwater, whereas the saline lake dolomites have hatshaped patterns resembling modern groundwater draining waterlogged, organic-rich areas. 65 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Role of felsic and feldpathic rocks in triggering subvolcanic emplacement of mafic intrusions: evidence from the Midcontinent Rift in northeastern Minnesota MILLER, James D. Dept. of Earth and Environmental Sciences and Precambrian Research Center, University of Minnesota Duluth, Duluth, MN 55812 Subvolcanic mafic intrusions make up over 40% of the igneous rocks of NE Minnesota associated with the 1.1Ga Midcontinent Rift. The greatest concentration of mafic intrusions comprise the enormous (~20,000 km3) Duluth Complex, which was emplaced into the basal section of a 5-10 km thick edifice of comagmatic volcanics. Most Duluth Complex intrusions, and mafic intrusions emplaced higher in the volcanic pile, occur as sheet-like bodies that commonly underlie either felsic rocks (rhyolite flows or granophyre intrusions) or feldspathic rocks (gabbroic to troctolitic anorthosite). In all cases, field relationships indicate that the felsic/feldspathic rocks are consistently older than the mafic rocks that underlie them. Three styles of mafic underplating are recognized in the Midcontinent Rift of northeastern Minnesota (Fig. 1): 1) Mafic layered intrusions beneath large granophyre bodies. Examples of this style of underplating are the Poplar Lake Intrusion beneath the Misquah Hill Granophyre in the Gunflint Trail area; the Sawbill Lake intrusion beneath the Eagle Mountain Granophyre, and the Sonju Lake Intusion beneath the Finland Granite. In all three cases, the mafic intrusion is well differentiated and in gradational contact with the overlying granophyre. The gradational contact is best explained by partial melting of the base of the granophyre and subsequent assimilation. 2) Mafic layered intrusions beneath anorthositic rocks. This style of mafic underplating is evident in almost all Layered Series intrusions of the Duluth Complex which are intruded beneath Anorthsotic Series rocks (Fig. 1). The Layered Series intrusions that will be highlighted in this talk include the Layered Series at Duluth, the Partridge River Intrusion, and the Tuscarora Intrusion. 3) Diabase sheets beneath thick rhyolite flows. Examples of this style of mafic underplating include the Endion Sill beneath the Tischer Creek Rhyolite and the Lester River Sill beneath the Lakewood Rhyolite, both in the Duluth area, the Beaver River Diabase beneath the Palisade Head Rhyolite in the Beaver Bay Complex, and the Mink Mountain diabase emplaced beneath and into the Grand Marais Felsites. Empirical evidence and density/viscosity/thermal considerations suggest that the felsic/feldspathic rocks served as density barriers to mafic magmas, which had risen into the upper crust to the point of neutral bouyancy. The felsic/feldspathic rocks not only triggered underplating of the mafic magmas, but also commonly served as thermal insulators to the underplated mafic bodies, thus resulting in their slow cooling and crystallization differentiation. 66 �Proceedings of the 61st ILSG Annual Meeting - Part 1 67 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geology of the North and South Temperance Lakes area of the Boundary Waters Canoe Area, Cook County, Minnesota - 2014 Precambrian field camp capstone mapping MILLER, Jim, BEAVER, Christopher, HAHN Timothy, MILLER Nikolas, PULIESE Joseph, and WRIGHT, Erick Precambrian Research Center, University of Minnesota Duluth, Duluth, MN 55812 As a capstone mapping project for the 2014 Precambrian field camp, a crew of five students under the supervision of Jim Miller conducted five days of field mapping bedrock geology in the South and North Temperance Lakes area. This area is located in the Boundary Waters Canoe Area west of Brule Lake in Cook County, Minnesota. The area is accessible from the Caribou Trail, which heads north from Tofte to a canoe landing on Brule Lake, and then via a 5-mile paddle to the west end of Brule Lake and a portage into South Temperance Lake. The main objective of this project was to conduct bedrock geologic mapping of rocks that previous capstone mapping has shown to comprise part of the footwall to the Sawbill Lake intrusion (Brooker and Miller, 2013). Previous studies of the Temperance Lakes area include reconnaissance mapping by Grout et al. (1959) and Davidson (1977). Grout’s (~1:100,000-scale) township maps of the area (Figures XXIII and XXIV, Grout et al., 1959) show it to contain granophyric granite and mafic volcanic rocks intruded by gabbro. Davidson’s (1:24,000-scale) reconnaissance map of the Cherokee Lake 7.5’ quadrangle show a similar mix of rock types in the Temperance Lakes area, but he subdivides the gabbroic rocks into an olivine gabbro unit and an anorthositic gabbro unit. The latter unit, Davidson correlates to the anorthositic series of the Duluth Complex. Capstone mapping conducted for UMD’s Precambrian field camp in the summers of 2007 (Frost et al., 2007), 2009 (Blakely et al., 2009), 2010 (Brooker et al., 2010), and 2011 (Asp et al., 2011) and field mapping conducted by Ben Brooker as part of his MS thesis at UMD in 2011 (MS, in preparation) revealed the existence of well differentiated, mafic layered intrusion which has been named the Sawbill Lake intrusion (SbLI; Brooker and Miller, 2012, 2013). Mapping of the lower contact of the SbLI showed is lower troctolitic cumulates were in contact with a footwall of evolved ferrodioritic cumulates that locally contain hornfels basalt inclusions. This observation and aeromagnetic data (Chandler, 1983) showing curvilinear anomalies in the footwall rocks that are conformable to similar anomalies internal to the SbLI suggest that another well differentiated mafic layered intrusion might exist beneath the SbLI. The 2014 capstone mapping project focused mapping shoreline exposures in the South and North Temperance Lakes which covers a 3 km wide by 4 km tall area extending north from the basal contact of the SbLI. The result of this mapping clearly shows that a well differentiated tholeitic mafic layered intrusion indeed is situated conformably beneath the SbLI. This as yet unnamed intrusion can be subdivided into seven distinct units based on dominant lithology, internal structure and position within the intrusion. Internal structure within the intrusion (layering and igneous foliation dips 20-40°) to the south. The base of the intrusion is exposed in the northern part of North Temperance Lake where a fine- to medium-grained, locally plagioclase-phyric, subophitic to ophitic olivine diabase is found in intrusive contact with granophryric leucogranite of the Misquah Hills granophyre – part of Duluth Complex Felsic Series. The diabase cuts the granophyre in an orthogonal pattern of N-S/EW dikes and shows chilled contacts. In the southern part of North Temperance Lake, the olivine diabase grades into an ophitic augite troctolite that locally displays moderate foliation and layering defined by augite oikocryst concentrations. This Pl+Ol cumulate persists upsection into the northern part of South Temperance Lake. In the mid-section of South Temperance Lake, a 68 �Proceedings of the 61st ILSG Annual Meeting - Part 1 troctolite-olivine gabbro transitional unit is defined by the intermittent (cyclical?) occurrence of cumulus augite and Fe-Ti oxide. The southern part of South Temperance Lake is dominated by a medium-grained, locally layered, moderately to well foliated, intergranular olivine oxide gabbro – a Pl+Cpx+Ox+Ol cumulate. Along the portage trail following the Temperance River south of South Temperance Lake, the olivine oxide gabbro transitions into an apatite ferrodiorite (Pl+Cpx+Ox+Ol+Ap cumulate), which in turn grades into a ferromonzodiorite with abundant basaltic hornfels inclusions. At the south end of the portage trail, the ferromonzodiorite abruptly transitions into a fine-grained, subophitic olivine diabase, which is the base of the overlying Sawbill Lake Intrusion. This olivine diabase is also observed to cross cut the upper three units of the Temperance Lakes sequence, clearly indicating that the Sawbill Lake was emplaced later by overplating the Temperance Lake sequence. Plans for the 2015 capstone mapping project are to follow the igneous stratigraphy defined in the Temperance Lakes area to the west and north into the Cherokee Lake area. REFERENCES Asp, K., Leu, A., Parisi, A., Sletten, D., Brooker, B., Miller, J., 2011, Bedrock geology of the Sawbill Lake area: University of Minnesota Duluth, Precambrian Research Center, PRC/MAP-2011-04, 1: 12,000. Blakely, S., Brown, A., Foley, D., Rowland, A., Stifter, E., and Miller, J., 2009, Bedrock geology map of Homer Lake and adjacent areas; Cook County, Northeastern Minnesota: University of Minnesota Duluth, Precambrian Research Center, PRC/MAP-2009-01, 1: 12,000. Brooker, B.P., and Miller, J.D., 2013, Bedrock geologic map of the Sawbill Lake Intrusion, Cook County, MN. Precambrian Research Center Map Series PRC/Map-2013-01, scale 1:24,000. Brooker, B.P., and Miller, J.D., 2012, Geology and petrology of a Mesoproterozoic layered mafic intrusion in portions of the Brule Lake and Cherokee Lake 7.5’ Quadrangles, northeastern Minnesota. Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 1 - Proceedings and Abstracts, v. 58, part 1, 15-16. Brooker, B.P.,Hadley, M.L., Markwood, L.W., Olson, J., Tomlinson, A.P., and Miller, J.D.,2010, Bedrock geologic map of the Jack Lake and Weird Lake areas, Cook County, northeastern Minnesota: University of Minnesota Duluth, Precambrian Research Center, PRC/Map-2010-05, 1: 12,000. Chandler, Val W, 1983, Aeromagnetic map of Minnesota, Cook and Lake counties: Minnesota Geological Survey, Aeromagnetic Map Series, Map A-1, scale 1:250,000 Davidson, D.M., 1977, Reconnaissance geologic map of the Cherokee Lake quadrangle, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series, M-30, scale 1:24,000 Frost, S.J., Juda, N.A., and Miller, J., 2007, Bedrock Geology Map of Homer Lake and Adjacent Areas; Cook County, Northeastern Minnesota: University of Minnesota Duluth, Precambrian Research Center, PRC/MAP-2007-02, 1: 12,000 Grout, F.F., Sharp, R.P., and Schwartz, G.M., 1959, The geology of Cook County, Minnesota: Minnesota Geological Survey Bulletin 39, 163 p. 69 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The mineralogy and petrology of a newly-discovered REE occurrence within the Coldwell Complex near Marathon, Ontario NIKKILA, D. and Zurevinski, S. Dept. of Geology, Lakehead University, Thunder Bay, ON P7B 5E1 The Coldwell Complex is situated within the Archean Schreiber-White River metavolcanicmetasediment of the Superior Province. Spanning over 25 km in diameter, it is the largest alkaline intrusion in North America (Figure 1). The 1108 +/- 1 Ma age of the Coldwell complex and close spatial proximity supports a strong relationship to the magmatism of the Keweenawan Midcontinent Rift (Heaman and Machado 1992). Early studies define three magmatic centers of the Coldwell Complex, which in order of intrusion are Center I, Center II and Center III (Mitchell and Platt 1982). Center I consists of an oldest phase gabbro, which borders a ferroaugite syenite to the east and north. Center II includes a nepheline-bearing biotite-gabbro and several intrusions of nepheline syenites, and Center III is composed of four syenites which in order of intrusion are: magnesiohornblende syenite, contaminated ferroedenite syenite, ferroedenite syenite, and quartz syenite. .000000 510000 .000000 520000 .000000 530000 .000000 540000 .000000 550000 .000000 .000000 5423000 .000000 $ 5414000 5414000 .000000 5423000 .000000 5432000 500000 .000000 .000000 5432000 .000000 490000 Radio Hiill .000000 Terrace Bay ! . ! . . !! . 5396000 .000000 Mink Creek ! P 490000 0 2.5 5 .000000 10 500000 15 .000000 20 510000 25 .000000 520000 .000000 530000 .000000 540000 .000000 Marathon 550000 .000000 Frog ! P 5396000 5405000 ! P 5405000 .000000 ! . .000000 30 Km Figure 1: Regional geology of the Coldwell Complex (purple). As part of an undergraduate thesis project, the focus of this study was to classify syenite rocks related to the three intrusive centers, and identify any REE-bearing minerals present. Fieldwork was completed along HW-17 roadcuts, and North of the highway on Canada Rare Earth claim blocks, termed the ‘Radio Hill’ occurrence (Figure 1). The syenites of the Radio Hill occurrence had not previously been identified due to limited access to the area. Through petrography, Highway-17 samples were classified as Center II Nepheline syenites, where the 70 �Proceedings of the 61st ILSG Annual Meeting - Part 1 textures of amphibole, biotite, pyroxene and natrolite compared well to the nephelene syenites previously described by Mitchell and Platt (1982). Syenites of the Radio Hill occurrence were classified as Center III, specifically, ferroedenite syenites. The Radio Hill syenites show an increase in the modal abundance of quartz, and a decrease in natrolite. Compositions of the amphiboles from the Radio Hill syenites compare well with the silicic ferroedenite and hastingsitic hornblende compositions, with a trend to Na, Si, and Fe enrichment with Ca and Al depletion. Radio Hill mica has been identified as annite end-member compositions, with Mg # ranging from 0.082 to 0.294. Radio Hill plagioclase feldspar compositions show An % from 0 to 12.04 %, representing albite to oligioclase end-members. Rare earth element minerals were described and identified from the Radio Hill occurrence using qualitative identification methods with the scanning electron microscope (SEM-EDX) at Lakehead University. Minerals found occurring in the Radio Hill syenites, in order of abundance, include apatite (elevated La, Ce, Nd, and Th), plumbopyrochlore ((Pb, Y, U, Ca)2-xNb2O6(OH)), ceriopyrochlore ((Ce, Ca, Y)2(Nb, Ta)2O6(OH, F)), monazite ((La, Ce, Nd)PO4), and fluorite. REFERENCES Heaman, L.M. and Machado, N. 1992. Timing and origin of midcontinent rift alkaline magmatism, North America: evidence from the Coldwell Complex; Contributions to Mineralogy and Petrology, v. 110, p. 289-303. Mitchell, R.H., Platt, R.G., Lukosius-Sanders, J., Artist-Downey, M. and Moogk-Pickard, S. 1993. Petrology of syenites from centre III of the Coldwell alkaline complex, northwestern Ontario, Canada; Canadian Journal of Earth Sciences, v. 30, p. 145-158. Mitchell, R.H. and Platt, R.G. 1982. Mineralogy and Petrology of Nepheline Syenites from the Coldwell Alkaline Complex, Ontario, Canada; Journal of Petrology, v. 23, p. 186-214. 71 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Petrology, geochemistry and mineral chemistry of the Crystal Lake and Mount Mollie mafic intrusions, northwestern Ontario O’BRIEN, Sean1, HOLLINGS, Peter1, and MILLER, Jim2 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada 2 Department of Geological Sciences, University of Minnesota Duluth, 1049 University Drive, Duluth, MN 55812, United States The Midcontinent Rift (MCR) extends ~2500 km through Canada and the United States, and comprises ~1,500,000 km3 of volcanic and intrusive rocks spanning four distinct stages of activity ranging from 1150-1087 Ma (Heaman et al., 2007). The 1108-1105 and 1100-1094 Ma periods has been interpreted by Heaman et al. (2007), to represent the main formation and maturation of the rift system and is associated with the majority of the igneous activity, producing mafic to ultramafic intrusions, basaltic sills, dikes and flows as well as alkaline rocks. The extent and volume of magmatic activity has led previous researchers to conclude that a plume was most likely the cause of the MCR (Miller and Nicholson, 2013). In this study, two mafic intrusions related to the MCR will be investigated using detailed petrography, geochemistry and mineral chemistry. The two intrusions, Crystal Lake and Mount Mollie, are located ~40 km south of Thunder Bay, Ontario and are located within a few km of each other. Crystal Lake is a Y-shaped layered intrusion with a north limb striking W-NW for 5 km and a south limb striking E-NE for 2.75 km. Mount Mollie varies from 60 to 350 m wide and extends for ~35 km, and is located just east of the Crystal Lake intrusion (Fig. 1). The intrusions have been targets for exploration for the past few decades as they both contain disseminated sulphides and are host to Ni-Cu-PGE mineralization (Smith and Sutcliffe, 1987; Lightfoot and Lavigne, 1995). The two units are very similar, consisting mainly of gabbro, troctolite and olivine gabbro with some disseminated sulphides, chromite and other spinels. The close spatial relationship led to the belief that the intrusions were related and contemporaneous (Lightfoot and Lavigne 1995), however, recent geochronology has shown that the Mount Mollie intrusion has an age of 1109.3 ± 6.3 Ma whereas Crystal Lake has been dated at 1099.6 ± 1.2 Ma (Hollings et al., 2010). Figure 1. Generalized map of Crystal Lake gabbro, Mount Mollie dike and surrounding rocks, adapted from Cundari (2013). 72 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Preliminary results suggest the Mount Mollie dike may in fact be contemporaneous with Crystal Lake. Core logging revealed that ~10 m of intrusive mafic rocks occur at the top of the drill core, overlaying ~20 m of sandstone, followed by gabbroic rocks. It is unclear whether the gabbros above and below the sandstone layer are directly related. With two generations of igneous activity a possibility, the temporal relationship between the two intrusions needs to be investigated further. Full descriptions of the mineralogy and textures of both intrusions, particularly the layering styles (i.e., modal, graded, and phase layering), will be completed through core logging and detailed petrography of thin sections. These results will be combined with whole-rock geochemical data to investigate fractionation trends, mixing and crustal contamination signatures to understand the evolution of the intrusions and the genesis of the mineralized horizons. Mineral chemistry of olivine and spinels will be used to determine cryptic layering. Olivine mineral chemistry is especially important as it can further constrain the evolution of the magmas, with an emphasis on the forsterite-fayalite and nickel contents, and to give insight into parental melt compositions. Spinel mineral chemistry will be used to help understand parental melt compositions as chromite is one of the first minerals to crystalize from the melt and is refractory. SEM analysis of the platinum group minerals (PGMs) will be conducted to determine textural and mineralogical associations as well as what compositional varieties are present, (i.e. alloys, sulphides, arsenides, etc.). This detailed study will allow us to build a model for how these two intrusions formed, how they fit into the Midcontinent Rift, and what the style of mineralization is. REFERENCES Cundari, R.S., Smyk, M., Campbell, D. and Puumala, M., 2013. Geology, Geochemistry and Cu-Ni-PGE Mineralization of the Crystal Lake Gabbro, 6th Annual PRC Professional Workshop Cu-Ni-PGE Deposits of the Lake Superior Region, Duluth, Minnesota. Heaman, L., Easton, R., Hart, T., Hollings, P., MacDonald, C. and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences, 44: 10551086. Hollings, P., Smyk, M., Heaman, L.M. and Halls, H., 2010. The geochemistry, geochronology and paleomagnetism of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada. Precambrian Research, 183: 553-571. Lightfoot, P.C. and Lavigne, Jr., M.J., 1995. Nickel, copper, and platinum group element mineralization in Keweenawan intrusive rocks: new targets in the Keweenawan of the Thunder Bay region, northwestern Ontario; Ontario Geological Survey, Open File Report 5928, 32p. Miller, J., and Nicholson, S., 2013, Geology and mineral deposits of the 1.1Ga Midcontinent Rift in the Lake Superior region: an overview, in field guide to copper-nickel-platinum group element deposits of the Lake Superior region. Precambrian Research Center, Guidebook, 13: 1-50. Smith, A.R., and Sutcliffe, R.H., 1987. Keweenawan intrusive rocks of the Thunder Bay area. in Summary of Field Work, 1987. Ontario Geological Survey, Miscellaneous Paper 137: 248-255. 73 �Proceedings of the 61st ILSG Annual Meeting - Part 1 So, an Environmental Impact Statement is required: Some Lake Superior area Geologic Parameters for Geologists, Consultants, Companies, and Regulators PETERSON, Dean1 1 Peterson Geoscience LLC, 306 West Superior Street, Suite 410, Duluth, Minnesota, 55802. Since the birth of the environmental movement in the early 1970s, modern society has evolved into an “environmentally concerned world”, where the public demands that industry (herein the mining industry) reduce its physical, social and environmental footprint have translated into specific national legislation. Under United States environmental law, an environmental impact statement (EIS) is a document required by the National Environmental Policy Act (NEPA) for certain actions "significantly affecting the quality of the human environment". Similar Canadian environmental law documents are required by the Canadian Environmental Assessment Act. An EIS type document is a project specific tool for decision making, and typically includes four sections: (1) an introductory statement of the Purpose and Need of the Proposed Action, (2) a description of the Affected Environment, (3) a Range of Alternatives to the proposed action, and (4) an Analysis of the environmental of each of the possible alternatives. In the Lake Superior area, certain environmental groups vehemently oppose all activities related to the mining industry and actively coordinate their opposition to the general public through meetings, web sites, social media, and the press. Most coordinated opposition ties directly into published EIS documents in general, and specifically on the Analysis of water quality for the proposed actions. To the author, environmental misinformation, half-truths, and outright lies constitute much of the anti-mining opposition. Have you read about “Sulfide Mining”, acid mine drainage, degraded water quality, highly fractured bedrock, canoeing to the white house, the 500 years, etc…..? So, what are field geologists to do? Geologists must first recognize and defend the fact that the basis for every type of geologic study related to an EIS is fundamentally rooted in observations made of rocks in their natural habitat, “in the field”. An intimate understanding of the projects geology gives company geologists an appreciation of the coherent and compelling field-based scientific arguments from which all other geology-based EIS interpretations grow. Company geologists working on mining related projects have to interact with a diverse group of people during the EIS process, and must vigorously defend the fundamental observations of the projects “in the field” geology. Geologists have to ask themselves this question: Do the lawyers, PR folks, environmental consulting firms, regulators, NGO’s, and the general public know the details of the projects geology as much as I do? As geologists, we must always remember Francis Pettijohn’s famous quote, “The rocks are the final court of appeal”. This talk will highlight some fundamental geologic parameters of the Lake Superior area that geologists, consultants, companies, and regulators need to understand to better design and complete mining related EIS documents. 74 �Proceedings of the 61st ILSG Annual Meeting - Part 1 New airborne geophysical data for the Lake Superior Region of northwestern Ontario: A new tool for the identification of Neoarchean to Mesoproterozoic structures and associated mafic-ultramafic intrusions PUUMALA, Mark1, CUNDARI, Rob1, CAMPBELL, Dorothy1, RAINSFORD, Desmond2, and METSARANTA, Riku2 1 Ontario Geological Survey, Ministry of Northern Development and Mines, Resident Geologist Program, Suite B002, 435 James St. South Thunder Bay, ON, P7E 6S7, Canada 2 Ontario Geological Survey, Ministry of Northern Development and Mines, Earth Resources and Geoscience Mapping Section, 933 Ramsay Lake Road, Sudbury, ON, P3E 6B5, Canada. Recent discoveries of mafic-ultramafic intrusion hosted Ni-Cu-PGE mineralization throughout the Midcontinent Rift (MCR) region (e.g. in Canada: Current Lake, Sunday Lake, Thunder, Steepledge, Marathon; in the U.S.A. Tamarack, Eagle) have revitalized exploration interest in the area surrounding Lake Superior. Airborne magnetic and radiometric surveys that were flown during 2014 (Figure 1; Ontario Geological Survey, 2015a, b) have provided new highresolution public domain airborne geophysical coverage that will assist in these exploration efforts. The new surveys cover a large portion of the northwest Lake Superior region in Ontario including areas underlain by MCR-related rocks, Paleoproterozoic and Mesoproterozoic sedimentary rocks of the Animikie and Sibley Groups, as well as parts of the Archean Quetico, Wawa and Wabigoon subprovinces. In the Lake Superior region, Ni-Cu-PGE mineralized mafic-to-ultramafic rocks were emplaced in a variety of settings during several tectonic events that occurred over a time period extending from the Mesoarchean to the Mesoproterozoic (Smyk et al. 2002, Smyk and Franklin 2007). As a result, these intrusions display a wide range of geochemical affinities and morphologies. In spite of these geochemical and physical differences, most of these mafic-toultramafic intrusions have a close spatial association with major crustal-scale structures (Rogers et al. 1995, Hart and MacDonald 2007), and many can be recognized by their distinctive magnetic signatures (i.e., positive or negative anomalies). When the new airborne survey magnetic data are combined with data from previous magnetic surveys (Ontario Geological Survey 2003, 2004), they highlight numerous structures that could have controlled the emplacement of mineralized MCR-related intrusions into Archean country rocks along the northwest margins of the rift north of Thunder Bay. One such structure is marked by several magnetic discontinuities and anomalies that can be traced for at least 145 km along an east-northeast (063°) trending line from the southeast end of Northern Light Lake (near the Ontario-Minnesota border), through to Greenwich Lake (50 km northeast of Thunder Bay). This structure is approximately parallel to Midcontinent Rift-related faults and dikes that are located farther to the southeast (Sutcliffe 1991) and Neoarchean faults that have been mapped to the northwest (Hart and MacDonald 2007). A second parallel structure is also evident in the magnetic data approximately 10 km further to the south-southeast. Ni-Cu-PGE mineralized mafic-ultramafic rocks of the Sunday Lake, Steepledge Lake and Current Lake intrusive complexes occur in a linear array that closely follows the Northern LightGreenwich Lakes structure (NLGLS), suggesting that it may have played a role in their emplacement. The NLGLS, which has been mapped as a fault over a portion of its length (Lodge et al. 2014), is also located in close proximity to Neoarchean gold (Tower Mountain) and komatiite-hosted Ni-Cu-PGE (Bateman Lake) mineralization in Conmee Township. This observation, together with its proximity to Mesoproterozoic and Neoarchean faults of similar orientations, suggests that the NLGLS may have been tectonically active both during the accretion 75 �Proceedings of the 61st ILSG Annual Meeting - Part 1 of the Superior Craton, and during the Midcontinent Rift event. As a result, it presents an attractive target for both Ni-Cu-PGE and gold exploration. Preliminary observations indicate that the NLGLS may continue further to the northeast, where it eventually merges with the Gravel River fault (GRF) approximately 60 km north of Terrace Bay). The GRF extends northeast along approximately the same trend to the James Bay Basin (Williams 1991), and it is also interesting to note its close spatial association with the Albany Graphite deposit. Figure 1. Location of new airborne geophysical surveys carried out over the Thunder Bay region during 2014 (geology from Ontario Geological Survey 2011). The data were released during the winter and spring of 2015 as Geophysical Data Sets 1077 (Mahon & Flatrock Lakes) and 1078 (Lac des Mille Lacs – Nagagami). REFERENCES Hart, T.R. and MacDonald, C.A. 2007. Proterozoic and Archean geology of the Nipigon Embayment: Implications for emplacement of the Mesoproterozoic Nipigon diabase sills and mafic to ultramafic intrusions; Canadian Journal of Earth Sciences v.44, p.1021-1040. Lodge, R.W.D., Ratcliffe, L.M. and Walker, J.A. 2014.Geology and mineral potential of Sackville and Conmee Townships, Wawa Subprovince; in Summary of Field Work and Other Activities 2014, Ontario Geological Survey, Open File Report 6300, p.9-1 to 9-17. Ontario Geological Survey 2003. Ontario airborne geophysical surveys, magnetic data, Shebandowan area; Ontario Geological Survey, Geophysical Data Set 1021 - Revised. Ontario Geological Survey 2004. Ontario airborne geophysical surveys, magnetic and gamma-ray spectrometer data, Lake Nipigon Embayment Area; Geophysical Data Set 1047. Ontario Geological Survey 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release-Data 126- Revision 1. Ontario Geological Survey 2015a. Ontario airborne geophysical surveys, magnetic and gamma-ray spectrometric data, grid and profile data (ASCII and Geosoft® formats) and vector data, Mahon Lake and Flatrock Lake areas; Ontario Geological Survey, Geophysical Data Set 1077. Ontario Geological Survey 2015b. Ontario airborne geophysical surveys, magnetic and gamma-ray spectrometric data, grid and profile data (Geosoft® format) and vector data, Lac des Mille Lacs–Nagagami Lake area; Ontario Geological Survey, Geophysical Data Set 1078b. 76 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Rogers, P.C, Thurston, P.C, Fyon, J.I, Kelly, R.I. and Breaks, F.W. 1995. Descriptive mineral deposit models of metallic and industrial deposit types and related mineral potential assessment criteria; Ontario Geological Survey, Open File Report 5916, 241p. Smyk, M.C. and Franklin, J.M. 2007. A synopsis of mineral deposits in the Archean and Proterozoic rocks of the Lake Nipigon region, Thunder Bay District, Ontario; Canadian Journal of Earth Sciences v.44, p.1041-1053 Smyk, M.C., Mason, J.K., Schnieders, B.R. and Stott, G.M. 2002. A synopsis of Archean and Proterozoic platinum group element mineralization in the Thunder Bay District, Ontario; in 9th International Platinum Symposium, Billings, Montana, July 25, 2002, Extended Abstract Vol., p.433-434. Stone, D. 2010. Precambrian geology of the central Wabigoon Subprovince, northwestern Ontario; Ontario Geological Survey, Open File Report 5422, 130p. Sutcliffe, R.H. 1991. Proterozoic geology of the Lake Superior region; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.627-658. Williams, H.R. 1991. Quetico Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.383-403. 77 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Spectrum of volcanogenic massive sulfide deposits in the Penokean Volcanic Belt, Great Lakes Region, USA QUIGLEY, Patrick* and MONECKE, Thomas Department of Geology and Geological Engineering, Colorado School of Mines, 1516 Illinois Street, Golden, Colorado 80401 The Paleoproterozoic (ca. 1880Ma) Penokean volcanic belt extends for over 250 kilometers across northern Wisconsin and western Michigan. The dominantly submarine volcanic rocks comprising the belt were formed in an island arc-related setting at the southern edge of the Superior craton. Despite relatively minor exploration, the majority of which occurred intermittently from 1960 to 1990, approximately 100 million metric tons of polymetallic massive sulfide ores have been delineated. Only the supergene enrichment zone of the Flambeau deposit has reached commercial production. The mined resource accounts for less than 2% of known mineral reserves, which makes the Penokean volcanic belt one of the most accessible, undeveloped, and underexplored volcanic terranes worldwide. The present study aims to characterize the volcanic setting, deposit characteristics, and alteration signature of significant deposits within the Penokean volcanic belt to provide a comprehensive metallogenetic model. To accomplish this goal, detailed core logging has been conducted at seven deposits across the belt, namely Back Forty, Bend, Flambeau, Horseshoe, Lynne, Reef, and Ritchie Creek. Representative sampling has been conducted at all deposits for detailed petrographic and geochemical investigation. Ongoing research has revealed a wide spectrum of volcanic environments and alteration styles across the Penokean volcanic belt. All major deposits occur within felsic-dominated volcanic successions and are hosted by vent-proximal volcanic facies associations. For example, the Back Forty deposit is hosted within a felsic succession (apparent stratigraphic thickness of 1,200 m) comprising coherent rhyolite units and associated volcanic breccias. Felsic volcanism was broadly contemporaneous with the deposition of mass-flow-derived volcaniclastic debris presumably generated by an explosive eruption of a rhyolite source. Mafic-dominated host rock successions are less common in the Penokean volcanic belt and appear to host some of the smaller tonnage deposits, including Horseshoe and Ritchie Creek. The styles of hydrothermal alteration vary between deposits, with sericite-chlorite-quartz alteration occurring at Back Forty, Bend, and Horseshoe. Acid-style alteration represented by andalusite-biotite-sericite schists has been noted at Flambeau and calc-silicate mineral assemblages are present at Lynne, Ritchie Creek, and Reef. Calc-silicate mineral associations have also been observed at the Pelican River and Spirit deposits, possibly suggesting that the volcanic host rocks were originally interbedded with limestone. Regional metamorphism varies from lower greenschist to amphibolite grade and has obscured relationships in some deposits. Most notable, primary volcanic textures are difficult to recognize at the Reef deposit, which is an unusual disseminated to quartz-sulfide vein confined Au-Cu deposit hosted by strongly deformed and recrystallized rocks. Recognition of significant variations in setting and deposit characteristics across the Penokean volcanic belt likely reflects first-order tectonostratigraphic controls during the development of the Penokean orogeny. 78 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Geochemical and petrologic characterizations of peridotite, Marquette County, Michigan SASSO, Andrew, and THAKURTA, Joyashish Department of Geosciences Western Michigan University1903 W Michigan Ave Kalamazoo MI 49008-5241 USA The discovery of the Eagle magmatic sulfide deposit, in 2002, sparked a renewed interest in exploration for magmatic sulfide mineral deposits associated with peridotite in Michigan’s Upper Peninsula. This study is a preliminary attempt to determine if other sulfide mineral deposits could potentially exist in association with the peridotites of Marquette County, Michigan. In order to achieve its goal, this study is attempting to determine if any petrologic or geochemical relationship exists between the peridotites of Marquette County, Michigan. As shown by Figure 1, peridotite has been mapped at four locations across the county: the Yellowdog Peridotite, located at the site of the Eagle Mine (Rossell and Coombes, 2005), the Presque Isle Peridotite, located in Marquette (Gair and Thaden, 1968), the Deer Lake Peridotite, located just north of Ishpeming (Clark, Cannon, and Klassner, 1975), and Black Rock Point, located north of Big Bay (Case and Gair, 1965). Field work was conducted in May, 2014 and was followed by petrographic analyses of collected samples. A peridotite rock unit could not be located at Black Rock Point by initial field work and petrographic studies. The area is composed of three major rock units. The southernmost unit is gabbro, the largest of the three units present. To the north there is an abrupt change to a heavily veined granite which is cut by at least two mafic dikes. The northern most unit of the area is a gneissic rock which has a much smaller outcrop than the other units. Thin section analysis of samples collected from Presque Isle revealed that the rock is a serpentenized lherzolite. Primary minerals include olivine, clinopyroxene, and orthopyroxene. A large portion of the rock has been altered to serpentine. Other secondary minerals such as chlorite and calcite are also present. Thin sections from samples of the Deer Lake peridotite show a fine-grained rock with a texture suggestive of hypabyssal origin, which has been heavily serpentenized. Serpentine is by far the most abundant mineral in all samples analyzed thus far. Primary minerals including olivine, clinopyroxene and orthopyroxene are present in small quantities. Other secondary minerals include chlorite, and calcite. New geochemical, petrologic, and structural data collected by this study will be included in detailed geologic maps of each site. The investigation will address the relationships of the peridotite units with respect to one another as well as their relationships with the surrounding rocks. The final results generated from this study will be useful in the creation of a new set of criteria to assess the mineralization potentials of peridotite units in Marquette County. 79 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1. Modified map of locations of the four research areas and their surrounding geology. Compiled by Simms, 1992. REFERENCES Case, James and Gair, Jacob., 1965Aeromagnetic Map of Parts of Marquette, Dickinson, Baraga, Alger, and Schoolcraft Counties, Michigan, and Its Geologic Interpretation. Clark, Lorin D., William F. Cannon, and J. S. Klasner., 1975, Bedrock Geologic Map of the Negaunee SW Quadrangle, Marquette County, Michigan. Reston, VA: Survey. Gair, Jacob Eugene, and Robert E. Thaden., 1968, Geology of the Marquette and Sands Quadrangles, Marquette County, Michigan. Rossell, Dean and Coombes, Steven., 2005, The geology of the Eagle Nickel-Copper Deposit Michigan, USA. Kennecott Minerals Co. Sims, P. K.., 1992, Geologic Map of Precambrian Rocks, Southern Lake Superior Region, Wisconsin and Northern Michigan. U.S. Geological Survey. 80 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Petrologic study of the “Chill” zone of the Layered Series at Duluth: Testing a possible plutonic-volcanic correlation within the Midcontinent Rift SAUER, Sarah and MILLER, Jim Department of Earth and Environmental Sciences, University of Minnesota Duluth, Duluth, Minnesota 55812 The Duluth Complex is a multiple intrusive mafic complex that represents the largest exposed plutonic component of the 1.1 Ga Midcontinent rift. Results from extensive field mapping and petrologic studies (Miller and Green, 2008a, 2008b; Green and Miller, 2008) of the mafic cumulates comprising the type locality of the Duluth Complex at Duluth have confirmed that it is composed of two fundamentally distinct rock series, and provided further detail regarding the igneous stratigraphy, internal structure of each series and thus the petrogenetic relationship between them. The Anorthositic Series (DAS), which forms a ~ 1-km thick cap to the Duluth Complex at Duluth, is a suite of structurally complex plagioclase-rich gabbroic rocks which are interpreted to have formed by multiple intrusions of a plagioclase crystal mush from a lower crustal magma chamber (Miller and Weiblen, 1990). Underlying the DAS cap is a 3-4.5 km thick, well differentiated, stratiform sequence of troctolitic to gabbroic cumulates forming the layered series at Duluth (DLS). The DLS is thought to have formed by open system crystallization differentiation (Miller and Ripley, 1997). Based on lithology and stratigraphic position, the DLS can be subdivided into five major zones: basal contact zone, troctolite zone, cyclic zone, gabbro zone and upper contact zone. The DAS had long been interpreted to be significantly older than the DLS based on the abundance of DAS inclusions in the DLS and, especially, on the occurrence of a fine-grained mafic rock that occurs at the sharp upper contact of the DLS with the overlying DAS, referred to as the DLS “chill”. However, high precision U-Pb ages from DAS and DLS samples (Paces and Miller, 1993) has shown that these two rock series are essentially identical in age at 1099±0.5 Ma relative to the 30m.y. window of MCR magmatism. This revelation warrants a reinterpretation of the relationship between the two series, along with reevaluating the origin of the DLS “chill.” Since the similar ages of the DLS and DAS preclude the DLS “chill” being a thermal quench of DLS parental magma against the DAS, Miller (Miller and Ripley, 1997; Miller, 2011) has suggested that quenching of DLS magma was caused by the decompression of a volatile-saturated magma accompanying volcanic venting from the subvolcanic DLS chamber. Several features lend evidence in support of a decompression quenching of hydrous magma interpretation including: 1) the evolved composition of the DLS “chill”, 2) the presence of biotite phenocrysts in the “chill”, and 3) the extensive hydrothermal alteration of overlying DAS rocks. Miller has further suggested that periodic venting of hydrous magma may have played an important role in the formation of the cyclic zone in the medial part of the DLS, particularly the occurrence of microgabbro cumulates in the upper parts of phase-layered macrocycles. This study seeks to test whether the DLS “chill” composition could be formed by decompression quenching of a volatile saturated magma and whether that magma is in equilibrium with the microgabbros of the Cyclic Zone and possible volcanic products represented in the NSVG overlying the Duluth Complex. To accomplish this, the lithological, petrographic and geochemical attributes of the DLS “chill”, microgabbros and flows from the NSVG were evaluated. Taking advantage of the fact that Brannon (1984) analyzed the chemistry of all mafic lavas occurring between 3-5 kilometers above the top of the Duluth Complex, the chemostratigraphy of the overlying NSVG were searched for compositions matching the DLS “chill”. If a correlative lava can be found, it would provide valuable constraints on the depth and pressure of the DLS magma 81 �Proceedings of the 61st ILSG Annual Meeting - Part 1 chamber. Once a pressure of crystallization is established, the DLS “chill” composition can be applied to PELE, a MELTS-based phase equilibrium modeling program developed by Boudreau (2006), to simulate whether pressure fluctuations caused by devolatilization and venting could effectively quench a volatile-saturated magma during decompression. PELE can also be used to evaluated the comagmatic relationship between the DLS “chill” and Cyclic Zone microgabbros by determining the mineral phases in equilibrium with the “chill” composition and comparing them to the phase compositions observed in the microgabbros. Preliminary results indicate two sequences of flows (Brannon’s (1984) flows 28-38 and 65) have the best match to the “chill” composition. Both series had the best fit for most major and trace elements, but most notable, however, was flow 65 which corresponds to the upper-most flow of an eight-flow sequence that defines an obvious differentiation trend. Flow 65 shows a good fit to the “chill” composition and parallel patterns of depletion of incompatible trace elements evident in the successively lower flows of the differentiation sequence. Treating these two sequences of flows as the potential volcanic products would indicate that the top of the DLS was emplaced within the volcanic edifice at a depth of ~4-5km. This depth is consistent with estimates for the formation of shallow reservoirs of magma (2-4km) beneath mafic volcanic centers like Kilauea in Hawaii (Ryan, 1987). Modeling with the PELE program is just underway. We hope to have results to present at the time of the meeting. REFERENCES Boudreau, A., 2006, Pele. (7.07). Computer modeling program. Duke University. (www.nicholas.duke.edu/eos/) Brannon, J.C. 1984. Geochemistry of successive lava flows of the Keweenawan North Shore Volcanic Group. Ph.D. thesis, Washington University, St. Louis, MO, . Green, J.C., and Miller, J.D., Jr., 2008, Bedrock geology of the Duluth quadrangle, St. Louis County, Minnesota. Minnesota Geological Survey Miscellaneous Map M-182, scale 1:24,000 Miller, J.D., 2011, Igneous stratigraphy of the Layered Series at Duluth – Type intrusion of the Duluth Complex. Institute on Lake Superior Geology, Proceedings Vol. 57, Part 2 - Field Trip Guidebook, p. 3-29. Miller, J.D., Jr.,and Ripley, E.M., 1996. Layered intrusions of the Duluth Complex, Minnesota, USA, in Cawthorn, R.G., ed., Layered intrusions: Amsterdam, Elsevier Science, p.257-301. Miller, J.D., Jr., and Green, J.C., 2008a, Bedrock geology of the Duluth Heights and eastern portion of the Adolph quadrangles, St. Louis County, Minnesota. Minnesota Geological Survey Miscellaneous Map M-181, scale 1:24,000 Miller, J.D., Jr., and Green, J.C., 2008, Bedrock geology of the West Duluth and eastern portion of the Esko quadrangles, St. Louis County, Minnesota. Minnesota Geological Survey Miscellaneous Map M-183, scale 1:24,000. Miller, J.D., Jr., and Weiblen, P.W., 1990, Anorthositic rocks of the Duluth Complex: Examples of rocks formed from plagioclase crystal mush: Journal of Petrology, v. 31, p. 295–339. Paces, J.B., and Miller, J.D., Jr. 1993. Precise U-Pb ages of Duluth Complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical Petrogenetic, paleomagnetic and tectonomagmatic processes associated with the 1.1 Ga Midcontinent rift system: Journal of Geophysical Research 98: 13,997-14,013. Ryan, M.P. 1987. Neutral buoyancy and the mechanical evolution of magmatic systems. In Magmatic processes physiochemical principles. B.O. Mysen (ed.) The Geochemical Society Special Publication 1, p. 259-287. 82 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Images on stone: Pictographs of the Ignace area, northwestern Ontario SMYK, Dennis W.1, ROSS, William.2 and SMYK, Mark C.3 1 P.O. Box 989,153 Balsam St., Ignace, ON P0T 1T0 2 William Ross Archaeological Research Associates, 189 Peter Street, Thunder Bay ON P7A 5H8 3 Resident Geologist Program, Ministry of Northern Development and Mines, Ontario Geological Survey, Suite B002, 435 James Street South, Thunder Bay, ON P7E 6S7 More than 400 rock paintings (pictographs) had been documented on outcrops of the Canadian Shield from Quebec, across Ontario and as far west as Saskatchewan (Rajnovich 1994). The senior author, an avocational archaeologist, has found and documented an additional 150 pictograph sites over the past 50 years. Most of them are situated within 160 km of Ignace, midway between Thunder Bay and Kenora in northwestern Ontario. Pictographs are the legacy of the Algonquian-speaking, early Cree and Ojibway peoples, whose roots may extend to the beginnings of post-glacial human occupancy in the area almost 10,000 years ago. This region of northwestern Ontario is underlain by Archean rocks of the Superior Province that are overlain by thin, discontinuous, unconsolidated glacial, lacustrine and organic deposits. The vast majority of pictograph sites are located on shoreline bedrock exposures of lakes and interconnecting rivers. Sites with relatively homogeneous and leucocratic bedrock faces (e.g., granitoids) are preferred, although more mafic rocks also host pictographs. Many sites consist of single elements (e.g., a canoe), while others are more complex with many figures and motifs. In many cases, there may only be one site on a lake, whereas in other examples, several sites may be scattered along a lakeshore. The largest concentration of pictographs found to date has 28 sites on cliffs on both sides of a 8 km-long, narrow stretch of lake. In almost all cases, the paintings are red, but one site with a gold handprint was found and documented by the senior author on Eagle Lake, southwest of Dryden. Discovered in the mid1960s, the unique Smyk Site (Figure 1), northeast of Ignace, is the only known local site with the three distinct colours of red, gold and dark purple. Although there is some scattered and anecdotal evidence of ochre quarrying, there has been only limited research into possible sources of natural pigmenting agents and the ochres themselves. Unlike sites in the Thunder Bay area, the geoarchaeology of the area west of the Lake Superior basin remains largely cursory. Much work remains in identifying and documenting pictograph sites and relating them in the larger context of spiritual places, habitation sites, transport and trading routes, local geologic materials and deglaciation history. REFERENCE Rajnovich, G. 1994. Reading Rock Art: Interpreting the Indian Rock Paintings of the Canadian Shield; Natural Heritage / Natural History Inc., Toronto, ON. 83 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Figure 1: Pictographs of three distinct colours (red, gold and dark purple) on Archean granitoid rocks at the Smyk Site, northeast of Ignace. Exposed panel is approximately 1 m across. 84 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The petrology, mineralization and regional context of the Thunder mafic to ultramafic intrusion, Midcontinent Rift, Thunder Bay, Ontario TREVISAN1, Brent, HOLLINGS1, Pete, AMES2, Doreen and RAYNER2, Nicole 1. Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada. 2. Geological Survey of Canada, Ottawa, Ontario, K1A 0E8, Canada The 1108 Ma Thunder mafic to ultramafic intrusion is a small, 800 x 100 x 500 m, Cu-PGE (platinum group element) mineralized body, located on the outskirts of Thunder Bay, Ontario. The intrusion was explored by Rio Tinto (formerly Kennecott Canada Exploration Inc.) in 2005 and 2007 (Fig. 1; Bidwell and Marino, 2007). It is associated with the early magmatic stages of the Midcontinent Rift (MCR) based on geochemical similarities to mafic and ultramafic rocks of the Nipigon Embayment and a 207Pb/206Pb zircon age of 1108.0 ± 1.0 Ma (Trevisan, 2014; Trevisan et al., 2015). Figure 1: Geological map of the greater Thunder Bay area including major road networks and outline of the current mineral claims that enclose the Thunder intrusion. The Thunder intrusion is located on the outskirts of the City of Thunder Bay, within Gorham Township and situated within the eastern limb of the Archean Shebandowan greenstone belt. Geospatial data from OGS (2011). The Thunder intrusion is similar to the other known mineralized early-rift MCR intrusions; however, it is the only known mafic/ultramafic intrusion of the MCR hosted in an Archean greenstone belt (Shebandowan). Major textural and geochemical differences can be used to subdivide the intrusion into a lower mafic to ultramafic unit and an upper gabbroic unit; the similar trace and rare earth element ratios of the two units suggest a single magmatic pulse that has undergone subsequent fractional crystallization and related cumulate phase layering. The estimated parental composition of the Thunder intrusion has a mg# (MgO/(MgO+FeOTot), mole %) of 57 which represents a more evolved magma than other early-rift mafic to ultramafic intrusions. This may indicate the involvement of multiple staging chambers during the ascent of the parent magma. 85 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Trace and rare earth element geochemical patterns are consistent with a mantle plume ocean island basalt-like source but, with high Th concentrations and the presence of a negative Nb anomaly, indicative of contamination . ԐNdt values from the intrusion range between -0.7 and +1.0, but lack trends indicative of progressive wall rock contamination, whereas the 87Sr/86Sri ratios range from 0.70288 to 0.70611 and trend towards wall rock values of 0.70712 and 0.70873. The weak correlation at Thunder between ԐNdt and 87Sr/86Sri is also a feature of the Nipigon Sills where it has been interpreted to be due to shallow-level crustal contamination whereas plots of MgO and SiO2 versus ԐNdt indicate contamination at depth by an older crustal source. Ni-Cu-PGE sulphide mineralization (20 m of 0.22 wt. % Cu, 0.06 wt. % Ni, 0.25 ppm Pt and 0.29 ppm Pd) is hosted by feldspathic peridotite in the lower mafic to ultramafic unit adjacent to the footwall of the Thunder intrusion. Sulphides typically occur from 1 - 5 modal %, rarely up to 30 modal %, with textures ranging from medium- to fine-grained disseminated, globular and rarely net-textured. Pyrrhotite, chalcopyrite and rare pentlandite with common secondary marcasite pyrite replacement occur along with the trace Pd, Ag, Au and rare Pt minerals, michenerite, kotulskite, merenskyite, sperrylite, hessite, electrum and argentian pentlandite. Whole-rock geochemical data display fractionated Ni-Cu-PGE patterns with depletion of iridium subgroup relative to the platinum subgroup of the platinum group elements. Sulphide δ34S values from the Thunder intrusion range from -2.0 to +3.8 ‰ and are similar to values for the metavolcanic host rock that range from -3.1 to +2.3 ‰. Two samples from the basal mineralization zone sulphides yield Δ33S values of 0.066 and 0.122 ‰ and one sample from the metavolcanic wall rock yields 0.149 ‰. The δ34S and Δ33S values for the Thunder intrusion fall within range of typical upper mantle compositions. The sulphur source appears to be of mantle origin; however, assimilation of crustal sulphur is a possibility but hard to resolve as the wall rock S isotope and S/SeTot signature is similar to that of upper mantle. REFERENCES Bidwell, G. E., and Marino, F., 2007, Thunder Project: 2007 Field program diamond drilling on the 1245457 claim: Thunder Bay Regional Geologist office, Assessment Files 2.34638. Trevisan, B.E., 2014, The petrology, mineralization and regional context of the Thunder mafic to ultramafic intrusion, Midcontinent Rift, Thunder Bay, Ontario: Unpublished M,Sc. Thesis, Thunder Bay, ON, Lakehead University, 299 p. Trevisan, B.E., Hollings, P., Ames, D.E., and Rayner, N., 2015. The petrology, mineralization, and regional context of the Thunder mafic to ultramafic intrusion, Midcontinent Rift, Thunder Bay, Ontario, In: Targeted Geoscience Initiative 4: Canadian Nickel-Copper-Platinum Group Elements-Chromium Ore Systems — Fertility, Pathfinders, New and Revised Models, (eds) D.E. Ames and M.G. Houlé; Geological Survey of Canada, Open File 7856. 86 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Bedrock and soil chemistry in paired watersheds in northeastern Minnesota WOODRUFF, Laurel G.1 and JENNINGS, Carrie E.2 1 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 2 Minnesota Department of Natural Resources, 500 Lafayette Road, St. Paul, MN 55155 Bedrock and soil have been collected and analyzed in two adjacent watersheds (Filson and Keeley) in northeastern Minnesota (Fig. 1). This sample collection effort is part of a three-year study to determine baseline geochemistry of solid media and water quality in this part of Minnesota. The Filson watershed, which includes Filson and South Filson Creeks, drains an area of about 26.3 km2 (~10.2 mi2) and discharges into the South Fork of the Kawishiwi River. The geology of the Filson watershed is complex. Bedrock includes the Archean Giants Range Batholith and the Duluth Complex. The Duluth Complex in the Filson watershed is mainly represented by the South Kawishiwi Intrusion, a thick sequence of troctolite and augite troctolite (with a heterogeneous sulfide-bearing, basal zone in contact with Archean quartz monzonite), the Anorthosite Series, and the Nickel Lake Macrodike, composed of oxide-rich gabbro and foliated troctolite (Fig.1). Two major mineral deposits in the Filson watershed are the Spruce Road deposit and the South Filson deposit (Fig.1). At the Spruce Road deposit, discontinuous sulfide-bearing heterogeneous troctolite is exposed at the surface across the northern part of the watershed. The South Filson deposit is a combination of primary disseminated Cu-Ni sulfide mineralization and secondary hydrothermal mineralization along fracture zones. Primary mineralization at South Filson occurs both at depth and sporadically in troctolite outcrop within a limited area; secondary mineralization occurs in fine-grained veinlets proximal to a northeast-southwest trending, highly altered fault zone. The adjacent Keeley Creek watershed, south and west of the Filson watershed, drains an area of about 28 km2 (~10.9 mi2) and discharges into Birch Lake. The geology of the Keeley watershed is fairly simple, with bedrock dominated by relatively homogeneous, typically unmineralized anorthositic troctolite to troctolite of the South Kawishiwi Intrusion (Fig.1). A thin zone of sulfidebearing melatroctolite has recently been mapped within the trace of the creek (D. Peterson, personal communication, 2014). In both the Filson and Keeley watersheds topography is largely controlled by the resistance to chemical weathering of underlying bedrock and subsequent removal of the saprolith by glacial erosion. Glacial cover is very thin. The landscape within the watersheds mainly consists of bedrock highlands surrounded by wetlands. In the Filson watershed soil was collected from 16 upland sites along two broad transects that cut across the general trend of the major bedrock types. In the Keeley watershed, the monotonous sea of anorthosite troctolite resulted in selection of 14 upland soil sites based on the rather problematic road access. At all soil sites, up to 3 samples were collected using hand tools, including the soil O horizon (where present – this area has abundant invasive earthworms that typically covert organic soil to mineral soil), the soil A horizon, and a deeper soil. Final sample depths were typically constrained by the stony substrate. Bedrock samples were collected from the abundant outcrop within each watershed. Bedrock sample sites were selected to be proximal to soil sample sites, if possible, or to capture the diversity of rock types. Bedrock was collected at 14 sites in the Filson watershed and 9 sites in the Keeley watershed. Soil and bedrock were analyzed for 44 major and trace elements following a near-total 4-acid digestion. 87 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Copper and Ni concentrations in mineralized bedrock in the vicinity of the exposed Spruce Road deposit are 4,600 ppm Cu and 1,130 ppm Ni. Non-mineralized South Kawishiwi Intrusion anorthosite contained Cu concentrations that range from 50 to 203 ppm and Ni concentrations that range from 119 ppm to 296 ppm; rocks of the Anorthosite Series in the Filson watershed have Cu concentrations from 67 to 235 ppm and Ni concentrations from 56 to 70 ppm. Soils collected over the exposed footprint of the Spruce Road deposit and in the down-ice direction from the Spruce Road have high Cu and Ni concentrations; a single soil sample collected in the vicinity of the South Filson deposit has relatively high Cu and Ni (Fig. 2A). The distribution of Ni and Cu in Fig. 2A is consistent with variable contributions to soil parent materials from sulfide (for example, high Cu and Ni) versus ferromagnesian silicate Figure 1. Location map showing the distribution of minerals (for example, high Ni but low Cu); soil and bedrock sample sites in the Keeley and Filson plots of other elements, such as Co, Cr, Fe, Mg, watersheds. and Mn provide similar evidence. Surface soils typically have metal values consistent with deeper soils. Because glacial transport distances are short and glacial cover thin, soil chemistry, for the most part, can be related back to bedrock contributions to soil parent material (Fig. 2B). Although these data are rather sparse, they describe the natural distribution of many elements within these two watersheds that can be attributed to geologic processes. The combination of these data with the on-going collection of water quality data in this three-year study will provide valuable information on the geochemical landscape in this region of potential mineral resource development. Figure 2. A) Box plot of Ni vs. Cu (in ppm) in soil; B) Ternary plot of Cu-Ni-Co in soil; bedrock data within solid similarly colored fields. 88 �Proceedings of the 61st ILSG Annual Meeting - Part 1 Exposure Surfaces of the Gunflint iron formation, northwestern Ontario YIP, Christopher, and FRALICK, Philip Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Exposure surfaces present in Precambrian rocks can be used as an environmental record of the conditions prior to their being covered by depositing sediments. The surfaces can show alteration of the preexisting rocks, which were exposed to the Precambrian atmosphere. The 1.88Ga Gunflint Formation in northern Ontario has two identifiable exposure surfaces found within its stratigraphy. The first one makes up the basal contact of the iron formation and is comprised of the basement rocks in places overlain by the thin Kakabeka Conglomerate all capped off by microbialite and/or grainstone of the Gunflint Formation marking the initial transgression of the sea. The second exposure surface is found approximately 45m above the basal contact. It records the regression of the ancient sea and is underlain by Gunflint grainstone and overlain by stromatolitic growth marking the shallowing of the sea. In the basal contact, the Archean rocks forming the exposed surface can show high levels of alteration. There are three outcrops present near Thunder Bay, Ontario, which contain complete sections of the basal exposure surface. The sample sites selected are two outcrops on the shoulder of Highway 11/17 and one on the shoulder of Highway 590. These three outcrops exhibit various alteration patterns within rocks near the paleosurface (Figure 1). Figure 1: Three examples of the Gunflint Iron Formation’s basal contact exposure surface showing the complete section through the exposure surface. The alteration horizons (1) are demarked. A) The formation of large core stones during alteration of the outcrop on the shoulder of Highway 11/17. B) The alteration of the KOA Hill outcrop showing the change in foliation from the vertical schistosity to the flaky altered horizon. C) The outcrop on Kakabeka Falls showing an indistinct difference between the unaltered and darker altered portion of the granodiorite unit. 89 �Proceedings of the 61st ILSG Annual Meeting - Part 1 The first outcrop on Highway 11/17 shows high levels of alteration through the formation of the large core stones as well as replacement of the original mineralogy of the granodiorite by mostly iron-rich chlorite. The outcrop at KOA hill on the shoulder of Highway 11/17 exhibits a change in the foliation in the Archean metasedimentary basement from a vertical schistosity to a flaky layer with no discernable pattern. The appearance of the Kakabeka Falls outcrop exhibits minor amounts of dark discolouration, but extensive replacement of the original mineralogy. The alteration history of these three outcrops and in particular the former, can be related to an earlier phase of surface weathering overprinted by massive diagenetic addition of Fe and Mn and extensive leaching of the initial constituents of the rocks. The exposure surface that is approximately 45m above the base of the Gunflint Formation consist of lithified grainstone blocks, some up to boulder size, that were in places rotated by current activity (Figure 2). This rubble zone and fractured basement below it contains small, wispy hematite dykes. Stromatolites developed on this lithified brecciated surface. Figure 2: Two representations of the exposure surface present approximately 45m above the base of the outcrop. A) The outcrop to the immediate north of Mink Mountain exhibits brecciation of the lithified grainstones with hematite dykes filling the fractures. B) A boulder removed from an outcrop present on Old School Road showing stromatolites forming above a brecciated grainstone boulder as well as hematite dykes filling the fractured boulder. 90 � https://digitalcollections.lakeheadu.ca/files/original/df52e8901482197925f050c5a95285c8.pdf 6927d891e486e9981bf1513e5ed0a06b PDF Text Text 61st ANNUAL MEETING InstItute on Lake superIor GeoLoGy Dryden, Ontario - May 20-24, 2015 Part 2 – Field Trip Guidebook �Sponsors The following organizations made generous contributions to the 61st Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. All of the funds contributed this year go toward travel awards for student registrants. For the past 60 years this organization has thrived as a result of the interest of individuals, corporations, universities and government agencies. The dedication to an exchange of scientific ideas and a passion for field trips has enabled the Institute to provide one of its primary objectives – to promote better understanding of the geology of the Lake Superior Region. Mary Arthur Steve Baumann Leonard Espinosa Gordon Medaris Jr. Allan MacTavish Jim Miller Paul Weiblen Canadian Institute of Mining and Metallurgy Thunder Bay Branch �61st annuaL MeetInG InstItute on Lake superIor GeoLoGy Supported by ONTARIO MINISTRY OF NORTHERN DEVELOPMENT AND MINES May 20-24, 2015 Dryden, Ontario HOSTED BY: Rob Cundari & Peter Hinz Co-Chairs Ontario Geological Survey Proceedings - Volume 61 Part 2 – Field Trip Guidebook Edited by Al MacTavish & Pete Hollings Cover photos: Top - Sakoose Mine, circa 1937 (from Humphrey and Tymura: The New Klondike to the Manitou), Middle - Central Patricia Headframe, Pickle Lake (courtesy Mark Smyk), Bottom - Pickle Crow Au mine (courtesy of Rob Cundari) �61st InstItute on Lake superIor GeoLoGy VoLuMe 61 consIsts of: part 1: proGraM and abstracts part 2: fIeLd trIp GuIdebook trIp 1: The CenTral red lake Gold BelT trIp 2: WesTern WaBiGoon suBprovinCe TranseCT, dryden To MeGGisi lake trIp 3: CanCelled trIp 4: Thunder lake (GoliaTh) projeCT trIp 5: ClassiC ouTCrops of The dryden area trIp 6: GOLD OCCURRENCES OF VAN HORNE TOWNSHIP, VAN HORNE GOLD PROPERTY FLAMBEAU EXPLOSURES trIp 7: unique MineralizinG evenT aT The pidGeon MolyBdenuM deposiT sTripped surfaCe exposure trIp 8: GeoloGy and Mineral deposiTs of The piCkle lake GreensTone BelT trIp 9: The GhosT lake BaTholiTh and relaTed peGMaTiTes trIp 10: MaTTaBi/sTurGeon lake hisToriC vMs CaMp Reference to material in Part 2 should follow the example below: Lichtblau, A., and Storey, C., 2015. Field trip 1 - The Central Red Lake Gold Belt. In; MacTavish, A. and Hollings, P., (Eds.), Institute on Lake Superior Geology Proceedings, 61st Annual Meeting, Dryden, Ontario, Part 2 - Field trip guidebook, v.61, part 2, 2-23. Published by the 61st Institute on Lake Superior Geology and distributed by the ILSG Secretary: Pete Hollings - ILSG Secretary Department of Geology Lakehead University 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Email: peter.hollings@lakeheadu.ca ILSG website: www.lakesuperiorgeology.org ISSN 1042-9964 �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table of Contents Introduction, safety considerations and acknowledgements ...............................................1 Field Trip 1 - The Central Red Lake Gold Belt ..................................................................2 Field Trip 2 - Western Wabigoon Subprovince Transect, Dryden to Meggisi Lake .........24 Field Trip 3............................................................................................................ cancelled Field Trip 4 - Thunder Lake (Goliath) Project ..................................................................40 Field Trip 5 - Classic Outcrops of the Dryden area ..........................................................46 Field Trip 6 - Gold Occurrences of Van Horne Township, Van Horne Gold property Flambeau explosures.................................................................................................51 Field Trip 7 - Unique Mineralizing Event at the Pidgeon Molybdenum Deposit Stripped Surface Exposure ......................................................................................................60 Field Trip 8 - Geology and Mineral Deposits of the Pickle Lake Greenstone Belt ..........67 Field Trip 9 - The Ghost Lake Batholith and Related Pegmatites ..................................112 Field Trip 10 - Mattabi/Sturgeon Lake Historic VMS Camp ........................................116 -i- �Proceedings of the 61st ILSG Annual Meeting - Part 2 Introduction, safety considerations and acknowledgements Rob Cundari and Peter Hinz Ontario Geological Survey, Thunder Bay, Ontario, Canada This volume is intended to serve not only as a guide for 61st ILSG field trip participants but also as a reference for those planning to revisit these areas at a later date. Consequently we have included UTM coordinates in the NAD 83 Zone 15 datum for the majority of stops, as well as instructions on how to reach them. For some of the stops on private land we have withheld the UTM coordinates to respect the privacy of the property owner. As some of the stops are on private and staked land, please be sure to obtain the land owners’ permission before entering their land. For up-to-date information on land ownership please contact the Kenora Resident Geologists’ Office at (807) 468-2819 or (for the Pickle Lake Area) the Thunder Bay Resident Geologist’s Office at (807) 4751331. Sample collection is prohibited at some stops on private land or within Provincial Parks. Many of the fieldtrips will be visiting stops along major and secondary highways or busy logging roads. Please take care when crossing or parking along these roads. For those field trips that are visiting active mine sites personal protective equipment will be required. For those field trips that are visiting past-producing mine sites please be very careful around visible open holes such as shafts, pits, or trenches and keep a wary eye out for hidden holes which may be overgrown with vegetation and therefore are very difficult to see. Please notify the field trip leaders if you have any medical conditions that may be of concern during the trip. Each trip leader is equipped with a first aid kit and satellite/ cell phone, so please notify them of any incident. We would like to thank all authors who contributed to this field guide and also all those who provided comments and assisted with the running of the field trips themselves. We appreciate the assistance and cooperation of the exploration and mining companies in providing us access and information concerning their properties. We are particularly grateful to Goldcorp Inc., Glencore plc, Cadillac Ventures Inc., PC Gold Inc. and Treasury Metals Inc. for running field trips on their properties. Figure 1. Map showing the general locations of field trips for the 2015 meeting. -1- �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 1 - The Central Red Lake Gold Belt Andreas Lichtblau and Carmen Storey Ontario Geological Survey, Red Lake, Ontario, Canada A Brief History of The Beginnings Of The Red Lake Mining Camp This excerpt from Horwood (1940) details the early history of the camp and ends with the first mine that went into production, the Howey, in 1930: A trading post, known as Red Lake House, was established by the North West Company some time prior to 1786 at Post narrows near the northeast end of the lake. Red Lake itself is shown on Arrowsmith’s map of 1801, which outlines the route travelled by Alexander Mackenzie in 1789. The post was taken over by the Hudson’s Bay Company in 1821 and has been in continuous operation ever since. In 1926 the site was changed to the west side of Howey bay opposite the Howey mine. company staked a group of claims near Slate Bay in McDonough township but abandoned the property after doing some development work and sinking a small shaft. In 1912, when the Provincial Geologist issued a report that the Keewatin rocks of the Patricia District should contain deposits of gold, some prospecting was done; but the work did not meet with success. In 1922, a party of prospectors became interested in the area. A press report of their discovery of a vein containing quartz and argentiferous galena attracted a number of other prospectors and resulted in some activity. Gus McManus, one of the prospectors, found some small gold-bearing quartz stringers near the outlet of Red Lake and staked several claims. Bruce, who was doing geological work to the south for the Ontario Department of Mines, heard of the finds and came in to make an examination. The area proved of such interest that the Department sent him back to complete his preliminary investigation. The map and report, published late in 1924, aroused considerable interest, and early in the summer of 1925 Lorne Howey and George McNeely, representing Haileybury interests, and Ray Howey and W. F. Morgan, representing McIntyre- Porcupine Mines, Limited, came in to prospect. Late in July, Lorne Howey and his partner found quartz stringers containing native gold and staked a group of claims on what is now the property of Howey Gold Mines, Limited. Ray Howey and his partner discovered gold mineralization to the southwest and; staked an adjoining group of claims, now part of the Hasaga property, for the McIntyre. News of the finds quickly reached the outside, and the Red Lake country had its first gold rush. The backers of Lorne Howey and George McNeely interested J. E. Hammell in their property, and he formed the Howey Red Lake Syndicate to develop the showing. In order to get men and supplies into the country before freeze-up seven Forestry planes, which were stationed at Sioux Lookout, were chartered. The success of this freighting venture was directly responsible for the development of commercial flying as an adjunct to mining From 1812 to 1872 no mention is made of Red Lake by the various exploring parties that penetrated the west. Its position near the heightof-land and away from the main canoe routes discouraged exploration, and it remained an outof-the-way but active outpost of the Hudson’s Bay Company. In 1872 A. R. C. Selwyn, who was making an exploratory trip down the English River from Lac Seul to Lake Winnipeg, heard of Onimini Sagaigan or Red Paint Lake from a group of Indians. They told him of the occurrence of slaty rocks, which, to him, suggested the presence of a belt of sedimentary and volcanic rocks to the north. He did not have time, however, to visit the area. In 1883, Bell made a track survey of the lake and confirmed Selwyn’s opinion of the presence of a wide belt of what he termed Huronian rocks. In 1893, Dowling examined the area and made a map, which till 1924 was the only source of geological information. Some prospecting was done between 1898 and 1924, but no mineral deposits of commercial importance were located. The first recorded discovery of gold was made in 1897 by the Northwestern Ontario Exploration Company, a prospecting venture headed by R. J. Gilbert. This -2- �Proceedings of the 61st ILSG Annual Meeting - Part 2 development in the more inaccessible parts of the north. By the spring of 1926 a fairly regular airservice had been established. In the fall of that year H. A. Oaks, one of the pioneer fliers, and James Richardson, Winnipeg financier, founded Western Canada Airways. This enterprise was known as Canadian Airways when taken over by Canadian Pacific Airlines in April, 1942. Red Lake District: 1. Uchi Sub-province rocks in the Red Lake District comprise the Red Lake and Birch-Confederation Lake greenstone belts in which the bulk of exploration and mining activity has taken place. The supracrustal rocks of the Red Lake greenstone belt can be subdivided into several assemblages with ages ranging from circa (ca.) 2990Ma to ca. 2700Ma (Table 1). Major granitoid intrusions show a range from ca. 2734Ma to 2699Ma (Table 2). Late in 1925 Mr. Hammell succeeded in interesting Dome Mines, Limited, in the Howey Syndicate. They took an option on the property and, by the summer of 1926, had completed an extensive diamond-drilling campaign. They decided, however, that the grade of the ore was too low to warrant the expenditure of more funds and in August of that year dropped their option. Many people became discouraged over the possibilities of the camp and left the area. Mr. Hammell, however, still had faith in the Howey property and continued work. With the financial assistance of W. S. Cherry a mill was built and the mine came into production on April 2, 1930. The undertaking proved to be a profitable venture and continued operations till November 3, 1941. 2. English River Sub-province rocks, south of the Uchi Sub-province, are predominantly metasedimentary and host minor intrusive rocks similar to those in the Quetico Sub-province. 3. To the north, the Berens River Sub-province formed the core of a microcontinent. This area is underlain by ca. 2750-2690Ma felsic plutonic rocks interpreted as a magmatic arc formed at an Andean-style margin that culminated in the Kenoran Orogeny. These plutonic rocks intruded an older substratum (North Caribou terrane) on which Mesoarchean volcanic rocks of the Red Lake belt are also interpreted to have formed. We will drive by an open stope at the Howey mine site during the tour. Regional Geology The Red Lake District (Fig. 1) is underlain by Archean rocks of the Superior Province of the Canadian Shield. Rocks of four sub-provinces are found in the Figure 1. Western Uchi Subprovince (modified from Percival et al., 2000). 4. The Sachigo Sub-province comprises crustal blocks ranging from Paleoarchean (>3.4Ga) to Neoarchean (ca. 2.7Ga) in age. Geology of the Red Lake Greenstone Belt (adapted from Sanborn-Barrie et al. 2001) The geology of the Red Lake greenstone belt (Fig. 2) is dominated by the (ca. 2990Ma) mafic-ultramafic Balmer assemblage, an oceanic plain sequence; minor calc-alkalic volcanic rocks of arc-like affinity terminate the assemblage. The majority of lode gold deposits in the camp are hosted by the basal mafic-ultramafic sequence. A later diverse lithologic association, the Ball assemblage, appears to represent a shallow marine, volcanic edifice built upon the Balmer substrate. Widespread circa (ca.) 2894Ma calc-alkaline volcanism is represented in Red Lake by the Bruce Channel assemblage. Overlying this is the ca. 2850Ma Trout Bay assemblage which includes substantial basaltic and gabbroic rocks in western Red Lake which are prospective for PGE mineralization, and which includes minor intermediate pyroclastic rocks throughout central Red Lake. The Trout Bay assemblage may correlate with Woman assemblage rocks of the Confederation Lake belt. -3- A regional angular unconformity is interpreted �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 2. General geology of the Red Lake greenstone belt (after Parker, 2000). Tour stops as indicated. to separate the Mesoarchean assemblages from the Neoarchean Confederation assemblages. Volcanogenic massive sulphide mineralization is associated with the younger sequence. A significant number of felsic units are classed as FII- and FIII-type rhyolites, considered highly prospective for large (Kidd Creek/Noranda type) massive sulphide deposits (Parker, 1999). A newly recognized component of the Neoarchean supracrustal package is the Huston sedimentary assemblage that includes surface exposures of polymictic cobble- to pebble-conglomerate and argillite; clasts include jasperoidal chert iron formation, massive sulfide pebbles, mafic flow rocks as well as well-bedded, graded turbiditic wacke and argillite. The Huston assemblage conformably to unconformably overlie the McNeely sequence (Confederation assemblage) and underlies the Graves assemblage (Sanborn-Barrie et al., 2004). The U-Pb age of detrital zircons give single age peaks of 2743 and 2746Ma at the cemetery and Madsen sites, respectively (Skulski et al., 2001), indicating single source derivation from erosion of pre-existing Confederation age rocks, and deposition after ca. 2743Ma. In the Campbell-Red Lake Deposit area the conglomerate defines the interface between the Balmer or Bruce Channel assemblages and the Confederation assemblage. Exposures on 16 Level of the Red Lake Mine reveal that (Dubé et al., 2004): “... the [Huston] conglomerate is a polymictic proximal conglomerate or breccia (debris flow) dominated by subangular to subrounded laminated cherty clasts with local jasper-rich and green mica fragments; it is similar to the Temiskaming conglomerate… it also contains clasts of andalusite-rich altered basalt and a few local clasts of layered (possibly sheeted) carbonate veins with small crustiform banding and cockade texture.” Detrital zircon analyses from the conglomerate on 16 Level of the Red Lake Mine define a single population with a mean age of 2747±4Ma (Dubé et al., 2004). This implies that the former two assemblages were exposed at surface by 2747Ma. The presence of local andalusite-rich clasts and clasts of carbonate vein material in an unaltered matrix demonstrates that there was a period of at least some -4- �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 1. Summary of supracrustal lithologies and radiometric ages in the Red Lake greenstone belt (modified from Parker 2000; with new ages from Sanborn-Barrie et al. 2001 and Skulski et al. 2001; final error estimates are not cited for the new unpublished ages of T. Skulski). Supracrustal assemblage English River? U-Pb Age (Ma) <2700±6 ConfederationGraves (north Red Lake) 2733±1.5 Huston (Cemetery) ≤2743 Huston (16 Level, Lake Mine) Rock types and descriptions References Polymictic pebble conglomerate. Thought to correlate with the Austin tuff, host to the Madsen gold deposit. Calc-alkaline andesitic to dacitic pyroclastic rocks Sanborn-Barrie et al. 2001 Well-bedded argillite and polymictic conglomerate Skulski et al. 2001 turbiditic wacke; Corfu and Andrews 1987 Dubé et al. 2004 Red ≤2747±4 Polymictic conglomerate; laminated cherty and jasper clasts; green mica fragments; aluminousaltered basalt and iron-carbonate vein clasts ConfederationHeyson (southeast Red Lake) 2739±3 Basal sequence is commonly tholeiitic to calcalkaline with lobe-hyaloclastite rhyolite flows; intermediate pyroclastic rocks; basalt; and feldspar-phyric andesite. Calc-alkaline rocks are more abundant at higher stratigraphic levels. Corfu and Wallace 1986 2742 2748+10/-5 Dominated by calc-alkaline, intermediate lapillituff breccia and lapilli tuff Sanborn-Barrie et al. 2001 2853 Lower tholeiitic basalt sequence with associated gabbroic rocks overlain by fine-grained clastic metasedimentary rocks (wacke, argillite) interlayered with subordinate intermediate pyroclastic rocks and chert-magnetite iron formation. Overlain by tholeiitic, pillowed basalts. Strongly calc-alkaline intermediate pyroclastic rocks overlain by pebble conglomerate, thinly bedded wacke and capped by chert-magnetite iron formation Interlayered, feldspathic wacke, lithic wacke and argillite; lenses of pebble and cobble conglomerates and quartz-rich pebble conglomerate and quartz arenite. Typically calc-alkaline intermediate pyroclastic rocks and rhyolite flows; komatiitic to tholeiitic basalt; overlain by chert-magnetite iron formation and dolomitic marble which contains stromatolites. Tholeiitic basalt, basaltic komatiite and komatiite interlayered with subordinate chert-magnetite iron formation; minor clastic metasedimentary rocks; minor intermediate to felsic pyroclastic rocks; and rhyolite. Sanborn-Barrie et al. 2001 ConfederationMcNeely (central and SE Red Lake) Trout Bay Bruce Channel Slate Bay Ball Balmer 2894±1.5; 2894±2 ≤2916 2940±2; 2925±3 2992+20/-9; 2989±3; 2964+5/-1 -5- Corfu and Wallace 1986; Corfu and Andrews 1987 Corfu et al. 1998 Corfu and Wallace 1986 Corfu and Andrews 1987 �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 2. Summary of lithologies and radiometric ages for major granitoid intrusions in the Red Lake greenstone belt (modified from Parker, 2000; new ages cited in Sanborn-Barrie et al., 2001 and elsewhere do not have final error estimates assigned, as this U-Pb data is not yet published). Granitoid intrusion Cat Island pluton Walsh Lake pluton Killala-Baird batholith Hammel batholith Dome stock U-Pb Age (Ma) 2699 2699 2704±1.5 Lake 2717±2 2718±1 McKenzie stock 2720±2 Red Crest stock 2729±1.5 Little Vermilion batholith Douglas Lake pluton Rock types and descriptions Potassium feldspar granodiorite Potassium feldsparand quartz-phyric monzogranite; xenolith-rich, diorite or granodiorite; possible oxidized phase at Ranger Lake with broad magnetic anomaly Potassium feldsparand quartz-phyric monzogranite; xenolith-rich, diorite or granodiorite, diorite or granodiorite; oxidized, magnetite-bearing marginal phase. Potassium feldspar and quartz porphyritic monzogranite; associated anorthositic intrusion. Granodiorite and augite porphyritic diorite/gabbro. Augite porphyritic diorite-gabbro; some ultramafic rocks; granodiorite Augite porphyritic diorite-gabbro 2731±3 Hornblende tonalite-granodiorite 2734±2 Biotite tonalite aluminous and carbonate alteration prior to deposition of the Huston conglomerate (Dubé et al. 2004). The position of colloform-crustiform iron-carbonate±quartz veins in the Campbell-Red Lake Deposit underneath the interpreted subaerial unconformity also leads Dubé et al. (2004) to interpret the veins as near-surface, epithermal-epizonal products, part of a protracted hydrothermal alteration event spanning pre- to postHuston assemblage time, a period of more than 35m.y. Recent age dating (Skulski et al. 2001) has also yielded multiple ages of detrital zircons from a fragmental unit thought to correlate with the Austin “tuff” ore zone at the former Madsen mine. Most of the Meso- and Neoarchean assemblages exposed in Red Lake are represented in this unit. Maximum age of deposition is consequently ≤2700±6Ma. Deformation (adapted from Sanborn-Barrie et al., 2001) The Red Lake greenstone belt has undergone at least three phases of deformation: 1. D0, a non-penetrative, early (pre-2748Ma) event involving overturning of the Balmer assemblage; References Sanborn-Barrie et al. 2001 Noble 1989 Corfu and Andrews 1987 McMaster 1987 Corfu 1987 Corfu 1987 Corfu 1987 Corfu 1987 Corfu 1998 and Andrews and Andrews and Andrews and Andrews and Stone 2. D1, (bracketed between 2733-2742Ma) resulted in a north-trending foliation that is axial planar to F1 folds and involved east-west shortening; and 3. D2, (ca. 2720-2700Ma) resulted in a dominantly east- to northeast-striking foliation that refolds F1 folds. A local ‘deflection’ of S2 around the McKenzie stock created an east-southeast-striking corridor of heterogeneous strain forming the “Mine Trend”, from Cochenour through the Balmertown area, hosting the major gold deposits of the camp. Hydrothermal Alteration Parker, 2000) (adapted from The Red Lake greenstone belt has been affected by a large-scale (10’s of kilometres) hydrothermal alteration system, resulting in approximately contemporaneous a) strong to intense, distal calcite carbonatization that affects rocks of all ages, and b) less extensive (kilometre), proximal, strong to intense ferroandolomite and potassic alteration, found in almost all areas hosting gold mineralization. Carbonate alteration affects both the Dome (2718±1Ma) and McKenzie (2720±2Ma) stocks and is overprinted by calc-silicate, skarn-like alteration formed during the intrusion of the -6- �Proceedings of the 61st ILSG Annual Meeting - Part 2 Killala–Baird batholith (2704±1.5Ma) and the Walsh Lake pluton (2699Ma). The significant carbonate alteration event is therefore bracketed between 2718 and 2704Ma, during D2. The main macroscopic features of carbonate alteration are pervasive replacement of rock matrix, open-space filling/replacement of primary porosity (vesicles, pillow selvages, hyaloclastite matrix), filling of extension veins with massive, colloform, crustiform and cockade breccia textures, networks of variably oriented veins, and “jigsaw puzzle” breccia veins. Multiple stages of carbonate alteration and veining have been recognized, indicating continuous carbonatization during D2 deformation. Potassic metasomatism takes the form of sericite/ muscovite alteration in greenschist-facies rocks; in ferroan-dolomite altered ultramafic rocks fuchsite occurs instead of sericite. Potassic alteration in amphibolite-facies mafic and ultramafic rocks takes the form of pervasive biotite±muscovite. Centimetreto m-wide, strong to intense, biotite±calcite±ferroandolomite±disseminated pyrite alteration halos often enclose ferroan-dolomite veins in amphibolite-facies mafic rocks. Biotite altered zones in amphibolite-facies rocks are characterized by a diverse assemblage of aluminosilicate minerals such as andalusite, staurolite and cordierite, with garnet, chloritoid, cummingtonite, and anthophylite. Barren, pervasive silicification within proximal alteration zones may be due to release and remobilization of silica during periods of pervasive carbonatization. The majority of gold deposits in the Red Lake belt are quartz and arsenopyrite rich selective replacement zones of colloform-crustiform ferroan-dolomite veins and breccia. Geology of the Campbell-Red Lake Gold Deposit (adapted from Dubé et al., 2002) Gold has been continuously produced from the Red Lake (formerly known as the Campbell-Dickenson) deposit since 1948. Current production levels and reserves are given in Table 3. Historic production figures for the Red Lake greenstone belt are shown in Table 4. Alteration facies in the High Grade Zone at Goldcorp Inc.’s Red Lake Mine have been described by Dubé et al. (2002) as: 1. An outer, metre-wide, garnet-chlorite-magnetite alteration with chlorite-amphibole-andalusite and locally associated centimetre- to metre-wide ‘bleached zone’ containing andalusite-muscovitequartz-ilmenite; 2. A proximal, centimetre- to metre-wide, massive to laminated, reddish-brown, biotite-carbonate alteration with disseminated pyrite (3-5%) and carbonate veinlets in well foliated basalt; and 3. A gold-rich, strongly foliated, silicified zone with abundant fine-grained arsenopyrite, sericite, and rutile, and lesser amounts of pyrite, pyrrhotite, magnetite, and stibnite (≤15%). This third alteration facies is adjacent to the silicified auriferous carbonate veins and replaces the biotite-carbonaterich alteration. The chronology of gold-rich replacement textures suggests a syn-D2 mineralizing event, dominated by silicification of carbonate veins, contemporaneous with boudinage of the veins. The silicified carbonate veins are hosted mainly by basalt; areas of high-grade gold mineralization are controlled by F2 fold hinges deforming the basalt-ultramafic contact. Multiple periods of silicification and gold deposition overprint and replace the carbonatization in these lower pressure hinge zones. The extremely high grade ore (>2.0oz/t Au) currently mined at Goldcorp Inc.’s Red Lake Mine, is possibly due to a combination of factors, including the presence of a low-permeability ultramafic cap, allowing the build-up of very high fluid pressure in the footwall basalt; the high iron content of the tholeiitic basalt, creating a chemical, as well as structural, trap for the Table 3. Current gold production and reserves, Goldcorp Inc., Red Lake Gold Mines Production in 2013 Mine Goldcorp Inc. Red Lake Gold Mines (1) Tonnage @ Grade 786 900 tonnes @ 20.33 g/t Au Total Commodity 493 000oz Au Production in 2014 Tonnage @ Grade 684 100 tonnes @ 19.47 g/t Au (1) Goldcorp Inc., news release MD&A, February 19, 2015. -7- Total Commodity 414 400 ounces Au Reserves (Proven and Probable) at end of 2014 Contained Ounces Gold 2 060 000 �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 4. Historical gold production, Red Lake greenstone belt GOLD PRODUCTION IN THE RED LAKE GREENSTONE BELT to December 31, 2014 GOLD PRODUCED RED LAKE BELT PRODUCER RED LAKE GOLD MINES CAMPBELL RED LAKE GOLDCORP (DICKENSON) MADSEN COCHENOURWILLANS MCKENZIE RED LAKE HOWEY HASAGA STARRATT OLSEN H.G. YOUNG MCMARMAC GOLD EAGLE RED LAKE GOLD SHORE BUFFALO ABINO LAKE ROWAN RED SUMMIT MOUNT JAMIE TOTAL YEARS OF PRODUCTION 2006 to present (1) ORE MILLED (SHORT TONS) TROY OUNCES OUNCES PER TON GRAMS PER METRIC TONNE 6,906,322 5,285,590 0.694 23.80 19,944,241 11,216,443 0.510 17.49 1948 - 2005 9,606,894 5,962,948 0.621 (4) 21.28 1938 - 1976, 1997 – (5) 1999 1939 - 1971 8,678,143 2,452,388 0.283 (6) 9.69 2,311,165 1,244,279 0.538 (7) 18.46 1935 - 1966 2,353,833 651,156 0.277 4,630,779 1,515,282 907,813 288,179 152,978 180,095 86,333 421,592 218,213 163,990 55,244 45,246 40,204 21,100 0.091 0.144 0.181 0.192 0.296 0.223 0.244 31,986 2,733 13,023 591 552 57,061,260 1,656 1,397 1,298 277 265 27,369,086 0.052 0.511 0.100 0.469 0.480 0.335 1949 – 2005 (2) 1930 - 1941, 1957 1938 - 1952 1948 - 1956 1960 - 1963 1940 - 1948 1937 - 1941 1936 - 1938 1981 - 1982 1985 - 1986 1986 - 1988 1935 - 1936 1976 (8) 9.48 (9) 1. Includes total production from the Red Lake complex from January 1, 2006, and production from the Campbell complex subsequent to May 12, 2006, the date of acquisition. 2. Includes production figures under Placer Dome (CLA) Ltd., to May 12, 2006. 3. For 1997, 1998 and 1999, no production due to strike by unionized employees. 4. From 1970, includes production from Robin Red Lake. 5. Includes clean up of ore and materials from the mine site. 6. Historic grade, actual grade for 1999 was 0.14 ounce per ton gold. 7. Includes production from Annco and Wilmar properties. 8. Continuous production 1930 to 1941; includes 268 ounces recovered from clean up in 1957. 9. The ore mined at Howey, before sorting totalled 5 158 376 tons. The average production from run-ofmine ore was therefore 0.0817 ounce per ton gold. -8- 3.12 4.94 6.19 6.57 10.14 7.65 8.38 1.78 17.53 3.42 16.07 13.30 17.4 �Proceedings of the 61st ILSG Annual Meeting - Part 2 auriferous fluids; multiple D2 strain events; repeated episodes of gold deposition and remobilization into a low pressure F2 fold hinge hosting the High Grade Zone. garnet-biotite-staurolite-amphibole; metre-wide stockwork amphibole veins and veinlets alternate with the pervasive alteration. Timing of this alteration is pre- to syn-D1, but its relationship to gold mineralization is not yet known; indeed, it could be classified as the amphibolite-facies equivalent of volcanogenic massive sulphide (VMS) type alteration, related to a Confederation age synvolcanic hydrothermal alteration system; Field Trip Stops The Field Trip starts in the west, at the Suffel Lake (Flat Lake) Road turn-off from Highway.618 and continues east and north, to end at the Redcon Prospect, approximately 17.2km north on Nungesser Road. Stop 1: Suffel Lake Road and Highway 618. Contact between Confederation and Balmer Assemblages (Fig. 3, Table 5) UTM Coordinates: NAD83; 15U 0434844E / 5645498N Exposures on the south side of the highway are part of the lowermost units of the Neoarchean Confederation assemblage, the dominantly calc-alkaline McNeely sequence. The outcrops here are amphibolite-facies, quartz-feldspar-porphyritic lapilli-crystal tuff, with thin, dark grey, collapsed pumice fragments; occasional lapilli sized lithic clasts are also observed. Strike of the rocks is generally northeast, facing and dipping steeply southeast. A sample from this unit, 800m northeast of the intersection, gave an age of 2744±1Ma (Corfu and Andrews, 1987). 2. An inner zone comprising a banded-laminated texture, characterized by bands of actinolitehornblende-microcline-calcite-tourmaline, alternating with biotite-rich bands. The amphibole is commonly randomly oriented. Diopside locally forms disseminated crystals up to 7-8cm long, or veinlets. Ore zones occur within the inner alteration zone, and comprise finely layered, sulphide-rich lenses up to a few metres wide. Sulphides (8-10%) comprise pyrrhotite, pyrite and/or arsenopyrite with trace chalcopyrite, and are found as disseminations or veinlets parallel to lamination/foliation. Gold occurs in the native state as inclusions in silicate minerals and locally as coatings on sulphide minerals. The highest grade is found in areas of most intense alteration, The north side of the road exposes highly altered tholeiitic, mafic volcaniclastic rocks of the Balmer assemblage. Abundant garnet and biotite rims clasts; minor andalusite is present. This outcrop, barren at this locality, forms part of the Austin “tuff” ore zone, described further below. Stop 2: Power Line Outcrops, Madsen Deposit UTM Coordinates: NAD83; 15U 0435235E / 5646040N Time limitations of the tour do not permit a complete visit of the Madsen deposit; a brief description of the deposit follows: Geology of the Madsen Deposit (adapted from Dubé et al., 2000) Madsen is a stratabound, replacement-style, disseminated gold deposit (Fig. 3), exhibiting two alteration facies, the mineralogy of which is now represented by two amphibolite-facies zones: 1. A pervasive aluminous, metre- to tens-of-metreswide, low-strain, outer zone, containing andalusite- Figure 3. Geology of the Madsen mine area (modified from Dubé et al., 2000) -9- �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 5. Major and trace element analyses Stop 1 lithologies (major oxides in %, trace elements in ppm). Sample Number UTM-Easting UTM-Northing Rock Type Sample No. SiO2 TiO2 Al2O3 Fe2O3 K2O MgO MnO CaO Na2O P2O5 LOI Total 2009-AL-02 434891 5645522 QFP 2009AL-02 76.19 0.1 12.06 1.58 4.86 0.36 0.03 0.94 2.53 0.01 0.82 99.48 2009-AL-03 434849 5645513 Basalt 2009AL-03 48.61 1.55 14.59 19.21 0.91 4.37 0.51 9.57 0.41 0.21 0.73 100.67 Au (opt) Ag (opt) <0.01 <0.1 0.01 <0.1 Rb Ba Sr Sc La Ce Nd Sm Eu Gd Tb 88.23 1044.1 76 3.5 47.57 86.49 27.27 3.78 0.47 2.46 0.303 19.66 84.2 44 33.9 7.84 18.69 12.85 3.85 1.26 5.23 0.936 represented by a quartz-biotite-muscovite-microcline assemblage in mm-cm bands or layers. Crenulation of alteration bands, sulphides and calcite veinlets by S2 as well as the large-scale deformation and folding of Austin ore lenses by F2 folds are consistent with pre- to early-D2 timing of gold mineralization. A minimum age on the deposit is 2699±4Ma (Corfu and Andrews, 1987), the age of a cross-cutting post-ore granodiorite dyke. Proximal alteration and style of mineralization may indicate the Madsen Deposit to be related to higher temperature (400º-600ºC) gold deposits and gold-skarn deposits hosted by mafic volcanic rocks (Parker, 2000). Sample Number Yb Lu Y Zr Th U Hf Nb Ta Cs Dy Er Ho Pr Tm Be Cd Ga Li Mo Sb Sn Tl V W Zn Pb Cu Cr Ni Co Bi Ti 2009AL-02 0.838 0.13 8.53 108 15.16 2.98 3.21 4.33 0.4 2.5 1.68 0.86 0.31 8.54 0.125 0.73 0.11 13.47 12.9 3 1.47 1.03 0.46 <10 1.32 30.56 12.6 22 36 5 1.7 0.094 605.94 2009AL-03 4.389 0.68 41.59 67 0.99 0.24 1.9 3.82 0.2 1.83 6.55 4.36 1.43 2.66 0.665 0.57 0.14 19.02 24.1 0.95 1.11 0.79 0.1 358.45 0.75 146.21 1.8 86 109 99 47.8 0.041 8772.63 South Austin Zone – Powerline Section The integrated mapping, geochronological, and lithogeochemical projects completed during the Federal–Provincial NATMAP program in the Red Lake greenstone belt from 1999 to 2004 significantly advanced the understanding of the geological history of the belt. Key outcrops elucidating the 300Ma history of the belt include the surface expression of the main ore zone (“South Austin Zone”) of Claude Resources Inc.’s past-producing Madsen Mine (produced 2.5 million ounces Au between 1938 and 1999; see Table 4), which is interpreted to occupy the position of the unconformity between the Mesoarchean Balmer Assemblage and Neoarchean Confederation - 10 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 4. Location of the South Austin Zone (adapted from Dubé et al., 2000). Assemblage (Sanborn-Barrie et al., 2001). Geological field tours of the Red Lake mining camp have been given since the start of the Red Lake office of the Resident Geologist Program in 1968. The South Austin Zone outcrop has become an important field trip stop since NATMAP re-interpretation of local geology, illustrating the complex interplay of lithology, alteration, deformation, and economic gold mineralization. The series of poorly exposed hillside outcrops are located approximately 600m southwest of the Madsen shaft (Fig. 4). They were mechanically stripped in the fall of 2009 and power-washed and channel-sampled by S. McDonald in the summer of 2010. The exposure is now complete from the Austin Zone footwall, through the Confederation Assemblage quartz-feldspar rhyolite, to the overlying Huston conglomerate (Fig. 5). Following is a preliminary description of the lithology and geochemistry of the exposure. The base of the exposure is a well banded/layered, locally contorted, example of Austin “tuff”, from which the bulk of the 2.5 million ounces Au of the Madsen deposit were mined. At this locality, the Austin Zone is a strongly altered (biotite, amphibole, garnet) mafic volcaniclastic/epiclastic rock, with up to 20% wacke/tuff (?) clasts, occupying the position of the unconformity between Balmer and Confederation assemblages. Major oxide chemistry (Table 8), uncorrected for alteration, indicates that 11 of 12 samples collected from this unit are tholeiitic (Irvine and Baragar, 1971, Fig. 2). A minimum age on the deposit is 2699±4Ma (Corfu and Andrews, 1987), the age of a crosscutting post-ore granodiorite dike (not exposed at this locality). Weakly to moderately foliated quartz-feldspar rhyolite porphyry (“QFP”) forms the structural and stratigraphic hanging wall of the deposit and marks the beginning of Confederation Assemblage time (McNeely sequence of Sanborn-Barrie et al., 2001). Contact with the underlying Austin Zone is sharp, with the Austin Zone characterized by an 80cm-wide transition zone of contorted and brecciated rock. Whole rock analyses of two samples of porphyry plot in the calc-alkaline rhyolite field of Jensen (1976; Fig. 1). Overlying the quartz-feldspar rhyolite is a conglomerate unit, part of the Huston Assemblage, which yielded a single peak in detrital U-Pb zircon ages of ≤2746Ma (Sanborn-Barrie et al., 2001). Fragments are generally intermediate to mafic, with a mafic biotite-garnet-andalusite altered matrix. Major - 11 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 5. Detail of stripping and sampling, South Austin Zone. Sample series 2010SM-01 through 2010SM-27. oxide chemistry (Table 8), uncorrected for alteration, indicates that the 4 samples collected from this unit are of tholeiitic affinity (Irvine and Baragar, 1971, Fig. 2). Tables 6 and 7 show the trace element analyses for samples collected during this study. Heyson felsic volcanic rocks, including spherulitic and lobe-hyaloclastite flows are seen in outcrops on both sides of Highway 618 on the way to Stop 3. Samples from all three rock units (Tables 6, 7, 8, and 9) were submitted to the Pacific Centre for Isotopic and Geochemical Research (“PCIGR”), Department of Earth and Ocean Sciences, University of British Columbia for geochronological analysis. The almost complete sequence across the interpreted unconformity is exhibited in these outcrops. A cross-section of age data was deemed important to understanding the relationships here, and in the >100km of unconformity interpreted throughout the Red Lake belt. TIMS dating employing the single grain chemical abrasion (CA-TIMS) technique (Plot 1). Five single grains analysed give results that are concordant and overlapping at 2741Ma. The best age estimate for the rock, 2741.0±0.8Ma, is based on the weighted average of all five 207Pb/206Pb dates. Sample 2010SM-031: The Huston Conglomerate was dated by the laser ablation (LA-U-Pb) technique, with 59 grains analysed (Plots 2 and 3). All analysed grains are less then ±5% discordant and 207Pb/206Pb A 20-litre sample was collected from each unit; sample preparation by PCIGR can be found at http://pcigr.eos.ubc.ca/services/samplepreparation. php#Geochronology. Description of results from PCIGR includes: Sample 2010SM-029: South Austin Zone-No zircons were found Sample 2010SM-030: QFP lapilli tuff; U-Pb Plot 1: Uranium-lead concordia diagram of QFP lapilli tuff. - 12 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 6. Trace element analyses (ppm) of South Austin Tuff samples (sample number prefix is 2010SM, i.e., 2010SM-01). Sample No. Rock Type -01 Austin -02 Austin -03 -04 Austin -05 Austin -06 Austin Austin -07 -08 Austin Austin -09 -10 Austin -11 Austin -12 Austin Austin Rb 100.79 64.46 73.7 108.19 115.67 104.61 70.52 122.2 109.1 96.68 82.53 142.82 Ba 259.6 199.3 357.5 282.5 264.8 275.1 226.6 249.2 245.9 125.7 166.2 190.8 Sr 31 80 126 98 77 65 41 44 41 6 9 9 Sc 43.3 34.7 42.2 40.1 53.7 51.8 42.9 50.9 53.2 47.9 32.3 31.1 La 5.64 5.4 7.39 6.8 7.66 7.31 6.26 6.58 7.24 11.35 7.01 9.96 Ce 14 13.14 18 16.56 18.96 17.77 16.23 17.16 17.23 24.24 15.08 20.37 Nd 9.5 8.75 12.05 11.61 12.64 11.81 11.37 11.16 11.77 15.18 9.49 12.27 Sm 2.9 2.82 3.8 3.72 3.98 3.58 3.72 3.44 3.55 4.14 2.57 3.11 Eu 1 1.19 1.29 1.25 1.4 1.33 1.45 1.24 1.18 1.37 0.81 0.79 Gd 3.71 3.73 5.03 4.96 5 4.67 4.99 4.1 4.46 4.95 2.89 3.7 Tb 0.686 0.655 0.889 0.896 0.895 0.821 0.925 0.738 0.839 0.863 0.482 0.629 Yb 3.412 2.886 4.092 4.013 3.822 3.41 3.959 3.702 4.158 4.356 2.054 2.415 Lu 0.52 0.43 0.61 0.59 0.57 0.5 0.6 0.57 0.63 0.69 0.31 0.36 Y 28.23 26.63 36.82 36.72 34.9 31.92 38.04 29.44 34.11 34.81 18.05 24.04 Zr 103 54 38 47 99 105 74 106 99 77 67 81 Th 0.99 0.71 0.99 0.96 1.01 1 0.94 1.03 0.95 1.01 0.7 0.97 U 0.25 0.22 0.26 0.25 0.22 0.24 0.24 0.28 0.23 0.25 0.25 0.37 Hf 2.97 1.61 1.3 1.41 2.8 3.05 2.13 3 2.78 2.19 1.99 2.36 Nb 4.13 2.91 4.16 4 4.25 4.1 3.78 4.24 3.92 3.88 2.86 4.28 Ta 0.3 0.2 0.3 0.2 0.3 0.3 0.2 0.3 0.3 0.3 0.2 0.3 Cs 6.25 2.95 2.16 4.1 6.12 5.41 2.99 6.22 5.39 6.27 3.7 6.08 Dy 4.74 4.53 6.15 6.14 6.04 5.58 6.34 5.17 5.87 5.94 3.14 4.15 Er 3.18 2.97 4.15 4.04 3.85 3.47 4.15 3.57 3.99 4.08 2.04 2.6 Ho 1.04 0.98 1.34 1.33 1.28 1.18 1.37 1.14 1.29 1.32 0.67 0.88 Pr 1.98 1.91 2.6 2.4 2.68 2.54 2.39 2.39 2.51 3.39 2.11 2.79 Tm 0.489 0.438 0.61 0.598 0.573 0.513 0.596 0.543 0.604 0.63 0.3 0.371 Be 0.69 3.06 2.23 1.94 1.24 0.87 0.62 0.52 0.49 0.23 0.28 0.75 Cd 0.11 0.1 0.13 0.07 0.07 0.08 0.21 0.12 0.16 0.28 0.07 1.86 Ga 20.34 18.27 20.93 20.51 22.75 22.26 19.38 22.42 20.46 18.14 14.51 18.55 Li 46 20.7 20.6 29.2 40.2 37.4 19.7 38.2 30 26.2 21.1 26.2 Mo 1.25 1.47 1.07 1.21 0.63 0.48 0.51 0.49 0.47 0.66 0.77 1.22 Sb 0.76 2.44 3.46 2.03 0.82 0.8 1.95 1.09 0.97 0.71 0.78 1.14 Sn 0.87 0.83 0.73 0.87 0.75 0.91 0.84 0.81 0.93 0.73 0.67 0.58 Tl 0.52 0.33 0.33 0.51 0.55 0.5 0.37 0.59 0.54 0.52 0.41 0.86 V 403.74 251.4 410.39 398.18 453.66 429 399.06 431.52 420.2 433.06 227.59 267.23 W 9.61 67.03 63.65 45.25 18.88 9.16 19.03 3.21 2.3 3.4 9.59 11.86 Zn 105.7 66.58 81.62 77.45 91.71 69.39 173.52 121.95 108.11 107.27 34.15 388.81 Pb 3.5 3.5 3.7 2.8 2.9 3.1 3.4 2.2 2.7 6.2 2 21.4 Cu 45 74 64 75 114 86 135 107 69 86 37 182 Cr 113 90 111 116 118 116 104 117 111 106 154 162 Ni 87 67 89 100 94 94 70 84 78 75 118 127 Co 44.7 34.9 50.9 52.7 52.7 57.1 43.7 51.9 44.5 39.5 33.1 38.8 Bi 0.031 0.065 0.105 0.086 0.032 0.033 0.083 0.051 0.077 0.066 0.036 0.128 Ti 9630.2 6115.15 9527.75 9381.44 10239.1 10082.4 8816.34 10119.2 9356.36 9089.29 5094.75 6837.98 Analyses by Geoscience Laboratories, Ministry of Northern Development, Mines and Forestry, Sudbury, Ontario. - 13 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 7. Trace element analyses (ppm) of QFP and Huston conglomerate samples (sample number prefix is 2010SM, i.e., 2010SM-13). Analyses for samples 2010SM-25 and -26 are pending. Sample No. Rock Type -13 QFP -14 QFP -15 Huston -16 Huston -17 Huston -18 Huston -19 Huston -20 Huston -21 Huston -22 Huston -23 Huston -24 Huston Rb 85.22 74.47 107.78 68.11 64.93 63.87 77.81 77.86 61.36 63.04 53.71 66.62 Ba 722 569.1 495.9 283.4 271.4 584.6 334.3 293.1 250.6 276.9 273.8 261.1 Sr 73 54 217 40 38 77 80 66 56 44 40 41 Sc 5 4 28.5 41.4 37.7 21 28.2 35.7 39.3 45.1 39.9 42.9 La 27.9 20.46 23.3 6.55 7.12 20.45 14.54 7.78 6.57 9.42 7.56 7.78 Ce 52.26 38.29 45.04 14.46 16.56 39.34 27.97 15.81 14.82 20.41 16.52 17.22 Nd 15.14 11.48 17.35 8.39 10.47 15.7 12.19 7.91 8.24 10.52 9.11 9.99 Sm 2.26 1.63 3.29 2.28 3.17 3.09 2.72 2 2.25 2.68 2.42 2.57 Eu 0.58 0.44 1.09 0.82 0.97 0.91 0.76 0.95 0.86 0.87 0.81 0.91 Gd 1.75 1.25 3.04 2.84 4.61 2.69 2.72 2.38 2.77 3.08 2.95 3.12 Tb 0.277 0.199 0.494 0.494 0.818 0.429 0.45 0.426 0.495 0.554 0.505 0.552 Yb 1.037 1.012 1.888 2.086 3.384 1.512 1.624 1.902 2.127 2.398 2.163 2.297 Lu 0.16 0.15 0.29 0.32 0.51 0.23 0.25 0.29 0.33 0.36 0.33 0.35 Y 10.83 8.66 18.68 18.81 35.54 14.91 16.06 16.41 19.34 21.67 19.85 21.14 Zr 125 122 109 72 73 132 87 86 81 98 75 74 Th 12.36 12.32 6.33 1.06 1 8.18 3.38 2.67 1.46 2.37 1.26 1 U 2.36 2.2 1.52 0.38 0.32 2.98 1.06 0.79 0.51 0.73 0.34 0.33 Hf 3.72 3.64 3.19 2.13 2.11 3.74 2.42 2.48 2.39 2.71 2.12 2.17 Nb 4.69 4.46 4.26 3.39 3.25 5.71 4.1 3.9 3.72 4.59 3.48 3.63 Ta 0.4 0.4 0.4 0.2 0.2 0.6 0.3 0.3 0.3 0.3 0.2 0.2 Cs 1.09 1.28 4.78 2.33 2.25 1.56 2.39 1.85 1.82 2.1 1.8 2.38 Dy 1.76 1.35 3.2 3.31 5.53 2.7 2.81 2.89 3.32 3.7 3.43 3.69 Er 1.08 0.93 1.97 2.1 3.66 1.57 1.69 1.88 2.13 2.4 2.21 2.33 Ho 0.36 0.3 0.67 0.7 1.22 0.55 0.59 0.61 0.72 0.8 0.73 0.78 Pr 4.68 3.59 4.84 1.93 2.29 4.43 3.22 1.94 1.93 2.57 2.13 2.28 Tm 0.157 0.144 0.285 0.31 0.518 0.229 0.248 0.288 0.324 0.354 0.326 0.349 Be 1.04 1.11 1.4 0.82 0.65 1.14 0.61 0.61 0.74 0.76 0.46 0.57 Cd 0.03 0.03 0.37 0.69 1.17 0.19 0.27 0.37 0.8 0.39 0.28 0.79 Ga 16.27 15.14 16.66 13.94 14.33 21.19 16.76 16.31 15.24 14.44 9.8 11.78 Li 17.9 18.6 26.7 24.9 26.7 33.5 30.1 25.5 27.9 44.6 25.6 24.3 Mo 1.52 1.05 1.51 1.14 0.93 1.33 1.66 1.81 1.43 1.11 1.06 0.96 Sb 0.34 0.22 0.93 1.25 1.15 0.62 0.61 1.13 1.42 1.49 1.02 0.74 Sn 0.73 0.73 0.8 0.58 0.54 0.87 0.8 0.7 0.56 0.62 0.69 0.59 Tl 0.31 0.29 0.55 0.36 0.32 0.28 0.39 0.35 0.3 0.33 0.29 0.37 V 18.54 12.18 189.3 284.59 263.86 145.66 202.73 249.17 270.06 307.75 270.75 288.7 W 1.76 3.9 2.3 2.43 1.98 3.05 1.39 1.62 1.57 1.46 1.34 2.16 Zn 21.37 26.71 139.65 105.4 131.72 51.19 77.43 99.6 101.71 50.85 30.67 72.94 Pb 6.5 4.8 8.1 9.7 8.8 5.6 5.1 6.6 8.2 4.7 5.7 6 Cu 4 3 382 235 199 44 172 191 191 225 145 197 Cr 32 <24 190 466 326 122 343 428 397 502 388 334 Ni 14 19 119 193 162 76 131 136 162 232 154 170 Co 3.7 5 38.6 62.7 55.8 25.5 39.4 37 53 62.7 47.3 58 Bi 0.024 0.061 0.024 0.088 0.104 0.028 0.088 0.106 0.123 0.086 0.102 0.05 Ti 1274.8 1039.0 4366.4 5646.3 5878.0 3920.6 4674.2 5442.6 6041.0 7120.2 5825.7 6060.0 Analyses by Geoscience Laboratories, Ministry of Northern Development, Mines and Forestry, Sudbury, Ontario. - 14 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 8. Major element chemistry of all lithologies, South Austin Tuff exposure. Sample No. Rock Type SiO2 (%) TiO2 (%) Al2O3 (%) Fe2O3 (%) K2O (%) MgO (%) MnO (%) CaO (%) Na2O (%) P2O5 (%) LOI (%) Total (%) 2010SM-01 Austin 54.33 1.67 15.68 15.38 4.66 3.55 0.37 1.72 0.55 0.15 1.65 99.71 2010SM-02 Austin 52.45 1.07 14.35 14.45 2.34 4.37 0.32 6.79 0.38 0.12 2.26 98.90 2010SM-03 Austin 46.23 1.64 15.22 14.31 3.34 4.97 0.40 10.25 0.20 0.18 2.51 99.23 2010SM-04 Austin 52.95 1.61 15.44 12.90 3.65 4.14 0.28 6.61 0.17 0.18 2.51 100.43 2010SM-05 Austin 53.84 1.68 16.44 11.79 4.37 3.93 0.21 4.47 0.29 0.20 2.39 99.60 2010SM-06 Austin 54.55 1.69 16.71 10.66 4.35 4.03 0.21 4.40 0.34 0.20 2.33 99.48 2010SM-07 Austin 47.39 1.53 14.45 16.10 2.61 5.92 0.45 8.57 0.36 0.16 2.12 99.67 2010SM-08 Austin 49.95 1.69 16.48 15.00 4.74 4.29 0.27 3.86 0.25 0.20 2.79 99.52 2010SM-09 Austin 53.44 1.59 15.49 14.73 4.36 3.74 0.27 3.33 0.21 0.19 2.37 99.71 2010SM-10 Austin 50.62 1.55 14.95 23.23 3.62 3.11 0.48 0.45 0.04 0.17 1.23 99.46 2010SM-11 Austin 64.42 0.86 14.08 12.20 3.57 1.89 0.27 0.24 0.05 0.10 1.76 99.44 2010SM-12 Austin 58.16 1.13 14.19 15.78 4.65 2.30 0.39 0.36 0.04 0.14 2.01 99.15 2010SM-13 QFP 74.89 0.21 14.08 1.64 3.79 0.74 0.04 1.53 1.01 0.03 1.83 99.80 2010SM-14 QFP 76.14 0.17 13.28 2.27 3.27 0.82 0.04 1.17 1.08 0.03 1.78 100.05 2010SM-15 Huston 62.92 0.72 14.84 8.94 3.45 2.79 0.15 2.65 1.88 0.05 1.42 99.83 2010SM-16 Huston 60.37 0.97 14.21 15.04 2.55 3.33 0.29 1.46 0.20 0.07 1.49 99.99 2010SM-17 Huston 59.94 0.99 13.16 16.81 2.34 2.98 0.42 1.58 0.21 0.06 1.04 99.54 2010SM-18 Huston 64.49 0.67 16.82 7.53 2.93 2.24 0.14 2.11 0.84 0.08 1.99 99.84 2010SM-19 Huston 64.50 0.80 13.64 9.40 2.80 2.85 0.17 2.62 0.90 0.07 1.26 99.01 2010SM-20 Huston 62.65 0.94 13.46 12.23 2.73 3.01 0.20 1.85 0.46 0.06 2.23 99.82 2010SM-21 Huston 60.95 1.05 13.82 14.09 2.22 3.10 0.22 1.65 0.33 0.06 1.40 98.89 2010SM-22 Huston 60.45 1.21 14.48 14.76 2.22 3.38 0.19 1.11 0.29 0.06 1.26 99.40 2010SM-23 Huston 63.56 1.04 13.74 14.01 2.05 2.93 0.28 0.98 0.28 0.06 0.99 99.92 2010SM-24 Huston 58.64 1.04 14.61 16.14 2.53 3.08 0.33 1.52 0.27 0.08 0.68 98.92 Analyses by Geoscience Laboratories, Ministry of Northern Development, Mines and Forestry, Sudbury, Ontario. than 2700Ma. Data from this sample suggest that the maximum age of the breccia is 2700Ma, and including the very youngest result the maximum age could be as young as 2655±9Ma. The 2741±0.8Ma age of the QFP lapilli tuff correlates Plot 2: Uranium-lead concordia diagram of Huston conglomerate. dates for nearly all are likely to record primary crystallization ages. Reversely discordant grains 25 and 33 are the only ones that give results younger Plot 3: Uranium-lead probability plot, Huston conglomerate. - 15 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 9. Precious metal analyses of Austin Tuff exposure samples Units Detect Limit 2010SM-01 2010SM-02 2010SM-03 2010SM-04 2010SM-05 2010SM-06 2010SM-07 2010SM-08 2010SM-09 2010SM-10 2010SM-11 2010SM-12 2010SM-13 2010SM-14 2010SM-15 2010SM-16 2010SM-17 2010SM-18 2010SM-19 2010SM-20 2010SM-21 2010SM-22 2010SM-23 2010SM-24 2010SM-25 2010SM-26 2010SM-15 SP Au ppb 6 37 431 73 35 30 9 <6 8 <6 29 9 36 <6 <6 22 11 16 6 12 15 9 10 8 <6 6 13 19 Pd ppb 1.3 <1.3 <1.3 <1.3 <1.3 <1.3 <1.3 <1.3 <1.3 <1.3 <1.3 3.9 7 <1.3 <1.3 3.4 7.3 5.8 2.2 4.3 5.6 6.1 6.4 6.1 8 6.2 6.1 3.2 Pt ppb 0.4 1 0.9 0.8 0.9 1 0.9 0.9 1 1 1 4.1 4.7 <0.4 <0.4 4.3 8.5 7.2 3 5.2 6.8 8.1 7.9 8 9.1 8.2 7.7 4.7 detrital ages. A similar unit (called the “Madsen Conglomerate”) was sampled by Sanborn-Barrie et al. (2004), approximately 1.2km northeast of the South Austin zone locality. A maximum depositional age of this unit was interpreted to be 2700±6Ma, based on a diverse detrital zircon profile (n=56), with ages that correspond to volcanic assemblage as old as ca. 2.9Ga and younger post-volcanic plutons ca. 2.7Ga. Its relationship to gold mineralization remains unknown. The questions arising from different age profiles obtained from the same outcrop of conglomerate, and the very young depositional ages of detrital zircons (English River assemblage) can be addressed by additional geochronology, detailed mapping of the units in the immediate area and inspection of core from diamond drilling of the Austin ore zone by Claude Resources Inc. and Pure Gold Mining Inc. Stop 3: Buffalo Deposit – Dome Stock mineralization UTM Coordinates: NAD83; 15U 0439890E / 5650800N The approximately 7km diameter hornblendebiotite granodiorite Dome Stock (Table 2) has been dated at 2718±1Ma (Corfu and Andrews 1987) and is interpreted to have been emplaced during D2 (SanbornBarrie et al. 2001). The stock is variably iron-carbonate, sericite, and chlorite altered and deformed. Exposures to be visited are at its southern contact; here it intrudes, and contains xenoliths of, foliated Balmer assemblage mafic volcanic rocks (Fig. 6). well with previous dates of 2744±1Ma obtained by Corfu and Andrews (1987) from the same lapilli tuff or immediately adjacent unit 150m to the southwest, and a felsic spherulitic flow approximately 120m in the hangingwall, dated at 2746+36/-17Ma (Corfu and Andrews, 1987). The same outcrop of interpreted Huston conglomerate at the present locality and one 15km to the northeast (referred to as the “Cemetery conglomerate”) were previously dated by T. Skulski at less than ca. 2746Ma, based on prominent (n=5060) single age peaks (cited in Sanborn-Barrie et al., 2004). They concluded a single-source derivation from Confederation assemblage volcanic rocks. The current data from Madsen suggests that this conglomerate unit may have sampled a very diverse substrate, and therefore displays a wider range of The Dome Stock hosts several gold occurrences and two small, past-producing mines 1) the Red Lake Gold Shore produced 21,100 ounces gold, and 2) the Buffalo Mine produced 1656 ounces gold. The Buffalo Mine was discovered in 1925 and explored several times since then. Note: the adit was reopened by Claude Resources Ltd. in October 1998 to further explore the Buffalo deposit. Rehabilitation of the adit started in 2004. Gold is hosted within two sets of quartz-tourmalinepyrite-calcite veins in conjugate orientation (cm-wide northeast-trending veins at 239°/73°N, and decimetrewide northwest-trending veins at 119°/76°S; Pettigrew, 1999). Their orientation may be a result of the intersection of two previously interpreted (Durocher and Hugon, 1983) deformation zones (St. Paul BayMartin Bay and Flat Lake-Howey Bay Deformation Zones). The dominant vein set strikes northwest and primary quartz vein fill was replaced by tourmaline, concomitant with bleached pink metasomatic halos - 16 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 6. Detailed geology of south side of Buffalo Pit (from Lavigne et al., 1986). developing around tourmaline-rich portions of the veins. Gold is concentrated in the calcite-albitesulphide halos, in particular at its outer fringe, where chalcopyrite and tellurides were deposited. A second stage of gold mineralization is associated with Bitellurides in fractures and cavity fillings in quartz and late fracture-filling pyrite, hosted within the quartztourmaline-pyrite-calcite veins. Stop 4: Flat Lake–Howey Deformation Zone Bay–Flat (similar in appearance to the Dome stock) containing mafic xenoliths and quartz-tourmaline veinlets. Table 10. Major and trace element analyses of mylonite, Stop 4 (major oxides in %, trace elements in ppm) Lake UTM Coordinates: NAD83; 15U 0441670E / 5651602N The Howey Bay–Flat Lake deformation zone was defined by Durocher and Hugon (1983), and was interpreted to be part of a belt-wide system of transcurrent shear zones hosting most of the major gold deposits. Recent detailed work has led to a reevaluation of this concept (Sanborn-Barrie et al., 2000). This stop is approximately 750m southwest of the Howey mine. Intense deformation at this stop has destroyed most primary textures that might be used to identify the rocks. They have been recently re-interpreted by Sanborn-Barrie et al. (2004) as mylonitized dioritic rocks (Table 10), representing a mafic border phase to the Dome stock. Pink felsic dikes that cut the mafic rocks are also mylonitized. Ironcarbonate veins are boudinaged and transposed into the sinistral shear direction. The far western extremity of the outcrops exposes a deformed quartz-feldspar dyke - 17 - Sample # UTM-Easting UTM-Northing Rock Type Sample No. SiO2 TiO2 Al2O3 Fe2O3 K2O MgO MnO CaO Na2O P2O5 LOI Total 2009AL-07 441670 5651602 Mylonite 2009AL-7 61.11 0.62 15.69 5.84 2.04 2.19 0.07 3.11 4.63 0.12 5.1 100.52 Au (opt) Ag (opt) <0.01 <0.1 Rb Ba Sr Sc La Ce Nd Sm Eu Gd Tb 51.6 333.5 518 14.1 20.47 37 19 3.57 1 2.95 0.41 Sample No. Yb Lu Y Zr Th U Hf Nb Ta Cs Dy Er Ho Pr Tm Be Cd Ga Li Mo Sb Sn Tl V W Zn Pb Cu Cr Ni Co Bi Ti 2009AL-7 1.154 0.18 11.25 116 3.12 0.88 3.03 1.74 <0.2 1.34 2.27 1.12 0.42 5.01 0.163 1.07 0.05 18.87 22.1 1.22 2.27 0.79 0.31 96.62 1.36 71.04 7.8 30 71 60 15 0.073 1730.68 �Proceedings of the 61st ILSG Annual Meeting - Part 2 Stop 5: Howey Mine (Fenced In Pit) – Drive-By by calcite. This represents the distal, outer halo of carbonate alteration, as discussed by Parker (2000). UTM Coordinates: NAD83; 15U 0442328E / 5652067N On the north side of Hammell Road a cement foundation marks the site the former Howey mine. Behind the fenced-off area is the site of the crown pillar mined out in the final stages of the mine. The Howey Mine was the first Au producer (1930-1941) in the Red Lake camp and remains the lowest grade profitable gold mine in Canadian mining history (final average grade 0.08oz/t Au, having produced 422,000oz gold). The Howey (and adjacent Hasaga) ore bodies occur in a boudinaged, variably sericitized and silicified quartz-feldspar porphyry dyke trending approximately 065°/80°S. Cm-wide, auriferous quartz veinlets trend 080°, making an angle of 15° with the contacts of the dyke and dip at 80°S. Gold-bearing quartz veinlets formed as the last of three episodes of quartz veining. Gold is associated with pyrite-sphalerite-galenatourmaline±tellurides. Small flat outcrops between the highway and the fence are highly deformed intermediate rocks of the Howey Mine hanging wall. This site lies within the northeast-trending Howey Bay–Flat Lake deformation zone and comprises Confederation age rocks. Stop 6: Chicken Chef Outcrop; Huston conglomerate UTM Coordinates: NAD83; 15U 0442396E / 5652067N This small outcrop is a clast-supported conglomerate, exhibiting cobble-sized, rounded granitoid clasts and stretched lapilli-sized intermediate to mafic clasts in a feldspar-phyric, foliated groundmass. Some of the granitoid clasts are moderately iron-carbonatized; a highly iron-carbonate-altered lense of detritus appears to be partially wrapped around a granitoid cobble. Similar clast lithologies, but also including rounded cherty fragments, are found in an outcrop 200m south along Highway 105. Stop 7: Outcrops on west side of Highway 125 and Sandy Bay Road; Calcite carbonatized pillowed flows (distal carbonate alteration facies). UTM Coordinates: NAD83; 15U 0446915E / 5654682N Slightly deformed pillows of the Confederation assemblage show pervasive calcite carbonatization, calcite veins, and pods (Table 11). Amygdules are also filled (replaced?) with calcite. Jig-saw puzzle breccias (created by fluid overpressure at depth) are cemented Stop 8: Woodland Cemetery Road and Highway 125; Meso–Neoarchean Contact UTM Coordinates: NAD83; 15U 0447235E / 5655572N These outcrops show altered, relatively low-strain, pillowed basaltic komatiite flows of the Balmer Assemblage (Table 12) unconformably overlain by polymictic conglomerate of the Huston assemblage. The exposures are in the transition from calcite carbonatization (distal alteration) to ferroan-dolomite (proximal) alteration. The pillowed and minor massive flows show extensive iron carbonate alteration as well as iron carbonate and quartz veins. Fuchsite is present in the central part of the outcrops on the west side of the highway (cemetery side). The Campbell Mine is approximately 1.5km to the north. While the mafic flows have not been directly dated at this locality, they are typically variolitic, and show a geochemical similarity with known Balmer-age rocks elsewhere; the massive and pillowed flows here can be traced to Balmertown, where an intercalated rhyolite at the Campbell mine was dated at 2989±3Ma (Corfu and Andrews, 1987). Variolitic flows occur in the northern part of the outcrops on the east side of the highway. However, they are unconformably overlain by Huston polymictic conglomerate further south along the outcrop. The conglomerate contains a large proportion of rounded cherty, jasperoidal and pyritic fragments. It represents an apron of Confederation assemblage (McNeely-age 2743Ma) detritus deposited at the break in paleoslope between the Confederation volcanic centre and its Balmer-age substrate. Stop 9: Cochenour Outcrop; Proximal ferroancarbonate alteration; carbonate veining; silicification; and gold UTM Coordinates: NAD83; 15U 0443689E / 5658640N This outcrop was mapped in detail by Williamson and Dubé (2003) at a scale of 1:150. It is situated directly above past-producing workings of the Cochenour Mine. The outcrop is underlain by maficultramafic Balmer assemblage volcanic rocks; the Meso- Neoarchean contact is exposed approximately 900m to the north. - 18 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 11. Major and trace element analyses, Stop 7 lithologies (major oxides in %, trace elements in ppm) Sample # UTM-Easting UTM-Northing 2009-AL-08 446914 5654682 Rock Type Andesite SiO2 TiO2 Al2O3 Fe2O3 K2O MgO MnO CaO Na2O P2O5 LOI Total 43.93 0.55 14.98 9.05 0.18 3.01 0.22 12.88 3.91 0.08 10.81 99.59 2009-AL-09 446914 5654682 Intermediate Dyke 56.29 0.84 16.27 6.81 0.41 4.28 0.09 5.64 4.74 0.42 4.21 100.02 Au (opt) Ag (opt) <0.01 <0.1 <0.01 <0.1 Rb Ba Sr Sc La Ce Nd Sm Eu Gd Tb Yb Lu 4.3 117.8 355 25.2 8.72 18.54 9.58 2.15 0.63 2.36 0.378 1.304 0.19 10.13 100.5 776 14.7 59.04 128.13 64.69 11.54 2.71 7.7 0.892 1.483 0.21 The outcrop contains exposures of iron-carbonate veins and breccia with spectacular colloformcrustiform and cockade textures, as well as a highgrade, auriferous, silicified zone (Williamson and Dubé, 2003). The latter is characterized by strong silicification, with varied proportions of sericite/green mica and carbonate. The presence of green mica hints at a mafic-ultramafic protolith; however, the presence of local subparallel layering in a feldspar- and quartzrich rock argues for a flow-banded felsic protolith. The silicified zone has been offset 50m to the west by latestage black line faults. Extensional iron-carbonate veins are generally barren of gold, except when cut by mm-scale, fine, Sample # Y Zr 2009-AL-08 13.43 47 2009-AL-09 21.16 177 Th 1.83 9.93 U Hf Nb Ta Cs Dy Er Ho Pr Tm Be Cd Ga Li Mo Sb Sn Tl V W Zn Pb Cu Cr Ni Co Bi Ti 0.47 1.19 2.98 0.2 0.62 2.46 1.45 0.51 2.35 0.208 0.45 0.14 14.54 24 1.01 1.58 0.54 0.03 148.81 <0.5 85.32 3.5 78 122 169 48.9 0.022 3160.87 2.27 4.53 8.15 0.5 0.63 4.42 1.9 0.75 16.21 0.246 1.46 0.05 21.74 33 0.79 1.72 1.26 0.06 126.67 <0.5 89.72 8.2 19 43 70 25 0.14 5021.85 auriferous arsenopyrite veinlets. Stop 10. Redcon Carbonate Zone; West and East sides of Nungesser Road; Proximal ferroancarbonate alteration; carbonate veining UTM Coordinates: NAD83; 15U 0448918E / 5661396N This area is approximately 4km north of the Campbell–Red Lake Deposit, still within the proximal, ferroan-carbonate alteration facies. The outcrops are weakly foliated (145°), dominantly massive to pillowed Balmer assemblage basalts (Table 13), occurring within the amphibolite-facies metamorphic aureole of the Walsh Lake Pluton. - 19 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 12. Major and trace element analyses, Stop 8 basalt (major oxides in %, trace elements in ppm). Sample # UTM-Easting UTM-Northing Rock Type Sample No. SiO2 TiO2 Al2O3 Fe2O3 K2O MgO MnO CaO Na2O P2O5 LOI Total 2009-AL-10 447235 5655572 Basalt 2009AL-10 61.58 0.5 6.7 18.09 1.28 3.85 0.36 3.13 <0.01 0.04 3.42 98.86 2009-AL-11 447235 5655572 Basalt 2009AL-11 49.76 0.81 10.58 10.92 1.89 4.05 0.33 7.9 0.29 0.05 13.02 99.6 Au (opt) Ag (opt) <0.01 <0.1 <0.01 <0.1 Rb Ba Sr Sc La Ce Nd Sm Eu Gd Tb Yb Lu Y 86.5 73.6 20 65.4 2.11 5.73 4.44 1.44 0.39 1.67 0.279 1.041 0.18 6.81 62.09 122.7 91 40.3 2.4 6.57 5.4 1.8 0.68 1.97 0.305 0.749 0.11 6.86 The stripped area on the east side of the Nungesser road was mapped in detail (Fig. 7) by Redcon Gold Mines in 1981 (assessment files) and now forms part of Goldcorp Inc.’s holdings. Here, a 1-2m wide carbonate vein is exposed near its southeastern termination. The vein can be traced in outcrop and drilling for approximately 750m to the west-northwest and will be seen at the next stop on the west side of the road. Gold occurs in north-northwest-trending, irregular, cm-thick quartz-actinolite stringers (tension gashes?) within the carbonate vein. After initial, pervasive potassic (biotite) Sample No. Zr Th U Hf Nb Ta Cs Dy Er Ho Pr Tm Be Cd Ga Li Mo Sb Sn Tl V W Zn Pb Cu Cr Ni Co Bi Ti 2009AL-10 32 0.19 0.06 0.75 1.3 <0.2 19.34 1.6 0.85 0.3 0.92 0.139 0.39 0.03 8.59 52.4 0.47 4.1 0.43 0.76 294.17 5.38 40.48 1 210 >600 655 100.5 0.037 2827.8 2009AL-11 30 0.28 0.07 0.83 1.76 <0.2 2.89 1.69 0.78 0.3 1.07 0.114 0.43 0.09 13.4 58.5 0.29 4.95 0.58 0.17 286.05 3.22 38.28 1 139 >600 493 83.5 0.06 4595.4 metasomatism of the basalt, the intrusion of the Walsh Lake Pluton produced a decimetre-scale, meshlike texture of amphibole-quartz veinlets that are prominently displayed due to their positive weathering features. Cross-cutting relationships suggest the following sequence of formation (from Lavigne et al. 1986): 1. quartz-calcite veins 2. ferroan-dolomite veins 3. mafic dyke 4. auriferous quartz veins - 20 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Silicification evident in the pillowed flow on the northern half of the outcrop is barren and apparently not related to the gold-rich silicification event; rather, it may be due to local silica dumping following pervasive carbonate metasomatism. Table 13. Major and trace element analyses, Stop 10 lithologies (major oxides in %, trace elements in ppm). Sample # UTM-Easting UTM-Northing 2009-AL-15 448918 5661396 2009-AL-16 448918 5661396 Basalt-Si+ Basalt Sample No. SiO2 TiO2 Al2O3 Fe2O3 K2O MgO MnO CaO Na2O P2O5 LOI Total 2009-AL-14 448918 5661396 Fe-Carb vein 2009AL-14 3.05 0.1 0.74 10.34 0.14 13.84 0.64 28.6 <0.01 0.02 41.85 99.18 2009AL-15 56.09 1.88 17.14 9.73 0.97 4.71 0.14 5.54 1.58 0.22 2.1 100.08 2009AL-16 50.04 1.69 14.65 15.95 0.86 7.5 0.15 5.74 0.77 0.17 1.89 99.42 Au (opt) Ag (opt) <0.01 <0.1 <0.01 <0.1 <0.01 <0.1 Rb Ba Sr Sc La Ce Nd Sm Eu Gd Tb Yb Lu Y 3 13.1 46 6.8 2.2 4.18 3.23 1.05 0.53 1.57 0.256 0.816 0.14 11.67 30.74 89.9 52 44.2 12.76 33.18 21.28 5.72 1.72 5.17 0.741 2.311 0.34 21.2 22.98 145.1 42 46.2 7.53 19.28 13.99 4.08 1.32 5.09 0.897 3.516 0.51 34.71 Sample No. Zr Th U Hf Nb Ta Cs Dy Er Ho Pr Tm Be Cd Ga Li Mo Sb Sn Tl V W Zn Pb Cu Cr Ni Co Bi Ti 2009AL-14 11 <0.09 0.03 0.26 0.21 <0.2 0.17 1.59 0.89 0.32 0.64 0.124 0.39 0.23 1.19 6.2 0.14 0.91 0.13 0.03 35.32 8.7 124.59 1.5 4 <24 44 9.4 0.015 441.71 2009AL-15 106 0.9 0.23 2.89 5.29 0.3 3.78 4.43 2.55 0.89 4.66 0.368 0.39 0.09 20.89 54.4 0.35 3.28 1.21 0.23 476.84 13.91 124.21 2.8 193 66 70 35.9 0.02 11217.71 2009AL-16 86 0.73 0.18 2.29 5.33 0.3 2.9 5.95 3.7 1.27 2.92 0.54 0.45 0.13 20.82 49.7 0.41 4.83 1.13 0.19 443.67 4.77 162.95 1.7 88 133 63 45.2 0.038 10004.7 Rock Type A “black line” fault occurs in the northern wallrocks of the main carbonate vein. A mafic (or lamprophyre) dyke (unit 4, above) cuts the vein, but is itself cut by late quartz-actinolite-gold stringers. The western outcrops are approximately 300m westnorthwest of the previous exposures. Things to note on the western series of outcrops: • differing colours of cross-cutting carbonate veins • colloform/crustiform textures in carbonate veins • andalusite-garnet-biotite alteration of pillows cut by calc-silicate veins (diopside±calcite, quartz, tourmaline; retrograding to epidote, tremolite, actinolite/hornblende, magnetite) • calc-silicate veins cross-cut by carbonate veins • folding of carbonate veins by the “Mine trend” S2 References (Used in Guide) Corfu, F. and Andrews, A.J. 1987. Geochronological constraints on the timing of magmatism, deformation and gold mineralization in the Red Lake greenstone belt, northwestern Ontario; Canadian Journal of Earth Sciences, v.24, p.1302-1320. Corfu, F. and Stone, D. 1998. Age, structure and orogenic significance of the Berens River composite Batholiths, western Superior Province; Canadian Journal of Earth Sciences, v.35, p.1089-1109. Corfu, F. and Wallace, H. 1986. U-Pb zircon ages for magmatism in the Red Lake greenstone belt, northwestern Ontario; Canadian Journal of Earth Sciences, v.23, p.27-42. Corfu, F., Davis, D.W., Stone, D., and Moore, M. 1998. Chronostratigraphic constraints on the genesis of Archean greenstone belts, northwestern Superior Province, Ontario, Canada; Precambrian Research, v.92, p.277-295. Dubé, B., Balmer, W., Sanborn-Barrie, M., Skulski, T., and Parker, J. 2000. A preliminary report on amphibolite-facies, disseminated-replacement-style mineralization at the Madsen gold mine, Red Lake, Ontario; in Current Research 2000-C17, Geological Survey of Canada, 12p. Dubé, B., Williamson, K., and Malo, M. 2002. Geology of the Goldcorp Inc. High Grade zone, Red Lake mine, Ontario: an update, in Current Research 2002-C26, Geological Survey of Canada, 15p. - 21 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 7. Detailed geology of the Redcon prospect (modified from Lavigne et al. 1986). Dubé, B., Williamson, K., McNicoll, V., Malo, M., Skulski, T., Twomey, T., and Sandborn-Barrie, M. 2004. Timing of gold mineralization at Red Lake, Northwestern Ontario, Canada: New constraints from U-Pb Geochronology at the Goldcorp HighGrade Zone, Red Lake mine, and the Madsen mine; Economic Geology, v.99, No.8, p.1611 to 1641. Durocher, M.E. and Hugon, H. 1983. Structural geology and hydrothermal alteration in the Flat Lake-Howey Bay deformation zone, Red Lake area, in Summary of Field Work, 1983, Ontario Geological Survey, Miscellaneous Paper 116, p.216 to 219. Horwood, H.C. 1940. Geology and mineral deposits of the Red Lake area, in Forty-ninth Annual Report of the Ontario Dept. of Mines, vol. XLIX, Pt. II, 231p. Irvine, T.N. and Baragar, W.R. 1971. A guide to the chemical classification of the common volcanic rocks; Canadian Journal of Earth Sciences, v.8, p.523-548. Jensen, L.S. 1976. A new cation plot for classifying subalkaline rocks; Ontario Geological Survey, Miscellaneous Paper 66, 22p. Lavigne Jr., M.J., Hugon, H., Andrews, A.J., and Durocher, M.E. 1986. Gold deposits of the Red Lake District, Relationships of gold mineralization to regional deformation and alteration in the Red Lake greenstone belt, Ontario, in Gold ‘86, Excursion Guidebook, ed. Pirie, J. and Downes, M.J., p.167 to 211. McMaster, N.D. 1987. A preliminary 40Ar/39Ar study of the thermal history and age of gold in the Red Lake greenstone belt; unpublished M.Sc. thesis, University of Toronto, Toronto, Ontario, 107p. Noble, S.R. 1989. Geology, geochemistry and isotope geology of the Trout Lake Batholith and the UchiConfederation lakes greenstone belt, northwestern Ontario, Canada; unpublished Ph.D. thesis, University of Toronto, Toronto, Ontario, 288p. Parker, J.R. 1999. Exploration potential for volcanogenic massive sulphide (VMS) mineralization in the Red Lake greenstone belt; in Summary of Field Work and Other Activities 1999, Ontario Geological Survey, - 22 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Canada, 14p. Open File Report 6000, p.19-1 to 22-26. Parker, J.R. 2000. Gold mineralization and wall rock alteration in the Red Lake greenstone belt: a regional perspective; in Summary of Field Work and Other Activities 2000, Ontario Geological Survey, Open File Report 6032, p.22-1 to 22-27. Chi, G., Dubé, B., and Williamson, K. 2003. Fluid evolution and pressure regimes in the Campbell-Red Lake gold deposit, Red Lake mine trend, Red Lake, Ontario; in Current Research 2003-C28, Geological Survey of Canada, 18p. Percival, J.A., Bailes, A.H., Corkery, M.T., Dubé, B., Harris, J.R., McNicoll, V., Panagapko, D., Parker, J.R., Rogers, N., Sanborn-Barrie, M., Skulski, T., Stone, D., Stott, G.M., Thurston, P.C., Tomlinson, K.Y., Whalen, J.B., and Young, M.D. 2000. An integrated view of Western Superior crustal evolution: highlights of 2000 NATMAP studies, in Summary of Field Work and Other Activities 2000, Ontario Geological Survey, Open File Report 6032, p.13-1 to 13-17. Dubé, B., Williamson, K., and Malo, M. 2001. Preliminary Report on the Geology and Controlling Parameters of the Goldcorp Inc. High Grade Zone, Red Lake Mine, Ontario; Geological Survey of Canada, Current Research 2001-C18, 13p. Pettigrew, N. 1999. Structural and alteration history of the Buffalo Gold Deposit, Red Lake, Ontario; unpublished B.Sc. Thesis, University of New Brunswick, 154p. Sanborn-Barrie, M., Skulski, T., and Parker, J. 2001. Three hundred million years of tectonic history recorded by the Red Lake greenstone belt, Ontario, in Current Research 2001-C19, Geological Survey of Canada, 32p. Sanborn-Barrie, M., Skulski, T., and Parker, J. 2004. Geology, Red Lake greenstone belt, Western Superior Province, Ontario; Geological Survey of Canada, Open File 4594, scale 1:50,000. Sanborn-Barrie, M., Skulski, T., Parker, J., and Dubé, B. 2000. Integrated regional analysis of the Red Lake greenstone belt and its mineral deposits, western Superior Province, Ontario, in Current Research 2000-C18, Geological Survey of Canada, 16p. Skulski, T., Sanborn-Barrie, M. and Sanborn, N. 2001 New U-Pb geochronology in the Red Lake greenstone belt, Western Superior NATMAP, unpublished poster. Williamson, P.K. and Dubé, B. 2003. Detailed geology, hydrothermal alteration and gold mineralisation of the Cochenour stripped outcrop, Red Lake gold district, Ontario; Geological Survey of Canada, Open File 1673. Bibliography of Recent Research (not used in Guide) Chi, G., Dubé, B., and Williamson, K. 2002. Preliminary fluid-inclusion microthermometry study of fluid evolution and temperature-pressure conditions in the Goldcorp High-Grade zone, Red Lake Mine, Ontario, in Current Research 2002-C27, Geological Survey of Dubé, B., Williamson, K., and Malo, M. 2003. Gold mineralization from the Red Lake mine trend: Example from the Cochenour-Willans mine area, Red Lake, Ontario, with new key information from the Red Lake Mine and potential analogy with the Timmins camp; in Current Research 2003-C21, Geological Survey of Canada, 15p. Gulson, B.L., Mizon, K.J., and Atkinson, B.T. 1993. Source and timing of gold and other mineralization in the Red Lake area, northwestern Ontario, based on leadisotope investigations, Canadian Journal of Earth Sciences, v.30, p.2366 to 2379. Parker, J.R. 2001. Intermediate to Felsic Plutons in the Red Lake Greenstone Belt: Relationship to Deformation and Gold Mineralization; in Summary of Field Work and Other Activities 2001, Ontario Geological Survey, Open File Report 6070, p.19-1 to 19-10. Penczak, R.S. and Mason, R. 1997. Metamorphosed Archean epithermal Au-As-Sb-Zn-(Hg) vein mineralization at the Campbell Mine, Northwestern Ontario, Economic Geology, v.92, p.696 to 719. Penczak, R.S. and Mason, R. 1999. Characteristics and origin of Archean pre-metamorphic hydrothermal alteration at the Campbell Gold Mine, Northwestern Ontario, Canada, Economic Geology, v.94. p.507-528. Pirie, J. and Downes, M.J., eds. 1986. Gold ‘86 Excursion Guidebook. Stone, D. and Hallé, J. 2000. Geology of the Blackbear, Yelling and Stull Lake areas, Northern Superior Province, Ontario, in Summary of Field Work and Other Activities 2000, Ontario Geological Survey, Open File Report 6032, p.15-1 to 15-9 Thompson, P.H. 2003. Toward a new metamorphic framework for gold exploration in the Red Lake greenstone belt; Ontario Geological Survey, Open File Report 6122, 52p. Twomey, T. and McGibbon, S. 2002. The geological setting and estimation of gold grade of the High-Grade zone, Red Lake mine, Goldcorp Inc.; Exploration and Mining Geology, Nos. 1 and 2, p.19 to 34. - 23 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 2 - Western Wabigoon Subprovince Transect, Dryden to Meggisi Lake Mark Puumala and Dorothy Campbell Resident Geologist Program, Ontario Geological Survey, Thunder Bay, Ontario, Canada Craig Ravnaas Resident Geologist Program, Ontario Geological Survey, Kenora, Ontario, Canada Introduction This field trip will provide participants with an overview of the geology of a portion of the western Wabigoon Sub-province along an approximately 52km long transect that follows the Highway 502 corridor between the City of Dryden and Meggisi Lake (see Fig. 1, A-B; Fig. 3). The trip will begin in metasedimentary rocks near the southern margin of the Sioux Lookout domain, crossing over the Wabigoon fault into the Atikwa-Manitou volcano-plutonic domain. The majority of the trip will be spent observing supracrustal rocks of the Eagle-Wabigoon-Manitou Lakes (EWM) greenstone belt and plutonic rocks of the Atikwa batholith and Taylor Lake stock. The area has long been known for its gold exploration potential, with modest levels of production (approximately 13,000oz. Au) occurring between 1895 and 1948. As a result, the trip will also include a stop to observe a gold occurrence located near Flambeau Lake. Regional Geology Blackburn et al. (1991) describe the Wabigoon Subprovince as a 900km long by 150km wide granitegreenstone sub-province in the northwestern Superior Province. It consists of metavolcanic and subordinate metasedimentary rocks that are surrounded and cut by granitoid batholiths. The Wabigoon Sub-province is bounded on the north by the Winnipeg River and English River Sub-provinces and on the south by the Quetico Sub-province (Fig. 1). The Wabigoon Sub- Figure 1. General geology of the western Wabigoon Sub-province, showing its extent, major supracrustal belts and structural features (Blackburn et al., 1991). The transect is defined by the A-B line. - 24 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 province has been divided into three regions, referred to as: i) the western region; ii) the central region; and iii) the eastern region. As noted above, this field trip will transect a portion of the western region. The following summary of the regional geology of the western Wabigoon Sub-province is excerpted from Beakhouse et al. (1995). “The western Wabigoon region is characterized by interconnected, arcuate, metavolcanic dominated ‘greenstone belts’ surrounding large elliptical batholiths. The metavolcanic component of greenstone belts includes minor ultramafic (komatiitic), through abundant mafic (tholeiitic, calc-alkalic and minor alkalic) to felsic (mostly calc-alkalic) varieties. Except locally, metasedimentary rocks are volumetrically minor but diverse including turbiditic, volcaniclastic deposits, alluvial fan-fluvial deposits and chemical (magnetite ironstone and chert) deposits. Stratigraphic sequences generally comprise basal, laterally extensive, mafic metavolcanic sequences overlain by laterally limited, diverse mafic to felsic sequences. Minor clastic metasedimentary deposits are associated with some of the intermediate to felsic volcanism. Very locally, coarse clastic-dominated metasedimentary sequences with subordinate chemically distinct metavolcanic rocks unconformably overlie the diverse volcanic sequences. The principal exception to the generalized lithologic proportions outlined above occurs in an area north of the Wabigoon fault in the Dryden-Sioux Lookout area where metasedimentary rocks predominate. A more detailed review of the stratigraphy and geochemistry of western Wabigoon greenstone belts is presented in the guide portion of this volume. generalization include the Dryberry and Ghost Lake batholiths which are younger than and compositionally distinct from metavolcanic rocks in this area. The smaller stocks are predominantly late- to post-tectonic and range compositionally from diorite to granite and syenite. Minor, late alkalic intrusions (e.g., Sturgeon Narrows) occur in the Sturgeon Lake area. The deformational style of much of the western Wabigoon region, and particularly that portion lying to the south of the Wabigoon fault, is dominated by structural domes cored by large batholithic masses giving rise to apparent synclinal keels of greenstone belts surrounding the batholiths. In detail, it is not possible to correlate units on either side of the apparent ‘synclinal axes’ and these zones of opposing stratigraphic facing correspond, in part, to faults that have juxtaposed segments of volcanic rock of contrasting ages. Laterally continuous deformation zones exhibiting complex kinematics typically occur along the central axis of the greenstone belts where greenstone sequences face one another and may be related to this faulting. The northern portion (north of the Wabigoon fault) of the sub-province has a distinct structural style reflected in linear, fault bounded panels trending parallel to the sub-provincial boundary that contrasts with that of the remainder of the western Wabigoon region. Here there is evidence for early recumbent folding and thrust faulting as well as a later phase of dextral, transcurrent shear. Greenschist-grade regional metamorphic mineral assemblages characterize much of the greenstone belts. The principle exceptions to this generalization are narrow amphibolite-grade zones that occur at the contact with granitoid batholiths and at sub-province boundaries. A particularly noteworthy exception occurs in the Dryden area where there is widespread evidence for in situ partial melting of pelitic metasedimentary rocks. Granitoid rocks within the western Wabigoon region include large elliptical to multi-lobate batholiths that define the architecture of the greenstone belts as well as smaller stocks. Most of the large batholiths (Aulneau, Atikwa, Sabaskong) range compositionally from ultramafic to granitic but are predominantly tonalitic to granodioritic. These are closely associated petrogenetically and temporally with the metavolcanic rocks of the greenstone belts and are interpreted to represent sub-volcanic chambers that have risen into their own volcanic ejecta. Exceptions to this U-Pb geochronological constraints indicate that metavolcanic rocks were deposited between 2775 and 2711Ma, and much of this in the narrow interval of time between 2740 and 2720Ma. Large granitoid batholiths occurring to the south of the Wabigoon fault were emplaced synchronously with adjacent volcanic rocks, whereas those to - 25 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 the north of the fault tend to be younger than 2710Ma. Small, post-tectonic plutons were emplaced over a 15Ma interval commencing at ~2699Ma.” different interpretations of their relative age of emplacement. Presence of layering in the Gabbro Lake body suggests that they are horizontally-emplaced sills, and regional crosscutting relationships suggest their emplacement subsequent to major folding. This argument led to the suggestion (Blackburn, 1980) that they were emplaced subsequent to overthrusting of the Boyer Lake volcanics. However, an age of 2722Ma later obtained (Davis et al., 1982) from the same sill is not consistent with a postthrusting emplacement. These problems remain unresolved.” Manitou-Stormy Lakes Greenstone Belt The first half of the field trip will be spent viewing rocks located in the Manitou-Stormy Lake portion of the EWM Lakes greenstone belt. Beakhouse et al. (1995) describe the Manitou-Stormy Lakes greenstone belt as follows. “Stratigraphic and structural relationships in this greenstone belt, supported by age dating (Davis et al., 1982; Davis, 1990; Parker et al., 1989: summarised in Blackburn et al., 1991), suggest a history in which two early mafic tholeiitic sequences, the Wapageisi Lake group (not dated directly, but in the range 2745 to 2730Ma) and the Boyer Lake group (not dated directly, but >2722Ma) have been juxtaposed against an intervening sequence of later calcalkalic pyroclastics and overlying clastic metasedimentary rocks (Manitou-Stormy groups: <2706Ma and >2696Ma). To the northwest of the Manitou Straits fault, supracrustal units (Upper Manitou Lake, Pincher Lake, Lower Wabigoon volcanics: >2732Ma) that extend northward to the Wabigoon fault are examples of volcanic ejecta that are invaded by their own magma chamber (Atikwa batholith, Dore Lake lobe: 2732Ma). The early mafic suites are distinctly different, and may represent the ensimatic mafic plane within which the emergent pyroclastic edifices rose in island arc-related environments. The mafic plane analogue is supported by the presence within Wapageisi volcanics of two 200m thick, laterally extensive (>10km strike length) plagioclase-phyric flow units, 1500m apart. Differing lines of evidence from two tabular gabbro bodies (Gabbro Lake, Mountdew Lake) within the Boyer Lake group lead to Economic Geology Exploration and Production History The Wabigoon Lake and Upper Manitou Lake areas have long been recognized as having significant gold exploration potential. During the time period between 1895 and 1943, several small gold deposits were developed and modest amounts of gold were produced in these areas. The majority of the historic gold production occurred in the Gold Rock mining camp near Upper Manitou Lake. Past-producers at Gold Rock included the Laurentian, Big Master (aka Kenwest), Elora (aka Jubilee) and Gold Rock mines (Parker, 1989). Total production from these four mines was 12,113.21oz Au. Production statistics for the individual mines are provided in Table 1. Gold production from the Wabigoon Lake area during the same time period was much less significant, with a total of 613.02oz Au mined from four deposits (Parker, 1989). The past producing mines and their production statistics are listed below (Table 2). Since 1948, the Wabigoon Lake and Upper Manitou Lake areas have both seen sporadic exploration activity, including a period of base metal exploration during the 1960s and 1970s (Parker, 1989). However, most exploration activity that has been completed in these areas since 1980 has continued to focus on gold. Table 1. Gold production statistics for the Gold Rock Mining Camp (from Parker, 1989). Mine Laurentian Big Master (Kenwest) Elora (Jubilee) Gold Rock Total Production (oz Au) 8143 2565.52 1369.69 35 - 26 - Years of Operation 1906-09 1902-05, 1942-43 1936-37, 1939 1929 �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 2. Gold production statistics for the Wabigoon Lake area (from Parker, 1989). Mine Redeemer Bonanza Rognon Van Houten Total Production (oz Au) 343.96 243.85 22.21 3 The most notable recent exploration activity in the Wabigoon Lake area has occurred on the Laurentian Goldfields Ltd. Van Horne Gold Property. This property will be visited during this field trip, and is described in detail as field trip Stop 11. In the Upper Manitou Lake area, Manitou Gold Inc. has been active recently on a number of claim groups, including their Elora and Kenwest properties. These properties include the past-producing Laurentian, Big Master, and Elora Mine sites. There are 61 known gold showings, occurrences, prospects, and deposits in the Upper Manitou Lake area (Maunula and Wilson, 2010). These showings include the Starr Gold Occurrence that is described below as an optional self-guided field trip stop (Stop 8). Years of Operation 1904-06, 1918 1920, 1923 1916-18 1940 Lake Volcanics, and underlies a pillowed mafic flow sequence occurring at the top of the volcanic succession, the Upper Wabigoon Volcanics.” The stratigraphic and structural controls on gold mineralization in the Upper Manitou Lake area are described as follows by Parker (1989). “It is apparent that stratigraphy plays an important role in the concentration of gold at Upper Manitou Lake. Gold deposits are situated near the contacts between the Blanchard Lake, Upper Manitou Lake, and Pincher Lake Volcanic Groups. The mines at Goldrock, and the significant The location of the Van Horne gold property and the historic Gold Rock mining camp are shown on Figure 2. This map also illustrates mining claims that were in good standing as of April 20, 2015 and the location of mineral exploration projects that were reported by Lichtblau et al. (2015) to have been active during 2014. Gold Mineralization Gold occurrences in the Eagle-Wabigoon-Manitou Lakes greenstone belt show strong spatial and genetic correlations with major deformation zones such as the Wabigoon and Pipestone Manitou Straits faults (Parker, 1989). Parker (1989) noted the following structural and stratigraphic relationships to gold mineralization in the Wabigoon Lake area. “Gold-bearing quartz veins west of Wabigoon Lake are controlled by northwest-trending tension fractures and east- and east-northeast-trending shear zones. The majority of known gold occurrences at Eagle and Wabigoon Lakes are situated within a mixed sequence of mafic to felsic metavolcanics grouped together as the Lower Wabigoon Volcanics: it overlies a thick sequence of massive and pillowed mafic flows, the Eagle Figure 2. Map illustrating locations of Van Horne gold property and Gold Rock historic mining camp relative to field trip stops. Mining claims and exploration properties active during 2014 are also shown. - 27 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 gold occurrences and prospects in the area, are situated within the Pincher Lake Group, near its contact with the underlying Upper Manitou Lake Group. The Pincher Lake Group may be identical to the Lower Wabigoon Volcanics, which host numerous gold deposits and past producing mines at Eagle and Wabigoon Lakes. The Pincher Lake Group and Lower Wabigoon Volcanics are mixed sequences of calc-alkaline to tholeiitic, mafic and felsic metavolcanic rocks. exposure of folded Thunder Lake metasedimentary rocks in outcrops located adjacent to the Dryden Walmart parking lot. As noted above, there are significant differences in structural style on the north and south sides of the Wabigoon fault, suggesting that these rocks represent domains that evolved separately and have been juxtaposed against each other along the fault. Rocks located to the north of the Wabigoon fault, including the Thunder Lake sediments, are collectively referred to as the Sioux Lookout domain (Beakhouse et al., 1995). The fact that gold deposits are concentrated northwest of the Manitou Straits Fault is partly due to the fact that geological successions on either side of the fault are substantially different. The rocks southeast and east of the fault consist of the dominantly mafic Boyer Lake Group, and the Manitou Lake Group, both of which host a few scattered gold occurrences. Rocks northwest of the fault are dominantly calcalkaline to tholeiitic and are more typical of gold-bearing environments elsewhere in the Dryden-Ignace area. Structural disruption is also more significant northwest of the fault, providing numerous dilatant zones for the emplacement of gold-bearing quartz veins. Extensive and intense northeast shearing controls the majority of goldbearing quartz veins. At Peak Lake, northeast of Upper Manitou Lake, gold-bearing quartz veins are controlled by northwest-trending fractures which crosscut east and northeast-trending shear zones. This style of deformation is similar to gold deposits at Dinorwic Lake, and differs from the dominant northeast controls at Goldrock. Blackburn et al. (1991) suggest that the Wabigoon fault displays an early history of thrusting (south side up), followed by dextral strike-slip movement. Evidence for the former is most evident to the south of the fault, where tight folds with sub-horizontal axes occur in the Wabigoon volcanic rocks. Evidence for strike-slip motion is most notable to the north of the fault, where the Thunder Lake metasedimentary rocks are folded about sub-vertical axes. The Thunder Lake metasedimentary rocks are described as follows by Beakhouse (2000). “The Thunder Lake sediments include two separate panels of rock separated along a portion of their strike-length by the Thunder Lake volcanics. Thin- to medium-bedded wackesiltstone characterized by even, continuous bedding is the predominant component in both panels. Thin magnetite ironstone layers are a conspicuous minor component within the Thunder Lake sediments north of the Thunder Lake volcanics but are rare within the southern panel. Minor, thin garnet-rich (<70% garnet) and calc-silicate layers may represent original more pelitic and marly layers, respectively. In one location, the calc-silicate material forms discordant veins, and suggests that some of this material is remobilized or originated by secondary alteration processes. The garnet-rich layers are often closely spatially associated with the ironstone layers. Competency and susceptibility to fracturing of the host rock is the controlling influence on the concentration of the quartz veins: ductility contrasts between felsic dikes and mafic metavolcanics have focussed zones of extension and compression, commonly resulting in fracturing of the more rigid felsic dikes.” Field Trip Stops Stop 1: Thunder Lake metasedimentary rocks UTM Coordinates: NAD83; 15U 0513498E / 5514739N The Walmart store property located on TransCanada Highway 17 (Government Street) at the east end of Dryden. This stop provides an opportunity to view an - 28 - A limited number of determinations indicate that tops, although locally reversed by tight to isoclinal minor folding, are generally to the south in both the north and south panels. Contact relations with the Brownridge volcanics have not been observed but the data are permissive of a conformable stratigraphic relationship. The contacts with the Thunder Lake volcanics �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 3. General geology with field trip stops (geology from Ontario Geological Survey 2011). - 29 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Lake pluton (~2698Ma). A laser ablation inductively coupled mass spectrometric (LAICP–MS) U/Pb age of 2727±5Ma for a gabbro occurring to the southeast of Stormy Lake (Stone et al., 2010) may approximate the age of the unit; however, the sampled unit occurs at the contact between the Wapageisi and Bending Lake volcanics that both contain gabbro intrusions. The basal contact of the Wapageisi volcanics is in part intrusive where younger phases (Islet Lake Stock, Stone, Hellebrandt, and Lange, 2011b; Meggisi pluton, Blackburn, 1982) are in contact with the basal volcanic units. However, a well-developed, commonly shallowly dipping, penetrative deformation fabric is restricted to the basal 500m of the Wapageisi volcanics with the intensity of fabric development increasing towards the contact suggesting that the local intrusive contacts may be superimposed on a tectonic contact. The character of the northern contact of the Wapageisi volcanics differs across the area.”x appear to also be conformable although these are commonly moderately highly strained and a loci of abundant quartz veining.” This outcrop exposure illustrates the effects of two phases of deformation that resulted in the development of refolded folds in the Thunder Lake metasedimentary rocks. The first phase of deformation resulted in the formation of isoclinal folds about steeply-plunging axes. These folds were subsequently refolded into open z-folds indicative of a late phase of dextral transcurrent shear along the Wabigoon fault (Beakhouse et al., 1995). The axial planes of the late folds strike approximately east-northeast. Parker (1989) also noted the presence of structures (e.g., northwest-trending tension fractures) on the south side of the Wabigoon fault that are consistent with a late phase of dextral transcurrent motion. We will now travel for approximately one hour along Highways 594 and 502 until we reach the southern end of the transect at Meggisi Lake. Stop 2: Wapageisi metavolcanic rocks UTM Coordinates: NAD 83; 15U 0524073E / 5463090N to 0524055E / 5463690N Turn left onto Highway 17 and travel 650m to the traffic lights at the intersection with Highway 594 (Duke Street). Turn left onto Highway 594 and travel 8km to the intersection with Highway 502. Turn left onto Highway 502 and travel 64.5km south to Meggisi Road. This field trip stop consists of a series of road cuts extending northward from Meggisi Road for approximately 600m along the west side of Highway 502. Uphill Road is intersected at a distance of 350m along the section. These bedrock exposures provide excellent examples of relatively undeformed pillowed and massive flows of the Wapagesi metavolcanic rocks. The Meggisi-Uphill Roads section through the Wapagesi metavolcanic rocks commences in plagioclase-phyric pillowed basalt (Fig. 4), typical of the two laterally extensive, plagioclase-phyric flow units that provide stratigraphic markers near the base of the Wapagesi metavolcanic rocks (Beakhouse et al., 1995). The flows at this location are upward-facing with a shallow dip (approximately 25°) toward the north. Moving north, the plagioclase-phyric unit is overlain by aphyric massive basalt that is overlain by pillowed flows and pillow breccia. A 10cm wide mafic dike can be observed cross-cutting the pillow breccia. The The Wapageisi metavolcanic rocks are described as follows by Beakhouse (2011). “The Wapageisi volcanics comprise a homoclinal, generally northward-facing panel of submarine mafic volcanic and related gabbroic rocks. Plagioclase glomeroporphyritic units are laterally continuous and make up approximately 10 to 15% of the unit. The age of the Wapageisi volcanics is constrained to being older than both the Manitou–Stormy supracrustal association (~2703Ma) and the late- to post-tectonic Taylor Figure 4. Plagioclase-phyric pillowed flows. - 30 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 5. Quartz-carbonate filled “eyebrow” structures. northeast-striking dike parallels shear fractures and dips approximately 70° toward the northwest. This unit is overlain by a second massive flow that passes into another sequence of pillowed material. The massive flow contains large quartz-carbonate filled “eyebrow” structure gas cavities (Fig. 5) whose long axes are oriented approximately parallel to stratigraphy. Figure 6. Roadside exposure of pillowed Wapageisi metavolcanic rocks. are scattered throughout. Narrow selvages and lack of vesicles suggest deep-water emplacement but the presence of carbonate-filled gas cavities appears to contradict this. Size of pillows ranges from 20cm or less to about 1m. On the east side of the highway, excavation for the road cut, followed by winter frost heave, has exposed exceptionally well preserved three-dimensional pillow shapes. In one specimen, the budding neck of the pillow can clearly be seen in three dimensions.” More pillowed flows and flow-breccia are present to the north of the Uphill Road turnoff. Cavities in the breccia have been in-filled with quartz, carbonate and epidote. Stop 3: Pillowed lavas in Wapageisi metavolcanic rocks UTM Coordinates: NAD83; 15U 0524129E / 5464528N Travel 1.8km north of Meggisi Road along Highway 502. Because these roadside outcrop exposures are located on a corner, park at a roadside stop located on the west side of the highway at UTM Co-ordinates: NAD83; 15U 0524385E / 5464742N. The outcrops are located approximately 350m to the southwest of the parking spot. This stop provides a good opportunity for participants to obtain photographs of excellent examples of pillowed mafic metavolcanic flows in the lower, tholeiitic mafic plane sequence of the Wapageisi metavolcanic rocks (Fig. 6). The following description of the outcrop exposures is excerpted from Beakhouse et al. (1995). “In the west face, pillow shape and packing indicates tops to the northwest. Good examples of bun-shaped to mattress-shaped to budding forms can be seen. Sparse feldspar phenocrysts Stop 4: Manitou Group conglomerate and quartzfeldspar porphyry UTM Coordinates: NAD83; 15U 0521390E / 5469239N Travel 5.4km north from Stop 3 along Highway 502 to the intersection with Mosher Road. Turn west on Mosher Road and travel 6.6km This stop provides an opportunity to view polymictic conglomerate of the Manitou group that has been intruded by a quartz-feldspar porphyry dike. Beakhouse et al. (1995) describe the Manitou group as follows. The correlatable Manitou and Stormy Lake groups are typical of upper, emergent, chemically and texturally diverse, predominantly calc-alkaline sequences that provided fill for marginal sedimentary basins. In other greenstone belts these have been termed “Temiskaming” type. The base of the sequence unconformably (in places with high angle) rests on Wapageisi Lake group tholeiites. Calc-alkaline pyroclastics are typically coarse, and intruded by comagmatic, dacitic, subvolcanic porphyry stocks, one of which has been dated at 2699Ma (Don Davis, personal - 31 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 communication, 1989). At the top of the pyroclastic sequence lies a mafic alkaline (trachybasalt) flow unit (Sunshine Lake formation) that is unique in the western portion of the Wabigoon Subprovince. Rubble at the top of the unit marks the passage into the epiclastic suite, in which there is vertical and lateral facies variation from alluvial fan and fluvial into resedimented conglomerates, sandstones and mudstones. The resedimented facies show all the characteristics of turbidite sequences: coarse heterolithic conglomerates, graded wacke to mudstone beds, and loading textures (e.g. flame structure) (Teal and Walker, 1977; Blackburn, 1981). The polymictic conglomerate seen here forms part of the Mosher Bay metasedimentary rock sequence. The Mosher Bay metasedimentary rocks are an approximately 2000m thick, north-northwest facing, steep northerly- to vertically-dipping sequence comprised of conglomerate, sandstone, mudstone and minor iron formation. The lower two thirds of the sequence are dominated by conglomerate, while the upper third is a mixture of conglomerate, sandstone and mudstone. The conglomerates include both volcanic clast-dominated and polymictic horizons (Blackburn, 1981). access road located just to the south of Rattlesnake Creek. The outcrops on both sides of the highway provide a good exposure of the late- to post-tectonic 2695Ma (Davis et al., 1982) Taylor Lake Stock. The stock is a relatively small intrusion (approximately 11km by 6km) that is completely enclosed by supracrustal rocks. The intrusion is inhomogeneous and ranges in composition from granodiorite to monzodiorite. However, the highway transects the relatively homogeneous, granodioritic, western portion of the stock (Blackburn, 1981). Interesting features that are visible in these outcrops include exfoliation fractures (Fig. 7), shear fractures, a pegmatite dike that parallels a shear, and numerous mafic xenoliths. Stop 6: Manitou Group alkaline metavolcanic rocks UTM Coordinates: NAD83; 15U 0526602E / 5471785N Travel 1.4km north from Stop 5 to outcrop exposures located on both sides of the highway, north of Rattlesnake Creek. The following description of this stop is excerpted from Beakhouse et al. (1995). The polymictic conglomerates have variable clast sizes (pebbles, cobbles and boulders) and include a mixture of lithologies including granitoid rocks, volcanic rocks, chert, jasper, and magnetite iron formation. The proportion of volcanic clasts found in the polymictic conglomerates decreases upward in the sequence (Blackburn, 1981). “The river valley follows the east-striking Mosher Bay - Washeibemaga (MBW) fault, along which the Boyer Lake volcanics may have been thrust over the Manitou and Stormy Lake groups. In the outcrop on the east side of the highway alkaline volcanics of the Manitou group are intruded by granitic rocks of the late tectonic Taylor Lake stock. Sigmoidal shears in the alkaline volcanics may be related to movement on the MBW fault. The Mosher Bay metasedimentary rocks are intruded by a concordant quartz feldspar porphyry intrusion that covers a strike length of more than 3 km and ranges in width from tens of metres to 300m. The Mosher Bay porphyry is petrographically identical to nearby porphyries at Sunshine and Thundercloud lakes and contains abundant 2 to 4mm phenocrysts of quartz and feldspar throughout. Several similar, much smaller porphyry intrusions occur in proximity to the Mosher Bay porphyry and are likely to be genetically related (Blackburn, 1981). Stop 5: Taylor Lake Stock UTM Coordinates: NAD83; 15U 0526461E / 5470489N Return to Highway 502 from Stop 4 and turn left back onto Highway 502. Travel 1.3km north along Highway 502 to the intersection with a Domtar forest Figure 7. Exfoliation fractures in the Taylor Lake stock. - 32 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 To the north across the valley is a narrow gabbro unit, similar to those at the last stop, emplaced into the Boyer Lake volcanics. Although not evident at this stop, the MBW fault is cut out for a portion of its length by the Taylor Lake stock, and has in turn been sinistrally offset along the north-northeast-striking Taylor Lake fault.” by Asamera Inc. failed to identify any significant mineralization at depth or along strike. The location of the original surface showing can be recognized as a notch in the exposure where shallow blasting was carried out following the initial discovery. Stop 7 (Optional Self-Guided Tour Stop): Mountdew Lake gabbro UTM Coordinates: NAD83; 15U 0525623E / 5474074N Travel 2.4km north from Stop 6 along Highway 502 to the intersection with a logging road. Turn right and park on the side of the logging road. Outcrops of the Mountdew Lake gabbro sill are visible here near Highway 502 and along a logging road on the east side of the highway. This gabbro body discordantly intrudes mafic metavolcanic rocks of the Boyer Lake group (Blackburn, 1981). The intrusion post-dates formation of the Kamanatogama syncline which has affected the surrounding metavolcanic rocks (Blackburn, 1982). Although the intrusion has not been mapped in detail, a number of intrusive phases were identified by Blackburn (1981). The dominant phase of the intrusion (visible at this stop) is a medium-grained, dark green gabbro exhibiting an ophitic to subophitic texture in which amphibole derived from pyroxene is intimately intergrown with saussuritized plagioclase feldspar (Blackburn, 1981). Other intrusive phases include coarse-grained to pegmatitic gabbro containing considerable quantities of magnetite or ilmenite, altered pegmatitic gabbro with colourless “eyes” of quartz and/or potassic feldspar, and granophyre comprised dominantly of pink potassic feldspar with minor quartz (Blackburn, 1981). Stop 8 (Optional Self-Guided Tour Stop): Starr Gold Occurrence UTM Coordinates: NAD83; 15U 0524624E / 5483569N Travel 10.6km north from Stop 7 along Highway 502. Located immediately north of a paved pull-over area on the east side of the highway. Visible gold was discovered in a narrow quartz vein located near the south end of the outcrop on the east side of the highway in 1982 by local area prospector E. Starr. Follow-up prospecting identified another gold showing approximately 15m east of the highway. However, follow-up drilling of the occurrence in 1983 Rocks at this location are deformed and altered pillowed mafic metavolcanic flows of the Pincher Lake group. The deformation, alteration and gold mineralization observed here are associated with a structure known as the Gold Rock splay (Beakhouse et al., 1995). It is a subsidiary structure to the Manitou Straits fault, which crosses the highway approximately 4km to the south. As noted previously, the most significant gold deposits in the Manitou-Stormy Lakes greenstone belt occur nearby. They are also located to the northwest of the Manitou Straits fault in rocks of the Pincher Lake group, near its southern contact with the Upper Manitou group. Alteration includes silicification and carbonatization with numerous quartz-carbonate pods, veinlets and stringers. In many locations, the alteration is accompanied by disseminated pyrite mineralization. Stop 9: Giant pillows in Pincher Lake metavolcanic rocks UTM Coordinates: NAD83; 15U 0521772E / 5485196N Travel 3.5km north from Stop 8 along Highway 502. Park vehicles at the entrance to the Ministry of Transportation gravel pit located on the east side of the highway. The outcrop is located opposite the Fort Frances 140km sign. Mafic metavolcanic rocks of the Pincher Lake group are exposed on this road cut located on the east side of the highway. These rocks are pillowed flows that show wide shape and size variations. This exposure is notable for the remarkable size of many of the pillows, some of which span the entire height of the road cut (a distance of at least 5m) and are up to 2m thick (Fig. 8). Pillow selvages are narrow, with a few much smaller pillows visible between the giant ones. In a well-exposed road cut such as this, these rocks can easily be classified as pillowed mafic metavolcanic flows. However, this type of exposure illustrates the potential pitfalls of trying to classify flow morphology in areas where only small surface exposures are available (i.e., giant pillows could easily be mistaken for massive flows). - 33 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 LILE-enriched mantle may be present in these plutons remains a question for further study. These rocks, chemically high-Mg andesites and originally referred to as sanukitoid, have many geochemically overlapping characteristics with basalt-derived tonalite but differ importantly in having high-Mg as well as Ni and Cr (e.g., Stern et. al., 1989; Sutcliffe et. al., 1990). An important difficulty in distinguishing these two origins, however, arises from the probability that even limited fractionation will obscure some of the diagnostic characteristics of primitive sanukitoid magmas.” Figure 8. Giant pillows of the Pincher Lake metavolcanic rocks. Stop 10: Atikwa Batholith UTM Coordinates: NAD83; 15U 0516387E / 5498799N Travel 15.7km north from Stop 9 along Highway 502. The Atikwa Batholith is one of a number of intrusions in the western Wabigoon Sub-province that may represent subvolcanic magma chambers that have risen into their own volcanic ejecta. The following description of these intrusions is excerpted from Beakhouse et al. (1995). “The Aulneau, Sabaskong and Atikwa batholiths, that are interpreted to represent the plutonic root to much of the volcanism in the western Wabigoon Sub-province, are dominated by rocks within the compositional spectrum tonalite-quartz diorite-granodiorite. The absence of inherited zircons (Davis et. al, 1988) and equivocal nature of field evidence for an unconformable relationship with potential underlying sialic crust, suggest development in an ensimatic regime. These rocks have distinctive geochemical characteristics including low K and Rb and high Ca, Sr and Na/K, moderately fractionated REE with HREE at 1-4x chondrite, LREE at 30-60x chondrite and negligible Eu anomalies and mantle-type radiogenic (Sr and Nd) isotopic signatures. These characteristics have led to the interpretation that many of these rocks were derived from the partial melting of tholeiitic basalt at mantle or lower crustal depths (Davis and Edwards, 1985; Beakhouse and McNutt, 1991). The extent to which a component derived from direct melting of The outcrop at this location consists largely of medium-grained, massive to weakly foliated tonalite of the Dore Lake lobe of the Atikwa Batholith. The tonalite contains numerous relatively small, partiallyassimilated mafic xenoliths. A prominent feature visible in the exposure on the eastern side of the highway is a 20cm wide mafic dike that strikes 165° and dips 70° to the west (Fig. 9). Such dikes are common near the eastern margin of the Dore lobe and may be indicative of the cooled carapace to an active magma chamber (Beakhouse et al., 1995). The eastern exposure also exhibits a prominent set of joints that strike approximately 135° and dip steeply (approximately 70°) toward the southwest. Stop 11: Van Horne Gold Property – Pritchard exposure trench UTM Coordinates: NAD83; 15U 0504890N / 5507870N Travel 17.2km north from Stop 10 along Highway 502 to the intersection with Flambeau Road. Note that this field trip stop is accessed via a private, gated road, Figure 9. Mafic dike cross-cutting tonalite of the Atikwa batholith. - 34 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 and that permission must be obtained from the property owner. The Pritchard trench is located approximately 850m southeast of the intersection with Highway 502. A summary of the exploration history and geology of the Van Horne gold property is provided below. (Fig. 10). The metavolcanic rocks occur together with synvolcanic diorite to quartz diorite intrusions and contain discontinuous horizons of interflow sedimentary rocks. Mafic and felsic dikes cut all of these rock types. The mafic dikes are mainly gabbro to diorite, while the felsic dikes include both feldspar porphyry and quartz-feldspar porphyry. All of these rock types have been cut by quartz veins. The youngest rock type on the property is diabase that occurs as a prominent west-trending dike that appears to have been emplaced after the formation of the quartz veins (Lengyel and Rennie, 2009). Exploration History Laurentian Goldfields Ltd. initiated an exploration program in 2008 on the Van Horne Gold Property, approximately 8km southwest of Dryden. There are16 historical shafts, 53 test pits, and 73 trenches located on the property. Some have underground workings and gold production has occurred at 4 of these excavations. There are approximately 24 historical gold occurrences on the property (Lengyel, 2007). The majority of this historical work occurred during a period of intense exploration and development activity that occurred between 1897 and the 1940’s. Structural Geology Rock Types All of the rock types on the property exhibit variable strain. Shearing and veining observed on the outcropscale in exploration trenches provide the basis for property-scale structural interpretations. With the exception of minor northeast-trending veins (D1 axial planar), most shear structures and veins lie along four dominant orientations associated with D2 dextral transpression. These structures are listed below: The Van Horne Property is situated in a mixed sequence of mafic to felsic metavolcanic rocks grouped together as the Lower Wabigoon volcanics • East-trending (90°) structures parallel to D2 regional Figure 10. Regional stratigraphy and location of Laurentian Goldfields Ltd.’s Van Horne gold property (from Lengyel and Rennie, 2009). - 35 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Disseminated euhedral magnetite occurs throughout the property in most rock types. However, greater magnetite concentrations occur on the western part of the property. Rennie et al. (2012) have proposed that the “magnetic anomaly underlying the Flambeau zone is associated with hydrothermal magnetite disseminated through the diorite and quartz-diorite dykes.” foliation; • West-northwest-trending (280° to 290°) D2 Riedel (R) shears; • North-northwest-trending (325° to 345°) D2 Riedel (R’) shears; and • Northwest-trending (300° to 310°) tension features. Most historical exploration and development on the property has targeted the gold potential of quartz veins. These quartz veins are structurally controlled and hosted in a variety of rock types (Lengyel and Rennie, 2009). Based on the results of a channel sampling program completed on the Van Horne Property exposures, Rennie and Chiang (2012) were able to identify the quartz vein systems that host gold mineralization and described them as follows: Sulphides “The vein array consists of west-trending shear veins and northwest-trending tension veins as well as a lesser array of sub-horizontal ladder veinlets. Shear veins tend to show dramatic pinch-and-swell but are locally continuous over up to about tens metres. Tension veins locally form left-stepping, echelon patterns and the individual veins tend to be narrow (up to 20 cm) with openly anastomosing planar forms. The shear veins and tension veins intersect without cross-cutting relation, suggesting that they are coeval and are interpreted to have formed during a single progressive strain event.” Sulphide minerals are found in all rock types, including the quartz veins. Samples collected during a second phase of sampling tested the sulphide-bearing rocks adjacent to gold-bearing quartz veins that were identified during the initial sampling. Some of the channel-cut samples collected from this sampling program returned anomalous gold values (Rennie and Chiang, 2012). Exposures – Trenching Alteration In 2009, mechanical removal of overburden and pressure washing was conducted at seven areas on the property (Fig. 11). The objective of the trenching was to provide a north-south cut across stratigraphy in areas of high outcrop density. Some of these areas are adjacent to historical excavations, whereas others were testing mineral potential sites identified from Mobile Metal IonTM (MMI) survey results and responses from a helicopter-supported airborne magnetometer geophysical survey. These exposures were geologically mapped and the mineral potential of selected parts of outcrops was tested by two phases of channel cutting and sampling. All rock types on the property exhibit variable hydrothermal alteration. The intensity of alteration in the exposures often increases where quartz veining is more abundant. Some parts of the exposures that lack quartz veins also display hydrothermal alteration (Lengyel, 2007). Several alteration features were identified during examination of the boundary between the differentiated phases within the mafic intrusive rock by Rennie et al. (2012). “... in many instances the primary mineralogy and textures in the rock are completely replaced by alteration minerals. Multiple phases of alteration are identified in drill core; these include an early phase of sericite and chlorite replacement, overprinted by carbonate and finally pervasively silicified. The main alteration minerals are carbonate, ankerite, chlorite, magnetite and silica with lesser amounts of albite and tourmaline.” Gold Mineralization Based on the assay results of samples collected by Laurentian Goldfields Ltd. since the initiation of activity on the property, Rennie and Chiang (2012) and Rennie et al. (2012) have proposed the following: • Both shear veins and tension veins contain anomalous concentrations of gold. However, the gold concentrations are highly variable; • Gold occurs in both barren and pyrite-bearing quartz veins; • Altered wall rocks with secondary pyrite locally contain high concentrations (up to 8g/t) of gold, but gold values are highly localized; • Lithological contacts and mechanically competent rock units are important host rocks for mineralization; - 36 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 11. Location of stripped exposures and historical excavation on part of the Laurentian Goldfields Ltd. Van Horne gold property (modified from Lengyel and Rennie, 2009). • The gold-mineralized zone is enriched in pathfinder elements, such as arsenic, silver and particularly tungsten; and • Hydrothermal magnetite is disseminated throughout the diorite and quartz diorite dikes. Future exploration efforts on the property could identify additional quartz-carbonate vein-hosted gold mineralization, and the sulphide-bearing rocks adjacent to known gold-bearing quartz veins may host additional gold-mineralized zones. Quartz veins and hydrothermal alteration zones localized along lithological boundaries between differentiated phases within the interiors of mafic intrusions should also be examined for gold mineralization. Pritchard Exposure Trench The Pritchard Exposure trench (Fig. 12) will be visited during this field trip. The following description of this gold occurrence is excerpted from Lengyel and Rennie (2009): “The Pritchard trench is underlain by felsic tuff and tuff-breccia, massive rhyolite, mafic intrusive, and one outcrop of [an] altered and brecciated mafic unit of unknown origin at the north end. Both mafic units and the felsic tuff adjacent to them at the south end of the trench contain strong pervasive ankeritization. Strong Figure 12. Pritchard trench geology map showing areas where anomalous gold assays have been obtained (modified from Lengyel and Rennie 2009) - 37 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 pervasive silica is ubiquitous in the trench. Sericitization is moderate to strong in the felsic units, and is also strong in the altered northern mafic outcrop. Beakhouse, G.P. and McNutt, R.H., 1991. Contrasting types of Late Archean plutonic rocks in northwestern Ontario: Implications for crustal evolution in the Superior Province. Precambrian Research, 49, p.141165. Structures observed in the trench were limited to faulting and strong fracturing in the felsic units. A north- to northeast-trending fault (D1 axial planar) in the felsic volcanics immediately south of the access road is defined by a 1-2m deep cleft in the outcrop. The fault extends across the entire unit, and associated tensional faults and quartz veins define a sinistral sense of movement. The rhyolite unit is more strongly faulted, with a 0.5m east-trending (D2) wide fault zone at its northern end and a 5m wide east-trending D2 fault zone at its southern end. A minor northwesttrending normal fault was also observed at the southern end of the rhyolite. Beakhouse, G.P., Blackburn, C.E., Breaks, F.W., Ayer, J., Stone, D., and Stott, G.M. 1995. Precambrian ’95 western Superior Province field trip guidebook; Ontario Geological Survey, Open File Report 5924, 94p. Blackburn, C.E. 1980. Towards a mobilistic tectonic model for part of the Archean of northwestern Ontario; Geoscience Canada, v.7, p.64-72. Blackburn, C.E. 1981. Geology of the Boyer Lake-Meggisi Lake area, District of Kenora; Ontario Geological Survey Report 202, 107 p. Accompanied by Maps 2437 and 2438, scale 1:31,680. Blackburn, C.E. 1982. Geology of the Manitou Lakes area, District of Kenora, stratigraphy and petrochemistry; Ontario Geological Survey, Report 223, 61p. Accompanied by Map 2476, scale 1:50,000. Despite the strongly-fractured, brittle nature of the felsic units, the bulk of the veining observed in the trench exists within the mafic units. The brecciated and strongly altered northern mafic outcrop contains four zones of <1cm wide tensional quartz-ankerite±tourmaline veins over 4m. The larger mafic intrusive unit at the south end of the trench is cut by numerous east-trending (D2) and northwest-trending (tensional), 2-26cm wide quartz-ankerite-tourmaline veins. The east-trending veins typically cut and offset the northwest-trending veins. Veining within the rhyolite is mainly associated with a zone of faulting at its south end, although the easttrending quartz-ankerite-tourmaline-muscovite vein is observed cross-cutting the faulting and local mafic dykes. Blackburn, C.E., Johns, G.W., Ayer, J., and Davis, D.W., 1991. Wabigoon Sub-province; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.303-381. Davis, D.W. 1990. The Seine-Couchiching problem reconsidered: U-Pb geochronological data concerning the source and timing of Archean sedimentation in the Western Superior Province; in Proceedings, 36th annual Meeting, Institute on Lake Superior Geology, pt.1, Abstracts, p.19-21. Davis, D.W. and Edwards, G.R. 1985. The petrogenesis and metallogenesis of the AtikwaLawrence volcanicplutonic terrain; in Geoscience Research Grant Program, Summary of Research 1984-85, Ontario Geological Survey, Miscellaneous Paper 127, p.101111. Davis, D.W., Blackburn, C.E., and Krogh, T.E. 1982. Zircon U-Pb ages from the WabigoonManitou Lakes region, Wabigoon Sub-province, northwest Ontario. Canadian Journal of Earth Sciences, 19, p.254-266. Significant assays (>1g/t Au) from the Pritchard trench include grab samples of 7.12g/t Au, 3.58g/t Au, 2.46g/t Au, 2.18g/t Au, and 1.55g/t Au. Significant channel samples include 2.21g/t Au over 0.6 m and 1.52g/t Au over 0.6m. Davis, D.W., Sutcliffe, R.H., and Trowell, N.F., 1988. Geochronological constraints on the tectonic evolution of a Late Archean greenstone belt, Wabigoon Sub-province, northwest Ontario, Canada. Precambrian Research, 39, p.171-191. References Beakhouse, G.P. 2000. Precambrian geology of the Wabigoon area; in Summary of Field Work and Other Activities 2000, Ontario Geological Survey, Open File Report 6032, p.20-1 to 20-8. Beakhouse, G.P. 2011. Western Wabigoon Sub-province Synthesis Project; in Summary of Field Work and Other Activities 2011, Ontario Geological Survey, Open File Report 6270, p.9-1 to 9-8. Lengyel, P. 2007. Compilation report on the Van Horne Area of Interest gold property; Kenora District Geologist’s office, assessment file 52E10NW BBB-1, City of Dryden. Lengyel, P., and Rennie, C. 2009. Assessment report on the Van Horne gold property; AFRO number 2.46195, Kenora District Geologist’s office, assessment file 52E10NW DDD-7, Laurentian Goldfields Limited. Lichtblau, A.F., Ravnaas, C., Storey, C.C., Tims, A., Debicki, R.L., Pettigrew, T.K., Wilson, A.C., and Wetendorf, J. - 38 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 2015. Report of Activities 2014, Resident Geologist Program, Red Lake Regional Resident Geologist Report: Red Lake and Kenora Districts; Ontario Geological Survey, Open File Report 6301, 83p. Maunula, T. and Wilson, J. 2010. Manitou Gold Inc., Kenwest property, national instrument 43-101compliant technical report, 55p. http://www.manitougold.com/_ resources/kenwest/43-101_KW.pdf Ontario Geological Survey 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release Data 126—Revision 1. Parker, J.R. 1989. Geology, gold mineralization and property visits in the area investigated by the Dryden-Ignace Economic Geologist, 1984-1987; Ontario Geological Survey, Open File Report 5723, 306p. Parker, J.R., Blackburn, C.E. and Davis, D.W. 1989. Constraints on timing and placement of gold mineralization in the Wabigoon Sub-province near Dryden, Ontario: evidence for synvolcanic through late tectonic emplacement; Program with Abstracts, Geological Association of Canada-Mineralogical Association of Canada, Annual Meeting, v.l4, p.A92. Rennie, C. and Chiang, M. 2012. Assessment report on the Van Horne gold property; AFRO number 2.52196, Kenora District Geologist’s office, assessment file 52E10NW DDD-9, Laurentian Goldfields Limited. Rennie, C., Chiang, M., and Meade, S. 2012. Assessment report on the Van Horne gold property; AFRO number 2.52620, Kenora District Geologist’s office, assessment file 52E10NW DDD-8, Laurentian Goldfields Limited. Stern, R.A., Hanson, G.N. and Shirey, S.B., 1989. Petrogenesis of mantle-derived, LILEenriched Archean monzodiorites and trachyandesites (sanukitoids) in southwestern Superior Province. Canadian Journal of Earth Sciences, 26, p.1688-1712. Stone, D., Davis, D.W., Hamilton, M.A., and Falcon, A. 2010. Interpretation of 2009 geochronology in the central Wabigoon Sub-province and Bending Lake areas, northwestern Ontario; Ontario Geological Survey, Open File Report 6260, p.14-1 to 14-13. Stone, D., Hellebrandt, B., and Lange, M. 2011. Precambrian geology of the Bending Lake area (south sheet); Ontario Geological Survey, Preliminary Map P.3624, scale 1:20 000. Sutcliffe, R.H., Smith, A.R., Doherty, W. and Barnett, R.L., 1990. Mantle derivation of Archean amphibolebearing granitoid and associated mafic rocks: evidence from the southern Superior Province, Canada. Contributions to Mineralogy and Petrology, 105, p.255-274. Teal, P.R. and Walker, R.G. 1977. Stratigraphy and sedimentology of the Archean Manitou group, northwestern Ontario; in Report of Activities, Part A, Geological Survey of Canada, Paper 77-1A, p.181184. - 39 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 4 - Thunder Lake (Goliath) Project Treasury Metals Incorporated Staff Wabigoon, Ontario Please Note: All information presented within this document was taken from the website of Treasury Metals Incorporated and all data is within the public domain. The project site is not open to the public and special permission must be obtained from Treasury Metals before access of any kind is granted. speculative geologically to have the economic considerations applied to them that would enable them to be categorized as mineral reserves, and there is no certainty that the preliminary economic assessment will be realized. Mineral resources that are not mineral reserves do not have demonstrated economic viability. Project Highlights/Introduction Location Treasury Metals Incorporated is a Canadian gold exploration and development company focused on its 100% owned high-grade Goliath Gold Project (Thunder Lake), which currently consists of an Indicated and Inferred resource of 1.7 Moz. The Project, which is near Dryden, Ontario and is slated for near-term Canadian gold production. Treasury Metals Inc.’s 50-km2 Goliath Property land package (Fig. 1) lies adjacent to the Trans-Canada Highway 17 and is located midway between Thunder Bay and Winnipeg. The closest city is Dryden, Ontario located 15km to the west. The community of Wabigoon is located six kilometres to the east. Treasury Metals Inc.’s Goliath Gold Project offices are located in former Ontario government industrial offices, workshops, and storage facilities within the community of Wabigoon, Ontario. Treasury Metals is advancing through the Canadian permitting process to begin production at its open-pit gold mine and 2,500tpd processing facility in the near future. Subsequent underground operations will be developed in the latter years of mine life and will be funded from the project’s initial cash flow. Beneficial features of the Project location are: • Year-round fieldwork can be comfortable working conditions • Excellent year-round road access • On-site power supply • Trans-Canada Natural Gas Pipeline • Neighbouring major industrial services in Dryden Low initial start-up CAPEX of C$92M with cashflows from initial 3 years of open-pit production funding underground development; • Nearby workforce in surrounding communities • In August 2012, Treasury Metals completed an updated preliminary economic assessment (PEA) on the project. Key highlights include: • After-tax NPV 5% of $144M and 32.4% IRR The project demonstrates numerous highly prospective targets that show potential to host gold mineralization including about six kilometres of prospective strike along trend from the Goliath deposit. • Producing 80,000 -- 100,000 oz annually over 10+ years of mine life • Payback Period of 2.8 years • Low initial start-up CAPEX of $92M • LOM average feed grade of 2.87 g/t Au and 9.30 g/t Ag 100%-owned high-grade, demonstrating indicated and inferred resource of 1.7 Moz; • Results of 2013 drilling program have defined high-grade near-surface intersections, indicating significant upside potential for both resource and project economics; • • an The PEA is preliminary in nature and includes inferred mineral resources that are considered too completed in Project Geology and Mineralization The Goliath Project is within the Wabigoon Subprovince of the Archean Superior Province, north of the Wabigoon Fault (Fig. 2). An amphibolite metamorphic grade volcanogenic-sedimentary complex topped by an upper layer of greenschist characterizes much of the project area. The assemblage itself comprises quartzporphyritic felsic to intermediate metavolcanic rocks represented by biotite gneiss, mica schist, quartzporphyritic mica schist, a variety of metasedimentary rocks, and minor amphibolites. Compositional layering - 40 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 1: Location of the Goliath Gold Project area. in metasedimentary rocks strikes ~70 to 90° and dips from 70 to 80· south-southeast. The Thunder River Mafic Metavolcanics underlie the southern part of the property. All rocks have been subjected to folding and moderate to intense shearing with local hydrothermal alteration, quartz veining and sulphide mineralization. muscovite schist, and metasedimentary rocks), with the footwall comprising metasedimentary rocks with minor porphyries, felsic gneiss, and schist. Gold within the central unit is concentrated in a pyritic alteration zone, consisting of quartz-sericite schist (MSS), quartz-eye gneiss, and quartz-feldspar gneiss. The main zones of mineralization within the Goliath Deposit consist of the Main Zone, Footwall Zone (B, C, and D subzones), and Hangingwall Zone (H and H1 subzones). Mineralized zones strike approximately east-west and dip 70° to 80° to the south-southeast. High-grade gold mineralization (>3g/t) is concentrated in several steeply dipping, west-plunging shoots with strike lengths (>50m) and significant down-plunge continuity. The high-grade shoots are interpreted to be the result of the intersection between tight F1 isoclinal folding of the mineralized horizon and regional F2 folding with gold often concentrated in fold noses. These high-grade “shoots”, which appear to occur in regular intervals, are separated by lower grade Au/Ag mineralized rocks. Goliath Mineralization The mineralized zones are tabular composite units characterized by anomalous to strongly elevated gold concentrations, increased lead and zinc sulphide content, and distinctive altered rock units. Stratigraphically, gold mineralization is contained in an approximately 100 to 150m wide central zone composed of intensely altered felsic metavolcanic rocks (quartz-sericite and biotite-muscovite schist) with minor metasedimentary rocks. Overlying hanging wall rocks consist of altered felsic metavolcanic rocks (sericite schist, biotite- Exploration Focus Treasury Metals has completed more than 100,000m of diamond drilling since 2008 (in addition to historic drilling by Teck Resources). The Company’s current exploration and drilling program has focused on targets located in the northeast and east of the Goliath - 41 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 2: Geology Map of the Goliath Gold Project. Gold Deposit within the >49 km2 property block. Significant gold values intercepted in previous drilling campaigns, as well as re-interpreted airborne EM and aeromagnetic geophysical surveys, are being used to guide the current drilling program. Diamond drilling exploration was completed during 2012 and 2013 to test additional targets projected over 11 km of potential ore zone strike extension. Recent exploration focus has been principally to the northeast and east of the Goliath Gold deposit, in the Project’s land package. Significant gold values intercepted in previous drilling campaigns, as well as re-interpreted airborne EM and magnetometer geophysical surveys were used to guide the drilling program. The program focused on pursuing strike extensions of previously identified mineralization as well as following potential new ore shoots down dip within the currently defined resource area. Most recently the program concentrated on delineating the C Zone mineralization, now largely in the Inferred category, both within and to the east of the proposed open pit boundary. The C Zone in-fill program focused primarily on defining near surface resources, converting inferred ounces to indicated ounces, and defining a recently discovered high-grade C zone ore shoot. A previously sparsely drilled area was delineated measuring approximately 1.2km in strike-length and infill drilling of select targets was also completed. The C Zone has the potential to represent an increase in the current open pit mineable resource and a reduction of the overall waste to ore stripping ratios. With the completion of the first phase of the C Zone in-fill drilling program, the company completed a gap analysis in 2014 to determine the location of holes required to upgrade all ounces within the current open pit design from inferred to the indicated category including both the Main and C Zones. The scope and size of the in-fill and expansion drill program was determined from the gap analysis. - 42 - Refer to Figure 3 below for the location of the �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 3: Goliath Deposit mineralized zones and 2103 diamond drill collars. Project History various Goliath mineralized zones and the locations of drill holes completed in 2013. Resource Estimate The Goliath Gold Project is an amalgamation of two historic properties: the Thunder Lake Property purchased from Teck Resources Limited and Corona Gold Corporation, and the Laramide Property, transferred to Treasury from Laramide Resources Ltd. upon the Company’s spin-out in 2008. The detail of the most recent resource estimate (August 2012) for the Project is presented in Table 1, below. At present the deposit consists of an NI-43-101compliant indicated open pit (OP) and underground (UG) resource of 9.14 million tonnes (Mt) grading 2.6 grams per tonne (g/t) Au and 10.4g/t Ag for 760,000 contained gold ounces and 3,070,000 contained silver ounces. Also included within the estimate is an inferred open pit and underground resource of 15.9Mt grading 1.7g/t Au and 3.9g/t Ag for 870,000 contained gold ounces and 1,990,000 contained silver ounces (AuEq refers to Au equivalent grade; i.e. the value of all payable metals calculated as Au grade). Treasury Metals has since brought the Project to a level sufficient to begin the company’s Feasibility Study. Successfully completed Goliath Gold Project programs include: • More than 160,000 metres of diamond drilling. • Two NI43-101 compliant Resource Estimates since 2008. • Full Environmental Baseline Study Program completed. • Robust PEA released in August 2012. - 43 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 • Figure 4: Aerial view of the conceptualized open pit and the main mineralized zones. • Feasibility level Metallurgical testing. The majority of the pre-2008 exploration work was completed by Teck Resources beginning in 1990 with Teck’s first discovery hole and continuing through to a 250m underground drift and 2,375 tonne bulk sample taken in 1998 when the project was put on hold before being acquired by Laramide in 2007. There is only limited documentation of the prospecting and early exploration activity conducted on the Project properties prior to 1989. Material exploration activity on the property completed by Teck, which ultimately defined the Thunder Lake mineralization now known as the Goliath deposit, began in 1989 after reconnaissance work. The discovery hole on the Main Zone of the Goliath Deposit was drilled in 1990 and intersected multiple horizons of gold mineralization. In 1994, Teck and Corona entered into an option agreement for the development of their property, pursuant to which they formed a Joint Venture partnership in 1996 and drilled together until 1998 at which time Teck collected the bulk sample. Table 1: Treasury Metals 2012 Resource Estimation for the Goliath Gold Deposit. Resource Category Cut-off Grade Average Tonnes (g/t Au) Grade (g/t Au) Contained Au (oz) Average Ag Grade (g/t) Contained Ag (oz) Ag Total AuEq AuEq oz oz (Au+Ag) Indicated Resource Estimate Surface 0.3 6,002,000 1.8 326,000 7.1 1,257,000 22,000 348,000 Underground 1.5 3,136,000 4.3 433,000 18.0 1,812,000 32,000 465,000 9,140,000 2.6 760,000 10.4 3,070,000 54,000 810,000 Total Inferred Resource Estimate Surface 0.3 11,093,000 1.0 352,000 3.3 1,184,000 21,000 374,000 Underground 1.5 4,789,000 3.3 514,000 5.2 807,000 14,000 528,000 Total 15,900,000 1.7 870,000 3.9 1,990,000 35,000 900,000   - 44 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 By 1999, Teck and Corona’s exploration led to the outlining of the Goliath deposit and the reporting of a historic (non-NI43-101-compliant) inferred mineral resource estimate of 2.974 million tonnes at 6.47g/t Au, using a 3.0g/t Au cut-off and a minimum thickness of 3.0 metres. the company’s latest metallurgy results, is scheduled for 2015, a major milestone in the mine permitting process. Present metallurgy by G&T Metallurgical Services demonstrates 95+% gold recovery, 60-70% of which is recoverable by gravity, with 8-12 hours leach time, and a medium hardness rock. Strategic land continued to be held by Laramide, an exploration and development company involved in advancing U.S. and Australian uranium projects to development and permitting. The Laramide Property, historically referred to as the Goliath Gold Project, was located immediately south of the western portion of the Thunder Lake Property (Teck and Corona). Laramide completed preliminary exploration in 1994 and 1996 that included 1,622m of diamond drilling in eight holes. In 2008, Treasury Metals acquired all properties surrounding the deposit. New discoveries at targets outside the Goliath Gold Project’s current resource area could add significant resources to the Goliath Gold Project. Engineering activities are progressing to support advancement of this full feasibility study. Metallurgical Studies Treasury has completed several detailed metallurgical tests including a 2375 tonne bulk sample and a 400kg sample. The studies were completed by ALS Metallurgy (G&T Metallurgical Systems) of Kamloops to Feasibility level standards. The project consistently returns extremely high recoveries that are easily scalable with low leach times and low reagent consumptions. • 2375 tonne sample previously completed with 97% recovery. 70% recovery from gravity alone. • G&T Metallurgical Services obtained 95+% Au recovery, 60-70% recovery by gravity, 8-12 hours leach time, medium hardness rock. • Gekko Systems Australia are currently testing their Python Process to confirm amenability of Goliath gold ore to treatment using VSI and continuous gravity concentration CGR. This would significantly reduce treatment and permitting requirements. Environmental Baseline Study Program Baseline studies are completed to gain an understanding of the current natural environment of the site, support mine development decisions and management plans, and to provide support to rigorous on-going monitoring and mine closure plans. Treasury Metals is completing baseline studies to provide the necessary data to support the Goliath Gold Project. Treasury Metals began a complete program of baseline studies in 2010 in support of the Goliath Gold Project with an operational team of professionals based locally and nationally. Baseline studies have continued to present day supporting current physical, biological, and socio-economic decisions. Baseline study results will be provided and reported as part of the federal and provincial permitting regulations. Permitting and Mine Development The permitting process has been under way since 2012 with the acceptance of Treasury Metals Inc.’s Goliath Project Description by the Canadian Environmental Assessment Agency (CEAA). The CEAA subsequently issued its Environmental Impact Statement (EIS) guidelines to the company in February 2013. A full bankable feasibility study (BFS), including - 45 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 5 - Classic Outcrops of the Dryden area Peter Hinz Ring of Fire Secretariat, Ontario Ministry of Northern Development and Mines, Thunder Bay, Ontario, Canada Foreword This trip will examine three stops of an eight-stop field trip which was developed for the Dryden High School Geography Department by the author in 2000. The three stops (Fig. 1) represent the most photoworthy of the eight stops and display some of the more significant depositional, structural, and metamorphic features in the Dryden Area. The field trip area lies within the Western Wabigoon Terrane formerly known as the Wabigoon Sub-province. The geological discussions related to tectonic setting, lithologies, and structure has been compiled from work completed by Beakhouse between 1999 and 2010. The field trip stop discussions include observations by the author and interpretations of host lithologies noted by Beakhouse. The photograph below of Dryden area pillowed mafic basalts is from Satterly (1943). with minimal superimposed strain. This area has a complex deformational history with an early, generally bedding-parallel fabric (D1) deformed into a series of megascopic to regional-scale, southwest-plunging, Z-asymmetric folds with the development of a second fabric (D2) parallel to the axial surface of these folds. Metamorphic grade varies regionally from upper greenschist to upper amphibolite with the lowest grade generally occurring nearest to the Wabigoon fault. The Atikwa domain occurs to the south of the Wabigoon fault and is characterized by dominantly volcanic sequences that face away from large, coeval batholiths (e.g., Atikwa Batholith, Aulneau Batholith). Within the map area, the Wabigoon Metavolcanics are typical of these sequences with a thick basal portion consisting almost entirely of mafic metavolcanic rocks overlain by a more heterolithic portion, which, although still dominantly mafic metavolcanic, includes minor intermediate to felsic metavolcanic rocks and rare metasedimentary rocks. Mineral assemblages indicative of regional middle to upper greenschist facies predominate with narrow amphibolitic contact metamorphic aureoles adjacent to some of the plutons. Lithologies (Beakhouse, 2000 & 2002) Wabigoon Metavolcanic rocks Introduction and (Beakhouse, 2005) Tectonic Setting The Wabigoon-Dinorwic area is transected by the Wabigoon fault, which is a major regional structure that separates two geologically distinct domains within the Wabigoon Sub-province. Distinct mineral deposit types and styles also characterize these domains. The Sioux Lookout domain, lying to the north of the Wabigoon fault, is characterized by a series of alternating sedimentary-dominated and volcanicdominated panels that consistently face to the south. Many of these panels are regionally interpreted to have fault contacts; however, some of the contacts appear to be conformable depositional contacts The Dinorwic area occurs within the Atikwa domain and, with the exception of several small, late intrusions, is underlain by the Wabigoon Metavolcanics. Mafic metavolcanic rocks dominate the Wabigoon Metavolcanics in the Dinorwic area. Massive and pillowed flows are approximately subequally abundant with minor, widely distributed flow breccia and hyaloclastite. Massive flows range from fineto medium-grained. Many of the pillowed flows and some of the massive flows are moderately vesicular. Equigranular flows are most abundant with conspicuous moderate to coarse plagioclase porphyritic flows occurring locally. Massive flows include both magnetite-poor (magnetic susceptibility ~0.5 to 1.0) and magnetite-rich (magnetic susceptibility commonly greater than 50) varieties, whereas the pillowed flows consistently have magnetic susceptibilities comparable - 46 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 1. Field trip stop locations (from Blackburn et al., 1981). to the magnetite-poor massive flows. The colour of both weathered and fresh surfaces is highly varied due to a range in intensity of a variety of types of alteration including carbonatization (both calcitic and ferroan carbonate), silicification, and epidotization. Thunder Lake Sediments The Thunder Lake Sediments include two separate panels of rock separated along a portion of their strike-length by the Thunder Lake Volcanics. Thinto medium-bedded wacke-siltstone characterized by even, continuous bedding is the predominant component in both panels. Thin magnetite ironstone layers are a conspicuous minor component within the Thunder Lake sediments north of the Thunder Lake Volcanics but are rare within the southern panel. Minor, thin garnet-rich (<70% garnet) and calc-silicate layers may represent original more pelitic and marly layers, respectively. In one location, the calc-silicate material forms discordant veins, and suggests that some of this material is remobilized or originated by secondary alteration processes. The garnet-rich layers are often closely spatially associated with the ironstone layers. A limited number of determinations indicate that tops, although locally reversed by tight to isoclinal minor folding, are generally to the south in both the north and south panels. Contact relations with the Brownridge Volcanics have not been observed but the data are permissive of a conformable stratigraphic relationship. The contacts with the Thunder Lake Volcanics appear to also be conformable although these are commonly moderately- to highly-strained and a loci of abundant quartz veining. Structural Geology (Beakhouse, 2000) The areas north and south of the Wabigoon fault have markedly different deformational styles. South of the Wabigoon fault, the dominantly volcanic sequence is a northward-younging homocline, but penetrative fabrics are sporadically and weakly developed. Primary form of pillows and other primary structures is well-preserved in horizontal sections, although there is currently little constraint on the extent of possible vertical stretching. - 47 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 The area north of the Wabigoon fault is characterized by alternating volcanic- or sedimentary-dominated units that are generally southward-facing and has heterogeneous, weakly- to strongly-developed penetrative fabrics. An earlier (D1), ubiquitous, weaklyto strongly-developed fabric is parallel to bedding with the exception of one outcrop where it is interpreted to be parallel to the axial surface of rootless, tight to isoclinal folds. The D1 fabric development heterogeneity is, in part, controlled by primary lithological characteristics (e.g., fabric in mafic volcanic sequences is better developed in interflow hyaloclastite/breccia units than massive, fine- to medium-grained units). The D1 fabric, along with bedding, is deformed on a variety of scales into southwesterly plunging (D2) folds having Z-asymmetry. These regional-scale folds, as well as the contrasting characteristics to the south of the Wabigoon fault, are readily apparent on regional aeromagnetic maps. On a detailed scale, most outcrops having suitable markers or well developed D1 fabrics display Z-asymmetric minor folds that commonly have a D2 fabric (spaced cleavage to pervasive penetrative schistosity) developed parallel to the axial surface of the minor folds. Alteration (Beakhouse, 2001) Noteworthy abundances of garnet, that may be indicative of alteration, occur in two settings. In the Thunder Lake Sediments, extremely garnet-rich (6080% garnet) rocks form layers that are commonly, though not exclusively, closely associated with magnetite layers and calc-silicate mineral assemblages. It is not clear if these layers are indicative of hydrothermal alteration or whether they may represent isochemical metamorphism of unusually aluminous sedimentary rocks. Field Trip Stops Figure 2. Pillows with younging direction to the upper right. pillowed mafic metavolcanic rocks which are part of the Boyer Lake group. U-Pb age determinations of 2719±3Ma and 2722 ±5Ma are reported for the Boyer Lake group by Davis (1990). The pillows display distinct thick selvages and provide excellent younging direction (Fig. 2). Small, well-developed vesicles are observed which suggest a high confining pressure and emplacement in deep water. Small sections of pillow breccia are also seen amongst the well developed pillows. Further north along the outcrops sections of interflow sediments may also be observed, suggesting a hiatus in eruptive activity. Moving northward a 2 to 3m thick unit of interflow sedimentary rocks is observed, suggesting a hiatus in extrusive activity. The sediments display bedding as well as cross-bedding. Overlying the interflow sedimentary rocks is a massive flow unit, the lower portion of which does not contain vesicles. Oriented vesicles appear in the upper portion of the flow unit (Fig. 3). From the MTO rest stop turn right and travel for 23.8km to Elm Bay Road, turn left and after 75m turn left again. Continue for approximately 350m and turn Depart Dryden from the traffic lights by McDonalds Restaurant on Trans-Canada Highway 17. Travel 39.7km east to the Snake Bay Road. Turn around and return 0.8km to the west. Park the vehicles at the MTO rest stop. Walk approximately 200m to the north to the low outcrops on the east side of the highway. Stop 1: Pillowed mafic metavolcanic and interflow sedimentary rocks of the Wabigoon Volcanics. UTM Coordinates: NAD83; 15U 0540211E / 5496631N The southern portion of this outcrop contains Figure 3. Oriented vesicles within the massive flow unit. - 48 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 have been sheared through, producing isolated fold noses. The form of these folds is a tight Z, in which axial planes are parallel to the local east-southeast formational strike. Differentially weathered beds in the metapelites and wackes serve to outline these folds, portions of some having been completely detached by shearing. A second phase of open folding refolds the first phase.” left. Drive an additional 100m and park. Walk 45m eastward along a trail to the large smooth outcrop. Stop 2 – Isoclinal folding in Thunder Lake Sediments UTM Coordinates: NAD83; 15U 0522545E / 5512553N The description below (Beakhouse et al., 1995) is for an outcrop located approximately 300m to the east along-strike and on Highway 17. The outcrop we are visiting was identified subsequent to Beakhouse et al. (1995) and was exposed as part of an aggregate operation. “Road-cut outcrops on either side of Highway 17 show style of folding and order of superposition of folding in this portion of Warclub group metasedimentary rocks. On the south side of the highway, on a smooth sloping glaciated surface, two phases of folding can be seen. Transposition of bedding into the plane of schistosity is seen where the limbs of first phase isoclinal folds The outcrop displays distinctive z-fold asymmetrical folding as well as parasitic folds. Detailed descriptions of the structural components that are visible at this stop are provided above in the Structural Geology section. Return to Highway 17, turn left and continue for approximately 10km, turn left into the Walmart parking lot and proceed to the low outcrop located on the west side of the building. Stop 3 – Folding and metamorphism/alteration of Thunder Lake Sediments UTM Coordinates: NAD83; 15U 0513496E / 5514716N At this stop pelitic sedimentary rocks and iron formation of the Thunder Lake sediments display a similar deformation history as observed at Stop 2. Thin beds of iron formation display striking Z-asymmetrical folding as well as parasitic folds on fold noses and limbs across the outcrop (Fig. 4). Garnet-rich beds are also observed throughout the outcrop. As noted in the discussion it is unclear whether the garnets are a result of hydrothermal alteration or isochemical metamorphism (Fig. 5). For detailed descriptions of the structural and alteration features at this stop, refer to the Structural Geology and Alteration sections above. Figure 4. Folding in iron formation and pelitic sedimentary rocks of the Thunder Lake sediments. - 49 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 References Blackburn, C.E., Beard, R.C., and Rivett, A.S. 1981, KenoraFort Frances, geological compilation series, Kenora and Rainy River Districts, Ontario Geological Survey, Geological Map 2443, scale 1:253,440. Beakhouse, G.P. 2000, Precambrian geology of the Wabigoon area; in Summary of Field Work and Other Activities 2000, Ontario Geological Survey, Open File Report 6032, p.20-1 to 20-8. Beakhouse, G.P. 2001. Precambrian geology of the Thunder Lake Segment, Wabigoon Area; in Summary of Field Work and Other Activities 2001, Ontario Geological Survey, Open File Report 6070, p.15-1 to 15-6. Beakhouse, G.P. 2002. Precambrian Geology of the Dinorwic Area, Wabigoon Subprovince; in Summary of Field Work and Other Activities 2002, Ontario Geological Survey, Open File Report 6100, p.10-1 to 10-6. Beakhouse, G.P. 2005. Precambrian Geology of the Dinorwic-Butler Lakes Area, Wabigoon Subprovince; in Summary of Field Work and Other Activities 2005, Ontario Geological Survey, Open File Report 6172, p.9-1 to 9-6. Beakhouse, G.P., Blackburn, C.E., Breaks, F.W., Ayer, J., Stone, D., and Stott, G.M. 1995. Precambrian ’95 western Superior Province field trip guidebook; Ontario Geological Survey, Open File Report 5924, 94p. Figure 5. Abundant garnet in pelitic sediments. Davis, D.W. 1990. The Seine-Coutchiching problem reconsidered: U-Pb geochronological data concerning the source and timing of Archean sedimentation in the western Superior Province; Institute on Lake Superior Geology, v.36, pt.1, p.19-21. Satterly, J. 1943. Geology of the Dryden-Wabigoon area, Kenora District; Ontario Department of Mines, Annual Report, 1941, v.50, pt.2, 67p. - 50 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 6 - Gold Occurrences of Van Horne Township, Van Horne Gold property - Flambeau explosures Steve Meade Ontario Geological Survey, Sudbury, Ontario, Canada Craig Ravnaas Resident Geologist Program, Ontario Geological Survey, Kenora, Ontario, Canada Location and Access The Van Horne Gold Property and the Flambeau Exposures are located 8km southwest of Dryden. The exposures are located on private lands and can be accessed from Highway 502 (Fig. 1). Exploration Activity Laurentian Goldfields Ltd. initiated an exploration program in 2008 on the Van Horne Gold Property (Hogg, 2009). There are 16 historical shafts, 53 test pits, and 73 trenches located on the property. Some have underground workings and gold production has occurred at 4 of these excavations. There are approximately 24 historical gold occurrences on the property (Lengyel, 2007). A majority of this historical work is the result of the intense exploration and development activity occurring in the area between 1897 and the 1940’s. Rock Types The Van Horne property is situated in a mixed sequence of mafic to felsic metavolcanics rocks grouped together as the Lower Wabigoon Volcanics (Fig. 2). These rocks have been intruded by synvolcanic diorite to quartz diorite intrusions. Discontinuous interflow sedimentary rocks occur throughout the property. Mafic and felsic dikes cut all of these rocks types. The mafic dikes are mainly gabbro to diorite. The felsic dikes include both feldspar porphyry and quartzfeldspar porphyry. All of these rock types have been cut by quartz veins. One of the youngest rock types on the property is the prominent west-trending diabase dike that appears to been emplaced after the formation Figure 1. Location and access to Van Horne Property and the Flambeau Exposures. - 51 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 2. Regional stratigraphy and location of the Van Horne gold property (from Lengyel and Rennie, 2009). of quartz veins (Lengyel and Rennie, 2009). Quartz Veins Strain A majority of the historical exploration efforts and development has targeted the gold potential of quartz veins on the property. These quartz veins are structurally controlled and hosted in a variety of rock types (Lengyel and Rennie, 2009). Based on the results from the Phase II channel sampling program completed on the Van Horne property exposures, Rennie and Chiang (2012, p.29) identified the quartz veins that hosted gold mineralization: All of the rock types on the property exhibit varying amounts of strain. Shearing and veining within the trenches are consistent with structural interpretations for the property. With the exception of minor northeasttrending veins (D1 axial planar), structures and veins lie along four dominant orientations associated with D2 dextral transpression. Orientation of structures and quartz veins: • east-trending (Az 090), parallel to D2 regional foliation; • west-northwest-trending (Az 280 to 290) D2 Riedel (R); • north-northwest-trending (Az 325 to 345) D2 Riedel (R’); and • northwest-trending (Az 300 to 310) tensional features. - 52 - “The vein array consists of west-trending shear veins and northwest-trending tension veins as well as a lesser array of sub-horizontal ladder veinlets. Shear veins tend to show dramatic pinch-and-swell but are locally continuous over up to about tens metres. Tension veins locally form left-stepping, echelon patterns and the individual veins tend to be narrow (up to 20cm) with openly anastomosing planar forms. The shear veins and tension veins intersect without cross-cutting relation, suggesting that they are coeval and are interpreted to have formed during a single progressive strain event.” �Proceedings of the 61st ILSG Annual Meeting - Part 2 Alteration All rock types on the property exhibit varying amounts of hydrothermal alteration. The intensity of alteration in the exposures often increases where quartz veining is more abundant. Some parts of the exposures which lack quartz veins also display hydrothermal alteration (Lengyel, 2007). Several alteration features were identified during examination of the boundary between the differentiated phases within the mafic intrusive rock by Rennie et al. (2012, p.31): “... in many instances the primary mineralogy and textures in the rock are completely replaced by alteration minerals. Multiple phases of alteration are identified in drill core; these include an early phase of sericite and chlorite replacement, overprinted by carbonate and finally pervasively silicified. The main alteration minerals are carbonate, ankerite, chlorite, magnetite, and silica with lesser amounts of albite and tourmaline.” higher magnetite content in the rocks underlying the western part of the property. Rennie et al. (2012, p.31) have proposed that the “magnetic anomaly underlying the Flambeau zone is interpreted to be associated with hydrothermal magnetite disseminated through the diorite and quartz-diorite dykes”. Sulphides Sulphides are found in all rock types, including the quartz veins. Samples collected during the Phase II sampling program tested the sulphide-bearing rocks adjacent to gold-bearing quartz veins identified from the Phase I program. Some of the channel-cut samples collected from this Phase II program returned anomalous gold values (Rennie and Chiang, 2012). Exposures - Trenching Disseminated euhedral magnetite occurs in a majority of the rock types on the property. There is a In 2009, mechanical removal of overburden and pressure washing was conducted at seven areas on the property (Fig. 3). The objective of the trenching was to provide a north-south cut across stratigraphy in areas of high outcrop density. Some of these areas are Figure 3. Location of stripped exposures and historical excavation on part of the Van Horne Gold Property (modified from Lengyel and Rennie, 2009). - 53 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 adjacent to historical excavations, whereas others were testing mineral potential sites identified from Mobile Metal IonTM (MMI) survey results and responses from a helicopter-supported airborne magnetometer geophysical survey. These exposures were geologically mapped and the mineral potential of selected parts of the exposed outcrops were tested by two phases of channel cutting and sampling. Gold Mineralization Based on the assay results of samples collected by Laurentian Goldfields Ltd. since the initiation of activity on the property, Rennie and Chiang (2012) and Rennie et al. (2012) have proposed: • Both shear veins and tension veins contain anomalous concentrations of gold; however, the gold concentrations are highly variable; • Gold occurs in both barren and pyrite-bearing quartz veins; • Altered wallrocks with secondary pyrite locally contain high concentrations (up to 8g/t) of gold, but gold values are highly localized; • Lithological contacts and mechanically competent rock units are important host rocks for mineralization; • The gold-mineralized zone is enriched in pathfinder elements, such as arsenic, silver, and particularly tungsten; and • Hydrothermal magnetite is disseminated throughout the diorite and quartz diorite dikes. Exploration efforts could identify additional quartz-carbonate vein-hosted gold mineralization. The sulphide-bearing rocks adjacent to known goldbearing quartz veins could represent additional goldmineralized zones. Quartz veins and hydrothermal alteration zones localized along lithological boundaries between differentiated phases within the interiors of mafic intrusions should also be examined for gold mineralization. Field Trip Stops The following tour stop exposures descriptions and figures are gleaned from Rennie and Chiang (2012). The locations of the exposures are shown in Figure 3. The geological legend for Figures 4 through 7 is shown Figure 4. Pritchard exposure – North part highlighting sample sites which returned assay results >1.0g/t Au. - 54 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 5. Pritchard exposure – Central part highlighting sample sites which returned assay results >1.0g/t Au. faulted, with a 0.5m wide east-trending (02) fault zone at its northern end and a 5m wide east-trending (02) fault zone at its southern end. A minor northwesttrending normal fault was also observed at the southern end of the rhyolite. in Figure 8. Stop 1: Pritchard Exposure (Figs. 4 and 5) UTM Coordinates: NAD 83; 15U 0504905E, 5507910N The Pritchard trench exposure is underlain by felsic tuff and tuff-breccia, massive rhyolite, mafic intrusive rocks, and one outcrop of an altered and brecciated mafic unit of unknown origin located at the north end of the trench. Both mafic units and the felsic tuff adjacent to them at the south end of the trench contain strong pervasive ankeritization. Strong pervasive silica is ubiquitous in the trench. Sericitization is moderate to strong in the felsic units and is also strong in the altered northern mafic outcrop. Structures observed in the trench were limited to faulting and strong fracturing in the felsic units. A north- to northeast-trending fault (01 axial planar) in the felsic volcanics immediately south of the access road is defined by a 1-2m deep cleft in the outcrop. The fault extends across the entire unit, and associated tensional faults and quartz veins define a sinistral sense of movement. The rhyolite unit is more strongly Despite the strongly-fractured, brittle nature of the felsic units, the bulk of the veining observed in the trench exists within the mafic units. The brecciated and strongly altered northern mafic outcrop contains four zones of <1cm wide tensional quartzankerite±tourmaline veins over 4m. The larger mafic intrusive unit at the south end of the trench is cut by Pritchard Exposure (assay results >1.0g/t Au) Channel Cut Samples (H 463_ _ _ sample series) Samples mainly from quartz veins and adjacent wall rock Cut # PR-02 PR-04 Sample # 140 147 g/t Au 1.52 2.21 Rock-types quartz vein and felsic massive flow quartz vein and mafic intrusive Grab Samples (H E44_ _ _ sample series) Samples mainly from quartz veins and adjacent wallrock Sample # 569 571 572 - 55 - g/t Au 3.58 2.00 2.46 Rock-types quartz vein and felsic volcanic lapilli tuff-breccia felsic volcanic lapilli tuff-breccia felsic volcanic lapilli tuff-breccia �Proceedings of the 61st ILSG Annual Meeting - Part 2 numerous east-trending (02) and northwest-trending (tensional), 2-26cm wide quartz ankerite-tourmaline veins. The east-trending veins typically cut and offset the northwest-trending veins. Veining within the rhyolite is mainly associated with a zone of faulting at its south end, although the east-trending quartzankerite-tourmaline-muscovite vein is observed crosscutting the faulting and local mafic dykes. Widow Showing Exposure assay results > 1.0 g/t Au Channel Cut Sample (H 463_ _ _ sample series) Samples mainly from quartz veins and adjacent wall rock Cut # WSB-01 WSB-02 WSB-03 WSB-04 g/t Au 1.77 2.08 2.50 1.12 Rock-type mafic volcanic quartz vein & mafic volcanic quartz vein & mafic volcanic quartz vein & mafic volcanic Grab Samples (G011_ _ _ sample series) Samples mainly from quartz veins and adjacent wall rock Sample # 822 823 824 827 828 830 Stop 2a: Widow’s Showing Exposure (Fig. 6) UTM Coordinates: NAD83; 15U 0504870E, 5507610N The Widow’s Showing exposure is entirely underlain by massive mafic volcanic flows with 4% 0.1-0.5mm, euhedral magnetite and local strong ankerite associated with veins and fault zones. The outcrop is cut by two main structures, a main east-trending 02 fault adjacent to the east-trending D2 vein that displays dextral Sample # 120 125 128 131 g/t Au 1.54 2.68 8.96 5.07 5.61 1.25 Rock-type quartz vein and mafic volcanic quartz vein and mafic volcanic quartz vein and mafic volcanic quartz vein and mafic volcanic quartz vein and mafic volcanic mafic volcanic wrench movement with later normal dip-slip, and an associated northwest-trending dilatational fault. Sense of motion on the east-trending structures is evidenced by northwest-trending tensional shears and drag-folded Figure 6. Widow’s Showing exposure –Highlighting sample sites which returned assay results >1.0g/t Au. - 56 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 tension veins. Widow Trench Exposure (assay results >1.0 g/t Au) Channel Cut Samples (H 463_ _ _ sample series) Samples mainly from quartz veins and adjacent wall rock The outcrop contains numerous veins, which are dominantly northwest- and north-northwest-trending, with one main east-trending vein. The northwesttrending tensional veins and north-northwest-trending R’ veins dip 60° to the northeast, and are observed being dragged into the east-trending fault with dextral sense of motion. Veins are 1-30cm wide and composed of quartz, ankerite, and tourmaline with 10% pinkishwhite feldspar and minor light brown feldspar. Cut # WS-02 Sample # 088 g/t Au 1.54 Rock-types quartz vein and mafic volcanic Grab Samples (G011_ _ _ sample series) Samples mainly from quartz veins & adjacent wall rock Sample # 805 808 809 810 811 812 815 818 g/t Au 9.37 1.29 4.22 1.59 2.50 3.06 1.18 1.29 Rock-types quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry quartz vein and mafic feldspar porphyry Stop 2b: Widow’s Trench Exposure (Fig. 7) diameter pyrite and trace chalcopyrite is observed The Widow’s Trench is underlain by three main lithologies, quartz diorite, quartz-K-feldspar crystal tuff, and diorite. The quartz diorite is cut by plagioclase-phyric mafic dykes. Alteration in the quartz diorite comprises strong pervasive ankeritization, and approximately 4% euhedral disseminated 1-2mm throughout. Alteration in the quartz-K-feldspar crystal tuff is dominated by strong pervasive silicification and sericitization, with local mo derate pervasive ankerite. The diorite is very weakly deformed and altered, and contains strong pervasive magnetite. Figure 7. Widow’s Trench exposure – North part highlighting sample sites which returned assay results >1.0g/t Au. - 57 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 The main structural features in the Widow’s Trench are northwest-trending faults and quartz veins that are restricted to the quartz diorite and quartz-K-feldspar crystal tuff. The two most dominant structures are a northwest-trending high-angle reverse fault in the northeast corner of the trench and a northwest-trending shear zone that crosses the baseline at 47m (5507635 mN). Several indicators of displacement indicate sinistral movement on the shear zone. The sense of motion and orientation of the shear zone are consistent with the interpretation that it is an R’ Riedel shear related to the main east-trending dextral shear observed in the Widow’s Showing. In addition to the dominant northwest-trending quartz veins, minor north-trending quartz veins occur and are observed being cross-cut by the northwesttrending veins. Tourmaline is observed locally in the northwest-trending vein set, but is absent in the northtrending set. northward to determine structural relationships between the large northeast-trending, faulted cliff face and any newly-exposed bedrock. Approximately 3ft (1m) of overburden was removed from the northern half of the trench exposing mafic flows, mafic lapilli tuffs, and a mafic dyke. The area quickly filled in with water, and was subsequently backfilled for safety reasons, therefore only the southern exposure is described in detail below. The trench is underlain by three lithologies, with sharp east-northeast-trending contacts except where they are in northwest-trending faulted contact. The units comprise felsic quartz crystal tuff at the northern and southern ends of the exposed trench, with finegrained mafic flows and medium-to coarse-grained mafic lapilli tuffs in the middle of the exposure. The The number of northwest-trending veins displayed on Map 7 is under-represented. Many 0.5-2cm quartz and quartz-feldspar veins exist within the northwesttrending exposure north of 40m and west of 68m. Stop 3: Flambeau Exposure (Fig. 8) UTM Coordinates: NAD83; 15U 0505100E, 5507620N The Flambeau trench totals 95m in length; however, the bulk of the exposed bedrock exists within the southernmost 45m. The trench was continued Flambeau Exposure (assay results >1.0g/t Au) Channel Cut Samples (H 463_ _ _ sample series) Samples mainly from quartz veins and adjacent wall rock Cut # FL-03 FL-05 FL-06 tuff FL-08 tuff FL-09 FL-10 tuff FL-11 tuff Sample # 181 187 192 g/t Au 1.20 1.92 1.72 Rock-types quartz vein and mafic lapilli tuff quartz vein and mafic lapilli tuff quartz vein and felsic quartz crystal 200 1.49 quartz vein and felsic quartz crystal 205 209 24.80 ** 1.49 quartz vein and mafic volcanic quartz vein and felsic quartz crystal 212 2.84 quartz vein and felsic quartz crystal Grab Samples (G011_ _ _ sample series) Samples mainly from quartz veins & adjacent wall rock Sample # 833 836 837 838 842 844 902 905 907 g/t Au 5.01 1.87 3.85 2.47 1.99 2.14 1.50 2.44 47.80 ** Rock-types mafic volcanic mafic lapilli tuff mafic lapilli tuff mafic lapilli tuff quartz vein and mafic volcanic quartz vein and mafic volcanic quartz vein and mafic volcanic quartz vein and felsic quartz crystal tuff quartz vein and felsic quartz crystal tuff Figure 8. Flambeau exposure – Highlighting sample sites which returned assay results >1.0g/t Au. - 58 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 9. Legend for all Flambeau exposure figures. fine-grained mafic dykes are observed cross-cutting the mafic lapilli tuff and the felsic quartz crystal tuff with sharp contacts. Structurally, the Flambeau trench has two distinct features. Firstly, a northwest-trending high-angle reverse fault that crosses the baseline at 34.5m has resulted in blocky/broken ground at the topographic apex of the trench. A conjugate fracture set measured around 29m on the baseline indicates pure dip-slip movement (reverse). Secondly, an east-northeasttrending (D2) well-developed 7m wide shear zone centered at 20m on the baseline (5507604 mN) has deformed the mafic and felsic host rocks into fissile schists. Strong chloritization and ankeritization are pervasive. The felsic tuffs are strongly sericitized and locally silicified, and the mafic units are moderately silicified toward the central fault zone. The large finegrained mafic unit and the adjacent felsic tuff unit to the north contain 10% fine-grained, euhedral magnetite. The Flambeau trench contains three distinct orientations of 1-10 cm wide quartz-ankerite veins: 1. West-northwest-trending, closely-spaced (1030cm) Riedel veins in the southern half of the exposed bedrock; 2. Northwest-trending tensional veins in the northern half of the exposed bedrock; and 3. Minor northeast-trending 01 axial-planar veins within the central shear zone. References Lengyel, P. 2007. Compilation report on the Van Horne Area of Interest gold property; Kenora District Geologist’s office, assessment file 52F10NW BBB-1, City of Dryden. Lengyel, P. and Rennie, C. 2009. Assessment report on the Van Horne gold property; AFRO number 2.46195, Kenora District Geologist’s office, assessment file 52F10NW DDD-7, Laurentian Goldfields Limited. Hogg, S. 2009. Assessment report on the Van Horne gold property; AFRO number 2.44181, Kenora District Geologist’s office, assessment file 52F10NW DDD4, Laurentian Goldfields Limited. Rennie, C. and Chiang, M. 2012. Assessment report on the Van Horne gold property; AFRO number 2.52196, Kenora District Geologist’s office, assessment file 52F10NW DDD-9, Laurentian Goldfields Limited. Rennie, C., Chiang, M., and Meade, S. 2012. Assessment report on the Van Horne gold property; AFRO number 2.52620, Kenora District Geologist’s office, assessment file 52F10NW DDD-8, Laurentian Goldfields Limited. - 59 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 7 - Unique Mineralizing Event at the Pidgeon Molybdenum Deposit Stripped Surface Exposure Craig Ravnaas Ontario Geological Survey, Kenora, Ontario, Canada Introduction The Pidgeon Molybdenum Deposit (Pidgeon Mo) mineralization occurs within a granodiorite body referred to as the Lateral Lake Stock (Figure 1). The known molybdenum mineralization is situated near the east-south-eastern margin of the stock and consists of molybdenite within quartz-pegmatite veins and as clots and disseminations within the rocks flanking the veins. Molybdenite is also present within unaltered and nonquartz vein parts of the stripped exposure. MPH Ventures Corp., who completed the last phase of exploration work on the property, have proposed that a 30m wide envelope containing molybdenite occurs within the eastern portion of the Lateral Lake Stock and that envelope appears to follow the boundary between the stock and adjacent supracrustal rocks (Figure 3). Mineralization has been traced for 2km in a northeasterly direction (30°) and has a dip of approximately 45° to the southeast. The field trip will examine the rocks, structural features, alteration, and mineralization exposed by a recently stripped exposure located above the historic Pidgeon Mo adit. Exploration History The exploration history for the Pidgeon Molybdenum Deposit described below was summarized from Colvine and McCarter (1977), Busch (2008), and Duke (2012). 1906: Molybdenum was discovered in a pegmatite south of Gullwing Lake by C.D. Coates. 1946: Molybdenite occurrences are reported by the Ontario Department of Mines at the east end of Lateral Lake. 1950: Initial claims are staked by G.L. Pidgeon. 1954: Claims are optioned to Detta Minerals Ltd., who drilled two sub-horizontal holes, totalling 107m. A 35m adit, located at the stripped exposure, was driven to collect a 115kg bulk sample for metallurgical testing. 1957: Pidgeon Molybdenum Mines Ltd. (PPLM) was incorporated. 1957 and 1958: Rio Tinto Canadian Exploration Ltd. optioned the property from PMML and drilled 21 holes, totalling 2348 metres. A possible, historic and non-NI43-101 compliant, resource of 568,000 tons grading 0.57% Mo was estimated. 1958: DeCoursey Brewis Minerals Ltd. completed 612m in 5 diamond drill holes along the south contact of the Lateral Lack Stock, north of Moly Lake. 1963: Denison Mines Ltd: completed 858 m in eight diamond drill holes. 1965 to 1966: Rio Algom completed a magnetometer survey and 29 diamond drill holes, totalling 3474m. The company released an internal, historic and nonNI43-101 compliant, mineral resource of 416,000 tons grading 0.57% Mo. 1977: The Lateral Lake stock was mapped by A.C. Colvine and P. McCarter of the Ontario Geological Survey and released in Miscellaneous Paper MP55. 1979 to 1980: Rio Algom completed 27 diamond holes, totalling 3710 metres. A capital and operating cost analysis was completed by Strathcona Mineral Services Ltd. who outlined an historic and nonNI43-101 compliant mineral resource of 9.0 million tonnes (Mt) grading 0.096% Mo. 1981: Rio Algom completed 3 diamond drill holes, totaling 352 metres. 2006 to 2011: MPH Ventures Corp. optioned the property in 2006. In 2007 the company completed 7 diamond drill holes, totalling 1210 m, and calculated an NI43-101 compliant revised Mineral Resources Estimate containing an inferred resource of 8.5 Mt grading 0.099% Mo. The next year (2008) MPH completed 31 diamond drill holes, totalling 2644 m, removed overburden to expose the mineralized zone above the historic adit, and completed a detailed mapping and sampling program on the stripped zone. 2012: NI43-101 Mineral Resource Estimate. During 2012 MPH Ventures completed another NI43-101 Technical Report and revised Mineral Resource - 60 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Estimate on the Mo mineralization (MPH Ventures Ltd. September 13 2012 Press Release; and Duke, 2012). The resource estimate using a cut-off grade of 0.04% Mo was calculated at: • 2.66 million tonnes (Mt) indicated @ 0.117% Mo; and • 12.39 Mt inferred @ 0.084% Mo. Location and Access The Pidgeon Molybdenum Deposit is located between the towns of Dryden and Sioux Lookout (Fig. 1). The stripped exposure to be examined is accessed by following Trans-Canada Highway 17 east of Dryden for a distance of 28km to the town of Dinorwic. Turn left (north) onto Highway 72 and drive for approximately 32km until you reach the junction with the Kathyn Logging road. Turn left onto the Kathyn Road and travel northwest for a distance of 8.8km to a trail that leads 250m west into the Field Trip Stop at the Pidgeon Mo stripped exposure. Regional Geology The Pidgeon Molybdenum Deposit is located within the eastern part of the Lateral Lake Stock (Figures 1 and 2). The Lateral Lake Stock is an elongate body extending approximately 12km from Gullwing Lake easterly to Lateral Lake and it is up to 2.8km in width. It consists predominantly of granodiorite with gradational contacts into a marginal quartz monzonite and larger country rock inclusions. Biotite is the principal mafic mineral and contains 1 to 2% epidote intergrowths. Foliation within the Lateral Stock varies from indistinct in the central part of stock to strongly foliated near the margins. The contact of the stock is concordant with the dip of the surrounding supracrustal rocks, varying from 45° on the southern side tod 30° on the northern side (Colvine and McCarter, 1977). The axial plane of a regional anticline structure trends along the central part of the Lateral Lake Stock. This regional anticline also extends into the supracrustal rocks (Fig. 2). The flanks of the anticline dip 40 to 60°, while the crestal zone plunges 25 to 30° in the east, 25 to 30° in the west. The fold is probably Figure 1. Kenora District geology map with location of Pidgeon Molybdenum Deposit (geology from OGS 2011). - 61 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 2. Regional geology and location of the Pidgeon Molybdenum Deposit and the stripped surface exposure (Blackburn, 1978). the combined result of gentle doming associated with the emplacement of the Lateral Lake Stock (Colvine and McCarter, 1977). exploration activity by MPH Ventures Ltd. provides a good exposure to view rock types, structure, alteration, and mineralization associated with the occurrence. The property and stripped exposure is situated at the eastern end of the Lateral Lake Stock, near the axial plane of the regional anticline (Fig. 2). Exposed Rock-types The metavolcanic rocks surrounding the stock consist of fine-grained chloritic units interbedded with medium- to coarse-grained amphibolitic units. Metaconglomerate units overlie the metavolcanic rocks both north and south of the stock; they contain well rounded trondhjemite (leucotonalite), chert, aplite, and mafic volcanic clasts in a quartz-feldspar-biotite matrix (Busch, 2008). The felsic intrusive rocks underlying the stripped exposure (Figs. 3 and 4) consist of coarse-grained granodiorite and quartz monzonite that are composed of quartz, albite plagioclase, orthoclase, microcline, and biotite with minor chlorite, muscovite, carbonate, sericite, sphene (titanite), epidote, and apatite (Busch, 2008). Structural Features Field Trip Stop All observed structural features are illustrated on Figures 4 and 5. UTM Coordinates: NAD 83; 15U 0545830E / 5533428N At the stripped exposure the contact of the Lateral Lake Stock with the supracrustal rocks is 70° and dips 45° SE (Fig. 3). Previous exploration has mainly focused on the mineral potential near the eastern boundary of Lateral Lake Intrusion (Fig. 3). The Pidgeon Molybdenum Deposit, exposed by stripping during the last phase of Colvine and McCarter (1977) mention that “foliation varies from indistinct in the centre of the Lateral Lake - 62 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 3. Rock-types and location of exposed rocks and mineralization and MPH Ventures Corp. diamond drill holes (Busch, 2008). Figure 4. Structural features, compositional banding, adit, and the large quartz vein in Pidgeon Mo stripped exposure (modified from Busch, 2008). - 63 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Stock to very strong on the margins, producing a finely banded to gneissic texture through separation of biotite, quartz and feldspar bands”. This compositional banding trends 42° and dips 80°E. The intrusion of several magmatic phases probably resulted in the formation of pink and grey-white banding (Busch, 2008). The pink banding is dominantly quartz and feldspar; whereas the gray bands are composed almost entirely of feldspar. The trend of this banding (42°) is not parallel to the contact of exposed supracrustal rocks (70°). The banding was overprinted by all of the other structural features present. There also appears to be increase in number of these 40°N-dipping fractures as distance increases from the contact and they overprint the 110°-trending fractures. The sloping edges of the outcrop, located in the western part of exposure near the adit entrance, appear to follow 30 to 80°-trending, 40° N dipping fractures. A similar feature comprising the slope of stripped exposure situated south of adit also appears to be associated with a 40° N-dipping fracture (Fig. 4). These slopes could comprise the footwall of a fracture. Prominent fractures observed within the stripped exposure have a strike of 110° and dip to the northwest. The dip of these fractures near the boundary with the supracrustal rocks is 45°NW, whereas that dip increases to 60°NW as the distance from the contact increases. The spacing between these fracture sets varies from 1 to 5 m. The 45-60° N-dipping and the 40° N-dipping fractures could be classified as joints and appear to have formed before the hydrothermal quartz veining event. There are also other randomly trending fractures such as a weak set of north-northeast-trending; vertically dipping fractures that overprint the quartz veins, the 45 to 60° N-, and the 40° N-dipping fracture/joint sets. A second prominent fracture set within the stripped exposure trends north-easterly between 30° and 80° and exhibits a consistent dip of 40°N irrespective of distance from the intrusion/supracrustal rock contact. It is interesting to note that the contact of the Lateral Lake stock and the supracrustal rocks dips at 45°SE, which is opposite to the dip direction of the all fracture sets observed within the stripped exposure. Figure 5. Fractures, quartz veining, alteration, and sample sites with high Mo values in the Pidgeon Mo stripped exposure (modified from Busch, 2008). - 64 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Quartz Veins present) in wallrocks; Quartz has filled some of the fractures present within the stripped exposure (Fig. 5). Quartz is not observed in the 110° trending fracture/joints but is often present within the northeast-trending 40°N dipping fracture/ joints. Quartz is not found in the northerly trending vertical dipping fractures. Quartz is also not observed along the contact of the intrusion nor extending from the intrusive rocks into the adjacent the supracrustal rocks. A majority of the quartz veins observed within the stripped exposure occur within the randomly trending fractures. These fractures could be related to strain associated with the regional anticlinal flexure. A large easterly-trending quartz vein, present in the central part of the stripped exposure, cuts the prominent 110°-trending fracture/joints (Figs. 4 and 5). Alteration Based on regional mapping and diamond drill intersections Busch (2008) proposed that “aplitic sills and potassic feldspar-bearing pegmatite are concentrated in the mineralized zone”. Busch (2008) also suggested that “based on mapping of the exposure these pegmatite and aplitic sills are [a] succession of coarse-grained alkali feldspar immediately adjacent to and within the quartz vein flanked by a uniform finegrained zone of alkali feldspar”. This alteration halo is up to 1.5m thick flanking the large easterly-trending quartz veins and is considerably narrower adjacent to other, smaller quartz veins (Figure 5). The compositional banding is commonly (but not always) destroyed in the vicinity of the alteration flanking the quartz veins. Biotite is irregularly distributed in the felsic intrusive rocks, sometimes as large patches, and is not abundant in the quartz-pegmatite veins. Mineralization Colvine and McCarter (1977) presented a summary of the molybdenite mineralization they observed at the Pidgeon Mo Deposit as follows: • Isolated grains in quartz veins and pegmatites; • Comminuted along vein margins; • Isolated grains and rosettes in wall rocks adjacent to pegmatite veins and quartz vein stockworks; • Narrow bands and lenses parallel to foliation (where • Comminuted along fractures and slippage planes in wallrocks not parallel to foliation (where present). Exploration activity, especially by MPH Ventures Corp, has outlined a 30m wide envelope containing molybdenite, extending for at least 1200 m, which occurs within the eastern part of the Lateral Lake Stock and appears to follow the boundary with the supracrustal rocks (Fig. 3). A majority of these exploration programs targeted the mineral potential of the Lateral Lake Stock along the boundary zone and associated quartz veins. The following summary of mineralization is based on examination of the exposures by Busch (2008), Duke (2012), and field visits by Kenora District OGS geological staff. Figure 6 shows the location of sample sites that returned anomalous Mo values. Molybdenite mineralization occurs in the quartz and pegmatite veins and is disseminated within the felsic intrusive rocks adjacent to veining. Molybdenite is very clotty in the quartz veins and pegmatites with a preference for the boundary between quartz and pegmatite and the boundary between the pegmatite and aplite. These clots comprise near solid molybdenite aggregates up to 3cm wide by 30cm long. Molybdenite also occurs as finer, more evenly distributed disseminations within the wallrock adjacent to the veins (Busch, 2008). Disseminated fakes of molybdenite are found in a majority of the 30 to 80°-trending, 40°N-dipping fracture/joints (Fig. 4). Molybdenite can also be found within micro-fractures cutting the felsic intrusive rocks adjacent to the 40°N-dipping fracture/joints. The mineralization present within the quartz veins can be disseminated, occur within fractures, or can concentrate along vein boundaries. Not all quartz veins present contain molybdenite (Fig. 5). Portions of the stripped exposure contain concentrations of disseminated molybdenite. These zones can contain up to 15% molybdenite, with rosettes ranging up to 1cm diameter, but do not occur within or adjacent to quartz veins, pegmatites, fractures, or fracture/joints sets (Fig. 4). Fluorite, amethyst, epidote, and tourmaline have been reported in drill logs, but were not directly related to molybdenum mineralization. Weak correlations associated with anomalous Mo mineralization are evident with B, Cd, S, SE, Th, and U values (Busch, 2008). - 65 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 6. Channel samples cuts, quartz veining, alteration, and sample sites with high Mo values in the Pidgeon Mo stripped exposure (modified from Busch, 2008) References Blackburn, C.E. 1978. Geological compilation, Kenora–Fort Frances; Ontario Geological Survey, Map 2243, scale 1:253,440. Ontario Geological Survey, 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release-Data 126 – Revision 1. Busch, D.J. 2008. Pidgeon Moly Assessment Report; Kenora District Geologist’s office, assessment file 52F16NW 028, AFRO# 2.41635, MPH Ventures Corp. Colvine, A.C. and McCarter, P. 1977. Geology and Mineralization of the Lateral Lake Stock; in Summary of Field Work 1977, Ontario Geological Survey, Miscellaneous Paper 75 No 47, p.205 to 208. Duke, C. 2012. Technical Report, Pidgeon Molybdenum project, mineral resource summary, prepared for MPH Ventures Corp, by Riverbend Geological Services Inc.; NI 43-101 Technical Report, filed September 13, 2012 with SEDAR®, see SEDAR Home Page. - 66 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 8 - Geology and Mineral Deposits of the Pickle Lake Greenstone Belt Mark Smyk Resident Geologist Program, Ontario Geological Survey, Thunder Bay, Ontario, Canada Pete Hollings Department of Geology, Lakehead University, Thunder Bay, Ontario, Canada Neil Pettigrew PC Gold/Fladgate Exploration Consuting Corporation, Thunder Bay, Ontario Introduction Exploration and Mining History Despite the fact that the Pickle Lake area has been the subject of geological mapping, mineral exploration, and mining for over a hundred years, it still remains a relatively unknown and poorly understood greenstone belt in northwestern Ontario. Renewed academic and exploration interest over the last two decades has provided new data and insights into its tectonic evolution and metallogeny in a modern context. The Uchi domain has a long history of economic mineral production, including several current and pastproducing gold mines in the Red Lake and Pickle Lake areas, and two past-producing base metal mines at Confederation Lake (South Bay volcanogenic massive sulphide (VMS) copper-zinc deposit) and Pickle Lake (Thierry copper-nickel-platinum group elements (PGE) deposit). The past-producing mines (Map 1) in the Pickle Lake belt are listed in Table 1 along with their production statistics and most recent published reserve figures. More detailed statistics are provided in the individual stop descriptions. This field trip will provide an overview of this greenstone belt, focusing on gold and base metal deposits in the vicinity of Pickle Lake (Pickle Lake Greenstone Belt Map below). To the authors’ knowledge, this is the first “formal” field trip that has been conducted in the Pickle Lake area. Many of these field trip stops are on mine properties which require access permission from the property owners. Resident Geologist Program staff (Ontario Geological Survey, Ministry of Northern Development and Mines) can provide current ownership and contact information for these properties. Field trip participants should adhere to all safety protocols and exercise caution around highways, mine workings and other potential hazards. Acknowledgments The authors have benefited from the guidance and support of a number of individuals and company personnel in developing this field guide and gaining access to mine properties: • Brian Newton (Billiken Management Services Inc.) • Norman Brewster (Cadillac Ventures Inc.) • PC Gold Inc. • Mike Aziz (Goldcorp Canada Ltd.) • Jim Hickey (Sigfusson Northern Ltd.) • Gerry White, Robert Cundari, Mark Puumala, Stuart Dunlop, Doug Lowman (Resident Geologist Program, MNDM). This synopsis of the exploration and mining history of the Pickle Lake camp has been modified from those of Thomson (1939) and Hennessey et al. (2011). A reconnaissance survey was made along the Crow River between 1903 and 1905 by McInnes (1906) of the Geological Survey of Canada. Prospecting in the Pickle Lake area commenced in 1926. In 1927, Louis Cohen of Haileybury formed a prospecting group and sent Alex and Murdock Mosher in to stake the first claims in December, 1927 on what ultimately became the Central Patricia Gold Mines Property. These claims were optioned by F. M. Connell & Associates in August 1928, and Central Patricia Gold Mines Limited was incorporated on February 19, 1929. Diamond drilling commenced at Central Patricia in February 1929 and production in March, 1930. In the spring of 1930 a mining plant was assembled at the Central Patricia and underground work was done during that summer. An ore body was outlined and a 50-ton mill recommended, but, because of financial conditions, the mill was ordered only in 1933 and was completely installed by May, 1934. In 1928, gold was discovered by Albany River Mines Ltd. at the No. 16 Vein on the Albany River claims to the east of the then Pickle Crow Property. The Crow River area had attracted little attention until the summer - 67 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Map 1: Pickle Lake greenstone belt geology and past-producing mine locations. of 1928, when these promising gold discoveries were announced and a gold rush ensued. During the following winter months an area extending 20km east of Pickle Lake was largely staked and subsequently prospected in the summer of 1929. Following the gold rush of 1928, M. E. Hurst (1931) made a preliminary geological examination of the Pickle Lake-Crow River area. Articles on the Central Patricia mine were written by J. M. Connie and J. A. Reid. Up to that time the most promising discoveries were at the Central Patricia Mine, the showings of Northern Aerial Minerals Exploration (NAME), a company set up in 1928 by J. E. (Jack) Hammell (later becoming the Pickle Crow Mine), and a vein on the Springer claims, which were under option to F. M. Connell and associates. Development on other groups of claims did not provide much encouragement, although on many claims very little rock was exposed and no real exploration could be done. The Central Patricia discovery paved the way for exploration in the region which led to the discovery and initial drilling (1929) of the first Pickle Crow ore body, the No. 1 (Howell) Vein, by NAME. The original discoveries on Central Patricia and NAME ground were drilled in 1929. Pickle Crow Gold Mines, Limited was organized and took over the property of Northern Aerial Canada Golds, Limited, successors to NAME. Underground work had commenced at the - 68 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 1: Past-producing mines in the Pickle Lake area (modified after Puumala, 2009). Mine Name UTM Zone Central 15 Patricia #1 15 Central Patricia #2 Easting (m) 697509 Northing (m) 5707914 Years of Production 1934–1951 703404 5708271 1938–1940 15 702034 5699957 1989–1994 Golden 15 Patricia Pickle Crow 15 629877 5692270 1988–1997 704366 5709757 1935–1966 684148 5708433 1976–1982 Dona Lake Thierry 15 Total Production Reserves 621 806oz (17,628kg) Au 13,158oz (373kg) Au None published 246 500oz Au 5080t @ 8.05g/t Au open pit reserves (1993)1 40,000t @ 5.95g/t Au (1994)1 None published 619,796oz (17,572kg) Au 2 10.150Mt @ 3.9g/t Au (2011) (~1.26M oz) 1,446,214oz (41,000 kg) Au 51,528,960kg Cu Measured and Indicated Resources: 8.131Mt @ 1.46% Cu, 0.18% Ni, 3.7g/t Ag; Inferred Resource: 11.507Mt @ 1.46% Cu, 2 0.15% Ni, 6.1g/t Ag (2012) K1-1 Deposit: Inferred Resource: 51.044 2 Mt @ 0.31% Ni, 0.08% Cu, 1.5g/t Ag (2012)   reserve estimates that cannot be confirmed to comply with the reporting standards of National Instrument 43Historical 101. 1 Resource estimate determined in compliance with the reporting standards of National Instrument 43-101. 2 original discovery (Howell vein) in the fall of 1933 and soon indicated an ore body. A 125-ton mill was ordered and production commenced on April 17, 1935. The immediate success attained at both properties led to a resumption of activity in the area. In the summer of 1936, 14 companies were at work. Owing to the large areas of overburden, detailed examination of the country was carried on by means of geological mapping, geophysical surveys, surface trenching, and diamond-drilling. Albany River Mines Ltd. sank the Albany Shaft to a depth of 190m (625ft) between 1933 and 1938 and completed extensive underground development. Winoga Patricia Gold Mines was created in 1936 and drilled 73 surface diamond drill holes on a pie-shaped property located between PCGM’s holdings and the Albany River Mines ground to the east. A shaft was subsequently sunk on the property in 1938. That same year, PCGM took over ownership of both Albany River Mines and Winoga Patricia Gold Mines through a new company called Albany River Gold Mines Ltd. It is believed that the Winoga Patricia Gold Mines shaft later became the No. 3 Shaft of the Pickle Crow operation. The Cohen-MacArthur zone, located 2km to the north of the developing Pickle Crow Mine, was discovered in 1933. A total of 14 surface diamond drill holes were drilled at Cohen-MacArthur in the winter of 1936. This property also was optioned by PCGM in 1938. With the acquisition of the CohenMacArthur claims, PCGM became one of the largest land holders in the Pickle Lake area. The Geological Survey of Canada completed a regional synthesis of the Pickle Crow greenstone belt during this period as well. Ground and airborne geophysical surveys have been completed over all or parts of the Pickle Crow property at various times during its early history. A dip-needle survey completed in 1936 on the Pickle Crow property was useful in tracing out the bands of iron formation. A detailed magnetic survey was carried out over the property by Teck (or its predecessor companies) around 1960. With the outbreak of World War II and the shortage of labour, mine operations slackened off considerably. Prospecting virtually came to a standstill, although both Pickle Crow and Central Patricia Mines continued production throughout the war. Pickle Crow was the last of the original gold mines to close, ending operations in 1966. Central Patricia Gold Mines began copper-nickel exploration in the Kapkichi Lake area in the mid1940s. Kapkichi Nickel Mines Limited followed, carrying out surveys and drilling from 1956 to 1966. Mining claims covering the Thierry copper-nickel deposit were optioned by Union Minière Explorations and Mining Corporation (UMEX) from Kapkichi Nickel Mines in 1969. UMEX’s follow-up work led to a decision to proceed with development of the deposit in 1974. The Thierry Mine was in production from 1976 to 1982 (Puritch et al., 2012). The property is currently held and is being explored by Cadillac Ventures Inc. In 2012, Cadillac revised the Resource Estimate for the underground portion of their Thierry Project, based upon the operating costs of conceptually combining the operations of the Thierry underground - 69 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 and the K1-1 open pit deposits. In early 1984, exploration by Dome Exploration (Canada) led to the discovery of the Dona Lake gold deposit, southeast of the main Pickle Lake camp (Cohoon 1986). Jointly owned by Dome Mines and Campbell Red Lake Mines, the Dona Lake Mine began production in 1988. After producing approximately 250,000oz of gold, it suspended operations in 1997; the property is now maintained by Goldcorp Canada Inc. The Golden Patricia Mine, 65km west of Pickle Lake, was discovered by St. Joe Gold Corp. Operated by Bond International Gold/LAC Minerals, the mine produced over 600,000oz of gold from 1988 to 1997. There has been ongoing exploration in the Pickle Lake camp since the suspension of mining operations. The most notable of the gold exploration projects has been that of PC Gold Inc. on the Pickle Crow Mine property. Since assuming ownership in 2008, the company’s exploration efforts have generated a 1.26 million-ounce, NI 43-101-compliant gold resource (10,150,000 tonnes averaging 3.9g/t gold). Mapping of the Pickle Lake area has been undertaken by the Ontario Department of Mines and its successor, the Ontario Geological Survey. The reader is referred to a number of maps and reports, including: Hurst (1931); Harding (1936); Thomson (1939); Evans (1941); Pye (1956; 1975; 1976); Ferguson (1966); Sage and Breaks (1982), Stott et al. (1989a, b) and Stott (1996). Regional Geology of the Pickle Lake Greenstone Belt The regional geology of the Pickle Lake belt has recently been elucidated by modern lithostratigraphic, geochemical, geochronologic and structural studies. These include research and mapping conducted by Stott (1996), Hollings (1998, 2002), Hollings and Kerrich (2004), Young (2003), Young and Helmstaedt (2001) and Young et al. (2006), among others. The Pickle Lake greenstone belt is located within the central Uchi Sub-province (Fig. 1). Stott and Corfu (1991) divided the belt into four volcanic assemblages based on stratigraphic relationships, isotopic age data, and aeromagnetic data (Fig. 1); in order of decreasing age these are the Northern Pickle (~2900Ma), Pickle Figure 1. Regional geology of the Pickle Lake Greenstone Belt (from Young et al., 2006). - 70 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Crow (2870Ma), Woman (2820Ma), and Confederation (2750Ma) assemblages (Fig. 1; cf. Sage and Breaks, 1982; Corfu and Stott, 1989). More recently, a revised stratigraphy has been developed by Young et al. (2006), from which the following synopsis is taken: “The Archean Uchi subprovince of the western Superior Province contains an imperfectly preserved record of over 300 million years of crustal evolution, characterized by episodic volcanism and plutonism initiated at ~3.0 Ga and culminating with subprovince-scale orogenesis at ~2.7Ga (Stott and Corfu 1991; Corfu and Stott 1993a, 1993b, 1996; Sanborn-Barrie et al. 2001; Thurston et al. 1991). The east–westtrending Uchi subprovince consists of narrow supracrustal belts surrounded by broad plutonic domains that occur at the southern margin of the North Caribou terrane (NCT, Fig. 1; Thurston et al. 1991). time of amalgamation of the Winnipeg River and Wabigoon subprovinces, respectively, against the NCT (White et al., 2003). These regional tectonic fabric-forming events are associated with gold mineralization throughout the Uchi subprovince (Stott and Corfu 1991). In particular, gold mineralization at the world-class Red Lake mining district is well studied and is most likely associated with ~2.718Ga regionally penetrative structures (Sanborn-Barrie et al. 2001; Dubé et al. 2003). In contrast, structures hosting gold mineralization in the Pickle Lake belt were thought to have formed prior to ~2.744Ga (Stott and Corfu 1991; Stott 1996); however, we can show that these structures are younger, approximately correlative with the timing of mineralization in Red Lake. The Pickle Lake greenstone belt comprises a 25km wide and ~70km long belt of Archean supracrustal rocks and internal granitoid plutons surrounded by large granitoid batholiths (Fig. 2). The supracrustal rocks have been metamorphosed to greenschist facies with amphibolite facies Neoarchean regional scale deformation and metamorphism is widespread along the southern margin of the NCT (Stott and Corfu 1991) at ~2.72 and ~2.70Ga, which probably identifies the Figure 2. Geology of the Pickle Lake Greenstone Belt (from Young et al., 2006). - 71 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 to 2740Ma trondhjemitic to granodioritic Ochig Lake pluton and Pickle Lake stock, as well as the ~2697 to 2716Ma Hooker–Burkoski stock. occurring within the narrow (<1km wide) thermal aureoles of younger plutonic bodies (Stott 1996). Although all of these rocks have been metamorphosed, in this paper the prefix “meta-“ is dropped for convenience. Up to three phases of deformation, variably developed, are recognized in the Pickle Lake belt and are described in detail later. The earliest recognized deformation (D1) is recorded by a local bedding-parallel foliation in the Pickle Crow assemblage. This foliation is truncated by the ~2735Ma Albany quartz– feldspar porphyry dyke and is not recognized in the volcanic rocks of the Confederation assemblage. The early deformation event is attributed to overturning of the Pickle Crow assemblage prior to deposition of the ~2744 to 2729Ma Confederation assemblage. Subsequent deformation and development of a regionally penetrative planar fabric (S2) postdates ~2729Ma volcanism, pre-dates the intrusion of the ca. <2716Ma Hooker–Burkoski stock and is host to gold mineralization. Corfu and Stott (1993a, 1993b) and Stott (1996) described a tectono-stratigraphy in the Pickle Lake belt. New field work, U–Pb ages, geochemical data, and Sm–Nd isotopic analyses have established the timing and determined the nature of volcanism, deformation, and tectonic assembly of the Pickle Lake greenstone belt [Young and Helmstaedt 2001, Young 2003, Young et al. 2006]. The >2860Ma Pickle Crow assemblage has been redefined to include the former Northern Pickle assemblage on the basis of stratigraphic continuity and similar volcanic geochemistry between the two units across a previously inferred fault contact. The Pickle Crow assemblage consists of tholeiitic basalt with thin, but laterally extensive, oxide-facies iron formation overlain by alkalic basalts and minor calcalkaline andesites to dacites with primitive Nd isotopic compositions (εNd 2.89 Ga = +2.1 to +2.4) suggestive of deposition in a sediment-starved oceanic basin. Integration of new and previously published field, geochronological, and geochemical data has permitted revision of the distribution and contact relationships of supracrustal assemblages in the Pickle Lake greenstone belt. The revisions can be summarized as follows: 1. Based on the lateral continuity of lithologic units and Nd isotopic compositions, the rocks formerly attributed to the inferred ~3Ga Northern Pickle assemblage have been assigned to the ~2.89Ga Pickle Crow assemblage. An earlier assumed accretionary boundary between these two assemblages is not supported by the present work. The ~2km thick 2836Ma Kaminiskag assemblage (former Woman assemblage) consists of tholeiitic basalt interbedded with intermediate and rare felsic pyroclastic flows with primitive Nd isotopic compositions (εNd 2.836 Ga = +2.4). Two samples of intermediate volcanic rocks interbedded with southeast-younging pillowed basalt, previously inferred to be part of the Pickle Crow assemblage, yielded U–Pb zircon ages of 2744+3/-2Ma and 2729±3Ma. These rocks are thus part of the younger Confederation assemblage, which consists of intercalated basalt and dacite (εNd2.74 Ga = +0.1 to +0.8) exhibiting diverse compositions probably reflecting eruption in a continental margin arc to back-arc setting. The contact between the Confederation and Kaminiskag assemblages is assumed to be a fault. 2. Based on new geochronology, the limits of the ~2.74–2.73Ga Confederation assemblage are known to extend further north and east in the belt and include rocks formerly assigned to the Pickle Crow assemblage. Opposite younging directions and enriched isotopic compositions suggest that the Confederation assemblage unconformably overlies the Pickle Crow assemblage. 3. The ~2836Ma Kaminiskag assemblage (formerly Woman assemblage), previously assumed to have developed autochthonously on the Pickle Crow assemblage, is thought to be in fault contact with the younger Confederation assemblage, having developed outboard of the North Caribou terrane. The greenstone belt is intruded by late synto post-tectonic plutons including the composite quartz dioritic to gabbroic July Falls stock with a U–Pb zircon age of 2749+4/-2Ma, and the ~2741 The proposed tectonic evolution incorporates the revised supracrustal assemblages and can be - 72 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 assemblages of the Uchi Sub-province are dominated by plume-related magmatism (c.f. Hollings et al., 1999; Tomlinson et al., 1998) whereas younger assemblages (<2.8 Ga) are characterized by arc related magmatism (c.f. Hollings et al., 2000; Hollings and Kerrich, 1999). The Pickle Lake belt (Fig. 2) offers the opportunity to investigate the transition between these regimes as he ~2.8Ga Pickle Crow assemblage offers a valuable opportunity to evaluate the tectonic processes occurring on the margins of the ancestral Superior Province during this significant transition. Hollings (2002) reported data on the Northern Pickle assemblage and distinguished three distinct volcanic suites (Fig. 3). Suite I, from the north of the assemblage, comprises tholeiitic basalts with unfractionated rare earth elements (REE). Suite II basalts are light REE (LREE) enriched with negative mantle normalised Nb anomalies relative to Th and La, similar to the CA suite of Sajona et al. (1996). Basalts from Suite III are also LREE enriched but lack negative Nb anomalies. Normalised Nb abundances are either similar to, or higher than Th and La, comparable to Nb-enriched basalts (NEB) from Phanerozoic arcs. Variable high field strength elements and heavy REE (HREE) systematics support the subdivision of Suite III basalts into two spatially distinct suites. Suite IIIa is interpreted to have melted at relatively shallow depths in the presence of Nb-Ti-bearing silicates whereas Suite IIIb melted in the presence of garnet. Variations in the geochemistry of the three suites can be accounted for by interaction between primitive mantle, adakite melts, and subduction modified mantle. The association of the three suites is interpreted to be the result of rifting of Suite I tholeiites in a backarc environment, characterized by Suite I tholeiites, permitting asthenospheric upwelling of subduction modified NEB and associated arc-like volcanic rocks (Suites II and III). The association of arc like basalts (Suite II) with NEB (Suite III) in the Northern Pickle assemblage extends the known occurrence of Archean NEB beyond the ~2.7 Ga age of previously recognised examples. summarized as follows: 1. Deposition of the Pickle Crow assemblage in a back-arc to emergent-arc setting prior to 2.86Ga. The isotopically enriched tholeiitic lower sequence may represent deposition on or near a thinned or juvenile continental margin; whereas, compositionally diverse and more evolved rocks of the upper sequence may have formed in a transitional arc to back-arc setting. 2. North- to northwest-vergent shortening (D1), widely bracketed between 2892 and 2744Ma, may have been responsible for overturning the Pickle Crow assemblage and producing a major recumbent fold. Prior to the deposition of the Confederation assemblage, the upper limb of this recumbent fold was eroded, preserving the downward-younging lower limb. This episode of deformation is one of the few documentations of pre-2.75Ga tectonism and, by implication, metamorphism in the western Superior Province other than in the North Caribou greenstone belt (Thurston et al. 1991) and cryptic evidence in the Red Lake belt (Sanborn-Barrie et al. 2001). 3. The isotopically juvenile ~2836Ma Kaminiskag assemblage developed parautochthonously with respect to the North Caribou terrane in an arc to back-arc setting followed by intrusion of isotopically enriched ~2821Ma tonalite (Quarrier tonalite gneiss). 4. The ~2744 to 2729Ma Confederation assemblage was deposited unconformably on the overturned Pickle Crow assemblage. Arc-derived plutonic rocks (the Ochig Lake pluton and Pickle Lake stock) intruded the Pickle Crow and Confederation assemblages during this interval of Neoarchean volcanism. 5. Subsequent to ~2739Ma Confederation assemblage volcanism, and prior to the intrusion of the ~2697–2716Ma Hooker–Burkoski stock, regionally penetrative deformation resulted in steep foliations and telescoping of the Kaminiskag assemblage against, and possibly tectonically interleaved with, the Confederation assemblage. Gold mineralization in the Pickle Lake belt is associated with these tectonic fabrics and is therefore contemporaneous with mineralization in the Red Lake camp, possibly reflecting a NCTwide mineralization event.” Geochemical evidence shows that the older (>2.9Ga) Hollings and Kerrich (2004) showed that the ~2.89Ga Pickle Crow assemblage comprises a basal unit of tholeiitic basalts with uniform SiO2 contents, and MgO spanning 12.4 to 3.1 wt.%. The basalts plot on a hyperbolic mixing array from a high Nb (4-8ppm), La/Ybn, but low Zr/Nb member to a low Nb (2-4 ppm), La/Ybn, but high Zr/Nb counterpart, indicative of a heterogeneous sub-arc mantle source (Fig. 3). The tholeiites were interpreted to have formed in a back-arc basin to the later calc-alkaline rocks found - 73 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 3. Chondrite- and primitive mantle-normalised plots for representative samples from the Pickle Crow Assemblage. Normalizing values are those of Sun and McDonough (1989). stratigraphically above them (Stott and Corfu, 1991; Young and Helmstaedt, 2001). The presence of NEB within the Northern Pickle assemblage, interpreted by Young and Helmstaedt (2001) to stratigraphically underlie the Pickle Crow assemblage is consistent with that model and overall the geochemical data are consistent with a paired back-arc and arc, where the high-Nb basalts are generated from back-arc mantle, and the lower-Nb basalts from sub-arc mantle. Field Trip Stops The field trip stops are shown in Figure 4 and a road log for the field trip is presented in Table 2. Stop 1: Pillowed Basalt, Pickle Lake Turn-Off UTM Coordinates: NAD83, 15U, 0696981E / 5706785N Two large, glacially polished outcrops just south of the junction of Highway 599 and Pickle Lake Road expose pillowed basalt (Fig. 5) of the tholeiitic Lower Sequence of the Pickle Lake Assemblage (Young, 2003). The tholeiitic rocks of the lower sequence display SiO2 contents of 44–51 wt. %, MgO of 4–11 wt. % and Fe2O3 of 9–15 wt. %. The majority of samples are characterized by unfractionated or weakly depleted light REE (LREE) and mildly fractionated heavy REE (HREE) (La/Smpm = 0.75–1.10; Gd/Ybpm = 0.97–1.35). The Th/La and Zr/Y ratios are generally less than the primitive mantle values of 0.11 and 2.44, reflecting the depleted character of the tholeiites. In addition, some samples show negative Zr and Ti anomalies relative to neighbouring elements on a primitive mantlenormalized plot (Young, 2003). Unfortunately, the westernmost of the two large outcrops has fallen victim to geo-vandals, who have applied white paint and localized spray paint to much of the outcrop. Despite this, recessively weathered sections and unpainted surfaces suggest that it is, indeed, pillowed basalt. The easternmost outcrop has escaped the fate of its neighbour and does offer a great exposure of the same pillowed basalt. Pockmarked by recessively weathered selvages, possible drainage cavities and fractures, these grey-weathering bun - 74 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 2: Road Log for Field Trip; trip stops are shown in Figure 4. All UTM co-ordinates on road log and field trip stops are stated as NAD83, UTM Zone 15. STOP NAME STOP NO. Starting Point Pillowed Basalt DISTANCE (km) EASTING NORTHING Junction Highway 599 and Pickle Lake Road 0.0 679055 5706910 Turn south on Hwy. 599; to Pull-off / parking area on right 0.05 696981 5706785 687610 5709480 1 THIERRY MINE AREA Sheared basalt LANDMARK Junction Highway 599 and Pickle Lake Road (reset) 0.0 Central Patricia Mine 0.9 Bridge across Crow (Kawinogans) River 1.0 Junction, Highway 599 and Thierry Mine Road 3.2 Reset, turn west onto Thierry Mine Road 0.0 Road to K1-1 Zone pit 10.3 2a 10.8 Mine Gate 14.4 Pillowed basalt, south of West Pit 2b 684010 5708540 Sheared basalt and gabbro, north of East Pit 2c 684107 5708717 Mafic dyke in sheared basalt, north of East Pit 2d 684140 5708686 “Ribboned” basalt, north of East Pit 2e 684122 5708670 “Breccia ore”, northwestern end of East Pit 2f 684086 5708638 PICKLE CROW MINE AREA Junction Highway 599 and Pickle Lake Road (reset) 0.0 Junction, Highway 599 and Pickle Crow Road 0.5 Reset, turn east onto Pickle Crow Mine Road 0.0 Mine Gate 7.1 703600 5709080 Field office and core yard 3a Trench C (No. 5 and No. 11 veins) 3b 7.5 703890 5709300 No. Vein stockpile and Shaft Pit 3d 8.2 704355 5709860 - 75 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 2: Continued Trench A (Conduit Zone) 3c DONA LAKE AREA 10.3 Junction Highway 599 and Pickle Lake Road (reset) 0.0 Junction Highway 599 and Dona Lake Mine Road 11.0 Reset, turn onto Dona Lake Mine Road 0.0 705820 5711200 Dona Lake Mine Portal 4a 702099 5699732 Sulphidic Banded Iron Formation 4b 701767 5699990 Pillowed Basalt 4c 701686 5700056 Banded Iron Formation 4d 701708 5699991 BIF, southern end of pit 4e 701764 5699994 693228 5693898 Junction Highway 599 and Dona Lake Road (reset) Ochig Lake Pluton 5 0.0 3.2   pillows are outlined by thin, dark selvages and locally developed pillow breccia. Pillow cusps and packing indicate “tops” to the southwest. A foliation, striking at ca. 125° may parallel flow contacts. Localized flexures and crenulations in this fabric were also noted. Stop 2: Thierry Copper-Nickel Mine (Sub-stops are described below) (N.B. Permission to enter the site must be granted by Cadillac Ventures Inc.) The exploration and development history of the Thierry Mine property has been taken and modified from Puritch et al. (2006). Despite Pickle Lake’s reputation as a gold camp, base metal occurrences were discovered in the late 1940s and became a significant focus of exploration and mining activity in the late 1970s. Central Patricia Gold Mines Limited completed diamond drilling from 1946 to 1950 on several gabbro-hosted copper-nickel prospects in the Kapkichi Lake area, west of Pickle Lake. Kapkichi Nickel Mines Limited conducted geophysical surveys and diamond drilling in this area between 1956 and 1958. The actual claims covering the mine site were optioned by Union Miniere Explorations and Mining Corporation (UMEX) from Kapkichi Nickel Mines in 1969. UMEX completed further exploration and drilling that eventually outlined four high-priority areas of copper-nickel mineralization: the K1-1, K2-1, G, and J anomalies, respectively. The K2-1 anomaly eventually became the Thierry Mine Deposit. In 1969, following a joint-venture agreement with Kapkichi, UMEX crews started ground electromagnetic, magnetometer and geologic surveys on the Kapkichi property. As a follow-up of the McPhar and UMEX geophysical surveys, drilling intersected low-grade (0.40% Cu, 0.11% Ni) mineralization in mafic and ultramafic rocks underlying Kapkichi Lake. This discovery was later named the “J” and “G” deposits. In September 1970, the first hole drilled outside the area covered by the joint-venture agreement with Kapkichi Nickel Mines, intersected 6m (20ft) of sulphides in biotite-chlorite schist containing 1.24% Cu and 0.14% Ni. This drill hole marked the actual discovery of the Thierry deposit, which was completely covered by overburden and thus had escaped discovery by earlier prospectors. Immediately following the discovery drill hole, the Thierry Deposit was drilled off, on a grid of cross sections 200ft apart, by 77 holes, totaling 45,000ft. In December 1971, at the end of the surface drilling campaign, the undiluted in-situ, drill-indicated reserves were estimated at 11,500,000 tons averaging 1.68% Cu and 0.18% Ni. In view of these results, UMEX decided to proceed with an underground development and - 76 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 4. Field trip stop locations, Pickle Lake Greenstone Belt geology (geology from Young et al., 2006). exploration program. Kilborn Engineering was awarded a contract to prepare a preliminary feasibility study of the deposit and to assume the project engineering. Figure 5. Pillowed basalt, Pickle Lake turn-off (Stop 1). After studying Kilborn Engineering’s feasibility report, it was decided to proceed with the development of the deposit. The official decision to place the property into production was made in 1974. Shaft Sinking started on June 27, 1972 to a planned depth of 2200ft. On December 8, 1973, after installing the cage and skip, cross-cut stations were started at the 600-, 1200-, and 1600-foot levels. At the end of 1974, the crosscuts on these levels were completed. Work progressed to complete the drifts in the hanging wall at the 600- 77 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 and 1600-foot levels and the drift in the footwall at the 1200-foot level. Horizontal and inclined holes were drilled from stations at 100-foot intervals in the drifts. Underground development, mill construction (started late 1974), and infrastructure development (town site, upgrading of access road, upgrading of the power line, etc.) were accelerated as much as possible to be ready for start-up by July 1, 1976. On July 1, 1976 the 4000 short tons per day mill was completed and production started at the Thierry Deposit in two separate small open pits (West and East) from the surface down to the 170-foot level. Dilution was very high (approximately 30%) due to the transverse mining method applied by the contractor. Ore was separated in two stock piles: high-grade (above 0.5% Cu), and low-grade (below 0.5% Cu). Underground production, utilizing the sub-level (blast hole) stoping method started in 1977. Underground work went on continuously but the milling rate was reduced from 4000 short tons per day to 2000 short tons per day in July 1977 given the low copper prices. At the same time, underground primary development was halted until October 1978. Production continued until April 1982 when the mine and mill were decommissioned. Production from the Thierry Deposit was accomplished by two open pits followed by underground operations. Underground development of the deposit included the development of a three-compartment shaft to a depth of 543m (1778ft) and 2890m of excavations at the 180, 360 and 850m levels. A total of 15,850m of underground diamond drilling was completed. Approximately 5.8 million tons of ore were mined and processed, with an average grade of 1.13% Cu and 0.14% Ni, between October 1976 and April 1982. This resulted in the production of approximately 217,750 short tons of concentrate that contained 113.6 million pounds of copper and 2.8 million pounds of nickel. Initially, only a copper concentrate was produced; however, by 1981 UMEX recognized the value of the nickel and a limited amount of nickel concentrate Table 3: Thierry Mine Estimated Historical Production (Novak and Mlot, 2004).   Year Production (tons) Cu (%) Ni (%) Pt+Pd+Au (ounces) Concentrate (tons) 1976 1977 1978 1979 1980 1981 1982 TOTAL 215,017 956,428 913,103 1,021,572 1,160,558 1,309,298 255,556 5,831,532 1.17 1.26 1.29 1.15 1.08 0.90 1.32 1.13 0.10 0.13 0.11 0.11 0.11 0.09 0.10 0.13 0.043 0.037 0.041 0.04 0.033 0.018 0.043 0.04 9,072.44 40,071.89 38,365.67 39,140.70 43,304.39 35,871.18 11,924.59 217,749 was produced which slightly enhanced the net smelter value of the ore. In addition, minor payable amounts of precious metals and PGE were also reported: 17,500 troy oz platinum; 47,000 troy oz palladium; 17,000 troy oz gold, and 900,000 troy oz silver. Mine production data were estimated from available UMEX production records and was tabulated by Novak and Mlot (2004). The estimate is summarized in Table 3 below. The mine was shutdown in 1982 due to low metal prices, and lower than anticipated ore grades. In June 1987, the mine was allowed to flood. After the mine closure in 1982, very little additional exploration was carried out; however, between 1987 and 1989 analyses were carried out for platinum and palladium. Between 1974 and 1975, 221 drill core samples from the Thierry Deposit were assayed by UMEX for Pt and Pd. In 1987, UMEX staff geologist D. Unger, implemented re-sampling and assaying of selected diamond drill holes. In 1987, R. Dahl was retained by UMEX to undertake a complete reevaluation of the PGE potential of the Thierry Mine and vicinity. An airborne geophysical survey (EM/ Resistivity/Magnetometer/VLF-EM) was flown by DIGHEM in 1988 over the Kibler Lake Stock. In 1989, UMEX Inc. re-evaluated the economic potential of the deposit with a view to reopening of the Thierry Cu-Ni Mine. UMEX, as a result of the PGE studies undertaken between 1987 and 1988, was aware that the mine contained nickel-copper zones that were coincident with anomalous PGE concentrations. UMEX also reported a large, low-grade zone of disseminated copper-nickel mineralization at the K1-1 anomaly. The K1-1 zone is tabular, 1500m long and up to several hundred metres thick. Test mining of this zone was undertaken in 1981. The average grade of the zone is 0.31% Cu and 0.1% Ni, at a cut-off of 0.2% Cu. Etruscan Resources Inc. purchased the property in 1990 with a view to placing the property into production. In 1991, Watts, Griffis and McOuat Limited (WGM) were engaged by Etruscan to prepare an economic analysis for the reactivation of the Thierry operation. WGM reported non-NI43-101-compliant diluted underground “mineral reserves” of 2.7 million short tons with average grades of 1.78% copper and 0.25% nickel. Etruscan completed reclamation of the mill and shaft facilities from 1993 to 1995. During this period, the entire drill core storage facility, including the contained core, was destroyed. PGM Ventures Corporation acquired the property from Etruscan under an asset purchase agreement - 78 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 in late 2000. PGM Ventures reviewed and initiated development of a digital database from available UMEX data. In 2002, PGM Ventures undertook a drilling campaign of 25 drill holes, totaling 8952m, to test mineralization at the Thierry Deposit and at other targets on the property. In total, PGM Ventures drilled 11 holes to confirm the presence of economically interesting mineralization at the Thierry Deposit. Drilling confirmed mineralization on the eastern and western portions of the deposit. Five drill holes intersected mineralization from depths of 1200ft to 1800ft below surface. PGM Ventures also took samples of ore representative of that previously mined from surface stockpiles to confirm the presence of consistently attractive PGE values in the well-developed breccias and massive sulphide ores. PGM engaged JVX Ltd. to complete a Time-Domain electromagnetic and magnetometer survey. The work was designed to identify geophysical targets along the main westtrending structure, which hosts both the Thierry and the Ros Zone (K1-1 zone). The survey results outlined coincident magnetic and EM anomalies, similar in intensity and extent to those at the Thierry Mine itself. Richview Resources Inc. acquired the property and commenced a diamond drilling program to explore the Thierry deposit and other areas of the property from October, 2004 through March 2005. Apart from the drilling program no other exploration work was complete by Richview. Puritch et al. (2012) outlined the subsequent work at Thierry. An NI43-101-compliant resource estimate with an effective date of February 1, 2006 consisted of 4,623,000 tonnes of Measured and Indicated material at a grade of 1.81% Cu, 0.20% Ni, along with 4,366,000 tonnes of Inferred material at a grade of 1.71% Cu and 0.18% Ni. Richview commenced its summer validation and exploration program on May 9, 2007 and completed a 14,000m drilling program. Surface drilling around the K1-1 open pit area to confirm and validate the historic drilling was also undertaken. A compilation of all mine data was conducted. A 3km corridor of unexplored ground between the Thierry Mine and the K1-1 deposit was cleared of overburden. A summer work program including excavation, geological mapping, prospecting, and geochemical sampling was completed by October 2008. A Mobile Metal Ion (MMI) geochemical survey of the Thierry Project was also conducted. The amalgamation of Cadillac Ventures Inc. and Richview Resources Inc., pursuant to a three-cornered agreement, became effective on Jan 15, 2010. Cadillac assumed 100% control of the Thierry Project. The following synopsis of the post-amalgamation period is taken from Cadillac’s Management’s Discussion and Analysis, January 26, 2015. Following the amalgamation, Cadillac obtained access to all the detailed geological and technical information on Thierry that was available. This data was then evaluated with the objective of developing a comprehensive work program. Cadillac revised the resource estimate for Thierry to include the additional data generated by the drilling of 20 holes in 2007 and 2008 by the former operators of the property as well as the drilling used in previous NI43-101-compliant resource estimates. Cadillac reported (press release, June 9, 2010) that the resource estimate update for the former Thierry Mine then consisted of a Measured and Indicated resource of 6,228,000 tonnes containing 1.92% Cu and 0.2% Ni and an Inferred resource of 8,379,000 tonnes containing 1.79% Cu and 0.16% Ni, using an NSR cut off of $46/tonne. This report enabled Cadillac to identify target areas for a surface diamond drilling campaign in conjunction with a dewatering initiative at Thierry. The objective of the planned drilling program was to infill drill an area at depth where there had been a lack of information in the Thierry Deposit model. Cadillac commenced the drilling program in May 2010 and in May 2011 reported the successful completion of all six holes, totalling 6800m. Assay results from each of the six holes drilled indicated a successful intersection of the Thierry mineralized zone at depth. Additionally, a diamond drilling program designed to test for shallower extensions of mineralization from the known and modeled deposit, commenced in January 2011. Cadillac subsequently reported that the assay results from these drill holes confirmed the extension of mineralization both to the west and to the east of the known deposit with both directions remaining open. The results of this drilling program, together with the extensions of the strike length to the west and east, formed the basis of an updated resource estimate (press release, September 1, 2011). This updated resource consisted of a Measured and Indicated resource of 8,281,000 tonnes containing 1.73% Cu and 0.2% Ni and an Inferred resource of 14,639,000 tonnes containing 1.70% Cu and 0.16% Ni, using an NSR cut off of $46/tonne. Cadillac subsequently completed an NI43-101-compliant Technical Report and Resource Estimate in December 2011. - 79 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Following a successful drilling program on the nearby K1-1 deposit in November and December 2011, Cadillac further revised the Thierry underground resource based upon conceptually combining the operations of the Thierry underground and the K1-1 open pit deposit (which resulted in a $5/tonne decrease in the NSR cut-off to $41/tonne). The further updated resource consisted of a Measured and Indicated resource of 8,815,000 tonnes containing 1.66% Cu and 0.19% Ni and an Inferred resource of 14,922,000 tonnes containing 1.64% Cu and 0.16% Ni, using the revised NSR cut-off of $41/tonne (press release, February 23, 2012). Cadillac subsequently completed an NI43-101compliant Technical Report and Resource Estimate in March 2012. Outside of the Thierry underground mine area, the eastern component of the property encompasses a project referred to as K1-1. This area, which lies 3 km east of the Thierry Mine, was drill-tested by UMEX who defined the mineralization in 1973 and 1981. UMEX had identified an estimated, non-compliant 75,000,000 tonne resource containing 0.38% Cu and 0.11% Ni. During February and March 2011, Cadillac completed three shallow drill holes on the K1-1 occurrence in order to test previous UMEX results and intersected wide zones of mineralization. During the first quarter of fiscal year 2012 Cadillac completed an additional 13-hole drilling program on the K1-1 deposit designed to confirm the historic data calculation and facilitate an NI 43-101-compliant resource. Cadillac subsequently reported the results of the 13 holes drilled, with each of the holes intersecting widespread shallow mineralization with localized higher-grade occurrences. The results of this drilling program enabled Cadillac to calculate the size of this open pit deposit and in October 2011, Cadillac reported the completion of a mineral resource estimate and exploration target for the K1-1 open pit project. This estimate was based on a combination of historic drilling by previous project operators and more recent drilling by Cadillac. The Inferred mineral resource estimate for K1-1 within a whittle pit shell consists of 19,897,000 tonnes containing 0.42% Cu and 0.10% Ni, using an NSR cut off rate of $ 15/tonne. The exploration target for K1-1 located outside and below the resource pit shell was also estimated to contain 45,000,000 to 55,000,000 tonnes grading 0.32% to 0.36% Cu and 0.08% to 0.12% Ni. Cadillac subsequently completed a NI 43-101-compliant Technical Report and Resource Estimate in December 2011. Cadillac then completed a further exploration program on the K1-1 deposit during November and December 2011 utilizing two diamond drill rigs and completing a total of 26 holes. The purpose of this program was to upgrade and expand the mineralization and models at K1-1 by infill drilling within the area of the pits and adjacent to the modeled pits, as well as targeting areas under the pits and along strike in the exploration program. Cadillac reported that the assay results of samples from the 26 holes enabled Cadillac to further update the initial K1-1 resource estimate. The updated Inferred mineral resource at K1-1 had been estimated within an economically optimized Whittle pit shell consisting of 53,614,000 tonnes containing 0.38% Cu and 0.10% Ni, using an NSR cut off rate of $11/tonne (press release, February 14, 2012). Cadillac subsequently completed an NI43-101compliant Technical Report and Resource Estimate in March 2012. The K1-1 deposit consists of a Global Mineralized inventory including the optimized pit and adjacent mineralization of 75,857,000 tonnes containing 0.38% Cu and 0.10% Ni, using an NSR cut off rate of $11/tonne (press release, February 14, 2012). This updated K1-1 resource estimate will impact on a future production decision at Thierry as the close proximity of both the K1-1 deposit and the Thierry underground mine to each other will allow cost efficiencies based on the sharing of the infrastructure and the processing plant capacity which should enable Cadillac to realize lower production costs and therefore process material of a lower grade than would be envisioned if either deposit were operating independently of each other. In May 2012, Cadillac received a positive Preliminary Economic Assessment (PEA) which demonstrated the technical and potential economic viability of reopening the Thierry Mine and processing material from both the underground Thierry Mine deposit and the K1-1 open pit deposit on a combined basis (press release, May 12, 2012). Cadillac had reported that it had commissioned a dewatering plan for the Thierry mine site. During fiscal year 2013, Cadillac completed more drilling at K1-1 (8 DDH, totalling 2200 m) to potentially increase the mineralized resource. Cadillac subsequently reported that each hole had successfully encountered mineralization outside of the current Whittle Pit-defined compliant resource and returned assay grades comparable, or better than, those contained within the current Whittle Pit-defined compliant resource (press release, September 13, - 80 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 2012). Cadillac also completed a field prospecting/ mapping program designed to provide field testing and sampling of correlated geophysical and coincident magnetic anomalies on the property outside of the area of the Thierry Mine and K1-1 during the year. Cadillac continues to evaluate the results of the aforementioned drilling, prospecting and mapping programs with a view to determining the next phase of the exploration program at the Thierry Mine and K1-1. Geology of the Thierry Mine The description of the geology of the Thierry Mine, below is taken from Anderson (2007). A vertical crosssection through the deposit showing the pit and some of the underground workings is present in Figure 6. Regional setting The Thierry deposit is situated in the northwest portion of the Archean Pickle Lake greenstone belt in the Uchi volcano-plutonic subprovince of the western Superior Province. Supracrustal rocks in this portion of the Pickle Lake belt are included in the Mesoarchean (>2860Ma) Pickle Crow assemblage, which consists of massive and pillowed tholeiitic basalt flows, with minor intercalated sedimentary rocks, iron formations, and calcalkaline andesite and dacite (Young et al., 2006). Along the northwest margin of the Pickle Lake belt, the Pickle Crow assemblage is intruded by stocks, plugs and sills of mafic to ultramafic composition. The largest of these intrusions, the July Falls stock, is composed of gabbro, diorite and quartz diorite, the latter of which has yielded a U-Pb zircon age of 2749Ma (Young et al., 2006). South of the Thierry deposit, these rocks are intruded by the 2740Ma Pickle Lake stock, which is interpreted to be broadly syntectonic Figure 6. Cross section of the Thierry deposit (Patterson and Watkinson, 1984). - 81 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 edge of the open pit, most outcrops consist of strongly foliated mafic tectonites, which appear to be derived from pillowed mafic flows and flow-breccia, with subordinate intercalations of sedimentary rocks. Individual pillows are typically strongly flattened and attenuated, although convincing, bun-shaped examples of pillows are observed locally. Intercalated mafic breccia layers tend to be more strongly tectonized than the pillowed flows, and locally exhibit moderate to strong epidote- calc-silicate alteration. Layering in these rocks dips moderately to steeply northwest. Mafic sills and dikes are abundant and consist mainly of medium-grained equigranular leucocratic and mesocratic gabbro. (Young et al., 2006). To the north, the Pickle Crow assemblage is intruded by the post-tectonic Bow Lake batholith. Metamorphic mineral assemblages indicate middle greenschist-facies regional metamorphism, although amphibolite-facies assemblages are developed in the thermal aureoles of the Neoarchean plutons. Overprinting relationships described by Young et al. (2006) indicate a relatively simple deformation history. A regionally pervasive and penetrative planar shape fabric, which is oriented subparallel to bedding and the general trend of the supracrustal belt, is assigned to the D2 deformation phase (Young et al., 2006). In the northern portion of the belt, the S2 fabric dips steeply northwest and is axial planar to tight to isoclinal, moderately to steeply-plunging, F2 folds. This fabric also contains a variably developed L2 stretching lineation that plunges steeply to the northeast or north-northeast. On the northwest flank of the East pit, the mafic volcanic and intrusive rocks are structurally overlain to the north by a 10–20m thick unit of well-layered chloritic tectonite, which appears to have been derived from stratified mafic to intermediate volcaniclastic rocks. These rocks are weathered dark green to light grey and are fine-grained, chloritic and feldspathic. Local geology Bedrock exposures in the footwall of the Thierry deposit consist mainly of pillowed mafic flows and intercalated flow-breccia, with subordinate mafic intrusions. The mafic flows weather greenish-grey to black and are finegrained, aphyric and non-amygdaloidal. Wellpreserved, bun-shaped to amoeboid pillows range up to 1.5m in maximum dimension, with 1–2cm thick selvages. Interpillow material consists mainly of strongly recrystallized and altered hyaloclastite. On the southwest margin of the East pit, individual pillowed flows range from 1 to 5m thick and are intercalated with similarlythick units of monolithic mafic breccia and pillow-fragment breccia. The subvertical flow contacts trend roughly north and a pillow cusp in one location indicates tops to the east. Mafic tectonites in the hanging wall of the Thierry deposit are intruded by fine-grained, lateto post-tectonite felsic dikes that weather pale brown and contain sparse quartz and feldspar phenocrysts. Structural geology Overprinting relationships in a single outcrop of pillowed basalt in the area roughly midway between the fenced shaft and West pit indicate at least three generations of ductile deformation fabric. In this outcrop, an early penetrative planar fabric that trends generally north is defined by foliated chlorite and amphibole, and a variably developed tectonite layering. This early fabric is overprinted by open to tight, symmetric to Z-asymmetric folds that trend northeast and plunge steeply to the north. These folds are, in turn, transected by a finely-spaced crenulation cleavage that dips subvertically and trends eastnortheast. In this location, the mafic volcanic rocks are intruded by minor dikes and sills of medium to coarse-grained gabbro/pyroxenite. In the largest of these intrusions, leucogabbro, melagabbro and pyroxenite exhibit highly contorted and locally gradational contact relationships that are suggestive of magmatic-mixing. Although these relationships clearly indicate three generations of ductile deformation, the available information is insufficient to determine the relationship (if any) of these structures to those observed in the East and West pits, or their significance with respect to the regional deformation history described by Young et al. In the hanging wall of the Thierry deposit, on the northwest margin of the East pit, bedrock exposures exhibit significantly higher finite strain than those in the footwall. Along the northwest - 82 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 common feature of the generally east-trending greenstone belts in the Superior Province. Such features can be interpreted in terms of two distinct generations of structure (e.g., early vertical stretching overprinted by transcurrent shearing), or a single generation of structure formed during progressive deformation within a transpressional shear zone (i.e., a shear zone with a significant component of zone-normal shortening). In the present case, the available data is insufficient to differentiate between these, or other, possibilities. (2006). It is noteworthy however, that the overall style, orientation and sequence of deformation structures observed in these outcrops correspond closely to structures assigned to the regional D1, D2 and D3 deformations by Young et al. (2006). As described above, finite strain appears to be considerably lower in the footwall of the Thierry deposit, as compared to the hanging wall. The stratification observed in the outcrops of pillowed mafic flows and breccia on the south margin of the East pit trends generally north, at a high angle to the strike of the mineralized zone. Open to tight, west-trending folds, such as those observed south of the West pit, could account for this marked obliquity. In this regard, it is possible that the localization and orientation of the shear zone that hosts the Thierry deposit was controlled by the axial surfaces or limbs of pre-existing, west-trending fold structures. Alternatively, the obliquity could simply result from progressive folding along the margins of the mineralized zone during shearing, in which case the strata further down in the footwall would be expected to gradationally change orientation into the more regional easterly trend. The tectonite fabric is overprinted by at least two generations of folds, which are particularly well-developed in the layered mafic to intermediate tectonite on the northwest margin of the East pit. The earlier folds are open to tight, strongly asymmetrical, similar-style structures. The fold hinges plunge moderately to steeply toward the northwest or north. Z-asymmetrical folds predominate over S-folds, although both are observed in single outcrops. This opposing vergence is consistent with the local occurrence of sheath-like fold closures. A sheath-fold interpretation for these closures is supported by the fact that the fold hinges are oriented subparallel to a locally observed quartz-ribbon lineation and the more extensively-developed L-fabric in the underlying mafic tectonites. As described previously, rocks in the hanging wall of the deposit consist mainly of variably tectonized mafic volcanic rocks. The tectonite fabric in these rocks consists of both planar and linear fabric elements. The planar fabric element dips steeply to the north-northwest or northwest and is defined by flattened primary features (e.g., pillows, clasts), transposed veins and primary layering, and foliated chlorite and amphibole. The presence of steeply plunging sheath folds and stretching lineations in the mafic to intermediate tectonite indicates a significant component of dip-slip shear, in apparent conflict with the abundance of dextral kinematic indicators on the horizontal outcrop surfaces. Although such features are not inconsistent with transpressional shear zones, they may be more readily interpreted in terms of a two-stage deformation history involving early dip-slip shear, overprinted by later strike-slip (dextral) shear. Such complex deformation paths would have important implications for the localization and geometry of Cu-Ni ore bodies in the Thierry deposit (see below). The linear fabric element is defined by aligned hornblende porphyroblasts, rare quartz ribbons and, in the mafic intrusions, aligned aggregates of amphibole and chlorite that likely represent pseudomorphic replacements after primary pyroxene phenocrysts. Some of these rocks approach pure L-tectonites. On the northwest margin of the East pit, the L-fabric plunges 65o towards the north. On the east flank of the East pit, the L-fabric plunges 40o toward the northwest. The tectonite fabric and early asymmetrical folds are overprinted by gentle to open, upright folds with steeply plunging hinges and northwesttrending axial planes. Asymmetric boudins and shear bands are abundant on horizontal outcrop surfaces and consistently indicate dextral shear. Tectonite zones exhibiting well-developed kinematic indicators on horizontal outcrop surfaces and steeply plunging stretching lineations are a In the mafic tectonite, the mineral assemblage appears to comprise chlorite, hornblende, feldspar, epidote, biotite and quartz, consistent - 83 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 have obscured or destroyed primary textures. Patterson and Watkinson (1984) describe the host-rocks as consisting of 75% metagabbro, 20% mafic metagabbro and 5% talc-carbonate schist, and provide convincing textural evidence of a peridotitic to dunitic protolith for the talccarbonate schist. Novak and Mlot (2004) cite previous drill-log descriptions of chemical sedimentary rocks in the mine sequence, which were not described by Patterson and Watkinson (1984). with upper greenschist- to amphibolite-facies metamorphism. Hornblende porphyroblasts are strongly aligned in the planar shape-fabric, and locally define a penetrative mineral lineation that plunges parallel to the stretching lineation. In one location, hornblende porphyroblasts are also randomly oriented within and across the planar shape-fabric, suggesting that hornblende growth outlasted penetrative ductile deformation. In the western portion of the large stripped outcrop on the northwest flank of the East pit, the mafic tectonites are cross-cut by a series of widely-spaced, discrete, brittle-ductile faults. These faults dip steeply toward the northnortheast and are defined by very planar slipsurfaces that sharply truncate the tectonite fabrics in the wall-rocks. The faults contain narrow (<20cm), laminated fault-fill quartz veins which locally widen into more irregular quartz breccia veins. Oblique internal fabrics in the laminated portions of the veins indicate dextral strike-slip shear. The most significant of these faults contains narrow, discontinuous veins and irregular blebs of remobilized chalcopyrite with minor pyrrhotite. The rock-types tentatively identified by this author in the mineralized drill cores from the Thierry deposit include possible chemical or siliciclastic sedimentary rocks, ultramafic intrusive rocks, and mafic intrusive/extrusive rocks. The possible ultramafic rocks are dark green to black, fine- to medium-grained and equigranular, and composed almost exclusively of chlorite (±carbonate, talc, serpentine and amphibole). The possible sedimentary rocks are light to dark grey or reddish-brown, fine-grained, siliceous and biotitic, and typically exhibit a finely laminated to layered structure. Evidence of intense transposition is provided by the common occurrence of tight to isoclinal, strongly asymmetric and rootless folds. From the descriptions in Patterson and Watkinson (1984, p.4), Novak and Mlot (2004), Keller (2005) and Puritch et al. (2006), it seems likely that most of these rocks were previously described as ‘mylonite’ or ‘chlorite-biotite schist’, with the implication that they were derived through intense deformation and hydrothermal alteration of mafic or ultramafic intrusive precursors. However, this interpretation appears inconsistent with the observation that these rocks commonly lie in very sharp contact with ultramafic and mafic intrusive rocks that lack mesoscopic deformation fabrics, which would require very abrupt strain-gradients from the mylonite into the precursor rocks. Discrete shear-fractures and faults locally form complex, diffuse networks within, and along the margins of, the zone of mafic-intermediate tectonite, and are particularly well-developed in the massive, equigranular gabbro sills. Asymmetric fabrics and minor offsets of the main tectonite fabric consistently indicate dextral strike-slip shear. The shear fractures dip steeply to the north and northeast, and appear to define a Riedel-type shear-fracture system to which the brittle-ductile faults described above may be related. Riedel-type systems comprise multiple orientations of synthetic and antithetic shear factures that, in a strike-slip system, will exhibit a common, sub-vertically plunging line of intersection. Orebodies overprinted by these types of shear-fracture systems would be expected to exhibit significant structural complexities. The gabbro is fine- to medium-grained and typically massive and equigranular, with a weak to moderate foliation defined mainly by chlorite. Leucogabbro grades continuously into melanogabbro, consistent with the descriptions of Patterson and Watkinson (1984), and the occurrence of possible magma-mixing textures in gabbro exposed on the south margin of the East Host-rocks to mineralization Identification of primary rock-types in the examined intervals of mineralized drill core from the Thierry deposit is significantly complicated by the combined effects of intense metamorphic recrystallization, penetrative deformation and local hydrothermal alteration, which - 84 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 irregular-shaped fragments or grains of silicate gangue. This style of mineralization typically contains 5 to 10% total sulphide. pit. Blocky, relatively undeformed inclusions of melanocratic gabbro in the layers of possible sedimentary origin are tentatively interpreted as gabbro dikes that underwent boudinage and segmentation during intense transposition of the host-rocks. Unless the host-rocks were significantly strained prior to emplacement of the gabbro dikes, the relatively undeformed aspect of the boudins would tend to support a sedimentary, rather than tectonic, origin for the fine-grained and laminated aspects of biotitic layers. Net-textured mineralization consists of irregular, inter-connected veinlets of [chalcopyrite] and [pyrrhotite] which surround fragments or grains of silicate gangue. Typically, this style of mineralization contains 10 to 25% total sulphide and, with increasing sulphide content, grades into breccia mineralization. Breccia mineralization consists of near-solid (50-80% total sulphide) to solid (>80% total sulphide) [chalcopyrite] and [pyrrhotite] that contains angular to sub-rounded inclusions of wall-rocks or other silicate gangue. Breccia mineralization tends to form discrete, relatively sharp-walled veins that are apparently controlled by irregular fracture arrays in relatively undeformed (i.e., weakly to non-foliated) wall rocks. These veins may result from ductile flow of sulphides during deformation (i.e., piercement veins) or fluid-state remobilization during, or subsequent to, deformation (e.g., Gilligan and Marshall, 1987). The breccia veins typically do not exceed 50cm in maximum thickness, and are composed mainly of [chalcopyrite], with subordinate to subequal [pyrrhotite]. Some veins exhibit a marked segregation of [pyrrhotite and chalcopyrite]. Styles of mineralization Patterson and Watkinson (1984) describe four types of sulphide ore in the Thierry deposit: breccia ore, mylonite ore, bornite ore and disseminated sulphides. Novak and Mlot (2004) propose a similar classification, but instead refer to the mylonite ore as ‘chlorite-biotite schist’ ore and include an additional, poorly defined, ‘oxidized’ ore-type. “… sulphide mineralization in the Thierry deposit is subdivided on the basis of texture into disseminated, matrix-textured, net-textured and breccia styles. It is recommended that the terms ‘mylonite’ or ‘chlorite-biotite schist’ should be avoided, as these describe only the hostrocks to the sulphide mineralization, not the mineralization proper…” Late-tectonic emplacement of the breccia veins is indicated by their relatively undeformed state, and the local presence of multiphase deformation structures in wall-rock inclusions. In addition, breccia mineralization is observed to discordantly cut tight to isoclinal folds in the laminated rocks of possible sedimentary origin, indicating emplacement subsequent to development of the intense transposition fabrics. These aspects, coupled with the relatively finegrained, non-annealed nature of the breccia mineralization indicate emplacement late in the tectono-thermal evolution of the host-rocks. Disseminated mineralization consists of isolated, monomineralic to polymineralic grains of pyrrhotite (po), chalcopyrite (cp) and/or pentlandite that typically do not exceed 2mm across and are evenly distributed throughout the host-rock. Typically, this style of mineralization contains less than 5% total sulphide. Disseminated mineralization is observed in ultramafic and mafic rocks of inferred intrusive origin, and has been interpreted to represent recrystallized primary magmatic sulphide (e.g., Patterson and Watkinson, 1984). It should be noted however, that disseminated sulphides are also locally observed in the finely laminated rocks of possible sedimentary origin, indicating that alterative mechanisms may be required to explain at least some of this mineralization. Nevertheless, some of the breccia veins do preserve evidence of a more protracted, multiphase emplacement history. One of the breccia veins intersected in DDH PGM 05-60, for example contains two generations of breccia mineralization. In this vein the early generation of breccia is inclusion-poor and distinctly layered and is sharply cross-cut by a later generation of Matrix-textured mineralization consists of blebs and irregular veinlets of [chalcopyrite] and [pyrrhotite] that form the matrix to equant to - 85 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 relatively massive and inclusion-rich breccia. chalcopyrite, pyrrhotite, pentlandite and pyrite. Although most examples of breccia mineralization are relatively undeformed, some exhibit evidence of ductile and brittleductile deformation, in the form of a moderate to strong planar fabric defined by aligned inclusions of silicate gangue. This mineralization is appropriately referred to as ‘durchbewegt’ (Marshall and Gilligan, 1989). In [one] example…the ductile fabric is cross-cut by enechelon arrays of [chalcopyrite]-filled extension veins (i.e., tension-gashes). Four principal types of sulphide mineralization are recognized at the Thierry Deposit (Patterson and Watkinson, 1984) with Patterson (1980) noting a fifth: Puritch et al. (2012) provided another synopsis of copper-nickel-PGE mineralization at the Thierry Deposit: • Chlorite-Biotite Schist Mineralization (mylonitic mineralization): 56% of all mineralized rock (CBS ore), containing 5-20% sulphide as stringers of chalcopyrite, pyrrhotite, pentlandite and pyrite; the stringers parallel foliation and where gradational with breccia mineralized rock, the breccia fragments are flattened and elongated. “Mineralization at the main Thierry and adjacent K1-1 deposits, is more or less coincident with what is best characterized as a chloritebiotite-hornblende altered mylonitic shear zone (the “CBS shear zone”). The shear zone extends across the ultramafic intrusive along a strike length of about one kilometre and a width up to 50m. Within the shear zone mineralization is hosted by highly schistose rocks containing stringer sulphides to less schistose ultramafic rocks containing massive stringers or veins and disseminated sulphides. Primary sulphides, listed in approximate order of decreasing abundance are pyrrhotite, chalcopyrite, pyrite and pentlandite. Cubanite, bornite, magnetite and minor ilmenite have also been identified. Violarite and mackinawite have developed from alteration of pentlandite. • Breccia Mineralization: 40% of all mineralized rock and composed of 20-30% sulphide, consisting of rounded to angular fragments of gangue in a matrix of chalcopyrite, pyrrhotite, pyrite and pentlandite. Breccia mineralized rock grades into CBS mineralized rock. • Bornite Mineralization: 2% of all mineralized rock, containing 1-5% sulphide as stringers and disseminations of chalcopyrite and bornite in carbonate veins associated with blocks of amphibolite schist in the main shear zone. • Primary Disseminated Sulphide Mineralization: 1% of all mineralized rock, occurring as blocks of chalcopyrite (with exsolution of bornite or cubanite) plus pyrrhotite and pentlandite between remnants of olivine. • Oxidized Mineralization: 1% of all ore comprised of several varieties of ore, characterized by violarite, millerite, bornite etc. Outside of the main mineralized zone, chalcopyrite and bornite occur as stringers as well as finely dissemination sulphides. Bornite is commonly associated with carbonate and quartz veins. Oxidized mineralizations are reported to contain violarite, millerite and bornite. The mylonite and breccia mineralization has a copper-to-nickel ratio of 8:1, compared to a 2:1 ratio in the disseminated sulphides. In addition, the chalcopyrite:pyrrhotite ratio is approximately 1:1 in the mylonite mineralization and 1:10 in the disseminated sulphides (Patterson, 1980). Copper-nickel-PGE mineralization at the Thierry and K1-1 deposits is hosted within a highly deformed and altered ultramafic sequence. Copper-nickel-PGE mineralization consists of: Precious metal minerals have been found in the Thierry Deposit in two distinct associations: • • In the breccia mineralization, the precious metal minerals merenskyite, moncheite, stutzite and an unnamed mineral Ag3BiTe3 occur with chalcopyrite, pyrrhotite. pentlandite, pyrite and violarite. Sulphide matrix breccia; • Blebs and small stringers, occasionally net textured sulphides; and • • In the bornite mineralized rock, the precious metal minerals, native silver, acanthite, stutzite and merenskyite are associated with chalcopyrite, Disseminated sulphides. The sulphide mineral assemblage consists of - 86 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 bornite and copper bismuth sulfosalt (wittichenite and emplectite). rocks along the Thierry Mine access road (STOP 2a) and at the mine site (Stops 2b to 2f) to the tholeiitic Suite I Upper Sequence of the Pickle Lake assemblage, correlative with Suite I of Hollings (2002). These basalts are characterized by unfractionated or weakly depleted light REE and mildly fractionated heavy REE (La/Smpm= 0.70–1.5; Gd/Ybpm = 1–1.2). The strongest positive correlation of metals is between silver and copper. There is a corresponding negative correlation between silver and nickel at values of nickel greater than 0.5%. A plot of Pt/(Pt+Pd) versus Cu/(Cu+Ni) shows average head grades of the Thierry Deposit to be enriched in copper and somewhat in platinum relative to other similar deposits (Naldrett and Cabri, 1976). From a PGE perspective the Thierry mineralization falls into two groups, both of which fall well off a characteristic trend line defined by Naldrett and Cabri (1976) for typical PGE ores. The first group is pyrrhotite-rich and correspondingly has a high Ni content. This group is platinum poor compared to the second Cu-rich, chalcopyrite rich fraction which has a high platinum content. Stop 2a: Mafic Metavolcanic Rocks, south of K1-1 Zone, south side of Thierry Mine access road UTM Coordinates: NAD83; 15U 0687610 E / 5709480N This stripped area exposes strongly sheared amphibolitized and epidotized mafic metavolcanic rocks south of the K1-1 mineralized zone (Fig. 8). Metrescale M- and S-folds are visible, as are boudinaged or ptygmatically folded felsic dykelets and quartz veins. A pervasive, sub-vertical foliation strikes 090° and is also locally folded. Millimetre-scale shear bands and spaced cleavages are ubiquitous. Disseminated pyrite generates localized gossan. No primary features are noted in these metavolcanic rocks. Ores at the Thierry Deposit underwent intense modification after their initial deposition as magmatic sulphides. Dynamic metamorphism has mobilized much of the breccia and mylonite mineralization. Stop 2b: Thierry Mine, open pits; pillowed basalt, south of the West Pit Sub-Stop Descriptions, Thierry Mine (Fig. 7) Geochemical sampling and analyses conducted by Young (2003) ascribed all the mafic metavolcanic UTM Coordinates: NAD83; 15U 0684010E / 5708540N Figure 7. Field trip stop locations, Thierry Mine Property. - 87 - Outcrops north of the exploration office and near �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 8. Sheared and folded mafic metavolcanic rocks, Stop 2a, Thierry Mine access road. the capped shaft and vent raises expose sheared and locally pillowed mafic metavolcanic rocks with pervasive epidote veins and patches and ptygmatically folded quartz-feldspar-epidote veins. Flattening and stretching have produced pillow aspect ratios of up to 10:1, molar-shaped pillows and preclude unequivocal “tops” direction determination. Pillows have dark, chlorite-, biotite- and amphibole-rich selvages and variably bleached cores (Fig. 9). Foliations in this area are more-or-less east-striking and have variable dips. Felsic dykes may be discordant to foliation or may be sheared and Z-folded, suggesting a protracted intrusion history. Recognizable pillowed basalt gives way to strongly foliated, phyllitic, amphibolitic equivalents closer to the southern margin of the west pit. The access ramp at THE southeastern end of the West Pit provides a crosssection through these phyllitic rocks, dipping circa (ca.) 50° north. Local malachite staining was noted in some of the feldspathic neosome of some of the banded amphibolite. An ~100m wide deformation zone (aka CBS Shear Zone) extends through the middle of the West Pit. In the West Pit, the main mineralized zone is situated on the southern margin of the deformation zone. Around the margins of the pit, contacts are exposed between phyllitic mafic metavolcanic rocks and weakly foliated, synvolcanic (?) gabbroic rocks. These gabbroic rocks are thought to host the bulk of the Thierry mineralized zone at depth, whereas chloritebiotite schist hosts the mineralization at surface to shallow levels of the mine (A. Carlson, Richview Resources Inc., pers. comm., 2008). Minor fold axes are generally northeast-plunging, as are ore shoots within the main mineralized zone. Steeply plunging stretching lineations are evident on foliation planes. On the eastern margin of the West Pit, massive mafic metavolcanic rocks locally host a chloritic, 110°-striking cataclasite unit which may offset the mineralized zone between the West and East pits. Biotite schist is locally developed along the contacts with gabbroic units. Stop 2c: Thierry Mine, open pits; sheared Basalt and Gabbro, north of East Pit UTM Coordinates: NAD83; 15U 0684107E / 5708717N Sharp contacts between sheared, locally laminated basalt and medium- to coarse-grained gabbro are wellexposed on these glacially polished exposures north of the western end of the East Pit and just east of the road that extends between the two pits. Stop 2d: Thierry Mine, open pits; mafic dyke in sheared basalt, north of East Pit UTM Coordinates: NAD83; 15U 0684140E / 5708686N Stop 2e: Thierry Mine, open pits; “ribboned” basalt, north of East Pit UTM Coordinates: NAD83; 15U 0684122 / 5708670N Figure 9. Deformed pillowed basalt flow, south of West Pit, Thierry Mine (Stop 2b). Sheared amphibolite/pillowed basalts predominate north of the East Pit and in some places are highly stretched and ribboned. Felsic dykes are folded and millimetre-scale shear bands are developed in sheared intermediate dykes. A distinctive, 25cm-wide dyke, hosted in amphibolite, is characterized by numerous, rounded granitoid xenoliths (Fig. 10). - 88 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 10. Mafic dyke with rounded granitoid clasts in deformed metabasalt, north of West Pit, Thierry Mine (Stop 2d). Stop 2f: Thierry Mine, open pits; “breccia ore”, northwestern end of East Pit UTM Coordinates: NAD83; 15U 0684086E / 5708638N At the northwestern end of the East Pit a rusty, malachite-coated exposure displays chalcopyrite- and pyrrhotite-rich, “breccia ore”-style-mineralized zones that crosscut folded metavolcanic rocks. Host rocks are well-foliated and resemble the main chlorite-biotite schist host to sulphide mineralization. Small, lenticular “islands” of more competent or xenolithic (?) material are enveloped by schistose rocks. Felsic dykes are dismembered. Stop 3: Pickle Crow Mine (N.B. permission to access the site must be granted by PC Gold Inc.) The Pickle Crow Mine was discovered in the early 1930’s by Northern Aerial Mineral Exploration which began sinking the No. 1 Shaft in 1933. Northern Aerial was acquired by Pickle Crow Gold Mines (PCGM) in 1934 and commercial production at the mine began in 1935. The Pickle Crow mine operated until 1966, during which time it produced 1,446,214 troy oz gold and 168,757 troy oz silver from 3,070,475 short tons of ore milled (at an average grade of 0.47oz/t or 16.14 g/t Au). After the mine closed in 1966 there was little work done until Highland Crow/Noramco Mining Corporation acquired the property in the 1980’s. Noramco completed over 46,000m of surface drilling, dewatered the mine workings to the 750 foot level, and completed an additional 9,000m of underground drilling. With the end of the flow-through era the property sat dormant until Wolfden Resources Inc. acquired the property in 1999. Wolfden subsequently entered into a surface mining agreement (top 100m of the deposit) in June, 2000 with privately held Cantera Mining Limited, resulting in fragmentation of the property ownership. Cantera constructed a 225 ton per day (tpd) extreme gravity mill on the site, submitted a partially completed production closure plan with MNDM, and began constructing a tailings management facility within the historic Pickle Crow tailings area. Cantera also commenced stockpiling of material mined from the historic No. 1 Vein shaft and crown pillar area. Cantera ceased work in 2002 and was placed into receivership in 2004. On November 5, 2007, Premier Gold Mines (successor to Wolfden) and Don Ross (successor to Cantera) announced the signing of a Letter of Intent (LOI) to sell their interests in the property to PC Gold, at that time a private company. A definitive agreement was signed on December 21, 2007. On May 13, 2008 PC Gold satisfied the terms of the definitive agreement and completed the acquisition by completing an initial public offering and listing on the TSX. PC Gold Inc. has been actively exploring the Pickle Crow property since 2008. The mineral concessions of the Pickle Crow project consist of 106 patented mining claims and 119 unpatented claims for a total area of approximately 21,690ha. Since acquiring the property PC Gold has completed a massive program of digitizing historical data and building a detailed 3D model of the mine and surrounding property. PC Gold has to date has completed 98,000m of drilling, as well as several trenching, mapping, and geophysical programs. Most recently, PC Gold has been pursuing financing for small-scale production using the onsite 225tpd mill with ramp access on the near surface, highgrade No. 22 and 23 veins. The Pickle Crow deposit currently hosts an NI43101-compliant inferred resource of 1.3 million ounces, including 637,000 high-grade vein-hosted ounces at an average grade of 9.1g/t Au. Local Geology Many geological investigations have been completed on the Pickle Crow Property (Fig. 11). Detailed rock descriptions on a property scale have been completed by Thomson (1939), Pye (1956, 1975), Ferguson (1966), and MacQueen (1987). Other workers, such as former PCGM employees R. J. - 89 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 11. General geology of the Pickle Lake Mine area, showing Field Trip Stop locations. Graham and L. D. S. Winter, wrote unpublished reports that contain much valuable information. The following descriptions of the geological units of the Pickle Crow Property are derived from the detailed, property-scale work referenced above, and placed into the tectonicstratigraphic framework of Young et al. (2006). The Pickle Crow Property in the immediate vicinity of the mine is underlain by rocks of both the Pickle Crow and Confederation assemblages; rocks of the Kaminiskag assemblage occur to the east. On the property, the Pickle Crow assemblage is dominated by tholeiitic basalts with intercalated sedimentary rocks, primarily banded iron-formation (BIF), and rare calc-alkaline volcanic and volcaniclastic units. The assemblage is interpreted to be unconformably overlain by the Confederation assemblage. An unnamed, Temiskaming-like sedimentary assemblage was identified and dated (<2752.2±2Ma) by PC Gold in 2009 and comprises polymictic conglomerate, sandstone, siltstone, argillite, and argillaceous iron formation. The assemblage occupies a small, fault-bound basin near the contact between the Pickle Crow and Confederation assemblages, and likely represents the erosional unconformity between the two assemblages. The assemblage appears to represent a similar event to that which produced the Houston Lake assemblage of the nearby Red Lake greenstone belt. Pickle Crow Assemblage Rare ultramafic rocks (>20 wt% MgO) were noted in drill core, particularly in deep drilling in holes PC08-014A and PC-10-085 at Shaft 1 and 3, respectively. These units are typically 5 to 20m thick and are intensely talc-altered, sometimes presenting a severe obstacle to drilling. Due to their alteration and deformation, the protolith cannot be accurately determined. No relict spinifex texture has been observed. These ultramafic units are always found intercalated with tholeiitic basalt and are currently interpreted to represent subvolcanic sills. The tholeiitic lavas are compositionally consistent; no flows of intermediate composition have been noted. - 90 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Greenish-grey-weathering, massive basalt is generally fine-grained. Although pillowed units are ubiquitous, it has not been possible to subdivide or map individual flows. Some flow contacts are marked by beds of iron formation. Pillows range in size from 0.5 to 1m and have narrow selvages ~1cm thick. Generally, there are no amygdules in the pillows. They display a variety of shapes and in only a few places on surface has it been possible to make top determinations. Flow-top breccias locally contain light-coloured, angular fragments up to 30cm in diameter. Parts of the flows are medium-grained with individual crystals up to 1.5mm in diameter. These rocks are of medium-grey colour with small light grey feldspars. All these medium-grained rocks have been included with the volcanic rocks and no dykes or sills have been mapped separately on surface. The mediumgrained metabasalts consist essentially of fibrous amphibole, chlorite, and highly altered plagioclase with small amounts of carbonate, epidote, saussurite, and quartz, and subordinate leucoxene, apatite, sphene, and sometimes pyrrhotite (Pye, 1956). One main band of banded iron formation (BIF) is known to be interbedded with the basalts adjacent to the workings of the Pickle Crow Mine but, in places, there are additional local bands. This BIF has been traced in the Shaft 1 area by surface mapping and by drilling for 2700m and ranges in thickness from ~1m up to 25m or, where it has been thickened by folding, to ~45m. In the No. 1 Shaft workings the BIF is known to extend down-dip for 1200m and is thought likely to persist to much greater depths. Magnetite-carbonate BIF is prominently banded with alternating layers, varying in thickness from thin laminae up to 5cm. The more siliceous layers may be light or dark grey, laminated chert. No jasper bands are known to be present. Some of the darker layers contain a high proportion of magnetite but the magnetite content varies along strike; consequently, magnetic surveys have been only partially successful in tracing these beds. The weathering of some BIF produces a rusty iron oxide alteration. This variety of BIF consists of bands of cryptocrystalline quartz and siderite, varying amounts of magnetite and pyrrhotite, and occasional streaks of chlorite (Hurst, 1931). Some of the bands are composed almost wholly of magnetite in small, angular grains. Pyrrhotite occurs as patches, streaks, or grains replacing iron carbonate and chlorite or as veinlets traversing the various bands. Streaks of chlorite, which probably represent inclusions of schistose greenstone, are often associated with the carbonate bands. Carbonate (ankerite) BIF is exposed in outcrops and old trenches on Pickle Crow claims PA774 and 777 in the Cohen-MacArthur area. This BIF is ~550m in length, with thicknesses up to 10m and has sharp contacts. Limonite from the weathering of this BIF has stained adjacent host rocks. At the northeastern end of this zone some outcrops are typically BIF but elsewhere, the more siliceous bands do not weather in relief and the weathered surface of the rock has a uniform surface. The fresh surface of the rock varies in colour from light to dark grey and some specimens of this BIF are very hard and siliceous. Sulphide-chert-rich argillaceous BIF is abundant in the Central Pat East Zone where it typically occurs as interbedded, magnetite-poor, chert-rich BIF and sulphide-rich (pyrite) argillite with minor intermediate tuff. Thin section work (Kolb, 2011) indicated that this BIF is also very carbonate-rich. Although extensive, this BIF displays great local variation in thickness and type (e.g. oxide- versus sulphide-facies). It is not wellexposed at surface and is known almost entirely from diamond drilling. The Pickle Crow assemblage contains significant amounts of calc-alkaline rocks. Although rare in the core mine trend, dacites are the most common rock type outside of it. Most units are lenticular and in cannot be easily correlated between adjacent outcrops. The presence of beds containing breccia fragments is a widespread and characteristic feature. There is no known interbedding between the basic metavolcanic rocks and the acid to intermediate metavolcanic rocks. In places, the contact is marked by interbedded sedimentary rocks. Feldspar-phyric dacite occurs in the northwestern part of the property, often in tuff-breccia units. Such rocks outcrop, or have been recorded in drill core, in the area between the No. 1 Shaft and Powderhouse Lake. Quartz–phyric rocks are very rare, but have been reported in the Central Pat East Zone. Although portions of the local Temiskaming-like sedimentary assemblage, notably the argillite and conglomerate, had been reported historically (Graham, 1965) it was not recognized as a separate unit from the Pickle Crow assemblage until 2009 (Lynch, 2010). The sedimentary basin is quite restricted in its dimensions extending from just northeast of Shaft 3 (hole PC-09050) to just north of the Springer Shaft. The unit appears to be a fault-bound basin that contains polymictic conglomerate at its northeastern end, grading into finegrained wacke, siltstone, argillite, and argillaceous - 91 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 BIF to the southwest. The assemblage is particularly pyrite-rich, specifically the conglomerate (interstitial pyrite) and argillite (bedded nodular pyrite). Locally, this pyrite has been remobilized and converted to pyrrhotite, especially in the conglomerates. The Confederation assemblage unconformably overlies the Pickle Crow assemblage on the property and includes: • The lowermost basaltic unit. • A thick succession of intermediate to felsic metavolcanic rocks in the southeastern part of the property. A thin, discontinuous unit of pillowed basalt occurs at the base of the Confederation assemblage and is distinguished from the lower sequence of the Pickle Crow assemblage on the basis of trace element geochemistry. This lower basalt unit, or geochemical Suite I, is characterized by elevated FeO contents (13 to 16 wt%) and displays weak LREE enrichment and HREE fractionation with minor negative niobium anomalies (Young et al., 2006). The southeastern part of the property is underlain by calc-alkaline felsic to intermediate volcanic and volcaniclastic material that is similar in most respects to the calc-alkaline rocks in the Pickle Crow assemblage. Feldspar-phyric dacite and related breccias predominate and are both widely distributed in this part of the Pickle Crow Property. The dacites contain feldspar, quartz, and sericite together with minor amounts of chlorite, epidote and leucoxene (Pye, 1956; Ferguson, 1966). The dacite breccias contain scattered, light-coloured angular felsic fragments, usually less than 5cm in maximum dimension. Near the eastern boundary of the property the fragments in the volcaniclastic rocks are rounded and were termed “agglomerate” or “volcanic conglomerate” by Thomson (1939). There are two porphyry stocks and several porphyry dykes within the property boundaries. The Pickle Crow porphyry stock occurs northwest of No. 3 Shaft and the Albany River porphyry stock outcrops near the Albany shaft. Dikes have been mapped on claims PA729, 1139 and 2011 (Ferguson, 1966). The Pickle Crow porphyry (2909+15Ma; Young et al., 2006) is elliptical in plan, 1.8km long and 200m wide. The major axis strikes 055° and appears to be generally conformable in strike and dip with the enclosing rocks. But, on the 229m (750ft) level plan it is shown to crosscut the trend of the volcanic units at a small angle. The complete outline of the stock has not been established on the 869m (2850ft) level, but over this vertical distance the porphyry appears to maintain its shape, becomes slightly wider, dips at 77° to the northwest, and does not appear to plunge. A few porphyry dykes or sills are present near the stock but the outline is regular, without apophyses extending outward from the main intrusion. On the 229m (750ft) level the southern contact of the intrusion with the adjacent country rocks is sharp. The Pickle Crow porphyry is distinguished by large (2 to 10mm) quartz phenocrysts which are rounded to oval in cross-section, but a few are rectangular with rounded corners. On the light grey weathered surface, quartz phenocrysts are enclosed in a matrix of kaolinized feldspar. Microscopically, the rock contains distinct, well-formed, but smaller, fractured crystals of albite (Ferguson, 1966). The matrix of the rock is an aggregate of tiny anhedral grains of quartz and altered plagioclase with accessory amounts of magnetiteilmenite, leucoxene, apatite, sphene, and rutile. The Albany porphyry (2735+10Ma; Young et al., 2006) is 670m long and 120m wide, striking 060°. This stock is somewhat irregular at the ends, with lobes and dyke-like apophyses. Some of the associated dikes are parallel with the trend of the enclosing rocks but others are crosscutting. From surface to the 625ft (190m) level, the stock dips 65° northwest. On surface, the major axis of the stock makes a small angle with the strike of the enclosing rocks. The stock appears to maintain a similar shape and extend down the dip with no known plunge. This faintly pink porphyry contains abundant feldspar phenocrysts, scattered quartz phenocrysts, and a few biotite crystals, all of which are about 2mm in size. In its most unaltered state it appears granodioritic in composition. In thin section the feldspar is considerably altered to white mica and in some local areas to saussurite. Some of the quartz phenocrysts are individual crystals, but other phenocrysts consist of clusters of crystals. Quartz and carbonate occur as trains of interstitial material between the larger feldspar phenocrysts. Small crystals of apatite are enclosed within the large biotite crystals. Miscellaneous feldspar porphyry dykes also occur on the mine property. A prominent northwest-trending diabase dyke cuts across the western portion of the property. Narrow, fine-grained diabase dykes have also been mapped in the workings and encountered in underground drilling at the Pickle Crow Mine. One dike occurs in the - 92 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 hanging wall of the No. 1 Vein and appears to be earlier than the vein, but another dike of similar appearance cuts this vein on the 411 m (1,350 ft) level (Pye, 1956). Diabase dykes also occur in the Pickle Crow porphyry. Although some of these dykes are parallel to the margins of the stock, others are at angles of from 20°to 40° to the contact. All of these dykes within the porphyry are cut by the mineralized veins. A biotite lamprophyre dyke outcrops along the southern side of the rock exposures on claim PA760, in the northwestern part of the Pickle Crow property and, a dyke of similar composition occurs in the Pickle Crow porphyry and cuts a mineralized vein. The lamprophyres are massive, dark grey to black, mediumgrained rocks with large (locally >5mm) biotite phenocrysts. Two varieties have been recognized. One, which cuts the Howell Vein, is composed chiefly of biotite, orthoclase, chlorite, and carbonate, and may be an altered minette. A second, post-ore dyke, which cuts the No. 2 Vein system, is made up of biotite, andesite, quartz, subordinate clinopyroxene, and accessory apatite and zircon (Pye, 1956). Lamprophyre dykes, similar to those exposed at Shaft 1, are also very common and are intimately associated with gold mineralization at the Central Pat East Zone. The Hooker-Burkowski quartz-phyric trondhjemite stock, southwest of the Pickle Crow property, is undeformed and intrudes all Pickle Lake greenstone belt assemblages. Structural Geology The structural history of the Pickle Crow Property, although extensively studied, is by no means definitely understood. The focus has been mainly on the Pickle Crow assemblage and is largely based on the descriptions of Ferguson (1966) and the deformation history of MacQueen (1987). Stratigraphic younging determinations and structural-facing directions have been based on rare, unequivocal pillow tops and graded argillaceous beds. One persistent magnetite-carbonate BIF forms an important marker unit within the Pickle Crow tholeiites. The general strike of rocks on the property is northeast and the dip is 75° to 80° northwest. The plunge of folds in the BIF near No. 1 Shaft is due north at 75° to 80°. The rake of the three productive veins in the No. 1 Shaft area is 70° at 020°. The Pickle Crow porphyry stock and the Albany porphyry stock both extend down the dip and do not appear to plunge. Several of the anticlines narrow and plunge beneath the younger rocks in a pattern that would be consistent with a plunge to the northeast. Some other anticlines maintain a constant width for considerable distances and some anticlines have a shape, in plan, which suggests a plunge to the southwest. Along some fold axes the stratigraphic sequence is repeated in reverse order which indicates plunge reversals. The major anticlines on the property are: the Pickle Crow; Albany Shaft; Pumphouse Lake; Sawmill; Pumphouse Creek; Powderhouse Lake south; Powderhouse Lake central; and Powderhouse Lake north. The major adjacent synclines are: the Albany Shaft; Township Line; No. 3 Shaft; Pickle Crow No. L; Pickle Crow No. 2; and Pumphouse Lake (Ferguson, 1966). It is important to note that this complex folding is essentially confined to the Pickle Crow mine area, and that the upper (northerly) part of the Pickle Crow assemblage outside of the property is a simple, north-facing, homoclinal sequence (Young et al., 2006). The general trend of the fold axes is northeast, but the Pickle Crow No. 2 syncline, the Pumphouse Creek anticline, and the Township Line syncline have fold axes which curve across the major fold axes. The folds strike northeast and dip steeply northwest, forming isoclinal folds with overturned southeastern limbs. A well-developed schistosity is present in the metavolcanic rocks in the mine area and on the limbs of folds. This schistosity conforms with the dip of the bedding and with the axial planar cleavage of the latest period of folding. The porphyries are not strongly sheared, but the platy minerals developed in the matrix are aligned in conformity with the schistosity of the adjacent volcanic rocks. Lithologic units in the Pickle Crow area have been metamorphosed to greenschist facies. The greenschistamphibolite facies isograd was defined below the mine workings in Shaft 1 (at approximately 1600m depth) and is identified by the appearance of hornblende and abundant garnet. Greenschist mineral assemblages in basaltic rocks were petrographically determined to be: chlorite+actinolite+epidote+quartz+albite; and chlor ite+sericite+quartz+albite. The chlorite and actinolite have a preferred orientation parallel to the second deformation foliation, suggesting that this was the peak metamorphic episode. MacQueen (1987) described a complex polyphase deformation history that includes four tectonometamorphic episodes summarized below (Figure 12): - 93 - “The earliest generation of structures (D1) is present as rare 1- to 3m isoclinal fold closures or �Proceedings of the 61st ILSG Annual Meeting - Part 2 hinge zones (F1) within BIF units that have been subsequently refolded inside second generation (F2) fold closures. Due to refolding, the F1 folds have axial planes with a mean strike direction between east-northeast to east and a steep dip to the north with hinge lines (L1) plunging steeply (60° to 70°) to the northeast (Pye, 1956). striking sets of shear fractures have, in some instances, developed into discrete 2 to 5m wide Type 2 shear zones that are strongly foliated and run between substantial Type 1 structures. These Type 2 shears are an important structural control on gold mineralization. D2 deformation was contemporaneous with gold-bearing vein emplacement and deformation definitely continued sometime after vein emplacement, as evidenced by Z-folding of the No. 1 Vein in Shaft 1 and even more intense folding and boudinaging of veins in Shaft 3. D2 structures are the most prominent in terms of metamorphic overprint and the current distribution of rock units and mineralization on the property. In the mine area, D2 is characterized by a penetrative axial planar schistosity (S2), parallel, or at a small angle, to bedding/S1, striking northeast and dipping 75° to 87° to the northwest. Stretching lineations (L2) in the S2 plane, defined by chert and magnetite in BIF, quartz phenocrysts in quartz feldspar porphyry and varioles in tholeiites, are steeply plunging (70° to 85°) to the north-northeast. The effect of this stretching lineation cannot be overstated, as pillows have been measured with stretching ratios in excess of 30 to 1. The lineation is typically the strongest and most consistent fabric, with the S2 foliation often a distant second. As a result, the most continuous direction of lithological continuity (and mineralization) on the property is vertical. D2 fold closures (F2) have axial surfaces that strike northeast and dip steeply (75° to 85°) to the northwest, and hinge lines (L2) that plunge 60° to 80° to the northeast. D2 closures are characterized by 1 to 200m wide, tight, to isoclinal, similar folds, (best developed in BIF). They have thickened hinge zones, attenuated limbs and wavelengths of 300m. The large D2 folds outlined by the BIF include the Pickle Crow Anticline, Pickle Crow Syncline, Sawmill Anticline, and Powder House Anticline (Ferguson, 1966). D2 shear zones occur throughout the Pickle Crow property as zones parallel to S2 surfaces (Type 1) and as discrete shear zones (S2’) that splay off the Type 1 shear zones in a eastnortheast direction, connecting Type l shear zones. Type 1 shear zones are strongly foliated zones greater than 30m wide which dip steeply (75° to 85°) to the northwest. They include the Pickle Crow Fault, Highland Crow Shear Zone, Pumphouse Shear Zone and Powderhouse Shear Zone. Shear fractures in outcrop surfaces (Stott, 1996) trend east-northeast at low to moderate angles (20° to 40°) to S2. These east-northeast- The third generation of structures on the Pickle Crow property (D3) consists of northwest- and north-striking conjugate faults, steeply dipping to the northeast, and crosscutting and displacing earlier structural fabrics. Late undeformed veins, felsic dikes and lamprophyre dikes have been emplaced along fractures parallel to northwestand north-striking conjugate sets of fractures which crosscut D2 fabrics. The final (D4) deformation generation is represented by the development of late, continuous, northwest-striking, en-echelon extensional fractures that crosscut earlier fabrics (Sage and Breaks, 1982). These fractures are up to 20m wide and have considerable strike lengths, up to hundreds of metres. The Pickle Crow diabase dike, and other diabase dikes hosted in these structures, crosscut all strata, including the Hooker-Burkowski stock and the granitic terrane to the north and south of the Pickle Lake greenstone belt. The Pickle Crow diabase dike has been displaced sinistrally along a northeast-trending fault, the only record of post-D4 deformation within the Pickle Crow area.” Gold Deposits PC Gold considers the gold occurrences in the Pickle Lake mining camp to be classical examples of deposits grouped under the descriptive model of Archean low-sulphide, gold-quartz veins. The major gold ore bodies at Pickle Lake are hosted by the Pickle Crow assemblage, which locally hosts at least two such large structural discontinuities or “breaks”: the core mine-trend break, which preserves Temiskaminglike sedimentary rocks; and the Cohen-MacArthur deformation zone, which is up to 100m wide. The gold-bearing veins at Pickle Crow fill pre- or syn-ore - 94 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 12. Structural elements of the Pickle Crow Mine area. faults, shears, and fractures in the various host rocks. Auriferous sulphide zones that are stratabound and contained within BIF occur adjacent to shear zones in some areas. The historically mined ore at Pickle Crow was contained in quartz veins that are generally banded (crack-seal) with tiny streaks of tourmaline, chlorite or sericite, and in fracture fillings. Quartz is by far the main vein mineral, with lesser carbonates, including siderite, ferruginous dolomite, and calcite. Minor albite, chlorite, sericite, and local traces of tourmaline, magnetite, and scheelite have been noted. Native gold was the main ore mineral. The main sulphide minerals are pyrrhotite and pyrite, which combined are usually less than 2% of the vein material, along with trace arsenopyrite, chalcopyrite, galena, and sphalerite. Some gold is closely associated with the sulphides and traces are found in the altered wall rock but, in general, the gold is free and occurs along sericitechlorite-fuchsite-lined fractures and seams in-filling minute fractures in the quartz. Spectacular samples of visible gold have been observed in a number of places in the mine. As a general rule, however, the gold is very finely divided and practically invisible to the naked eye (Pye, 1956). At Pickle Crow, alteration minerals include silica, sericite, chlorite, carbonate, and pyrite. Host rocks for the gold mineralization at Pickle Crow include tholeiitic basalts, BIF, intermediate volcanic/volcaniclastic rocks, and quartz feldspar porphyry. Gold occurrences are associated with four styles of mineralization: • Narrow, high-grade, gold-scheelite-bearing quartz veins, which were the main source of gold produced at the Pickle Crow Mine from 1935 to 1966. • BIF-hosted gold mineralization adjacent to vein structures. BIF contains stringers and discontinuous lenses of quartz and the iron-bearing minerals have been replaced by sulphides. Both quartz and sulphides are gold-mineralized. Only a limited amount of this type of material was processed at the Pickle Crow Mine. However, BIF-hosted gold was - 95 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 the main ore type at the adjacent Central Patricia Mine to the southwest. • Shear zone-hosted gold mineralization, consisting of complex, wide zones of intense shearing and alteration which are intimately associated with the intrusion of the Albany porphyry and characterized by disseminated pyrite, discontinuous quartz veining, and sulphidation of interflow BIF. • Arsenopyrite-associated gold mineralization which typically occurs as disseminated to semi-massive arsenopyrite and quartz-arsenopyrite stockworks hosted by BIF, but can be also found, to a lesser extent, in shear zones and/or quartz veins in volcanic rocks. Similar arsenopyrite-rich BIF-hosted gold was the main ore type at the adjacent Central Patricia Mine. A substantial number of auriferous quartz veins have been located on the property, along with several occurrences of BIF-hosted mineralization. The following quartz vein descriptions are mainly from the work of Thomson (1939), Pye (1956), and Ferguson (1966), while the subsequent BIF mineralization descriptions also include information from MacQueen (1987) and Winter (1988). Gold was produced from the No.1, No. 2, No. 5, No. 6, No. 7, No. 8, and No. 9 Veins during the life of the Pickle Crow Mine. The only additional mineralization of this type that was processed at the Pickle Crow mill came from exploration drifts at the Albany Shaft area. The most productive of the quartz vein ore bodies was the No. 1 Vein (Fig. 13). This vein has been traced on surface over a strike length of 900m and extends below the lowest level of the mine, or beyond a depth of 1500m (almost 700m below the lowest historically mined level of the deposit). The average thickness of this vein is 0.9m. The eastern part of the vein is highly contorted with an overall strike of 083°, cuts across the lithologic units at an angle of 30° to 40°, and has an average dip of 73° north. The western part of the vein has an overall strike of 058°, cuts the formations at about 10°, and has an average dip of 75° northwest. The No. 1 ore body is a ‘shoot’ within the No. 1 Vein, with the eastern boundary determined by gold content and the western boundary by diminishing vein thickness (Pye, 1956). Although the No. 1 Vein is typical of the veinstyle mineralization at the Pickle Crow Mine there are variations throughout the mine property. Veins in the Shaft 1 area (Fig. 14) are relatively undeformed, and more laminated, with more fine-grained gold. Figure 13. Cross-section through the No. 1 Shaft area, Pickle Crow Mine. They have very little shearing or wall rock alteration and have significant down-dip continuity. Veins in the Shaft 3 area (e.g., No. 2), are more deformed, have few laminations with more coarse-grained gold, and possess wider zones of shearing and alteration in the wall rock. They are more en echelon in nature and have less down-dip continuity. Veins in the Albany Shaft area (e.g., No. 16) are even more deformed, with few laminations and generally rare visible gold, possess wide zones of shearing and intense alteration of the wall rocks, and generally poor continuity. The No. 1 Vein consists largely of white or greyish, coarse- to fine-grained, almost sugary quartz, a little ferro-dolomite, tourmaline, scheelite, and subordinate amounts of metallic sulphides. The ankerite, tourmaline, and scheelite, although locally occurring as patches completely enclosed by the quartz, generally occur in thin seams replacing chloritized basalt in ‘book and ribbon’ structures or in the walls, and so - 96 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 14. Structural relationship between mineralized zones, No. 1 Shaft area (Coates and Anderson 2008). help accentuate the banded character of the vein. Two distinct generations of quartz are recognizable; one gold-bearing and making up the body of the No. 1 Vein, the other barren and occurring, along with calcite, in narrow transverse veinlets that cut sharply across the earlier type. Most of the other mineralized veins at the Pickle Crow property have similar characteristics to the No. 1 Vein. Quartz is by far the most abundant mineral of all the veins and the two generations of quartz are found in many of the veins. The quartz in other veins in the metabasalt is a light grey colour and is banded due to the presence of inclusions of schist and dark coloured minerals. Similarly, the second generation of quartz consists of veinlets, generally less than a centimetre in thickness, which are approximately at right angles to the veins and extend completely across the veins but rarely extend into the wall rock. These veins consist of quartz with abundant white or pink calcite (Ferguson, 1966; Pye, 1956). Pyrite and pyrrhotite are the most abundant sulphides and both are about equally common in the veins in the metabasalt, but pyrite is more abundant in the No.2 Vein, and only pyrite occurs in the No.16 Vein. In the metabasalt and iron formation wall rocks, pyrrhotite occurs as irregular masses, disseminated grains, and narrow seams. It occurs in fractures in the veins, or the wall rock, or both, and as grains healing broken crystals of arsenopyrite and pyrite. The pyrrhotite is later than the quartz and the late variety of pyrite, but is - 97 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 replaced by sphalerite, chalcopyrite and galena. The second style of mineralization at Pickle Crow is the gold-bearing BIF type. Considerable mineralization of this type was identified by PCGM on the property during its exploration and development work, mainly in the No. 1 Shaft area. These were locations adjacent to the No. 1 ore body (the No. 1 Iron Formation Zone or Eastern Iron Formation), the No. 5 ore body (the No. 5 Iron Formation Zone), and an area approximately midway, at depth, between the No. 1 and No. 5 Veins known as the Central Iron Formation Zone. After running iron formation material from test stopes in all three zones, PCGM found that it was unable to satisfactorily process both vein and iron formation mineralization styles due to their different metallurgical characteristics (Winter, 1987). It has also been reported that the average auriferous iron formation grade, believed to be about 6.85g/t Au (0.20oz/ton Au), was below the then cut-off grade of 8.57g/t Au (0.25oz/ton Au) (MacGregor, 1989). The Eastern, Central, and No. 5 iron formations or quartz-sulphide zones comprise stringers and discontinuous lenses of quartz within BIF. The iron-bearing minerals of the BIF have been replaced by sulphides adjacent to the veins and gold is present in the veins and the associated sulphides. Approximately 15,000t of this type of mineralization, taken from test stopes in each of the above occurrences, were processed during the historic mining operations. The BIF-hosted mineralization consists of bleached and altered iron formation with variable amounts of pyrite, pyrrhotite, and occasionally arsenopyrite, often with heavy secondary magnetite and quartz, and carbonate veins and veinlets. The gold mineralization at Pickle Crow (both quartz vein- and iron formation-hosted) is localized along, or adjacent to, Type 2 shear zones, i.e., shear zones that are developed oblique to greenstone belt lithological trends and cross between adjacent Type 1, lithologically concordant shear zones, or faults. The shear zone-hosted type of mineralization is restricted to the Albany Shaft area, and is referred to as Conduit style mineralization after Conduit Zones 1, 2, and 3 (formerly the A, B, and C zones). The mineralization is characterized by wide, highly complex zones (both lithologically and structurally) of shearing with discontinuous quartz veining, sulphidized interflow BIF, and disseminated pyrite. All rock types can be mineralized, with a preference for the interflow BIF where the highest grades occur, often in association with a pronounced crenulation fabric and abundant Z-folds. Lithological complexity is a key component, providing abundant small-scale competency contrasts. Conduit-style mineralization, when present in homogeneous rocks such as massive basalt, is much less intense and lower-grade. Alteration mineralogy includes widespread carbonate (some calcite, but primarily ankerite), strong sericitization, chlorite, silicification, quartz veining, and abundant disseminated pyrite. Visible gold was not observed although grades greater than 1oz/ton have been recorded. Minor alteration minerals include tourmaline, hematite, and fuchsite. The geometry of the mineralization is poorly understood. In the case of Conduit Zone 1 it has been defined by drilling to be an approximately 40m wide, northerly-plunging (~55°), pipe-shaped body; however Conduit Zones 2 and 3 do not mimic this geometry. There is also strong evidence that the Conduit style of mineralization is simply a much stronger manifestation of the shearing that surrounds the high-grade quartz veins at the Pickle Crow Mine. For instance there is evidence that Conduit Zone 2 is the southwest extension of the No. 16 Vein. The mineralization is often moderate- to low-grade and possibly amenable to open pit or bulk underground mining methods. Arsenopyrite-bearing gold mineralization was described early on in the history of the Pickle Crow property when it was discovered at the CohenMacArthur Zone and MacArthur Vein in the 1930s. These are located north of the core mine trend and just south of the Kawinogans (Crow) River. While historically it was not a significant style of mineralization at Pickle Crow, it was the principal ore at the nearby Central Patricia Mine. Subsequent work by PC Gold has found this style of mineralization to be much more widespread on the property than previously thought. The Cohen-MacArthur deformation zone is a wide (up to 100m) zone of intense shearing and carbonate (ankerite)-sericite alteration that roughly parallels the core mine trend and runs the entire length of the property. It was identified through geophysics and drilling in 2010. The Cohen-MacArthur structure is the strongest structure present on the property and forms a dividing line between mineralization styles. All mineralization north of this structure is associated with arsenopyrite whereas the mineralization south of it (with the rare exceptions of the Arsenide Vein and No. 21 Vein) are associated with minor scheelite and low arsenopyrite contents. Arsenopyrite mineralization can occur in several forms such as localized shear zones, quartz veins, quartz stockworks, and disseminations in - 98 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 both volcanic and chemical sedimentary rocks. The most widespread mineralization and highest grades occur within iron formation such as the Central Pat East Zone, where the mineralization is also spatially associated with several late unmineralized lamprophyre dikes which presumably used the same structures as the gold mineralization originally exploited. The tenor of the gold mineralization is very closely tied to arsenopyrite content, as well as the degree of silicification, quartz flooding and/or veining. Alteration minerals associated with arsenopyrite mineralization are widespread and include often intense carbonate (ankerite) alteration and strong silicification, quartz flooding and/or veining, and moderate sericite alteration. Minor alteration minerals include tourmaline, pyrrhotite, and chalcopyrite. Petrographic studies by Kolb (2011) on the Central Pat East indicate that the gold is free, located within fractures or next to arsenopyrite crystals. An important structural component of the arsenopyrite-associated gold mineralization is the presence of fractures and/or veinlets at high angles to bedding. At the Central Patricia Mine, the host BIF possessed an easterly strike; however, the mineralization was present in sulphide-rich fractures oriented perpendicular to the strike of the BIF. On the 750 level, mineralized fracture patterns occur at a high angle to the strike of the BIF, generating ore bodies (Tigert, 1949). At Central Patricia miners exercised care to drill holes at approximately right angles to the locally prevailing fracture pattern (i.e. parallel to bedding). Normal drilling, at right angles to the drift (i.e. perpendicular to bedding), would produce very false results (Tigert, 1949). These fractures plunge approximately 55° to the east and formed very consistent high-grade shoots within the BIF. Similar structures have been observed at the Central Pat East Zone, where arsenopyrite-filled fractures and arsenopyrite-quartz veinlets cut bedding at high angles. Similar high-angle quartz veinlets, plunging to the northeast and perpendicular to the strike of the host iron formation, have been observed in the arsenopyritepoor iron formation mineralization at the Sawmill Vein and may play an important role in controlling the geometry of this style of gold mineralization as well. Mineral Resources The NI43-101-compliant inferred mineral resource estimate was compiled by Fladgate Exploration Consulting Corporation and audited by Micon Internation Limited in April 2011. It was amended by Fladgate in August 2014 with the inaugural estimates on the No. 22 and 23 veins. The Pickle Crow project resource estimate is divided into three distinct areas within the core mine trend: the Shaft 1 area; the Shaft 3 area; and the Albany Shaft area. The drill hole database used for the resource estimation comprised drill holes, underground chip samples, and surface trench channel samples. In total, 1597 drill holes, totalling ~138,000m, were used, of which 167 drill holes (~50,000m) belonged to PC Gold drilling campaigns. A total of 27,826 chip samples taken by PCGM and 45 surface trench channel samples, taken by PC Gold, were used for estimation purposes. Tables 4 and 5 detail the resource estimate and Figure 15 is a longitudinal section showing resource areas for the Pickle Crow property. Table 4: Updated Pickle Crow Inferred Mineral Resources – August, 2014 Category Cut-off Grade (g/t Au) Tonnes Grade (g/t Au) Ounces Percentage of Total Ounces High-Grade Vein Underground 2.8 2,165,000 9.1 637,000 49% Bulk Underground* 2.0 4,510,000 3.7 536,000 41% Total Underground 2.25** 6,680,000 5.5 1,173,000 90% 0.35 3,628,000 1.1 126,000 10% 10,303,000 3.9 1,299,000 100% Total Open Pit Grand Total *Bulk Underground resources comprises primarily Banded Iron Formation (BIF) hosted mineralization.   **Represents a combination of potentially bulk mineable underground resources (2.0 g/t cut-off) and cut-and-fill underground resources (2.8 g/t cut-off, with vein intersections diluted to a minimum of 1 metre). - 99 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 5: Pickle Crow Updated Inferred Mineral Resources – August, 2014; Detailed Breakout by Shaft and Zone Area Shaft 1 Shaft 3 Albany Shaft Total Pickle Crow Zone Host Mining Technique Grade Au g/t Metric Tonnes Contained Ounces Au Cut-off Grade g/t BIF BIF & Vein Open Pit 1.1 3,628,000 126,000 0.35 BIF BIF & Vein Bulk Underground 3.7 4,320,000 508,000 2.0 No. 1 Vein Vein Underground 10.1 718,000 233,000 2.8 No. 5 Vein Vein Underground 5.2 141,000 24,000 2.8 No. 9 Vein Vein Underground 5.4 203,000 35,000 2.8 No. 11 Vein Vein Underground 6.5 18,000 4,000 2.8 No. 19 Vein Vein Underground 14.0 381,000 171,000 2.8 Shaft 1 Total Ounces 3.6 9,409,000 1,100,000 No. 2 Vein Vein Underground 9.1 96,000 28,000 2.8 No. 6 Vein Vein Underground 8.2 156,000 41,000 2.8 No. 7 Vein Vein Underground 5.8 49,000 9,000 2.8 No. 8 Vein Vein Underground 7.9 64,000 16,000 2.8 No. 12 Vein Vein Underground 11.9 14,000 5,000 2.8 No. 13 Vein Vein Underground 6.5 103,000 22,000 2.8 No.22 Vein* Vein Underground 5.9 28,000 5,000 2.8 No.23 Vein* Vein Underground 7.9 125,000 32,000 2.8 Shaft 3 Total Ounces 7.7 635,000 158,000 CZ1 Conduit-Style Bulk Underground 4.9 168,000 27,000 2.0 CZ3 Conduit-Style Bulk Underground 2.7 22,000 2,000 2.0 No. 15 Vein Vein Underground 4.7 42,000 6,000 2.8 No. 16 Vein Vein Underground 6.3 28,000 6,000 2.8 Albany Shaft Total Ounces 4.9 260,000 41,000 Total Open Pit 1.1 3,628,000 126,000 0.35 Total Bulk Underground 3.7 4,510,000 536,000 2.0 Total Underground 9.1 2,165,000 637,000 2.8 Grand Total 3.9 10,303,000 1,299,000   *2014 Resource Estimates Notes: 1. The mineral resource estimate is entirely classified as inferred mineral resources. 2. CIM Definition Standards were followed for mineral resources. 3. The cut-and-fill (highgrade vein) underground component of the mineral resource has been estimated at a cut-off grade of 2.8 g/t Au over a minimum width of 1 metre. Vein widths less than 1 metre were diluted to 1 metre prior to application of the 2.8 g/t Au cut-off grade. Grade and tonnes for the cut-and-fill component of the mineral resource are reported as diluted grade and tonnes. 4. The long-hole bulk underground (moderate-grade) component of the mineral resource has been estimated at a cut-off grade of 2.0 g/t Au. 5. The open pit (lowgrade) component of the mineral resource has been estimated at a pit discard cut-off grade of 0.35 g/t Au, using a preliminary Whittle pit shell to constrain the resource estimate and other assumed pit parameters. 6. The open pittable mineral resource extends to a depth of approximately 150 metres below surface. Only mineralization located within the pit shell has been reported at open pit cut-off grades. 7. The mineral resource has been estimated using a gold price of US$1,100 per ounce. 8. High-grade assays have been capped. Each domain was capped with respect to their unique geology and statistics. Caps for cut and fill (high-grade vein) underground resources range from 35 g/t to 145g/t. 9. Bulk density of 3.14 t/m3 was used for BIF and 2.70 t/m3 was used for veins. 10. The mineral resource was calculated via block model. Three dimensional wireframes were generated using geological information. A combination of Kriging and inverse distance estimation methods were used to interpolate grades into blocks of varying dimensions depending on geology and spatial distribution of sampling. 11. Mineral resources that are not mineral reserves do not have demonstrated economic viability. 12. Mineral resources have been adjusted for mined out areas. Small rib and sill pillars around old stopes have not been considered. 13. Numbers may not add due to rounding. - 100 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 15. Longitudinal section showing resource areas, Pickle Crow property Sub-stop Descriptions, Property (Fig. 16) Pickle Crow Deposit Stop 3c: Trench A (Figure 18) UTM Coordinates: NAD83; 15U 0705820E / 5711200N Stop 3a: PC Gold Inc. field office and core yard UTM Coordinates: NAD83; 15U 0703600E / 5709080N The geological setting of the Pickle Crow deposit along with various styles of mineralization will be discussed in detail at this stop. The property’s core library, which covers all rock types, alteration, and mineralization styles found on the property will be laid out for display and discussion. Stop 3b: Trench C UTM Coordinates: NAD83; 15U 0703890E / 5709300N This trench is an excellent example of the Core Mine Trend BIF-style mineralization (Fig. 17). The trench exposed the No. 5 vein in the northeastern corner and the No. 5 BIF zone in the centre of the trench, as well as extensive low-grade BIF mineralization. The highgrade, but narrow No. 11 vein is also exposed in the southeastern corner of the trench. It should be noted that although these structures appear to be relatively narrow and have a short strike-length, they have incredible down-plunge continuity, having been traced to more than 1150m at depth. This trench is the best exposure of what is referred to as the conduit-style mineralization. This mineralization is part of the high-grade vein Au-W style, but with some notable differences. In the Albany shaft area, the deformation of the vein is much more intense and hence continuity is diminished. In addition to this, the degree of alteration is also much more intense, with wide (tens of metres) alteration halos surrounding the veins. In the Albany area, the veins have been deformed to the point that they often form boudinaged fold hinges. These hinges have limited strike extent and complex internal structure, but with strong down-plunge continuity. Stop 3d: High-grade No. 1 Vein stockpile and Shaft Pit (Fig. 16) UTM Coordinates: NAD83; 15U 704355E / 5709860N This high-grade stockpile is composed of vein material from the No. 1 Vein crown pillar (Fig. 19). The larger stockpile behind it is composed of BIF-style mineralization which surrounds the No. 1 Vein, also from the crown Pillar. Abundant fine free gold can be observed along the chlorite-sericite-filled seams in the high-grade veins. Sheared, isoclinally folded (F2) BIF - 101 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 16. General geology of the southern half of the Pickle Crow property, showing veins, trenches and Field Trip Stop locations. is well-exposed in the nearby No. 1 Shaft Pit (Figure 20). Carbonate-altered lamprophyre dykes are also present in the No. 1 Vein surface stockpile and exposed in the No. 1 Shaft Pit. The 225tpd extreme gravity mill build by Cantera in 2002 lies a short distance away. STOP 4: Dona Lake Mine UTM Coordinates: NAD 83; 15 U listed with sub-stop descriptions (N.B. permission to enter the site must be granted by Goldcorp Canada Inc.) Figure 17. Stripped area at Trench C (Stop 3b), Pickle Crow property, showing contact between No. 5 Zone BIF and metavolcanic rocks. The past-producing Dona Lake Mine deposit, approximately 9km southeast of Pickle Lake, was discovered by Dome Exploration Limited in 1980 (Cohoon, 1986). Between 1980 and 1985 geophysical and geological surveys and diamond drilling were carried out by Dome Exploration. Advanced exploration, including sinking of a 176m exploration shaft, was conducted by Dome Mines and Campbell Red Lake Mines between 1985 and 1987. Developed jointly by Dome Mines and Campbell Red Lake Mines, it was put into production by Placer Dome Canada Limited - 102 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 18. General geology of the southern half of the Pickle Crow property, showing veins, trenches and Field Trip Stop locations. Figure 19. Broken chunks of No. 1 vein on stockpile at Stop 3d, showing typical laminated (“crack-seal”) structure (gloves for scale). Mill is in background. Figure 20. F2 fold in BIF, No. 1 Shaft pit, Stop 3d. Field of view is 2 metres. - 103 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 in February 1989 at a rated concentrator capacity of 550tpd, with Proven and Probable reserves totalling 754,000 tons (684,000 tonnes), averaging 0.24oz/t Au (8.23g/t Au; Coates and Anderson 2008). Placer Dome Canada operated the site until July 1993, at which time ownership was transferred to Ross-Finlay who operated the mine until closure in August 1994. The mine closed in 1994 due to exhaustion of viable reserves. During this time period, it is estimated that a total of 939,237 tonnes of ore was milled to produce approximately 6988kg (246,500 oz) of gold (Puumala 2009). This estimate is based on grade and tonnage estimates obtained from The Northern Miner (article, May 30, 1994), Janes et al. (1994) and Jen and McCutcheon (1995). The first source summarizes production under the ownership of Placer Dome from 1989 to 1993, and the second and third sources summarize production under the ownership of Ross Finlay from 1993 to 1994. At the time of mine closure, it is estimated that reserves of approximately 40,000 tonnes grading 5.95g/t Au remained at depth (Chronicle-Journal, article, June 24, 1994). Responsibility for the tailings reclamation remained with Placer Dome Canada which became Goldcorp Canada on May 12, 2006. Ochig Lake Pluton. The entire assemblage has been metamorphosed to amphibolite grade, as indicated by the presence of garnet porphyroblasts, biotite, local amphibolites with blue-green hornblende, and the relative lack of chlorite. The Dona Lake deposit is located in iron formation near the east-central portion of the property. The general geology in the immediate vicinity of the deposit, as derived from scattered outcrops, drilling and ground magnetic surveys, consists of tholeiitic basalt separated by several major units of iron formation, and intruded by felsic dykes and albite porphyry. The volcanics and sediments strike north-south to locally northwest-southeast and dip to the east and northeast at 60°. Tops, as determined from wellpreserved pillows, are also to the east. Several stages of deformation are evident in the area of the deposit: • The basalts, which are normally pillowed and massive, are very schistose and foliated in the vicinity of the iron formations. Some, but not all, of the felsic dykes are also affected by this foliation event. The Dona Lake Deposit is predominantly BIFhosted, geologically similar to the auriferous iron formation-hosted zones at the Central Patricia and Pickle Crow mines. The most comprehensive description was provided by Cohoon (1986): “The Dona Lake property is south of the previous producers [i.e. Pickle Crow, Central Patricia, etc.] in a separate greenstone sequence that trends south and merges with the OsnaburghPickle Lake belt [sic]. The main trend on the property is described by the nearly circular, 11kmlong arc of high magnetics which wraps around the tongue of the Ochig Lake Pluton. The high magnetics are caused by a major, semi-continuous unit and numerous minor discontinuous units of oxide iron formation. • Virtually all of the iron formation has been isoclinally folded. The fold planes are parallel to overall stratigraphy and the fold axes plunge east down the dip of the iron formation. These folds have wavelengths of about 1m and amplitudes of up to 10m. • Superimposed on the isoclinal folds with schistosity are low-amplitude cross folds with wavelengths of about 200m and amplitudes of 20m. • A stratigraphically controlled shear zone has been identified over a distance of 1km in the footwall of the iron formation. It is composed almost entirely of chlorite and presumably postdates the metamorphic event. The iron formations occur within a package of tholeiitic, usually pillowed basalt and amphibolite with local tuffs and minor felsic volcanics and clastic sediments. These units dip away from the pluton at a very consistent 60° and also young away from the pluton, suggesting a pre-erosion domal structure over the intrusive. All of the volcanic and sedimentary units have been intruded by sodium-rich felsic dykes and albite porphyry with a composition similar to the It is within the fold axis of one of these broadwavelength cross folds, in iron formation, that the Dona Lake gold deposit occurs. The spatial relationship is illustrated in [the Figure referred is not included in this guide; refer to Cohoon, 1986] which depicts the surface projection of the mineralization and the host iron formation. The other structural feature apparent in the diagram is a horizontal overlap of the iron formation of - 104 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 approximately 150m, caused by a north-south sinistral fault at a low angle to the stratigraphy. Vertical movement on this fault is unknown but is likely to exceed 300m. This feature may be related to the chloritic shear mentioned above. altered to pyrrhotite and/or grunerite. • Visible gold is common on the polished outer surface of core but has rarely been seen on broken surfaces. Visible gold occurrences usually correlate well with assay results. Gold mineralization occurs in both of the iron formations illustrated but the majority of the established reserves are in the one to the east. Mineralization has been defined over a strike length of 200m, but more than 80% of the reserves are in the central 100m in the core of the crossfold axis. With width exceeding 25m in the core, the overall geometry is that of a flattened pod, which plunges east down the axis of the crossfold, entirely within iron formation. • Gold mineralization is accompanied by traces of microscopic chalcopyrite and sphalerite. However, neither mineral was noted microscopically in drill core. Several features are notable by their absence: • There are virtually no quartz veins and the few which do occur seldom contain gold. There is also no visible or geochemically evident siliceous alteration, at least not within the basalts. Within the iron formation, variable quartz content and the possibility of remobilization of original chert makes identification of siliceous alteration difficult in drill core. ... The iron formations occur within the tholeiitic pillow basalt and amphibolite, which is strongly foliated in proximity to the mineralization. Intruding these units are numerous irregular albite-rich felsic dykes, albite porphyries and lamprophyre dykes. • There is no arsenopyrite and no evidence of geochemically anomalous arsenic. This situation should be contrasted with the other gold deposits in the Pickle Lake camp where vein quartz was usually the immediate host and arsenopyrite was often the main sulphide. The iron formations are usually classic oxidefacies iron formation, composed of finely bedded magnetite, chert and hornblende, with local grunerite, garnet, calcite and sulphides. A finely bedded chert-sulphide unit in the hanging wall has been genetically grouped with the iron-rich minerals other than sulphides and might be more properly termed a chert, since the sulphides may be secondary. Significantly, no carbonate iron formation, nor iron-rich carbonates, have been located on the property. Detailed studies concerning the genesis of the deposit are just beginning but, in general, it is assumed that the important factors relating to localization of mineralization are: • The iron formation which provided a competency contrast with the adjacent basalts and a chemically attractive site for the replacement of iron-rich minerals by iron sulphide. Limited polished section work on mineralized iron formation indicates that the magnetite beds are finely fractured and brecciated, leaving what would presumably be a very porous network for the movement of fluids. All of the gold occurs in oxide-facies iron formation. Mineralized sections display the following characteristics: • Between 5% and 15% pyrrhotite virtually always accompanies gold. Notably, this relationship does not apply to pyrite; even when pyrrhotite is abundant, if the pyrite content exceeds 3%-4%, gold values are usually low. The pyrrhotite is fine-grained and wispy, cross cutting bedding and apparently replacing or displacing other minerals. • Cross-folds which induced the fracturing in the iron formation and provided a locus for the entry and deposition of mineralized solutions. • In most cases, evidence of original bedding in drill core has been partially or completely destroyed. • The sodium-rich felsic dykes and albite porphyry, whose exact role is uncertain, but whose association with most gold deposits in the Canadian Shield empirically suggests that they are a necessary factor. • There is some evidence to suggest that the magnetite content is considerably reduced in mineralized sections, perhaps having been • Amphibolite-grade metamorphism, which - 105 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 may have resulted in ductile, as opposed to brittle, deformation. It is assumed that this is the reason for the amorphous zone of sulphide replacement rather than a distinct break or shear zone with quartz filling. All of the above hypotheses assume that the deposit is epigenetic, that the pyrrhotite mineralization is secondary and partially or completely replaced original iron-rich minerals, and that the gold mineralization was penecontemporaneous with the pyrrhotite. However, the sequence and mechanisms of deformation, and even the degree and nature of sulphide replacement, are open to further study. Outcrop is very sparse in the vicinity of the deposit, except for some of the hanging wall basalts and one small exposure of the unmineralized chert. [Exposures were subsequently created during mining and development activities.] Consequently, considerable reliance was placed on geophysics and geochemistry in the early stages of exploration, prior to detailed diamond drilling. Not surprisingly, magnetics were the most useful tool in defining the distribution and gross structural characteristics of the iron formation. However, very closely spaced readings on a 50m box grid were required to define the subtle variations in structure caused by the crossfold hosting the gold mineralization. Horizontal loop EM surveys with a 100m coil separation gave a weak but discernable response over the deposit. It is assumed that the response is weak due to the wispy and patchy nature of the sulphide mineralization, despite sulphide concentrations in excess of 15%. Surface soil geochemistry, involving both Cand A-horizon soils, failed to reveal the deposit. This failure is attributed to the uniform blanket of till which rests directly on [bed]rock. Rock geochemistry was somewhat more useful during initial stages of exploration prior to diamond drilling. Samples of basalt in the hanging wall returned up to 82ppb gold and more than half of the samples, in a radius of 200m, contain more than 10ppb gold. Background elsewhere on the property, including in iron formation, is in the order of 1ppb gold. The Dona Lake gold deposit is intimately associated with a folded, oxide-facies iron formation. It is the iron formation which played a crucial role in providing a structural and chemical trap for mineralizing solutions.” Structural setting for gold mineralization East of the Ochig Lake pluton, in the vicinity of the Dona Lake Mine, Stott and Corfu (1991) described shear zones with normal sense movement that are parallel and dip away from the margin of the pluton and interpreted these shear zones as related to pluton emplacement. East of the northeastern lobe of the Ochig Lake pluton, within a mapped contact strain domain, in the vicinity of the Dona Lake area, Stott (1996) documented margin-parallel shear zones dipping away from the intrusion, showing pluton-side-up kinematics. A discussion of gold in high-strain zones associated with pluton emplacement was given in Smyk et al. (2011). Stott and Biczok (2010) had related the structural geology of Musselwhite Mine (125km north of Pickle Lake) to the emplacement of the crescentshaped North Caribou pluton and revisited the idea (Stott et al., 1989) that gold mineralization there may also be related to the intrusion of this pluton. In this model, lateral compression imparted by the intrusion of the pluton produces shallowly plunging, isoclinal and locally transposed folds, and high-strain or shear zones in relatively incompetent lithologic layers that fold more tightly. These narrow high-strain/shear zones reflect coaxial strain and would not be related to any through-going, transcurrent, regional shear zone. Any non-coaxial shear features on folded limbs can be accounted for by subjecting rocks to the limit of flattening and by accommodating the extreme shortening by rotation and lateral shear. These zones, however, may focus hydrothermal fluid flow and host fault-fill veins (cf. Dubé and Gosselin, 2007). Many such zones are characterized by replacement and/or alteration zones, rather than discrete vein systems. The mineralized zones that result from such focussed fluid flow are localized and may only represent a very small proportion of a favourable host lithology (e.g. banded iron formation). Gold mineralization at Dona Lake is manifested as a flattened, plunging pod, hosted by a folded, sulphidized banded iron formation, largely in the core of a crossfold axis (Cohoon, 1986). There is a notable lack of brittle features, quartz veining and a mineralized shear zone or ‘break”, ascribed in part to the relatively high grade (amphibolite facies) of metamorphism. The deposit is situated in the amphibolite-facies contact - 106 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 metamorphic and strain aureole surrounding the Ochig Lake pluton (Stott, 1996). This pluton (2741Ma) has post-dated and overprinted the regional penetrative deformation fabrics with its own contact strain fabric (cf. Young et al. 2006). Sub-stop Descriptions, Dona Lake Gold Mine (Fig. 21): Stott (1996) and Stott and Smith (1988) listed several different settings for gold deposits directly attributed to felsic pluton emplacement: UTM Coordinates: NAD83; 15U 0702099E / 5699732N • Type A: A set of shear or high-strain zones bordering an uplifted and rotated wedge of supracrustal rock; • Type B: Shear or high-strain zone defining the outer margin of a pluton-induced strain aureole; • Type C: Shear or high-strain zone focussed along a contrast in rock ductility (e.g., BIF-basalt; Dona Lake Mine, Musselwhite Mine); • Type D: Shear or high-strain zones in conjugate sets at small angles to the schistosity; this strain environment may correspond to regional orogenic shortening or to a pluton-induced strain aureole; (e.g., Pickle Crow Mine); and • Type E: Shear or high-strain zones, tangential to a strain aureole, which may have been initiated during regional orogenic shortening and locally reactivated. Stop 4a: Dona Lake Mine Portal area The rehabilitated site of the former Dona Lake Mine is now marked by a grassy field along the sides of the access road, south of Sika Pond. The mine portal, now sealed, lies just south of the road. Foliated and folded mafic metavolcanic and gabbroic rocks outcrop in the vicinity of the site and along shoreline exposures on Sika Pond. The exposures near the reclaimed mine site are dominated by gabbro and pillowed basalt of the lower sequence of the Pickle Crow Assemblage with local banded iron formation (Young 2003). Confederation assemblage mafic volcanic rocks occur east of the Dona Lake Mine area. Well-preserved pillowed flows east of the mine area indicate east-southeast younging (Young 2003). Young (2003) has noted that in the Dona Lake area, gabbro is common and has a weathered surface characterized by a ‘knobby’ texture caused by coarse-grained hornblende and anastomosing fractures superficially resembling pillowed basalt (Fig. 22). Figure 21. Dona Lake Mine area with Field Trip Stop locations. - 107 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 basalt (Fig. 24) is situated on the west side of the mine road, across from where a pit had been developed in BIF. Bun pillows and brecciated pillow fragments have thin, dark, amphibole-rich selvages. Figure 22. Coarse-grained, amphibolitized gabbro, south of Sika Pond (Stop 4a). Stop 4b: Banded Iron Formation UTM Coordinates: NAD83; 15U 0701767E / 5699990N; 0701708E / 5699991N, and 0701764E / 5699994N Banded iron formation outcrops (Fig. 23) in a number of locations (see three sets of UTM coordinates above) along a mine access road that skirts the western shore of Sika Pond and accesses a flooded mine pit. A southeasterly trending foliation and M- and Z-folds are developed in these BIF units. A rusty BIF sample collected by OGS staff returned <0.01 z/ton Au and 0.3 oz/ton Ag (unpublished data, Resident Geologist’s Files, Thunder Bay) Figure 24. Pillowed basalt, west side of road, near pit, west of Sika Pond (Stop 4c). Stop 5: Ochig Lake Pluton Stop 4c: Pillowed Basalt UTM Coordinates: NAD83; 15U 0693228E / 5693898N. UTM Coordinates: NAD83; 15U 0701686E / 5700056N A small outcrop of relatively undeformed pillowed Figure 23. Rusty, folded BIF near pit, west of Sika Pond (Stop 4b). One of the few exposures of the Ochig Lake Pluton occurs on the eastern side of Highway 599, south of Fault Lake. These low-lying outcrops expose grey, medium-grained, equigranular biotite granodiorite with local, lenticular quartz-filled gashes (Fig. 25). One of the internal granitoids in the Pickle Lake belt, the Ochig Lake Pluton is semi-circular in plan, with a domical internal structure defined by outward-dipping foliation (Stott 1996). A northeasterly trending, steeply southeast-dipping foliation was noted by Stott et al. (1989a) just northeast of this location. Most of the pluton consists of homogeneous, medium- to finegrained granodiorite to trondhjemite. The Ochig Lake pluton yielded a U-Pb zircon age of 2741±2Ma (Corfu and Stott 1993a). - 108 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 mineralization within the Red Lake mine trend: example from the Cochenour–Willans mine area, Red Lake, Ontario, with new key information from the Red Lake mine and potential analogy with the Timmins camp. Geological Survey of Canada, Current Research 2003-C21. Evans, J.E.L. 1941. Geology of the Eastern Extension of Crow River area; Ontario Department of Mines, Annual Report, v.48, pt.7, 9p. Ferguson, S. A., 1966. Geology of Pickle Crow Gold Mines Limited and Central Patricia Gold Mines Limited, No. 2 Operation”; Ontario Department of Mines, Miscellaneous Paper 4. Figure 25. Quartz veins and lenses in equigranular granodiorite of the Ochig Lake Pluton, Highway 599 (Stop 5). References Anderson, S.D. 2007. Thierry Mine project, Pickle Lake belt, Ontario: Observations pertaining to the structural geology of the Thierry Cu-Ni-(PGE) deposit; unpublished report, Richview Resources Inc., 39p. Coates, H. and Anderson, W., 2008. Technical Report on the Pickle Crow Gold Property, Patricia Mining Division, Ontario, Canada; unpublished report, PC Gold Inc. Cohoon, G.A. 1986. Gold in an iron formation, the Dona Lake deposit; The Northern Miner Magazine, v.1, no.8, p.16-20. Corfu, F. and Stott, G.M. 1989. U-Pb geochronology of the central Uchi Subprovince, northwestern Ontario; p.A55 in Geological Association of Canada, Program with Abstracts, v.14, 140 p. Corfu, F., and Stott, G.M. 1993a. Age and petrogenesis of two late Archean magmatic suites, northwestern Superior Province, Canada: zircon U–Pb and Lu–Hf isotopic relations; Journal of Petrology, v.30, p.1179– 1196. Corfu, F., and Stott, G.M. 1993b. U–Pb geochronology of the central Uchi subprovince, Superior Province. Canadian Journal of Earth Sciences, v.78, p.53–63. Corfu, F., and Stott, G.M. 1996. Hf isotopic composition and age constraints on the evolution of the Archean central Uchi subprovince, Ontario, Canada. Precambrian Research, v.78, p53–63. Dubé, B. and Gosselin, P. 2007. Greenstone-hosted quartzcarbonate vein deposits; in Mineral deposits of Canada: A synthesis of major deposit-types, district metallogeny, the evolution of geological provinces, and exploration methods, Geological Association of Canada, Mineral Deposits Division, Special Publication No. 5, p.49-73. Dubé, B., Williamson, K., and Malo, M. 2003. Gold Gilligan, L.B. and Marshall, B. 1987. Textural evidence for remobilization in metamorphic environments; Ore Geology Reviews, v.2, p.205–229. Graham R. J., 1965. Unpublished letter to Dr. S. A. Ferguson, Ontario Department of Mines, Pickle Crow Gold Mines internal report. Harding, W.D. 1936. Geology of the Cat River–Kawinogans Lake area; Ontario Department of Mines, Annual Report, 1935, v.44, pt.6, p.53-75. Hennessey, B.T., San Martin, A.J. and Shoemaker, S.J. 2011. A Mineral Resource Estimate for the Pickle Crow property, Patricia Mining Division, northwestern Ontario, Canada; unpublished report, PC Gold Inc., 209p. Hollings, P. 1998. Geochemistry of the Uchi subprovince, northern Superior Province: an evaluation of the geodynamic evolution of the northern margin of the Superior Province ocean basin. Unpublished Ph.D. thesis, University of Saskatchewan, Saskatoon, Saskatchewan. Hollings, P., Stott, G. and Wyman, D., 2000. Trace element geochemistry of the Meen-Dempster greenstone belt, Uchi subprovince. Canadian Journal of Earth Sciences, 37, 1021-1038. Hollings, P. 2002. Archean Nb-enriched basalts in the northern Superior Province; Lithos, v.64, p.1–14. Hollings, P. and Kerrich, R. 1999. Trace element systematics of ultramafic and mafic volcanic rocks from the 3Ga North Caribou greenstone belt, northwestern Superior Province. Precambrian Research, v.93, p.257–279. Hollings, P. and Kerrich, R. 2004. Geochemical systematics of tholeiites from the 2.86Ga Pickle Crow assemblage, northwestern Ontario: arc basalts with positive and negative Hf–Nb anomalies; Precambrian Research, v.134: p.1–20. Hollings, P., Wyman, D.A., and Kerrich, R. 1999. Komatiitebasalt-rhyolite volcanic associations in northern Superior Province greenstone belts: Significance of plume-arc interaction in the generation of the protocontinental Superior Province; Lithos, v.46: p.137– 161. Hurst, M.E. 1931. Pickle Lake-Crow River area; Ontario - 109 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Department of Mines, Annual Report, 1930, v.39, pt.2, p.1-35. Janes, D.A., Seim, G.W., Hinz, P. and Storey, C.C. 1994. Sioux Lookout Resident Geologist’s District—1993; in Report of Activities 1993, Resident Geologists, Ontario Geological Survey, Open File Report 5892, p.106-128. Jen, L. and McCutcheon, B. 1995. Principal Canadian nonferrous and precious metal mine production in 1994; Canadian Minerals Yearbook, 1995, p.68.1 to 68.7. Keller, G.D. 2005: Technical Report, Thierry Deposit, Pickle Lake, Ontario; SRK Consulting (Canada) Inc., 150 p. Kolb, M. J., 2011. PC Gold thin section report, Central Patricia East, unpublished, internal company report, PC Gold Inc. Lynch, T. 2010. Technical Report for MNDM Assessment Purposes: PC Gold - Pickle Lake Property. Ontario Ministry of Northern Development, Mines and Forestry, unpublished Assessment Report, Resident Geologist’s Files, Thunder Bay. MacGregor, H. 1989. Report on Mineral Reserves at the Pickle Crow Property dated January 20, 1989; unpublished company report for Noramco Mining Corporation, by Watts Griffis and McOuat Limited. MacQueen, J. K. 1987. Stratigraphy, structure and gold mineralization of the No.5 Vein/Iron Formation Zone, Pickle Crow Gold Mines, Pickle Lake, Ontario. Unpublished M.Sc. thesis, Carleton University, Ottawa, Ontario. Marshall, B. and Gilligan, L.B. 1989. Durchbewegung structure, piercement cusps and piercement veins in massive sulfide deposits: formation and interpretation; Economic Geology, v.84, p.2311–2319. McInnes, W. 1906. The headwaters of the Winisk and Attawapiskat rivers; in Summary report of the Geological Department of Canada for the calendar year 1905, Geological Survey of Canada; Geological Survey of Canada, Summary Report (1905), 1906; p.76-80. Naldrett, A.J. and Cabri, L.J. 1976. Ultramafic and related mafic rocks; their classification and genesis with special reference to the concentration of nickel sulfides and platinum-group elements. Economic Geology, v.71, p.1131-1158. Novak, N. and Mlot, S. 2004. Technical report on the geology and mineral resources of the Thierry Copper-Nickel project, Pickle Lake, Ontario; Internal Consultants Report, 38 p. Patterson, G.C. 1980. The Geology of the Kapkichi Lake ultramafic-mafic bodies and related Cu-Ni mineralization, Pickle Lake, Ontario; unpublished Ph.D. thesis, Carleton University, Ottawa, Ontario. Patterson, G.C. and Watkinson, D.H. 1984. The geology of the Thierry Cu-Ni mine, northwestern Ontario; Canadian Mineralogist, v.22, p.3–11. Puritch, E., Armstrong, T., Burga, D., Routledge, R., Pearson, J.L., Hayden, A., Orava, D. and Rodgers, K. 2012. Technical report and Preliminary Economic Assessment of the Thierry and K1-1 Cu-Ni-PGE deposits, Thierry Project, Pickle Lake area, Patricia Mining District, northwestern Ontario, Canada; unpublished company report, Cadillac Ventures Inc., 230p. Puritch, E., Ewert, W.D., and Armstrong, T. 2006. Technical report and resource estimate on the Thierry Cu-NiPGE mine property, Pickle Lake area, Patricia mining district, northwestern Ontario, Canada; P& E Mining Consultants Inc., 161 p. Puumala, M.A. 2009. Mineral occurrences of the central and eastern Uchi domain; Ontario Geological Survey, Open File Report 6228, 294p. Pye, E.G. 1956. “Geology and Mineral Deposits of the Crow River Area”, Ontario Department of Mines, p.1-239. Pye, E.G. 1975. Crow River area, District of Kenora (Patricia Portion); Ontario Geological Survey, Preliminary Map P.1009, scale 1:12,000. Pye, E.G. 1976. Geology of the Crow River area, District of Kenora (Patricia Portion); Ontario Geological Survey, Open File Report 5152, 264p. Sage R.P. and Breaks, F.W. 1982. Geology of the Cat Lake -Pickle Lake area, Districts of Kenora and Thunder Bay; Ontario Geological Survey, Report 207. Sajona, F. G., Maury, R. C., Bellon, H., Cotten, J. and Defant, M. J. 1996. High field strength element enrichment of Pliocene-Pleistocene island arc basalts, Zamboanga Peninsula, western Mindanao (Philippines). Journal of Petrology, v.37, p.693–726. Sanborn-Barrie, M., Skulski, T., and Parker, J. 2001. Three hundred million years of tectonic history recorded by the Red Lake greenstone belt, Ontario. Geological Survey of Canada, Current Research 2001-C19. Smyk, M.C., White, G.D., Lockwood, H.C., and Bennett, N.A. 2011. Report of Activities 2010, Resident Geologist Program, Thunder Bay North Regional Resident Geologist Report: Thunder Bay North District; Ontario Geological Survey, Open File Report 6262, 47p. Stott, G.M. 1996. The geology and tectonic history of the central Uchi subprovince; Ontario Geological Survey Report 5952, 178p. Stott, G.M. and Biczok, J. 2010. North Caribou greenstone belt: gold and its possible relation to the North Caribou pluton emplacement—a belt-wide contactstrain aureole?; in Summary of Field Work and Other Activities 2010, Ontario Geological Survey, Open File Report 6260, p.22-1 to 22-12. Stott, G.M., Corfu, F., Breaks, F.W., and Thurston, P.C. 1989. Multiple orogenesis in northwestern Superior Province; Geological Association of Canada– - 110 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Mineralogical Association of Canada, Joint Annual Meeting, Program with Abstracts, v.14, p.A56. Stott, G.M., Brown, G.H., Coleman, V.J., Green, G.M., and Reilly, B.A. 1989a. Precambrian geology of the Pickle Lake area, western part; Ontario Geological Survey, Preliminary Map P.3056, scale 1:50,000. Stott, G.M., Brown, G.H., Coleman, V.J., Green, G.M., and Reilly, B.A. 1989b. Precambrian geology of the Pickle Lake area, eastern part; Ontario Geological Survey, Preliminary Map P.3057—Revised, scale 1:50,000. Stott, G.M., and Corfu, F. 1991. Uchi subprovince. Geology of Ontario, Ontario Geological Survey Special Vol. 4, Part 1. Ontario Geological Survey, Ontario, p.144– 236. Stott, G.M. and Smith, P.M. 1988. Development of goldbearing structures in the Archaean: the role of granitic plutonism; in Bicentennial Gold 88, extended abstracts, poster programme, v.1, Geological Society of Australia, Abstracts Series no.23, p.48-50. Sun, S.S. and McDonough, W.F. 1989. Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes; in Magmatism in the ocean basins. Saunders, A.D. and Norry, M.J. (Editors), Geological Society of London, London; v.42, p.313-345. Thomson, J.E. 1939. The Crow River area; Ontario Department of Mines, Annual Report, 1938, v.47, pt.3, p.1-65. Thurston, P.C., Osmani, I.A., and Stone, Northwestern Superior Province: review analysis. Geology of Ontario, Ontario Survey, Special Vol. 4, Part 1. Ontario Survey, Ontario, p.80–142. D. 1991. and terrane Geological Geological Tomlinson, K.Y., Stevenson, R.K., Hughes, D.J., Hall, R.P., Thurston, P.C., and Henry, P. 1998. The Red Lake greenstone belt, Superior Province: evidence of plume-related magmatism at 3Ga and evidence of an older enriched source; Precambrian Research, v.89, p.59–76. White, D.J., Musacchio, G., Helmstaedt, H., Harrap, R.M., Thurston, P.C., van der Velden, A., and Hall, K. 2003. Images of a lower crustal slab: Direct evidence for tectonic accretion in the Archean western Superior Province; Geology, v.31, p.997-1000. Winter, L.D.S. 1987. Geological Reserve (Mineral Inventory) Estimate, Pickle Crow Property, Pickle Lake, Ontario; unpublished company report for Highland Crow Resources. Winter, L.D.S. 1988. Geological Reserve (Mineral Inventory) Estimate, Pickle Crow Property, Pickle Lake, Ontario; unpublished company report for Noramco Mining Corporation. Young, M. 2003. New structural, geochronological and geochemical constraints on the tectonic assembly of the Archean Pickle Lake greenstone belt, Uchi subprovince, western Superior Province; Unpublished M.Sc. thesis, Queen’s University, Kingston, Ont. Young, M. and Helmstaedt, H. 2001. Tectonic evolution of the northern Pickle Lake greenstone belt, northwestern Superior Province, Ontario. Geological Survey of Canada, Current Research 2001-C20. Young, M.D., McNicoll, V., Helmstaedt, H., Skulski, T., and Percival, J.A. 2006. Pickle Lake revisited: new structural, geochronological and geochemical constraints on greenstone belt assembly, western Superior Province, Canada; Canadian Journal of Earth Sciences, v.43, p.821–847. Tigert, T. T. 1949. Geology of the Central Patricia Mine; Canadian Mining Journal, p.72-75. - 111 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 9 - The Ghost Lake Batholith and Related Pegmatites Shannon E. Zurevinski Dept. of Geology, Lakehead University, Thunder Bay, Ontario, Canada Introduction The Late Archean (2685 Ma) Ghost Lake Batholith (GLB) is a roughly 280km2 elongated intrusion which trends east to northeast, following regional structures (Fig. 1). The GLB is exposed from Eagle River, northeast to Ghost Lake, and is classified as a two mica granitoid (biotite and muscovite). Mineralogy of the peraluminous fertile S-type granitoid includes cordierite, sillimanite, Mg-garnet, tourmaline, beryl, and rare dumortierite (a fibrous borosilicate). The batholith shows trends of increasing geochemical fractionation, and as a result, is host to exotic rare element pegmatite occurrences, such as the Mavis Lake pegmatite group. pegmatites can be difficult when pegmatites are not found in association with the fertile primary parent granite. This is further hindered when the melts separate and mix with the crustal material during ascent and emplacement, producing hybrid geochemical and mineralogical signatures. In contrast, the Ghost Lake Batholith represents an intrusive complex where the rare element pegmatite facies exists within its consanguineous, primitive peraluminous granite parent (Breaks and Moore, 1992). This relationship shown in outcrop presents an opportunity to assess mechanisms that concentrate rare elements in pegmatites. For this reason, the Ghost Lake Batholith has been the subject of much past academic research, and likely will continue to be for years to come. Identifying the initial sources and processes that concentrate lithophile metal enrichment in rare element This ½ day field trip will examine granite and pegmatitic outcrop from the Ghost Lake Batholith, Figure 1. Generalized map showing the Sioux Lookout Terrane boundary and the location of the Ghost Lake Batholith and the Mavis Lake pegmatite group (from Brand et al., 2009). - 112 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 and two associated pegmatites, both of the Mavis Lake Pegmatite Group. This field guide builds upon those previously written and compiled by Breaks (1982), Breaks (1985), Breaks and Janes (1991) and Beakhouse et al. (1995). One stop is a road cut, so caution must be exercised. Never stand on the paved portions of the road, and be aware of the traffic at all times. Road intersection. Regional Geology This road cut is located in the western part of the batholith and features different plutonic phases. Most of the outcrop at this location was termed GLB-1 by Breaks and Janes (1991). GLB-1 is one of the most common granitic units of the batholith. It is an inequigranular, coarse grained, massive to weakly foliated biotite and cordierite-biotite granite. Zones of inhomogenous biotite granite with inclusions of metasedimentary rocks and mafic segregations occur throughout GLB-1. Biotite and cordierite-biotite pegmatitic leucogranite segregations occur sporadically throughout the unit. Breaks and Janes (1991) document the pegmatitic masses as consisting of blocky, graphic K-feldspar-quartz masses up to 60cm wide, grading into smoky quartz segregations in association with blocky K-feldspar, biotite, and apatite. Cordierite can be seen occurring as square megacrysts, but it appears more often as dark masses that are up to 1.5cm in diameter. It can be replaced by chlorite, muscovite, and andalusite. Accessory minerals include garnet, apatite, zircon, monazite, and sillimanite. Secondary muscovite occurs after K-feldspar, plagioclase (An10-30), biotite and cordierite. Other plutonic phases at this location are dikes of fine-grained, grey, muscovite-biotite granite and coarse-grained, white biotite-muscovite leucogranite and pegmatite segregations. The GLB and related Mavis Lake Pegmatite group occur within the Winnipeg River–Wabigoon Subprovince boundary zone, referred to as the Sioux Lookout Terrane (SLT) (Beakhouse, 1989). The SLT is characterized by metasedimentary and metavolcanic assemblages exhibiting a wide range of metamorphic grade, including zones of migmatized metasedimentary rocks (Breaks and Moore, 1992). This terrane is host to about 150km of peraluminous granite plutons, with the GLB being the largest of the group. Geochemically, the plutons of the SLT exhibit enrichment in lithophile elements and rare-earth elements (REE’s), including Cs, Be, Li, Rb, Nb, Ta, Ga, U, Th, Mo, W, and Sn. The GLB is broadly concordant to the eaststriking foliations and is weakly foliated. Evidence of the foliations is shown by preferred orientations of phyllosilicates, and biotite-cordierite-sillimanite segregations. Foliation diminishes towards the eastern segment of the batholith. Multiple events of ductile deformation in the GLB are described by Breaks and Moore (1982), and will be discussed at each field stop. The GLB is subdivided into two portions comprising the Western GLB and the Eastern GLB. The Western GLB is also known as the lower intrusion, where the granitoids are less homogenous, and form transitional contacts with the migmatized sedimentary host. In the Eastern GLB, also known as the upper intrusion, the contacts with the metasedimentary rocks are abrupt and contain fewer sedimentary inclusions. Traversing from the lower intrusion through to the upper intrusion, there are marked trends of geochemical fractionation involving enrichment of B, Be, Ga, Li, Nb, and Rb; and depletion of Ba, Sr, Zr, and total REE. Ghost Lake Batholith stops Starting at the parking lot of the Best Western Hotel and Conference Center, Dryden, drive 13km west along the Trans-Canada Highway 17 to the village of Oxdrift. Park vehicles in the old Oxdrift School parking lot and walk to the road cut located just west of the Corner Stop 1: GLB-1, Oxdrift School Stop, Muscovitebiotite granite, biotite-muscovite granite, rare granodiorite (± cordierite). UTM Coordiantes: NAD 83 15U 0499697E / 5518078N Mavis Lake Group Pegmatites stops Drive east from Stop 1 on Highway 17 back through Dryden for approximately 16.4km until you reach Airport Road. Turn left onto Airport Road and travel north for 4.8km to Ghost Lake Road. Turn right on Ghost Lake Road and follow the main road to the junction with a logging road at UTM coordinates 527090E, 5522510N, a distance of approximately 11.6km from Airport Road. Turn south and drive along this main road to an intersection with a secondary logging road at UTM coordinates 526970E, 5521410N, a distance of about 1.4km. Veer west along this secondary road and up the hill where several vehicles can park. Most of the better outcrops at this stop are located south of the logging road. - 113 - Beakhouse et al. (1995) describe the Mavis Lake �Proceedings of the 61st ILSG Annual Meeting - Part 2 Pegmatite Group as an east-striking, 8km by 1.5km, concentration of rare-element pegmatites and related metasomatic zones. With increasing distance from the parent Ghost Lake Batholith, the pegmatite mineral assemblages exhibit well-defined zonations. More than 12 distinctive spodumene pegmatites are described from the Mavis Lake area and each ranges in size from 3 by 15m to 15 by 280m. Occurring parallel to the foliation of the host metavolcanic rocks, the pegmatites intrude as roughly lensoidal bodies. The rare-element pegmatites are classified as the albite-spodumene type (after Černy, 1982). Mavis Lake Group Pegmatites are structurally confined to the northern limb of the westplunging Thunder Lake Syncline and show localized effects of late-tectonic deformation, for example: strained contacts; healed fractures involving tourmaline and spodumene buckling; and boudinage of pegmatitic granitic dykes near the contact zones. STOP 2: Dryden Airport Pegmatite UTM Coordinates: NAD 83 15U 0526720E / 5521360N This stop shows pegmatitic-aplite dykes intruding metasedimentary rocks (Fig. 2). This dyke is described as a barren potassic pegmatite with a characteristic mineral assemblage including garnet, muscovite, biotite, abundant coarse tourmaline, albite, quartz, blocky microcline (Or70-77Ab23-30), and rare, large (up to 10cm in length) lime-green beryl. Remnant rock saw cuts represent poached rare beryl. Road. Turn down Mavis Forest Road and travel 2.6km and park to your right at a small BMX trail entrance (just after the powerline). Follow the trail in and veer to your first left. Follow the trail until you reach a fork in the trail and then veer right and follow the intersection of the rutabaga trail and out onto a large ridge. Follow this around to the pegmatite clearing. STOP 3: Fairservice Spodumene-Beryl-Tantalite Pegmatites UTM Coordinates: NAD 83 15U 0523789E / 5518050N The Fairservice pegmatite intrudes foliated and gneissic mafic metavolcanic rocks and fine-grained laminated metagreywackes. There is some vague internal zonation - in which this particular occurrence is representative of a quartz-rich core and spodumenerich, albite-quartz pegmatite. This pegmatite hosts a randomly oriented, green primary spodumene, large coarse quartz pod-like segregations, thick yellowgreen muscovite, albite, apatite, lesser white and blue beryl, large black tourmaline (Fig. 3), blue apatite, and orange garnet in a matrix of light grey massive quartz. A sharp contact with laminated metagreywackes is noted. Healed fractures of tourmaline are shown at this occurrence. References Return to the vehicles and drive back to Highway 17. Turn right (east) onto the highway and drive 3.7km until you reach Thunder Lake Road. Turn left onto Thunder Lake Road and travel 2.7km to Mavis Forest Beakhouse, G.P., Blackburn, C.E., Breaks, F.W., Ayer, J., Stone, D., and Stott, G.M. 1995. Western Superior Province Fieldtrip Guidebook. Precambrian 1995 Meeting. Geological Survey of Canada Open File 3138/Ontario Geological Survey Open File Report 5924. Figure 2. Dryden Airport Pegmatite in contact with metasedimentary rocks. Figure 3. Fractured tourmaline from the Fairservice Occurrence, Mavis Lake Pegmatite Group - 114 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Brand, A.A., Groat, L.A., Linnen, R.L., Garland, M.I., Breaks, F.W. and Giuliani, G. 2009. Emerald mineralization associated with the Mavis Lake pegmatite group, near Dryden, Ontario. Canadian Mineralogist, 47, p.315-336. Breaks, F.W. and Moore, J.M, Jr. 1992. The Ghost Lake Batholith, Superior Province of Northwestern Ontario: A fertile, S-type, peraluminous graniterare-element pegmatite system: The Canadian Mineralogist, v.30, p.835-875. Breaks, F.W. 1989. Origin and evolution of peraluminous granite and rare-element pegmatite in the Dryden area of Northwestern Ontario; unpublished PhD thesis, Carleton University, Ottawa, Ontario, 594p. Breaks, F.W., Selway, J.B., and Tindle, A.G. 2001. Fertile peraluminous granites and related rare-element mineralization in pegmatites, Superior Province, Northwest and Northeast Ontario in Summary of Field Work and Other Activities, 2001, Ontario Geological Survey, Open File Report 6070, p.39-1 to 39-9. Breaks, F.W. and Janes, D.A. 1991. Granite-related mineralization of the Dryden area, Superior Province of Northwestern Ontario. GAC-MAC-SEG Joint Annual Meeting, 1991, Field Trip B7 Guidebook, 71p. Breaks, F.W. and Kuehner, S. 1984. Precambrian geology of the Eagle River-Ghost Lake area, Kenora District; Ontario Geological Survey, Map P.2623, 1:31,680. Černy, P. 1982. Petrogenesis of granitic pegmatites. In Granitic Pegmatites in Science and Industry (P. Černy editor) Mineralogical Association of Canada, Short Course Handbook 8, p.405-461. - 115 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Field Trip 10 - Mattabi/Sturgeon Lake Historic VMS Camp George J. Hudak Minerals Division, Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN Introduction showings. Detailed field, petrographic, and lithogeochemical studies performed since the mid-1980’s have indicated that the south Sturgeon Lake region of northwestern Ontario is underlain by an extremely well-preserved, though partially eroded, Neoarchean subaerial to submarine volcanic caldera complex (Fig. 1: Morton et al., 1991; Hudak et al., 2003). This caldera complex, the Sturgeon Lake Caldera Complex (SLCC), hosted six massive sulphide orebodies which produced nearly 20 million tons of polymetallic volcanogenic massive sulfide (VMS) ore during mining operations between 1972 and 1991 (Table 1), as well as numerous sub-economic massive and semi-massive sulphide The combination of well-preserved volcanic and hydrothermal alteration textures, the variable 55° to 90° dip of a north-facing, essentially homoclinal volcanic sequence, and more than 600,000m of diamond drilling over an apparent 4,500 meter stratigraphic interval has enabled geologists the opportunity to examine the lithological, lithogeochemical, and metallogenic evolution of the SLCC. This includes the synvolcanic subvolcanic intrusions (Biedelman Bay Intrusive Complex), initial subaerial to shallow subaqueous precaldera volcanism (Pre-caldera Sequence, PCS), early subaerial to submarine explosive silicic volcanism (Early Caldera Sequence, ECS) that is associated Figure 1. Location map (inset) and regional geological map of the south Sturgeon Lake region (modified after Trowell, 1983; Morton et al., 1991; Morton et al., 1999; Galley et al., 2000). - 116 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Table 1. Grade and tonnage statistics from the VMS ore deposits of the south Sturgeon Lake region (after Franklin, 1996). Caldera sequences include the Early Caldera Sequence (ECS) and the Late Caldera Sequence (LCS; Hudak, 1996; Morton et al., 1999; Hudak et al., 2003). F-Group Caldera Sequence ECS 106 Tons 0.34 Cu % 0.64 Zn % 9.51 Pb % 0.64 Ag (g/t) 60.4 Ag (g/t) 2.13 Mattabi ECS/LCS 12.55 0.74 8.28 0.85 104.0 3.67 Sturgeon Lake LCS 2.07 2.55 9.17 1.21 164.2 5.79 Creek Zone/Sub-Creek Zone* Lyon Lake* LCS 0.91 1.66 8.80 0.76 141.5 4.99 LCS 3.95 1.24 6.53 0.63 141.5 4.99 19.82 1.06 8.50 0.91 119.7 3.85 Deposit Total/Average * VMS deposits interpreted to have originally formed as part of Sturgeon Lake deposit with the development a multi-cyclic, piecemeal caldera complex up to 25km in strike length, and later intracaldera submarine mafic to felsic, dominantly effusive submarine volcanism and clastic and chemical sedimentation (Late Caldera Sequence, LCS). Postvolcanic structural deformation has not only led to the apparent transposition of non-caldera associated strata (Lyon Lake Fault Sequence, LLFS) up-section from the caldera-associated volcanic sequence, but has locally led to remobilization of massive sulphide mineralization along a major northwest-southeasttrending fault zone to form smaller, but locally economic, massive sulphide deposits (e.g., the Lyon Lake, Creek Zone, and Sub-Creek Zone deposits). The stratigraphic evolution in the SLCC represents one of the few examples of Neoarchean calderas which illustrate the modern “caldera cycle” (Smith and Bailey, 1968; Hudak et al., 2003; Mueller et al., 2004; Mueller et al., 2008). This field trip has several purposes: 1) to illustrate the textures, lithologies, and geological structures resulting from various volcanological and sedimentological processes associated with various stages of caldera complex development; 2) to illustrate the hydrothermal alteration mineral assemblages within the caldera complex, and their spatial relationships to volcanogenic massive sulphide (VMS) mineralization; and 3) to illustrate methods by which exploration geologists can use physical volcanology, hydrothermal alteration mineral assemblages, and structural geology, along with other lithogeochemical and geophysical data, to effectively explore for VMS deposits in the Sturgeon Lake camp and elsewhere. This guidebook represents an updated version of previous field trip guidebooks for the Sturgeon Lake region that were associated with the Geological Association of Canada – Mineralogical Association of Canada meeting during May, 1996 (Morton et al., 1996), as well as the Precambrian Research Center Professional Workshop on VMS and lode gold deposits in Archean greenstone belts (Hudak et al., 2008). In addition to this guidebook, the reader is referred to two papers which describe the physical volcanology and mineralization present within the SLCC. These papers provide detailed overviews of the regional geology as well as the literature related to subaerial and submarine volcanic processes and the genesis of VMS mineralization. Morton et al. (1991) presents an early interpretation of the stratigraphic sequence within the SLCC, and discusses the interrelationships between volcanological and mineralizing processes that formed the Mattabi VMS orebody. This classic paper is the first to discuss the development of a submarine caldera complex in the south Sturgeon Lake region, and how processes associated with caldera development led to conditions favourable for massive sulphide mineralization. Hudak et al. (2003) is a detailed discussion of submarine explosive silicic volcanological processes within the SLCC, and how these processes dictated the stratigraphic intervals, as well as the intracaldera locations, where economic VMS mineralization occurs within the SLCC. In addition, Hudak et al. (2003) includes the most upto-date stratigraphic nomenclature and interpretations of the volcanological and ore-forming processes associated with the genesis of this exceptionally wellpreserved Neoarchean subaerial to submarine caldera complex. In essence, Hudak et al. (2003) indicate that - 117 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 larger VMS deposits correlate with larger explosive eruption events within the SLCC. Regional Geologic Setting The SLCC is located within the South Sturgeon Sequence of the Savant Lake-Sturgeon Lake Greenstone belt that occurs in the Wabigoon Greenstone Belt of the Superior Province (Fig. 1; Sanborn-Barrie and Skulski, 1999). The SLCC is up to 25 km in strike length, and contains an approximately 3000 meter stratigraphic thickness of north-facing, vertical to moderately northdipping (55°) intracaldera fill (Morton et al., 1991; Morton et al., 1999; Hudak et al., in press) composed of greenschist to locally amphibolite facies (Trowell, 1974, 1983; Groves et al., 1988) metavolcanic, metasedimentary, and meta-intrusive rocks of Neoarchean age (Table 2). The eastern margin of the SLCC has been interpreted by Morton et al. (1991, 1999) to be the Lac David Fault. The western margin of the SLCC is poorly constrained by limited outcrop and diamond drilling, but based on the western limit of intracaldera strata, numerous dikes, and extensive hydrothermal alteration, is believed to be located beneath Biedelman Bay at the West Sturgeon Fault (Fig. 1). The caldera complex and associated ore deposits formed within an evolved, continental margin oceanic arc (Sanborn-Barrie et al., 2001; Galley, 2002) which contained magmas derived from back-arc basalts (Galley, 2003) that were not contaminated to any large extent by older 3.0 Ga Wabigoon Province continental crust (Bernier et al., 1999). Two coeval intrusive complexes intrude into the pre-caldera supracrustal volcanic strata in the Sturgeon Lake region, and have been extensively studied by Galley et al. (2000), Galley (2002), and Galley (2003). The Pike Lake Layered Complex (PLLC) consists of massive to crudely layered ferrogabbro, melanogabbro, and gabbro that, combined, are up to 10km in strike length and up to 2500m thick. The PLLC is intruded along its eastern margin by the Beidelman Bay Intrusive Complex (BBIC), which is approximately 20km in strike length and up to 2000m thick (Trowell, 1983; Galley et al., 2000). The BBIC is composed of at least six separate intrusive phases, from oldest to youngest including xenolithic tonalite, quartz-porphyritic pre-main phase dikes, main phase leucotonalite, post-main-phase leucotonalite dikes, quartz-plagioclase-phyric porphyritic dikes, and late phase quartz monzonite and granodiorite stocks (Galley, 2002; Galley, 2003). Wide variations in geochemical signatures and age dates for these intrusive rocks indicate a complex magmatic history encompassing at least three separate magma sources over a time span of up to 20 my (Table 2; Galley, 2002). The pre-main phase dikes, main phase leucotonalite, and post-main phase leucotonalite dikes are interpreted to be syncaldera magmatic phases based on similar U/Pb zircon Table 2. Summary of geochronology in the south Sturgeon Lake region (modified from King et al., 2000). Sample No. Source Rock Unit King et al., 2000 Rock Unit (this study) Date (Ma) PN76-13 Davis & Trowell, 1982 Galley et al., 2000: Galley, 2002 Post-caldera felsic volcaniclastic N/A 2717.9 JH82-2 Davis & Trowell, 1982 Davis et al., 1985 Swamp Lake rhyolite 2735.2 JH82-1 Davis et al., 1985 Swamp Lake rhyolite 2734.8 JH82-5 Davis et al., 1985 Lyon Lake Andesite and Rhyolite Lyon Lake Andesite and Rhyolite Lyon Lake Andesite and Rhyolite Not Reported Post-caldera felsic volcaniclastic Biedelman Bay quartz-plagioclase porphyritic dike Swamp Lake rhyolite 2735.0 JH82-4 Davis et al., 1985 Lyon Creek dacite lava Middle L tuff PN76-15 JH82-3 Davis et al., 1985 Davis et al., 1985 Middle L tuff Mattabi (?) tuff 2734.7* 2736.3 DD78-17 Davis & Trowell, 1982 Davis & Trowell, 1982 Galley, 2002 Pike Lake Complex 2732.7 Beidelman Bay biotite trondhjemite Beidelman Bay biotite leucotonalite 2733.8 96-GIA-328 PN76-14 DD78-18 Not Reported Lyon Creek Lava Dome Breccias Mattabi Ash Flow Lyon Creek Lava Dome Breccias Pike Lake Pluton Gabbro Beidelman Bay biotite trondhjemite N/A * Davis et al. (1985) note this is a minimum age. - 118 - 2720 2736.0 2735.5 2734.0 Analytical Error (Ma) + 2.7 - 1.5 +3.5 - 3.0 +1.8 -1.8 +6.9 -3.2 +2.8 -2.5 +1.7 -1.7 +3.0 -1.9 ±1.6 +9.3 -3.9 +3.6 -2.0 +1.4 -1.3 +3 -3 �Proceedings of the 61st ILSG Annual Meeting - Part 2 dates (2738.5-2732.5 Ma; Davis and Trowell, 1982; Davis et al., 1985; Galley, 2002) and trace element geochemical characteristics (Campbell et al., 1981; Galley et al., 2000; Galley, 2002) that correspond with SLCC intracaldera tuffs. These intrusive phases are believed, in part, to have provided thermal energy to drive regionally extensive hydrothermal systems which formed semi-conformable zones of sericite ± dolomite alteration that were subsequently cross-cut by diffuse, semi-conformable to disconformable alteration zones comprising iron carbonate (ferruginous dolomite, ankerite or siderite), iron-chlorite, chloritoid, and aluminum silicate (pyrophyllite, andalusite and/or kyanite) (Franklin et al., 1975; Campbell et al., 1981; Hudak, 1989; Jongewaard, 1989; Morton et al., 1991; Galley et al., 2000; Galley, 2002; Holk et al., 2002, Galley, 2003) that are genetically associated with VMS formation. Intrusive and hydrothermal breccias, and quartz-plagioclase-phyric dikes associated with porphyry-style Cu-Mo mineralization described by Poulsen and Franklin (1981) are approximately 15 my younger than the synvolcanic intrusive phases, indicating that porphyry mineralization in the south Sturgeon Lake area post-dates caldera-associated VMS mineralization (Galley et al., 2000; Galley, 2002). Three distinctive types of faults have been recognized in the south Sturgeon Lake area, as well as at numerous other VMS deposits (Gibson et al., 1999). These include: a) synvolcanic faults; b) post-volcanic highangle faults; and c) post-volcanic shear zones which may occur at a relatively low angles to the stratigraphy and may represent thrust faults. Synvolcanic faults are recognized using the following criteria: a) apparent offset of a layered unit, with subsequent units not offset; b) abrupt thickening or thinning of a volcanic, volcaniclastic, or sedimentary unit; c) rapid changes in alteration intensity or abrupt changes in alteration mineral assemblages; and d) the presence of apophyses or dikes associated with synvolcanic intrusions (Gibson et al., 1999; Hudak et al., 2003). Post-volcanic high angle faults have the following characteristics: a) offset of all stratigraphic units; b) abundance of brittle fracturing and/or zones of lost core; and c) abundance of quartz ± carbonate ± tourmaline veins filling fractures. Post-volcanic shear zones (thrust faults?) can be identified using a combination of the following criteria: a) intense shearing within adjacent units; b) displacement of older stratigraphic units into positions up-section from younger stratigraphic units; c) presence of “lamprophyre” and/or massive diorite dikes; and d) locally, the presence of graphite-rich breccia zones containing angular quartz-rich fragments (Hudak, 1996). Several episodes of faulting and folding have taken place in the south Sturgeon Lake region. Numerous synvolcanic fault zones were identified by Morton et al. (1991, 1999) based on abrupt thickness changes in volcanic strata, terminations of fault zones by overlying volcanic strata, and proximity to intense hydrothermal alteration and/or VMS mineralization. Two episodes of post-volcanic faulting are preserved in the south Sturgeon Lake area. The earliest event is a low angle shear zone (possibly a thrust fault) which juxtaposed caldera-associated strata and non-caldera associated strata in the eastern and northeastern parts of the south Sturgeon Lake area (Hudak, 1996: Morton et al., 1999). The fault zone interpretation is supported by a) the presence of highly strained rocks along the contact between the two sequences (Dube et al., 1989; Koopman, 1993; Hudak, 1996); b) the lack of consistent stratigraphy along a 20-50 m thick zone adjacent to the contact in the vicinities of the Sub-Creek Zone, Creek Zone, and Lyon Lake VMS orebodies (Hudak, 1996); c) abrupt terminations of alteration mineral assemblages across the contact (Hudak, 1996); and d) sharp oxygen isotopic gradients across the contact (Moss, 1992; Holk et al., 2002). This fault zone has been successfully geophysically imaged by Nedmović and West (2002). A subsequent faulting event formed a series of NNE-trending faults that offset the intracaldera supracrustal strata and the early low angle fault (Figure 1: Koopman, 1993; Hudak, 1996; Morton et al., 1999). The major fold event in the region resulted from a ca 2.7 Ga north-south compression event associated with the collision of Neoarchean rocks in the south Sturgeon Lake region and a Mesoarchean volcanic rift sequence to the north (Sanborn-Barrie and Skulski, 1999). The north-facing and north-dipping stratigraphic sequence in the south Sturgeon Lake assemblage represents the southern limb of an E-SE-trending, shallow eastplunging F1 fold axis located in the Post-Lake – Barge Lake region (Sanborn-Barrie et al., 1998). A major 116°-trending, 16°E-SE plunging fold documented by Dube et al. (1989) and Koopman (1993) near the Lyon Lake and Creek Zone orebodies is consistent with the orientation of this F1 syncline. Generalized Stratigraphy and Physical Volcanology of the SLCC Thirteen supracrustal stratigraphic successions have been grouped into four stratigraphic sequences based on their temporal and genetic relationships to caldera - 119 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 development, as well as their stratigraphic position relative to the Mattabi VMS orebody (after Hudak, 1996). Progressing stratigraphically up-section, these sequences are: a) the Pre-caldera Sequence (PCS); b) the Early Caldera Sequence (ECS); c) the Late Caldera Sequence (LCS); and d) the Lyon Lake Fault Sequence (LLFS). A major zone of structural deformation interpreted as either a shear zone (Koopman, 1993) or a thrust fault (Hudak, 1996) marks the contact between the PCS, ECS, and LCS in the south, and the LLFS to the north. Figures 2, 3, and 4 illustrate the geology in the western, central, and eastern parts of the SLCC, respectively. Figure 5 contains stratigraphic columns in various parts of the caldera complex. The PCS (Figs. 6 and 7) comprises a 200-2100m thick succession of subaerial and shallow subaqueous basalt lava flows, scoria-rich volcaniclastic deposits, and minor associated rhyolite lava flows (Groves et al., 1988; Morton et al., 1991; Morton et al., 1999). Pillow lavas and hyaloclastite are conspicuously absent in this sequence, except locally in the easternmost regions of the south Sturgeon Lake area (Jongewaard, 1989). Groves et al. (1988) and Morton et al. (1999) have interpreted the pre-caldera volcanic environment as a subaerial to shallow subaqueous shield volcano with local fields of scoria cones and tuff cones. The ECS (Figs. 6, 8, 9, and 10) contains a 650-1300m thick succession of volcanic and volcaniclastic strata. Up-section, these include: a) subaerially deposited ash fall tuff deposits (Jackpot Lake Succession); b) interstratified polymict breccias and subaerially and subaqueously deposited quartz-phyric lapilli tuff and tuff deposits (High Level Lake Succession); c) subaerial felsic lava flows, mafic-intermediate lapilli tuffs and volcaniclastic deposits (Bell River Succession); d) interstratified subaqueously deposited polymict breccias, volcanic sandstones and mudstones, and dacitic to andesitic lava flows and tuffs (Tailings Lake Succession); and e) subaqueous, massive to locally well-bedded, quartz-phyric lapilli tuff and tuff deposits (Mattabi Succession). Hudak (1996) interpreted this stratigraphic sequence as indicative of the early stages of caldera development (c.f. Smith and Bailey, 1968; Busby-Spera, 1984; Lipman, 1976; Lipman, 1997). The distribution of voluminous felsic volcaniclastic units and associated polymict breccias suggest that Figure 2. Geological map of the western one-third of the Sturgeon Lake Caldera Complex (modified after Morton et al., 1999). Lines A – A’ and B – B’ correlate to stratigraphic sections in Figure 5. - 120 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 3. Geological map of the central one-third of the Sturgeon Lake Caldera Complex (modified after Morton et al., 1999). Line C – C’ correlates to stratigraphic section in Figure 5. Figure 4. Geological map of the eastern one-third of the Sturgeon Lake Caldera Complex (modified after Morton et al., 1999). Lines D – D” and E-E” correlate to stratigraphic sections in Figure 5. - 121 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 5. Stratigraphic sections across the Sturgeon Lake Caldera Complex (Hudak et al., in prep.). Section is hung on top of High Level Lake Rhyodacite-Rhyolite Tuff/Lapilli-tuff unit. - 122 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 6. Immobile trace element Geochemical classification (after Winchester and Floyd, 1977) of least-altered volcanic and volcaniclastic rocks associated with the Sturgeon Lake Caldera Complex. Pre-caldera strata are illustrated in diagram A, early caldera strata are illustrated in diagrams B and C, Late-Caldera strata are illustrated in diagrams D and E, and Lyon Lake Fault Sequence strata are illustrated in diagram F. Data from Groves (1984), Hudak (1989, 1996), Jongewaard (1989), Walker (1993) and Franklin (unpublished data). - 123 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 7. Pre-caldera (Darkwater Succession) strata associated with the SLCC. A) rare pillow lavas near Lac David; B) massive basalt-andesite tuff with carbonate-altered scoria lapilli; C) photomicrograph of scoria in basalt -andesite tuff (field of view 8mm); D) fusiform- and spindle-shaped bombs in basalt-andesite tuff; E) cored bomb in basalt-andesite tuff; and F) flow-banding in rhyodacite – rhyolite lava flow. - 124 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 8. Strata from the Jackpot Lake and High Level Lake Successions. A) Jackpot Lake Succession lapilli tuff; B) photomicrograph of recrystallized ash matrix and pumice from Jackpot Lake Succession tuff (field of view 8mm); C) High Level Lake Succession polymict breccia (mesobreccia); D) High Level Lake Succession polymict breccia (megabreccia); E) pumice lapilli in massive High Level Lake Succession tuff south of the Mattabi orebody; F) interbedded High Level Lake Succession lapilli-tuffs and tuffs near the F-Group VMS deposit. - 125 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 9. Strata from the Bell River and Tailings Lake Successions. A) photomicrograph of amygdaloidal Bell River Succession rhyodacite-rhyolite lava flow; B) massive Bell River basalt lapilli tuff with carbonate- and quartz-altered scoria; C) polished drill core sample of Bell River basalt lapilli-tuff (scale bar = 2cm); D) Tailings Lake Succession polymict breccia from south of the Mattabi orebody. initial caldera development occurred simultaneously with the deposition of the High Level Lake Succession. A synvolcanic basin, bounded to the west and east by the Darkwater and Lac David Faults respectively (Fig. 1), appears to represent a nested caldera produced simultaneously with the eruption of the Mattabi tuffs (Hudak et al., 2003; Hudak et al., in prep). The LCS (Figs. 6, 11, 12, 13, and 14) comprises a 500-1500m thick succession of quartz- and quartz + feldspar-phyric tuff and lapilli tuff deposits, volcaniclastic sedimentary rocks, andesitic to dacitic lava flows, domes, and cryptodomes, and Algomatype banded iron formations (Koopman, 1993; Hudak, 1996). Hudak (1996) interpreted this sequence as characteristic of a maturing, late-stage caldera complex (c.f. Smith and Bailey, 1968; Lipman, 1997), and Hudak et al (in prep.) have utilized GIS analysis of the Sturgeon Lake geology that reflects relative percentages (by area and volume) of caldera collapse breccias, eruption-fed tuffs, lava flows, and sedimentary rocks that is consistent with this interpretation (Table 3). The most voluminous explosive eruptions in the LCS produced the Middle L lapilli tuffs and tuffs which are the host rocks for all the VMS orebodies in the LCS. Several intra-caldera intrusive rocks have been identified and include: a) the Beidelman Bay Intrusive Complex (BBIC), a multiphase, dominantly trondhjemitic intrusive complex that petrochemical and lithogeochemical data indicate was, in part, the subvolcanic intrusion associated with the SLCC; b) massive to amygdaloidal rhyolite which occurs as feeder dikes to the Bell River Succession rhyolite lava flows; c) coarse-grained massive feldspar-phyric diorite which occurs as feeder dikes to the Lyon Creek Succession lava domes and cryptodomes; and d) finegrained massive to amygdaloidal diorite and quartz - 126 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 10. Mattabi Succession strata: A) massive lapilli tuff (compass is 55cm wide); B) pumice block in massive Mattabi Succession lapilli tuff from immediately southwest of the Mattabi VMS orebody; C) laminated to thinly-bedded tuff and subeconomic VMS horizon north of the F-Group VMS orebody (scale bar divided into centimeters); D) drill core appearance of bedding replaced by sulphides (mainly pyrite; scale bar is 2cm). Table 3. Geographic information system (GIS) analysis of Pre-Caldera, Early Caldera, and Late Caldera sequences in the Sturgeon Lake Caldera Complex (Hudak et al., in prep.) Distribution of Lithology Types in the Sturgeon Lake Caldera Complex by Area % Lithology Type Pre-Caldera % Early Caldera % Late Caldera % Caldera Collapse Breccia 0.0 49.6 0.0 Lava Flows 94.1 8.6 41.0 Eruption-Fed Tuffs 4.8 41.0 12.3 Sedimentary Rocks 1.1 0.8 46.7 Distribution of Lithology Types in the Sturgeon Lake Caldera Complex by Volume %* Lithology Type Pre-Caldera % Early Caldera % Late Caldera % Caldera Collapse Breccia 0.0 57.6 0.0 Lava Flows 98.4 3.5 42.1 Eruption-Fed Tuffs 1.6 38.7 6.3 Sedimentary Rocks 0.1 0.2 51.6 *Volume estimates of intracaldera strata completed by assuming that the Sturgeon Lake Caldera Complex was a rcular caldera structure similar to Cenozoic subaerial ash flow calderas such as Valles (Smith and Bailey, 1968), Santorini (Druitt and Francaviglia, 1992), and Crater Lake (Bacon and Druitt, 1988), and submarine ash flow calderas such as Healy (Wright et al., 2003) and Myojin Knoll (Fiske et al., 2001). Similar assumptions have been used to estimate eruption volumes in other ancient volcanic sequences (Busby-Spera, 1984; Kokelaar and Busby, 1992). - 127 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 11. Lower L and Middle L strata within the SLCC: A) bedded Lower L rhyodacite-rhyolite tuff and lapilli tuff (lens cap is 55mm); B) Middle L Succession laminated to thinly bedded rhyo-dacite-rhyolite tuffs overlain by very thickly bedded lapilli tuff (scale bar divided into centimeters); C) massive Middle L tuff with 1-3 mm subhedral to euhedral quartz phenocrysts (scale bar is 2cm); D) normal graded Middle L tuff (up is to right of photo, scale bar is 2cm); E) outcrop appearance of Middle L rhyodacite-rhyolite tuff breccia (scale bar divided into centimeters); F) close-up view of Middle L rhyodacite-rhyolite tuff-breccia (note jigsaw puzzle-fit lapilli and bombs, scale bar divided into centimeters). - 128 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 12. Upper L and Bell River Lake succession strata: A) Upper L Succession rhyodacite-rhyolite tuff and lapilli tuff (scale bar divided into centimeters); B) Upper L Succession crystal-rich reworked tuff (large divisions on scale bar are centimeters); C) massive, crudely graded, reworked Upper L tuff (scale bar divided into centimeters); D) Bell River Lake Succession quartz- and plagioclase-phyric rhyodacite – rhyolite lava flow (compass at left for scale). - 129 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 13. No Name Lake succession strata: A) pillow lava with concentric cooling joints; B) pillow lava with multiple selvedges (large divisions on scale bar are centimeters); C) laminated tuff (resedimented hyaloclastite; D) amoeboid basalt andesite dike with surrounding peperite and resedimented hyaloclastite; E) close-up of relationship between basalt-andesite dike and surrounding peperite; F) close-up of close-packed blocky to irregular peperite. Compass in all photos is 55 mm wide. - 130 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 14. Lyon Creek Succession strata: A) vertical facies from margin (left) to center (right) of andesite-dacite cryptodome; B) andesite-dacite lithic tuff-breccia that occurs along margin of cryptodome; C) andesite-dacite lithic tuff breccia containing both fine- and coarse-grained feldspar phyric lava clasts; D) resedimented crystal-lithic tuff; E) laminated to very thinly bedded graphitic mudstone; F) folded magnetite-chert banded iron formation interbedded with chlorite-rich mudstones. Scale bar in photos 14A-14E is 2 centimeters. Scale bar in photo 14F is divided into centimeters. - 131 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 diorite which dilate the stratigraphy across the entire caldera complex and may have been a subvolcanic sill complex which fed the No Name Lake Sequence basaltic and andesitic lava flows (Morton et al., 1999). Post-caldera intrusive rocks also occur throughout the SLCC. These include: a) fine-grained, massive, locally feldspar-phyric diorite; b) coarse-grained amphibole-phyric diorite; and c) thin (<1-13m wide) “lamprophyre” dikes that occur within high strain zones associated with the structural deformation between caldera-associated strata and the non-caldera associated Lyon Lake Fault Sequence. Hydrothermal Alteration Assemblages in the SLCC Mineral Hydrothermal solutions, both ore-forming and nonore forming, have altered the volcanic and intrusive rocks associated with the Sturgeon Lake Caldera Complex. Following alteration, rocks in the region were metamorphosed to greenschist and locally amphibolite facies. Therefore, the alteration minerals now present are, in part, metamorphosed equivalents of alteration minerals formed during the synvolcanic hydrothermal alteration events. Hydrothermal alteration is widespread within the complex and in the uppermost parts of the Beidelman Bay Intrusive Complex (Franklin et al., 1975; Jongewaard, 1989; Hudak, 1989; Walker, 1993; Hudak, 1996; Galley et al., 2000; Galley, 2002; Galley, 2003). Discrete assemblages of metamorphosed alteration minerals form zones that are a) widespread and largely conformable to the volcanic stratigraphy (semiconformable alteration); and b) locally lens-shaped or pod-like beneath massive sulphide occurrences and deposits (semi-conformable alteration); and c) narrow and elongate that cross-cut stratigraphy and are proximal to synvolcanic structures or synvolcanic faults (disconformable alteration; Franklin et al., 1975; Groves, 1984; Morton and Franklin, 1987; Jongewaard, 1989; Hudak, 1989; Walker, 1993; Hudak, 1996; Gibson et al., 1999). Hydrothermal fluids formed five distinct, mapable massive sulphide ore-associated alteration mineral assemblages in the area (Fig. 15). From least- to most-altered, these assemblages are: a) widespread, semi-conformable carbonatization and silicification; b) widespread, semi-conformable iron carbonate ± iron-rich chlorite; c) widespread, semi-conformable chloritoid ± iron-carbonate and/or iron-rich chlorite; d) localized lens-shaped to pod-like, locally linear zones of aluminum silicate (pyrophyllite, andalusite, and/or kyanite) + chloritoid; and e) localized linear disconformable and semi-conformable, generally stratiform zones of aluminum silicate. Algoma-type iron formations associated with the LCS are associated with semi-conformable to disconformable veins, patches, and lenses of iron-carbonate + iron- rich chlorite + magnetite ± Mn-rich almandine garnet ± grunerite. Late sericite and/or magnesium-rich chlorite alteration locally overprints these five alteration mineral assemblages. Iron-carbonate ± iron-rich chlorite assemblage rocks contain at least 10% iron-carbonate + iron-rich chlorite with less than 5% chloritoid or aluminum silicate minerals. Outcrops containing this alteration assemblage can generally be easily identified by their orange to orange-brown stained, commonly pitted surfaces. Staining varies from irregular patches up to 15cm in diameter, to veins and veinlets 1-15mm in width that are aligned parallel to the foliation. Pumice/ scoria that has been replaced by iron-carbonate can be recognized as rounded to oval, orange-brown stained pits up to 5cm in diameter which commonly contain rounded- to lens-shaped quartz amygdules (10-60%). In thin section, this assemblage contains iron-carbonate ± iron-rich chlorite (10-60%), sericite (up to 30%), magnesium-rich chlorite (up to 50%), and locally, traces of chloritoid and/or aluminum silicate minerals. Chloritoid ± iron carbonate and/or iron-rich chlorite assemblage rocks contain greater than 5% chloritoid, and are characterized by the presence of 1-3mm dark green to greenish-black chloritoid prisms and rosettes. The presence of chloritoid commonly gives the rocks a “salt and pepper” appearance. Locally, chloritoid porphyroblasts occur with sericite (10-55%) in 1-5mm wide grey green veins that vary from semi-conformable to disconformable in orientation. Other minerals present include iron-carbonate (up to 60%), ironrich chlorite (2-20%, locally as a retrograde product of chloritoid), magnesium-rich chlorite (1-20%) and pyrite (up to 12%). Aluminum silicate + chloritoid bearing rocks typically contain 1-3mm chloritoid porphyroblasts (up to 33%) in a grey to grey-pink matrix composed of massive pyrophyllite (5-20%), 1-3mm blocky pink andalusite (up to 10%), and/or blue tabular to bladed kyanite porphyroblasts (up to 8%). Chloritoid occurs as 1-3mm prisms or rosettes disseminated throughout the rock or in chloritoid + andalusite veins up to 1cm in width. In addition, iron-rich chlorite (up to 7%), - 132 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 15. Field and drill core hydrothermal alteration mineral assemblages in the SLCC. A) Iron-carbonate assemblage in the F-Group area; B) chloritoid + chlorite alteration assemblage from Area 17; C) aluminum silicate alteration (kyanite + sulfides) from footwall to the Mattabi VMS deposit; D) patchy aluminum silicate (andalusite) alteration proximal to VMS mineralization in Area 17; E) iron-formation associated alteration (iron-chlorite + iron carbonate + chloritoid + magnetite) in Area 23; F) vein of iron formation-associated alteration (iron carbonate + magnetite + iron chlorite) in Area 23. - 133 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 magnesium-rich chlorite (up to 35%), iron carbonate (3-35%), pyrite (up to 30%) and sphalerite (up to 5%) are also associated with this alteration assemblage. Figure 16 illustrates the distribution of the VMS ore-associated alteration mineral assemblages in the vicinity of the F-Group orebody. Iron carbonate ± iron-rich chlorite assemblage rocks form a large, semiconformable zone along the southern, eastern, and northern boundaries of the area. Locally, pod-like regions (up to 50m in diameter) of this assemblage occur within the chloritoid assemblage zone, which exists as a large, semi-conformable zone up to 100m thick in the hanging wall to the F-Group orebody, and is present as pipe-like alteration in the F-Group deposits footwall. More proximal to mineralization, semi-conformable and locally linear pipe-like zones containing the aluminum silicate ± chloritoid alteration assemblage are present. Aluminum silicate assemblage rocks are most closely associated with the mineralization at the F-Group deposit. These rocks are distributed in two distinct patterns: a) in disconconformable, pipe-like linear, 5-30m wide zones that trend NNE in proximity to synvolcanic fault zones; and b) in a broad semiconformable zone (700m by 500m at surface) located both in the footwall and the hanging wall rocks to the F-Group orebody. Figure 17 illustrates the distribution of the ore- associated alteration at the Mattabi deposit. Semiconformable iron-carbonate (ankerite- and/or sideritebearing) and silicified rocks form a broad zone in the lower footwall rocks. This zone is cross-cut by several westward-dipping tabular zones which contain an aluminum silicate-rich core and aluminum silicate ± chloritoid margin. These zones occur in close proximity to synvolcanic fault zones which lead upward to, and cross-cut, a broad semi-conformable zone of aluminum silicate + chloritoid altered rocks that crudely surrounds the Mattabi orebodies. The aluminum silicate-rich tabular zones, associated with the synvolcanic fault zones, spread out into a semiconformable alteration zone directly stratigraphically below and lateral to the deposit. Stratigraphically overlying the deposit is a semi-conformable zone of chloritoid ± iron-rich chlorite ± iron-carbonate up to 150m thick (Walker, 1993). Figure 18 illustrates the distribution of iron formation-associated alteration mineral assemblages that occur in Areas 17 and 23 southwest of the Lyon Lake and Creek Zone VMS ore deposits. It is important to keep in mind that based on stratigraphic and structural data, the iron formation-associated alteration assemblage has no genetic relationships to the Lyon Lake and Creek Zone deposits, as these deposits are believed to have been structurally remobilized into Figure 16. Distribution of hydrothermal alteration mineral assemblages associated with the F-Group VMS orebody. - 134 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Figure 17. Distribution of hydrothermal alteration mineral assemblages associated with the Mattabi VMS orebody. Figure 18. Distribution of hydrothermal alteration mineral assemblages associated with the Algoma-type iron formation and VMS mineralization in Areas 17 and 23. - 135 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 erupt simultaneously with caldera collapse that generates the High Level Lake Succession polymict breccia deposits (megabreccia and mesobreccia). Facies analyses suggest initial trap-door like caldera collapse where the floor of the caldera complex collapses into a subaqueous environment west of a synvolcanic fault located approximately at the boundary between the F-Group claims and Area 15 (refer back to Fig. 5). The eastern two-thirds of the caldera complex, from the Darkwater Fault to the Lac David Fault, remains in a subaerial depositional environment. their present locations (Koopman, 1993; Hudak, 1996; Morton et al., 1999). Summary of the Geologic History of the SLCC The following summary represents the temporal sequence of geological, volcanological, and oreforming processes that occurred during the genesis of the Sturgeon Lake Caldera Complex. Pre-Caldera Geological History 1. Subaerial basalt shield volcanism (Darkwater Succession basalt-andesite lava flows) with minor effusive felsic volcanic activity (Darkwater Succession rhyodacite-rhyolite lava flows and associated flow breccias). 2. Intrusion of spherulitic rhyolite dikes in southern Area 16 - Area 17 which feed the eruption of Bell River Succession dacite-rhyolite lava dome in a subaerial environment in the eastern two-thirds 2. Subaerial-shallow subaqueous scoria cone and tuff cone genesis (Darkwater Succession basalt-andesite tuffs, lapilli tuffs, and tuff-breccias) as the end stage in the development of a shield volcano – tuff-cone – scoria-cone complex. 3. Initial intrusion of the ancestral Beidelman Bay Intrusive Complex and the Pike Lake Layered Complex as a high-level magma chamber. This causes regional tumescence which results in the formation of pre-caldera structures that evolve into ring fractures. 4. Subaerial explosive felsic volcanism, probably erupting through vents proximal to pre-caldera structures caused by extension related to regional tumescence, deposits the Jackpot Lake Succession rhyodacite and rhyolite tuffs. This eruption can be interpreted as minor, pre-caldera explosive activity which commonly precedes caldera formation in the Smith and Bailey (1968) caldera cycle model. Further development of pre-caldera faults continues during this time. Early Caldera Geological History 1. Formation of the ~25km strike length Sturgeon Lake Caldera Complex occurs as a result of voluminous subaerial explosive volcanism (at least 16km3, possibly on the order of ~500-900km3 based on observed caldera diameter and eruption volume observation (Cas and Wright, 1987; Hudak et al., 2003; Fig. 19) coupled with simultaneous caldera collapse. Explosive volcanism produces the High Level Lake rhyodacite-rhyolite lapilli tuffs which Figure 19. Comparison of estimated eruption volumes and caldera diameters for modern subaerial and subaqueous caldera eruptions, and the caldera-forming eruptions associated with the SLCC (Hudak et al., 2003, modified from Cas and Wright, 1987). Point A represents relationships for the initial SLCC event which resulted from the combined eruptions of the Jackpot Lake and High Level Lake succession rhyodacite-rhyolite lapilli tuffs and tuffs. Point B represents relationships for the Mattabi eruptive event, which formed the Mattabi Succession rhyodacite-rhyolite lapilli tuffs and tuffs, and a nested caldera between the Darkwater and Lac David faults. Point C represents the Middle L eruptive event, which formed the Middle L rhyodacite-rhyolite breccias, tuff-breccias, lapilli tuffs and tuffs, and possible two nested calderas of 1.5km and 3.1km in diameter (equivalent to a single caldera of ~3.4km diameter) in the eastern and western parts of the SLCC in the vicinities of the F-Group deposit and Areas 17 and 23. - 136 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 of the SLCC. Post caldera-collapse submarine hydrothermal activity is centered in the F-Group and Darkwater regions of the SLCC, and forms the F-Group VMS orebodies. Based on observations of modern VMS mineralization in terms of tectonic setting, metal contents, physical volcanology and hydrothermal alteration mineral assemblages (Gibson et al., 1999; Morgan and Schulz, 2012; Monecke et al., 2014), it is likely that these VMS orebodies were formed as replacement-style VMS deposits (c.f. Doyle and Allen, 2003) at water depths less than 1500 meters (Hudak, 1996); Subaerial hydrothermal activity occurs in the eastern twothirds of the caldera complex, and forms subaerial epithermal-like stringer mineralization in the High Level Lake lapilli tuffs and tuffs, as well as the Bell River Succession lava dome. 3. Eruption and deposition of the Bell River Succession basalt lapillistones and lapill tuffs from Area 16Area 17 in the vicinity of the Bell River lava dome. These deposits pond to the west along the Darkwater fault, which marks the western limit of a paleo-basin that extends eastward to the Lac David Fault. 4. Intracaldera clastic sedimentation forms the Bell River Succession laminated to thickly bedded polymict tuffs and tuffs. 5. Continued subsidence of the basin between the Darkwater and Lac David faults changes the depositional environment from a subaerial to a shallow subaqueous setting. Scalloping of the basin margins leads to deposition of the Tailings Lake Succession massive to stratified polymict tuffs and lapilli tuffs. Periodic effusive eruption of andesitic (Tailings Lake andesite lava flows) and small volume explosive eruptions of dacite-rhyolite (Tailings Lake dacite-rhyolite tuff) also occurs. 6. High temperature hydrothermal activity forms the Mattabi “E” ore lens within the Tailings Lake dacite-rhyolite tuff. 7. Pulsating, voluminous (~30km3: Hudak et al., 2003) subaqueous explosive eruptions form hot (?) highand low-concentration eruption fed density currents within the basin between the Darkwater and Lac David faults, and cool high- and low-concentration, eruption-fed density currents as outflow sheets in the western one-third of the SLCC (Mattabi bedded quartz-phyric rhyodacite-rhyolite lapilli tuffs and tuffs). A nested caldera is formed during these eruptions between the Darkwater and Lac David faults, enabling the submarine environment to subside to depths in which VMS may form. 8. Regional low temperature submarine hydrothermal activity alters the volcaniclastic strata to regional sericite and carbonate alteration assemblages. This regional alteration is responsible for semiconformable zone of Na-depletion that extends for 25km along strike within the Sturgeon Lake camp. The regional carbonate present is generally dolomite. Post-eruptive, high temperature synvolcanic hydrothermal activity in the vicinity of synvolcanic structures forms the Mattabi “B”, “C”, and “D” lenses, primarily as sheet-like replacement VMS deposits (c.f. Doyle and Allen, 2003) at water depths <1500 meters. Hydrothermal alteration associated with this mineralization leaches alkali elements and locally adds iron and manganese to form the premetamorphic precursors to the iron carbonate + iron-rich chlorite assemblage, the chloritoid ± ironrich chlorite assemblage, the aluminum silicate + chloritoid assemblage, and the aluminum silicate assemblage. Regional dolomite reacts with the hydrothermal solutions to form ferrodolomite, ankerite, and siderite as one approaches VMS mineralization (Franklin et al., 1975). Downwelling of cool seawater deposits magnesium and potassium to form late magnesium-rich chlorite and sericite veins. 9. Subaqueous deposition and post hydrothermal reworking of Mattabi Succession massive rhyodaciterhyolite tuffs. Continued regional submarine hydrothermal alteration produces regional semiconformable alteration zones dominated by sericite. Late Caldera Geological History 1. Subaqueous eruption, deposition, and post-eruptive reworking of the Lower L Succession rhyodaciterhyolite tuffs and lapilli tuffs. 2. Continued erosion of caldera walls, subaqueous reworking of intracaldera deposits, and low temperature hydrothermal activity generates the Lower L Succession interbedded graphitic mudstones, tuffs, and lapilli tuffs). 3. Initial eruptions of No Name Lake Succession basalt-andesite lava flows east of the Bell River, and Bell River Lake basalt-andesite lava flows adjacent to the Darkwater fault. 4. Eruption of rhyodacite-rhyolite lava dome, and formation of VMS deposits on the dome, approximately 2km west of the present location of - 137 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 the Sturgeon Lake VMS deposit (indirect evidence from rhyodacitic to rhyolitic lava lapilli and blocks in the Middle L breccia and tuff-breccia deposits). 5. Subaqueous disintegration of rhyodaciterhyolite lava dome and associated VMS deposits, simultaneously with subaqueous rhyodaciticrhyolitic plinean (phreato-plinean?) volcanism form high- and low-concentration eruption fed mass flows which deposit Middle L Succession rhyodacite breccia, tuff-breccia, lapilli tuffs, and tuffs within a nested caldera that has a well-defined western margin near the Bell River, and a poorly defined eastern margin (the location of the eastern margin is no longer able to be recognized at surface due to the structural deformation associated with the Lyon Lake Fault Zone). Simultaneous submarine plinean eruptions appear to have occurred north of the F-Group VMS orebody, and deposited relatively thick sequences of quartz-phyric rhyodacite-rhyolite lapilli tuffs and tuffs in a localized synvolcanic basin (Middle L Succession rhyodacite-rhyolite breccia, tuff-breccia, lapilli tuff, and tuff). 6. Post-eruptive high temperature hydrothermal activity forms the Mattabi “A” ore lens, the Sturgeon Lake VMS deposit, and subeconomic VMS deposits hosted in the Middle L Succession lapilli tuffs and tuffs in Area 17 and near the border between the F-Group claims and Area 15. Regional low temperature submarine alteration of the rocks yields semi-conformable sericite-rich and carbonate-rich alteration zones. 7. Intracaldera low temperature hydrothermal activity, sedimentation and reworking of intracaldera volcanic and volcaniclastic rocks (Upper L Succession interbedded graphitic mudstones, tuffs, lapilli tuffs). 8. Eruption, subaqueous deposition, and subaqueous reworking of plagioclase- + quartz-phyric volcaniclastic rocks (Upper L rhyodacite-rhyolite tuffs and lapilli tuffs). 9. Post-eruptive synvolcanic hydrothermal activity forms subeconomic VMS deposits in the Upper L Succession lapilli tuffs and tuffs in the eastern onethird of the SLCC. 10. Continued intracaldera sedimentation and submarine reworking of intracaldera deposits (Upper L Succession interbedded graphitic mudstones, tuffs, and lapilli tuffs). 11. Intermittent subaqueous mafic-intermediate and felsic volcanism (No Name Lake Succession and Bell River Lake Succession basalt-andesite lava flows, and Bell River Lake Succession rhyodaciterhyolite lava flows). 12.Continued intracaldera sedimentation and reworking of intracaldera deposits (Upper L Succession interbedded graphitic mudstones, tuffs, lapilli tuffs). 13. Intrusion of plagioclase-phyric diorite-dacite dikes and sills, and the formation of the Lyon Creek Succession andesite-dacite lava flows, lava domes and cryptodome. Subsequent erosion of flows and domes yields sediment which lithifies into Lyon Creek Succession feldspathic tuffs, lapilli tuffs, and tuff breccias. 14. Development of a localized basin within the eastern part of the Lyon Creek Succession dome/cryptodome complex, and the deposition of volcaniclastic sediments with the basin (Lyon Creek Succession feldspathic tuffs, lapilli tuffs, and tuff breccias). 15. Low temperature hydrothermal activity within the Lyon Creek basin forms Lyon Creek Succession Algoma-type iron formation, chert, and graphiterich mudstones. 16. Continued intracaldera sedimentation (Lyon Creek Succession feldspathic tuffs, lapilli tuffs, and tuff breccias). Post-caldera Geological History 1. Regional compression, and formation of the Lyon Lake Fault Zone, possibly as a low-angle thrust fault; shearing of the Sturgeon Lake VMS orebody, remobilization of sulphides, and emplacement of Lyon Lake, Creek Zone, Sub-Creek Zone VMS orebodies within the high strain zone associated with the Lyon Lake Fault 2. 2.7Ga regional north-south compression yields north-facing homoclinal sequence, regional greenschist and lower amphibolite facies regional metamorphism. Field Trip Stops This field trip is designed to be completed in one to two days, depending upon the time spent, and the depth of discussions, at each field trip stop location. The field trip stops have been numbered so that the participants observe rocks from the base toward the top of stratigraphic sequence (older rocks to younger rocks) throughout the excursion. In addition to moving - 138 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 more or less up-section, the field trip will successively investigate volcanic rocks, hydrothermal alteration, and mineral deposits in the western, central, and eastern parts of the SLCC. N.B. All field trip stops are located on GlencoreCanada Property; therefore, permission must be obtained before attempting access to the field trip stops. The first day of the field trip will concentrate on rocks that comprise the subvolcanic Beidelman Bay Intrusive Complex, the Darkwater Succession (pre-caldera rocks upon which the Sturgeon Lake Caldera Complex was developed), and the High Level Lake, Tailings Lake, and Mattabi Successions which comprise the Early Caldera Sequence, as well as host the majority of the VMS mineralization with the Sturgeon Lake camp. The second day of the field trip will focus on rocks that comprise the Late Caldera Sequence, including variably (and locally intensely) altered Middle L Succession tuffs and tuff-breccias (which host the VMS deposits at the Sturgeon Lake Mine about a kilometer east-southeast of our field trip stops), coherent dacites that comprise the Lyon Creek Succession cryptodome, and exceptionally well-preserved Neoarchean pillow lavas, volcaniclastic sediments, and locally, peperite deposits that make up the No Name Lake Succession. Please note, there are no co-ordinates available for the stops described within this field trip guide. UTM Co-ordinates; however, will be supplied to participants at the beginning of the field trip. Day 1 – Synvolcanic Intrusion, Precaldera Rocks, and Early Caldera Rocks Associated with the Sturgeon Lake Caldera Complex The Beidelman Bay Intrusive Complex Stop 1: Beidelman Bay Intrusive Complex Two coeval intrusive complexes intrude the precaldera supracrustal strata in the Sturgeon Lake region (refer to Fig. 1), and have been extensively studied by Galley et al. (2000), Galley (2002) and Galley (2003). The Pike Lake Layered Complex (PLLC) is composed of massive to crudely layered ferrogabbro, melanogabbro, and gabbro, which combined have a strike length of approximately 10km and a composite thickness of approximately 2500m. The PLLC is intruded along its eastern margin by the Beidelman Bay Intrusive Complex (BBIC), which has a strike length of approximately 20km and a composite thickness of approximately 2000m (Trowell, 1983; Galley et al., 2000). The BBIC is composed of at least six separate intrusive phases. From oldest to youngest, these include xenolithic tonalite, quartz-porphyritic premain phase dikes, main phase leucotonalite, post-main phase leucotonalite dikes, quartz- and plagioclasephyric dikes, and late-phase quartz monzonite and granodiorite stocks (Galley, 2002, 2003).Wide variations in geochemical signatures and age dates for these intrusive phases indicate a complex magmatic history encompassing at least three separate magma sources over a time span of up to 20 million years (refer to Table 2). At this field trip stop, we will investigate the product of the middle intrusive event, the main phase biotite leucotonalite of the BBIC. According to Galley (2002), the main phase biotite leucotonalite makes up the bulk of the BBIC. It comprises a massive, mediumgrained hypidiomorphic to seriate-textured rock that displays few internal variations in texture or grain size. Galley (2002) notes that this phase of the BBIC has lithogeochemical characteristics that matches favourably with those of the High Level Lake, Mattabi, and Middle L pyroclastic rocks within the SLCC, in agreement with earlier interpretations (Davis and Trowell, 1982; Trowell, 1983) that the BBIC was the subvolcanic magma chamber that erupted to form the intracaldera felsic pyroclastic rocks within the SLCC. Precaldera Strata – The Darkwater Basalts and Darkwater Rhyolites Stop 2: Darkwater Basalts and Darkwater Rhyolites The Precaldera Sequence extends along strike across the entire south Sturgeon Lake region, and is composed of a 200-2100m thick succession of subaerial and shallow submarine basalt and andesite lava flows (Darkwater basalt-andesite lava flows), volcaniclastic deposits (Darkwater basalt-andesite tuffs and lapillituffs), and minor rhyodacite-rhyolite lava flows (Darkwater rhyodacite-rhyolite lavas; Figs. 2, 3 and 4). At this field trip location, we will have the opportunity to investigate Darkwater basalt-andesite lava flows as well as Darkwater rhyodacite-rhyolite lavas. The Darkwater basalt-andesite lava flows have been studied in detail by Groves (1984) and Groves et al. (1988). These lava flows comprise a 200-1800m thick sequence of aphyric to plagioclase-phyric, massive to amygdaloidal (2-25% amygdules) basalt – andesite lava flows. Groves (1984) and Groves et al. (1988) indicate that individual lava flows vary in - 139 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 stratigraphic thickness from 3-40m, and are commonly characterized by billowy flow tops containing 45-70% amygdules and locally flow marginal autobreccias that contain 40-70% poorly sorted, massive to scoriaceous basalt-andesite lapilli and blocks. It is only on the far eastern end of the south Sturgeon Lake region that pillow lavas, sheet flows, and associated hyaloclastite deposits have been recognized (Jongewaard, 1989), suggesting that these pre-caldera strata were formed largely in a subaerial environment. At this location we will investigate the physical characteristics of Darkwater Succession basalt-andesite lava flows. Moving to the north, we will also investigate aphyric to quartz-plagioclase-phyric, typically massive but locally spherulitic and/or flow banded rhyodaciterhyolite lava flows that occur within the Darkwater Succession. Individual lava flows are up to 80m thick, and are separated by clast-supported volcanic breccias that have been interpreted by Groves (1984) and Groves et al. (1988) to be autobreccias. Similar felsic lava flows with stratigraphic thicknesses up to 10m locally overlie thin basalt-andesite lava flows in the eastern part of the south Sturgeon Lake region near Lac David (Jongewaard, 1989). The F-Group VMS deposit was discovered from airborne and ground geophysical surveys combined with exploration diamond drilling conducted by Mattagami Lake Mines in 1969. The orebody was mined by open pit methods from 1981-1984, and produced approximately 377,564 short tons of ore which contained 0.64% Cu, 9.51% Zn, 0.64% Pb, and 60.4g/ton Ag (M. Patterson, personal communication, 1990; Franklin, 1996). Stop 3: High Level Lake Tuffs and Polymict Breccias (Mesobreccia) F-Group Region Intracaldera Strata At this location (Fig. 20), approximately 75m southeast of the F-Group pit, one can observe the contact between massive to sparsely graded polymict breccia deposits (the High Level Lake Succession polymict breccias) and overlying rhyolitic massive lapilli tuffs and tuffs (the High Level Lake Succession Tuffs). The green to green grey polymict breccia contain a finegrained matrix compose of chlorite, sericite, and locally iron-rich carbonate (ferrodolomite and/or ankerite). Lapilli- to small block-sized clasts vary in abundance from 30-70%, and include: a) easily recognizable, light grey to pale white angular to subangular, locally flowbanded rhyolite lava flow fragments which have been In the F-Group and Darkwater regions, participants will observe Early- and Late Caldera Succession strata which illustrate the three stratigraphic horizons which host VMS orebodies in the SLCC. These stops will include: a) interbedded polymict breccias (High Level Lake Succession polymict breccias) and rhyolite tuffs and lapilli tuffs (High Level Lake Succession rhyodacite-rhyolite lapilli tuffs and tuffs) that formed during simultaneous explosive felsic eruptions and foundering of the region above the shallow ancestral Beidelman Bay intrusion during formation of the Sturgeon Lake Caldera Complex; b) extreme aluminum silicate alteration within the High Level Lake Succession breccias and tuffs that occur down-section and along strike from the F-Group VMS deposit; and c) well-bedded quartz-phyric lapilli tuffs and tuffs that comprise the Mattabi VMS ore horizon (including a small replacement-type sphalerite-rich VMS deposit) in the western part of the caldera complex, and the hanging wall rocks to the Mattabi VMS deposit in this part of the caldera complex, the Middle L Succession rhyodacite-rhyolite lapilli tuffs and tuffs. The Middle L Succession lapilli tuffs and tuffs are the host rocks for the Sturgeon Lake VMS deposit in the eastern part of the SLCC. Figure 20. Geological sketch map of the F-Group area, with field trip stop locations (from Morton et al., 1996). - 140 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 derived from the underlying Darkwater Succession rhyodacite to rhyolite lava flows; b) light grey to pale white, angular to subangular, locally silicified rhyolite tuff lapilli (presumably derived from the underlying Jackpot Lake Succession tuffs); c) green, commonly difficult to recognize, angular to rounded amygdaloidal basalt lapilli (derived from the underlying Darkwater Succession basalt-andesite lava flows); and d) green, commonly difficult to recognize, generally rounded to subrounded, locally pale-brown iron carbonate altered scoria lapilli (derived from the underlying Darkwater Succession andesite tuffs, lapilli tuffs, and tuff breccias). As one moves from the southwestern part of the outcrop toward the northeast, the abundance of lapilli and blocks within the breccia deposits decreases, and the abundance of <1mm quartz phenocrysts increases to 1-2% over a zone which varies from 1-5m in width. This zone represents the gradational contact between the High Level Lake mesobreccia deposits and the overlying and intercalated High Level Lake quartzphyric lapilli tuffs and tuffs. This gradational contact formed during mixing of the two units as they were deposited simultaneously (Hudak, 1989). Similar intercalations of felsic pyroclastic deposits and polymict breccia deposits have been documented in calderas within the San Juan Mountains in Colorado by Lipman (1976), and have been interpreted to represent deposits which result from simultaneous explosive volcanism and caldera collapse. Continuing to move to the north on the outcrop, the gradational contact zone gives way to massive deposits of the High Level Lake Succession lapilli tuffs and tuffs. The light grey quartz- and sericite-rich, recrystallized ash matrix of these deposits contains 5-25% euhedral to subhedral, locally resorbed 1mm quartz phenocrysts and up to several percent, typically difficult to see, subangular to subrounded pumice lapilli. Irregular patches and lenses of red-brown iron carbonate alteration vary from 1-10cm in length, and locally comprise 10-15% of the outcrop. In thin section, many of the altered 1-3cm carbonate patches appear to be altered pumice lapilli. The polymict breccia deposits at this outcrop, and numerous other outcrops in the Sturgeon Lake camp, are interpreted to be mesobreccias and megabreccias. Mesobreccias and megabreccias form from material that slumps off oversteepened walls of a caldera during and after caldera collapse. By definition (Lipman, 1976; Lipman, 1997; Lipman, 2000), megabreccia deposits contain blocks which are dominantly greater than 1m in diameter, whereas mesobreccia deposits contain lapilli and blocks which are dominantly less than 1m in diameter. Megabreccia deposits are generally formed proximal to caldera walls. Mesobreccia and megabreccia deposits which occur in the footwalls to the Mattabi and Sturgeon Lake Mine orebodies are stratigraphically equivalent to the polymict breccia deposits which are observed at this location. Stop 4: Altered High Level Lake Polymict Breccias, Lapilli Tuffs, and Tuffs (optional) Intensely aluminum silicate- and aluminum silicate + chloritoid-altered, intercalated High Level Lake Succession breccias and rhyolite lapilli tuffs / tuffs occur at this location, approximately 300m southwest of the F-Group pit (Fig. 20). Although difficult to recognize, these rocks comprise the same stratigraphic units exposed at the first field trip stop. Here, the High Level Lake Succession polymict breccia deposits vary from green to grey-green to pale pinkish-grey, and contain up to 50% lapilli and blocks. The fragments are of three principal types: 1) 5% rounded, 3-10mm in diameter, intensely amygdaloidal quartz- and sericite-rich fragments which petrographic observations indicate are altered pumice; b) 5-10% light grey lapilli- to small block-sized felsic lava fragments (Darkwater Succession rhyodacite/ rhyolite lava flows); and c) up to 30% subround to oval, 3-10mm diameter chlorite-rich amygdaloidal fragments which petrographic observations suggest are scoria and amygdaloidal basalt. The matrix of this unit is generally composed of magnesium-rich chlorite, quartz, and sericite, but locally, where pale pinkish grey, andalusite and/or pyrophyllite also occur. High Level Lake Succession rhyolite lapilli tuffs and tuffs overlie, and are locally intercalated with, the High Level Lake polymict breccia deposits in this area. The felsic lapilli tuffs and tuffs contain 1-20% 1mm diameter euhedral to subhedral quartz phenocrysts in a recrystallized ash matrix, and vary in colour from greygreen to greyish-pink depending upon the alteration mineral assemblage present. Grey-green regions contain an alteration assemblage of quartz, sericite, and locally magnesium- and/or iron-rich chlorite. Pinkish-grey regions are composed of an alteration mineral assemblage composed dominantly of quartz and up to 1mm diameter ragged anhedral to blocky subhedral andalusite. Light grey, subrounded to oval pumice lapilli vary from 3-20mm in diameter, and - 141 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 are composed of recrystallized quartz. Petrographic observations indicate that the pumice lapilli contain 30-50% <1mm round quartz-filled vesicles. Pervasive aluminum silicate alteration, greenschist facies metamorphism, and subsequent retrograde processes have led to the development of three different aluminum silicate phases within this group of outcrops. Andalusite is the most common aluminum silicate phase present (5-50%), and occurs as 1-6mm ragged anhedral to equant subhedral pink porphyroblasts. Thin section examinations indicate that the ragged andalusite porphyroblasts have locally undergone retrograde reactions and now contain inclusions and rims composed of sericite and/or pyrophyllite. Kyanite is present in two distinctive forms: a) as ragged tabular porphyroblasts up to 5mm in length within the altered matrix of the tuffs; and b) as pale blue blades ranging from 3-20mm in length within white to reddish-brown quartz-iron carbonate veins that are up to several centimeters in width. Pyrophyllite can commonly be found along the margins of both the andalusite and kyanite porphyroblasts, and may also occur in veins up to several centimeters in width as soft, pale greenishwhite radiating micaceous aggregates. Pyrophyllite is commonly present where quartz veins intersect kyaniterich veins. Minor amounts of chloritoid (generally <5% but locally up to 10%) are locally associated with these aluminum silicate minerals. This series of outcrops is interpreted to be proximal to, and in part, within, a synvolcanic fault zone. These faults provided cross-stratal channel ways in which high temperature, acidic metalliferous hydrothermal fluids traveled upward through the subseafloor to the seafloor. The aluminum silicate alteration has been shown via mass balance analyses (Jongewaard, 1989; Hudak, 1989; Hudak, 1996) to have developed as the acidic fluids leached cations from the rocks (for example, during the alteration of feldspar or volcanic glass), leaving them rich in aluminum and silica (and presumable with a clay-rich pre-metamorphic alteration mineral assemblage). Andalusite and kyanite certainly formed during the greenschist facies metamorphism of the strata; however, studies of both modern and ancient hydrothermal alteration assemblages (White and Hedenquist, 1990; White and Hedenquist, 1995) associated with epithermal mineral deposits indicate that andalusite may also form as a primary, high temperature alteration mineral phase. Pyrophyllite may have formed as a primary mineral as well, but based on textural evidence, appears primarily to be due to retrograde metamorphism of kyanite and andalusite. Stop 5: The F-Group Trench Excavated in 1989, the F-Group reclamation trench was designed to channel runoff waters from the F-Group waste dump into the F-Group pit (Fig. 20). Four different lithologies are exposed in this trench: 1) a dark grey to green, locally amygdaloidal gabbro to quartz-diorite sill-like synvolcanic intrusion; 2) light grey to pink, aphyric to locally quartz-phyric, bedded to massive Mattabi Succession lapilli tuffs and tuffs; 3) semi-massive to massive, replacement-type lenses of pyrite ± sphalerite which occur at the equivalent stratigraphic horizon to the “B”-lens of the Mattabi VMS orebody; and 4) grey to grey-green, bedded to massive, quartz- ± feldspar-phyric lapilli tuff and tuff deposits of the Middle L Succession. Several alteration mineral assemblages can also be recognized at this locality. These include: a) iron carbonate ± iron chlorite assemblage; b) the chloritoid assemblage; c) the chloritoid + aluminum silicate ± sericite assemblage; d) the aluminum silicate ± sericite assemblage; and e) locally, silicification. Note the cross-cutting relationships of the various alteration assemblages which can be observed by close examination of the trench wall rocks. Structurally, two different generations of faults have been identified by Walker and Hudak (1989) within the F-Group trench. Synvolcanic structures, which locally led to minor differences in the thicknesses of the volcanic units within the trench, produced cross-stratal permeability that focused upwelling metalliferous fluids in pathways to the paleoseafloor. Metasomatism resulting from interactions between the volcaniclastic strata and these synvolcanic metalliferous fluids produced the extensive hydrothermal alteration in the area. Cooling and neutralization of these fluids by cooler seawater or lower temperature hydrothermal fluids within the tuffs within the shallow seafloor led to the development of replacement-style (Doyle and Allen, 2003) massive sulphide occurrence within the Mattabi tuffs. Locally, strongly sheared, commonly sericite-rich, east-northeast trending high strain zones represent post-volcanic structural deformation which is believed to be related to splays off the northeasttrending Sturgeon Narrows Shear Zone (Figure 1) that is located to the north and west of this location beneath Sturgeon Lake. Stop 6: Bedded and Graded High Level Lake Mesobreccia Deposits (optional) - 142 - As indicated previously, it appears that the initial �Proceedings of the 61st ILSG Annual Meeting - Part 2 development of the Sturgeon Lake Caldera Complex involved trap-door like caldera collapse in which the most significant displacement occurred in the western part of the caldera. Here, in the western one-third of the caldera complex (refer to Fig. 1), we can observe planar thin- to thick-bedded, matrix-supported, normal and locally reverse-graded polymict lapilli-tuffs and tuffs. The finer-grained deposits on this group of outcrops display Ta, Tb, and Te-type bedding units (c.f. Bouma, 1962), whereas coarser deposits display R2R3 and S1-S3 bedding units (c.f. Lowe, 1982). Mattabi Region (Area 16) Intracaldera Strata In the vicinity of the Mattabi VMS deposit, field trip participants will observe: a) the High Level Lake Succession polymict breccias and rhyolite lapilli tuffs and tuffs in the central part of the SLCS: b) thick, massive, pumice-rich, quartz-phyric lapilli tuffs and tuffs that comprise the immediate footwall and host strata to the Mattabi VMS deposit (Mattabi Succession lapilli tuffs and tuffs); and c) hanging wall basalt to andesite pillow lavas, pillow breccias, and locally, bedded volcaniclastic sedimentary strata and peperite (the No Name Lake Succession lava flows and associated interflow sedimentary strata). Previous field trips to the Sturgeon Lake area also included a stop at an exceptional exposure of well-bedded polymict breccia deposits and associated normal graded volcaniclastic sandstones and mudstones which comprise the Tailings Lake Succession sedimentary strata; however, during reclamation, this exposure, which occurs immediately east of the former Mattabi headframe, was buried by approximately 3-5m of fill. Due to the importance of this exposure in terms of understanding the development of the SLCC, we have chosen to include its field trip description although we will not be able to observe the outcrop. The Mattabi VMS deposit was discovered by Mattagami Lake Mines in 1969 from follow-up drilling of airborne geophysical anomalies. The orebodies were mined by open pit and underground mining methods, and comprised five stratiform lenses of massive sulphide ore separated by stringer base-metal mineralization or barren host rock which occur in three distinct stratigraphic successions (Tailings Lake, Mattabi, and Middle L successions, respectively). Massive sulphide lenses which cropped out and extended to approximately 250m below the surface were mined via open pit methods between 1972 and 1980. The deeper VMS deposits were mine using underground mining methods until reserves were depleted in 1988. Combined, the five lenses comprising the Mattabi VMS deposit produced approximately 12.55 million tons of VMS ore grading 8.28% Zn, 0.74% Cu, 0.85% Pb, and 104g/ton Ag (M. Patterson, personal communication, 1990; Franklin, 1996). Stop 7: Mattabi Footwall - High Level Lake Polymict Breccias and Rhyolite Lapilli Tuffs/Tuffs This series of outcrops is located approximately 500m stratigraphically below the lowermost lens of the Mattabi VMS orebody. The southern portion of these outcrops comprises coarse polymict breccias which are interpreted to be mesobreccia deposits formed during caldera collapse. The far northeastern outcrops comprise quartz-phyric rhyolite tuffs which overlie, and are intercalated with, the polymict breccias (the field relationship seen previously in the vicinity of the F-Group orebody). Figures 21 and 22 illustrate the locations of the outcrops described below. Stop 7a: Polymict Breccia Deposits (High Level Lake Succession Mesobreccia) These outcrops comprise coarse, polymict, breccia that contains up to 50% 1-25cm light-coloured subangular felsic lithic fragments (Darkwater Succession rhyolite lava flow lapilli), up to 10% lapillito block-sized pumice fragments, which commonly have silicified rims, and <5% amygdaloidal mafic lapilli, which are up to 5cm in diameter (Darkwater Succession amygdaloidal basalt lapilli). Petrographic observations indicate that the matrix is composed of a mixture of fine-grained recrystallized quartz (40%), chloritoid (12%), magnesium-rich chlorite (30%), white mica (sericite ± pyrophyllite, 15%), and opaque minerals (pyrite and/or magnetite, 3%). The northeastern section of the outcrop contains a bussized felsic block (>10m in diameter) with similar composition to the smaller felsic lapilli which occur throughout the breccia deposits (Walker, 1993). Stop 7b: High Level Lake Succession Tuffs High Level Lake Succession tuffs at this location are light tan to grey in colour and contain 2-5% 0.5-1.5mm diameter subhedral to euhedral, locally resorbed quartz phenocrysts. In thin section, only “ghosts” of 2-10mm diameter pumice lapilli can be observed within a matrix composed of recrystallized, inequigranular polygonal quartz (up to 65%), fine-grained sericite (25%), magnesium-rich chlorite (6%), 3% pyrite, and 1% magnetite and/or ilmenite. Further to the - 143 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 east, the unit contains a coarse, fragmental texture. Here, the rock contains 60-70% lapilli- to blocksized subangular to angular felsic lithic fragments in a felsic matrix. The texture has been interpreted by Walker (1993) to represent deposits formed by postdepositional slumpage of partially consolidated High Level Lake Succession tuffs. Only minor variations in the alteration mineralogy can be found between the matrix and the fragments present at this location. Stop 7c: Chloritoid–Aluminum Silicate-Chlorite Altered High Level Lake Mesobreccia This exposure is composed of High Level Lake polymict breccia deposits which contain 20-30% rounded, 2-30cm, highly amygdaloidal scoria lapilli and blocks, as well as up to 10% 1-15cm subangular felsic lapilli and blocks which are similar in composition to those observed in outcrop M-1a. In thin section, the matrix comprises 30% quartz, 20% magnesium-rich chlorite, 10% iron carbonate, 15% sericite ± pyrophyllite, 10% chloritoid, up to 10% ragged andalusite, and 5% opaque minerals. Chlorite alteration increases toward the center of the outcrop where massive chlorite veining up to 50cm in width occurs. Figure 21. Geological plan map of the Mattabi area, with field trip stop locations (after Walker, 1993; Morton et al., 1996). The chloritic veining at this location is interpreted to have resulted from recharge of cool, fresh seawater into a hot hydrothermal system. This appears to have resulted in magnesium-dumping and subsequent chlorite alteration (Walker, 1993; also see Seyfried et al., 1999). The quartz-filled tension fractures which Figure 22. Geological sketch map of the High Level Lake Succession outcrops south of the Mattabi VMS orebody (after Walker, 1993; Morton et al., 1996). - 144 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 occur at this location may be the result of volume changes due to hydrothermal alteration. Kyanite can locally be found in these fracture-filling veins on the southern part of the outcrop. A 10-20cm wide band of silicified rock, striking North-South, bisects the outcrop. Samples from this silicified rock contain 50% quartz, 15% sericite ± pyrophyllite, 8% iron-rich carbonate, 7% magnesium-rich chlorite, 2% andalusite, and 2% opaque minerals. This band of more intense alteration may represent a conduit responsible for upward, hightemperature, acidic metalliferous fluid movement which was later overprinted by lower temperature, more neutral hydrothermal fluids that produced sericite and chlorite alteration. Stop 7d: Chloritoid – Aluminum Silicate - Fecarbonate Altered Mesobreccia This series of outcrops illustrates a definable, confined and symmetrically zoned increase in alteration intensity within a synvolcanic fault zone in the High Level Lake polymict breccia unit. Original volcanic textures within the breccia unit within this synvolcanic structure are strongly overprinted, but still recognizable when carefully inspected. In particular, one can still relatively easily recognize the abundance of relatively unaltered felsic lithic lapilli. Alteration in the small outcrops on the road and immediately north of the road consist of 1-4cm clots of iron-carbonaterich material surrounded by anastomosing veinlets of quartz, chloritoid, and andalusite. The dominant change in rock mineralogy from the previous outcrop (Stop 7c) is an increase in the amount of Fe-carbonate and chloritoid. Stop 8 (No Longer Available): Mattabi Footwall Tailings Lake Succession Bedded Sediments Note: Reclamation in the vicinity of the former Mattabi Mine headframe has unfortunately resulted in the burying of the classic outcrop of the Tailings Lake Succession. Although this outcrop can no longer be observed, I have included its description below, as the field relationships and textures observed in this former outcrop were extremely important for the development of the Sturgeon Lake Caldera Complex volcanological model described by Morton et al. (2001), Morton et al. (1999), Walker (1993) and Hudak et al. (2003). The Tailings Lake Succession consists primarily of highly variable polymict breccias, sandstones, and mudstones, with minor intercalated felsic tuff horizons and intermediate to mafic lava flows. The polymict breccias vary in their clast composition, clast abundance, and bedding characteristics. Three clast types are most common: 1) mafic lithic clasts, which are commonly replaced by chlorite and/or ironcarbonate (up to 50%); 2) fined-grained cherty felsic lithic lapilli (up to 30%); and 3) rounded pumice lapilli (up to 20%). Bedding is uncommon, but where present, is usually defined by sorting of the clasts, as well as changes in the compositions of the clasts. The Tailings Lake polymict breccia deposits are lithogeochemically indistinguishable from the High Level Lake polymict breccia deposits, and suggests their provenance is similar (e.g., from infilling of a basin by clastic material derived primarily from Pre-caldera strata). Many of the characteristics of the Tailings Lake polymict breccias are conspicuous in the exposure located beside the Mattabi Mine ventilation shaft. This exposure is atypical in its cross-sectional view, and because of the presence of well-defined bedding within the breccia unit. Here, the unit contains 5-30% 2-30mm diameter mafic lapilli (which contain 5-30% oval to rounded iron-carbonate- and quartz-filled amygdules or weathering pits), as well as 2-30mm felsic lithic or pumice lapilli (5-30%). In thin section, the matrix comprises quartz (35%), iron-carbonate(20%), and andalusite (10%), as well as late patches and veins of magnesium-rich chlorite (20%) and sericite/ pyrophyllite (10%). Bedding in the unit is defined by various abundances of mafic and felsic fragments. It appears mafic fragments are generally normal graded and felsic pumice lapilli are reverse graded. Such reverse grading of pumice clasts is commonly attributed to the slow settling of cold, highly vesicular fragments as water infiltrates vesicles in an aqueous environment (Whitham and Sparks, 1986). Bedding trends approximately 100° and dips 60-70° to the north. Stop 9: Mattabi Footwall – Aluminum SilicateAltered Mattabi Lapilli Tuffs and Tuffs This large outcrop on the southwestern side of the Mattabi open pit consists of moderately- to stronglyhydrothermally altered, massive Mattabi Succession quartz-phyric rhyolite lapilli tuffs and tuffs (Fig. 21). This exposure contains 10-50% subrounded to rounded juvenile felsic lapilli, 5-15% 5-30cm (although locally, up to 70-80cm diameter) delicate amoeboid to subrounded pumice, and 2-7% 0.5-1.5mm diameter subhedral to euhedral quartz phenocrysts. The recrystallized altered ash matrix consists of finegrained polygonal quartz (40-50%), chloritoid (1030%), sericite ± pyrophyllite (15-35%, after andalusite), - 145 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 andalusite (up to 5%), and opaque minerals (generally pyrite, 1-5%). Lithic fragments consist of quartz ± sericite ± pyrophyllite and the pumice are replaced, and locally contain amygdules with, quartz ± sulphide minerals. Bounding facies, as well as the presence of massive sulphide mineralization, indicate that these strata were deposited in a subaqueous environment. To the north (up-section), this unit grades into a series of bedded and locally normal graded aphyric tuffs and quartz-phyric tuffs. Petrographic studies indicate the presence of heat retention structures (lithophysae) and spherulites (indicative of the original glassy nature of the deposits) within the massive lapilli tuffs at this location (Walker, 1993: Hudak et al., 2003). White (2000) has shown that massive lapilli tuff deposits, which contain delicate, amoeboid-shaped pumice as well as heat retention textures, which are overlain by bedded tuffs of similar composition represent primary, hot, eruption-fed high concentration mass flows which are essentially true submarine pyroclastic flow deposits (e.g. originally deposited from collapse of hot, gas charged flows of juvenile volcanic material). Hudak et al. (2003) have, based on textures at this outcrop as well as in several diamond drill core intersections, shown that the Mattabi tuffs were, at least locally in the vicinity of the Mattabi VMS orebody, deposited as submarine pyroclastic flows from voluminous submarine explosive rhyolitic volcanism. End of Day 1 of Sturgeon Lake Field Trip Day 2 – The Late Caldera Sequence Lyon Lake and Sturgeon Lake Mine Regions (Areas 17 and 23) In this part of the caldera complex, we will observe two important units within the Late Caldera Sequence: 1) the Middle L Succession tuff breccias, lapilli tuffs, and tuffs (which are the host rocks to the Sturgeon Lake VMS deposit approximately 1.5km to the east of the field trip stop location); and 2) the Lyon Creek Succession dacite cryptodome–lava dome complex, which are interpreted to comprise the final igneous products associated with the genesis of the Sturgeon Lake Caldera Complex. Four VMS orebodies occur in the eastern part of the SLCC: 1) the Sturgeon Lake Mine, a 2.07 million ton ore deposit mined by open pit methods which contained 2.95% Cu, 9.17% Zn, 1.21% Pb, and 164g/ ton Ag; 2) the Lyon Lake deposit, a 3.95 million ton ore deposit mined by underground methods which contained 1.24% Cu, 6.53% Zn, 0.63% Pb, and 142g/ ton Ag; 3) the Creek Zone deposit; mined via open pit methods; and 4) the Sub-Creek Zone deposit, mined by underground methods. Combine, the Creek Zone and Sub-Creek Zone deposits contained 0.91 million tons of ore which graded 1.66% Cu, 8.80% Zn, 0.76% Pb, and 141g/ton Ag (Franklin, 1996). The genesis of these four VMS deposits has historically been controversial in the Sturgeon Lake Camp. The Lyon Lake, Creek Zone, and Sub-Creek Zone deposits were originally known as the NBU ore deposits. Harvey and Hinzer (1981) interpreted abrupt facies changes, coarser volcaniclastic rocks, increased alteration intensity, and a greater MnO/FeO ratio in the Sturgeon Lake, Creek Zone, and Lyon Lake deposits to indicate formation along the same stratigraphic horizon at various distances from high temperature hydrothermal vents. Harvey and Hinzer (1981) suggested that these VMS orebodies occurred stratigraphically up-section from the Mattabi and F-Group VMS deposits. Severin (1981) postulated that the Sturgeon Lake deposit occurred on the same stratigraphic horizon as the Mattabi orebody, and that the Lyon Lake, Creek Zone, and Sub-Creek Zone deposits formed in topographic lows from hydrothermal activity which post-dated the formation of the Sturgeon Lake deposit. More recent detailed mapping (Dube et al., 1989: Koopman, 1993: Hudak, 1996; Morton et al., 1999), petrographic studies, and lithogeochemical evaluations now indicate that the Sturgeon Lake deposit is cut by a major postvolcanic fault zone, and that the Lyon Lake, Creek Zone, and Sub-Creek Zone deposits represent parts of the Sturgeon Lake deposit which were moved into their present locations by this structural deformation. Stop 10. Aluminum Silicate Altered Quartz-Phyric Middle L Tuffs The Middle L Succession comprises a sequence up to 150m thick composed of quartz-phyric rhyolite tuff breccias, lapilli tuffs, and tuffs which can be followed along strike for at least 15km across the SLCC. VMS orebodies occur within this sequence of rocks at both the Mattabi (Mattabi ore lens “A”) and Sturgeon Lake Mine deposits. In addition, anomalous Cu and Zn concentrations occur in the Middle L tuffs under Sturgeon Lake near the F-Group – Area 15 property boundary. - 146 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 This small outcrop (Fig. 23), located immediately south of a drill road which runs approximately eastwest through Area 17, is one of the few exposures of the Middle L Succession in the eastern part of the caldera complex. At this location, the matrix varies from pale grey to pinkish-grey in colour and contains 1-2% <12mm subhedral to euhedral quartz phenocrysts. Pumice and felsic lithic lapilli are not obvious on this exposure, but occur in minor amounts (1-2%) in thin sections and from diamond drill core intersections of the unit. The colour of this outcrop is largely due to the presence of 5-15% 1-3mm blocky subhedral andalusite and finegrained quartz. Immediately down-dip and up-section from this exposure, a thin (1m thick) massive sulphide horizon occurs within bedded ash deposits in diamond drill hole 17-64. Due to a lack of exposure, it is difficult to determine whether the aluminum silicate alteration at this location represents disconformable alteration within a synvolcanic structure, or semi-conformable aluminum silicate alteration along the stratigraphic horizon which hosts the Sturgeon Lake deposit to the east. Stop 11: Middle L Succession Tuff Breccia This stop is located immediately north of the Area 17 drill road approximately 30m north of the previous stop (Fig. 23). The spectacular volcanic breccia (tuffbreccia) is composed of a chlorite-, sericite-, iron carbonate-, and locally biotite-altered matrix which contains 30-60% subangular to angular, 1-35cm light grey felsic lava lapilli and blocks which contain 1-5% 1mm quartz phenocrysts, as well as rare massive sulphide lapilli and iron-carbonate-altered pumice lapilli. The felsic lava flow fragments are composed of spherulitic quartz- and K-spar-phyric rhyolite lava, are vaguely reverse graded, commonly have fine-grained (apparently chilled) rims, and often exhibit jigsaw puzzle-fit with adjacent fragments. The tuff-breccia deposits can be followed several hundred meters down dip in numerous diamond drill holes within Area 17. Bounding facies indicate that these deposits were formed in a submarine environment. Hudak (1996) and Hudak et al. (2003) believe that these tuff-breccia deposits resulted from the collapse of a Middle L Succession submarine lava dome. This lava dome was likely located near a synvolcanic fault within a few hundred meters of this location. This tuffbreccia essentially represents deposits from block and ash flows that occurred in a submarine environment (Gibson et al., 1999). Massive sulphide lapilli within the tuff-breccias suggest that massive sulphides were being deposited on or within the lava dome prior to its collapse. Stop 12: Lyon Creek Succession Cryptodome The Lyon Creek Succession is composed of the youngest strata clearly associated with the development of the Sturgeon Lake Caldera Complex. As indicated above, these andesitic to dacitic lava flows, lava domes, cryptodomes and associated clastic and chemical sedimentary strata have been interpreted by Hudak (1996), Morton et al. (1999), and Hudak et al. (2003) to represent lava dome building and associated intracaldera clastic and hydrothermal sedimentation associated with the terminal stages of caldera development within a Valles-like caldera cycle (Smith and Bailey, 1968). Figure 23. Geological plan map in Areas 17 and 23 (after Walker, 1993; Morton et al., 1996). This stop (Fig. 23), located near the top of a small hill immediately west of the Lyon Lake mine road, is one of a handful of small exposures of the Lyon Creek dacite cryptodome. At this location we observe light grey massive plagioclase-phyric dacite lava which comprises the central part (and most - 147 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 coarsely porphyritic) of the cryptodome. Hudak et al. (in prep.) have shown a systematic increase in the maximum size of plagioclase phenocrysts from the margins toward the center of the cryptodome. Tan to orange-brown, locally carbonate-altered tabular plagioclase phenocrysts (10-15%) generally vary from 1-3mm in diameter. Locally, one may observe plagioclase phenocrysts as large as 6mm in diameter. Hydrothermal alteration at this location comprises anastomosing 2-10mm wide veins containing ironcarbonate, chlorite, magnetite, and locally, burgundyred Mn-rich almandine garnets. Locally, 1-2mm tabular deep green chloritoid porphyroblasts occur in the cryptodome immediately adjacent to the veins, and suggest hydrothermal processes which included the chemical breakdown of feldspar combined with iron metasomatism to form chloritoid. Alteration in this outcrop is genetically associated with the formation of Algoma-type banded iron formation which occurs upsection on the northeastern margin of the cryptodome– dome complex. matrix consist of 0.5-2cm angular andesite lapilli within a blocky, fine- to coarse-ash-sized matrix composed of delicately preserved hyaloclastite with convex fragment edges. Interflow sedimentary strata are finely bedded, and are locally intruded by amoeboid to relatively straight, locally discontinuous amygdaloidal andesite dikes which are up to 50cm in width. Look for blocky peperite where the wet sediments interacted with magma along the edges of the dikes. Stop 13: No Name Lake Andesite and Interflow Sediments and Peperite At this, the final stop of our field trip (Fig. 21), we will observe “classic” pillow lavas associated with the No Name Lake Succession. At this location, the extremely well preserved, moderately- to highly amygdaloidal “bun-” and “mattress-”shaped pillows (nomenclature of Dimroth et al., 1978) illustrate exception concentric cooling cracks and locally, what may be multiple pillow selvedges. Stratigraphic topping directions for these pillow lavas are consistently to the north. The tannishgreen color of these submarine lava flows is indicative of moderate intensity iron carbonate ± iron chlorite alteration that has been observed petrographically. The No Name Lake Succession comprises basaltic to andesitic sheet flows, pillow lavas, pillow breccias, interflow sedimentary rocks, and locally, peperites. Only the uppermost section of this succession is exposed at this location (Fig. 21); the lower sections can be observed only in diamond drill core and appear to consist primarily of thick, amygdaloidal sheet flows and pillow lavas. The first outcrop (behind the core racks) consists of thin (30-70cm thick) sheet flows with 5-25% oval-shaped carbonate-filled amygdules (2-30mm in diameter) that are generally aligned parallel to strike and the dominant east-west-trending rock foliation. In thin section, these rocks are composed of 20-30% fine laths of plagioclase and 15% quartz in a secondary groundmass composed of chlorite (40%), biotite (10%), and epidote (2%). Pillow breccia, hyaloclastite, interflow sedimentary strata, and peperite are exposed in the outcrop immediately east of the water tower. These rocks consist of approximately 25% amygdaloidal pillow lapilli and blocks (5-50cm) in a matrix comprising hyaloclastite. The pillow breccia fragments contain 10-20%, 1-4cm diameter carbonate-filled amygdules (these are commonly weathered-out to form small pits on the outcrop surface). The hyaloclastite portion of the Pillowed andesite flows are exposed in the outcrop north of the water tower. This rock consists of wellformed 1-4m long amygdaloidal pillows with 25-30% 1-2mm carbonate-filled amygdules. The amygdules illustrate a bimodal size distribution; most amygdules are 1-4mm, with another distinct group being larger and up to 30mm in diameter. Massive pillow selvedges vary from 5-15cm thick. Such thick selvedges may indicate proximity to an eruptive vent (Kennish and Lutz, 1998; Hudak et al., 2002). Stop 14: No Name Lake Andesite Pillow Lavas End of Day 2 of Sturgeon Lake Field Trip Acknowledgements The author would like to thank the many people and organizations that have made his research in the Sturgeon Lake region possible, and have made this field trip a reality. First and foremost, my mentors Ron Morton and Jim Franklin gave me numerous opportunities during my graduate and post-doctoral research to map, analyze, interpret, reinterpret, and publish many of our findings during our 20+ years of research in the Sturgeon Lake area. As well, Ron and Jim, along with the University of Minnesota Duluth and the Geological Survey of Canada, provided me not only research funding, but found ways to fund my travels to many key geological locations and - 148 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 important geological conferences that enhanced my ability to understand the geological environment and mineralization associated with the Sturgeon Lake region. I benefitted immensely from my years of research with my Sturgeon Lake research colleagues Jamie Walker and Peter Jongewaard. As well, Dean Peterson played a major role in developing the GIS studies that helped to further understand the evolution of the SLCC. Having the ability to toss around and develop ideas with these three exceptional field-oriented economic geologists was instrumental in developing many of the key concepts needed to understand the volcanological, hydrothermal, and mineralization processes that took place within the Sturgeon Lake Caldera Complex. As well, funding from Noranda Exploration, Mattabi Mines Ltd., Rio Algom, and Minnova during the late 1980’s and early 1990’s was instrumental in being able to conduct the extensive field mapping programs, petrographic research, and geochemical studies necessary to evaluate this spectacular volcanic feature and its associated mineral deposits. Key industry personnel that contributed greatly to further understanding the SLCC included Wally Gibb (Mattabi Mines, Ltd.), Mike Patterson (Mattabi Mines, Ltd.), Al Smith (Noranda Exploration) and Ron Kennedy (Mattabi Mines, Ltd.). Lucy Potter and Aaron MacDonell (both with Glencore Canada Corporation) are thanked for their key roles in allowing access to the Sturgeon Lake VMS camp for this field trip. Finally, I am humbled that the Institute on Lake Superior Geology requested that I lead this field trip to this exceptional geological area. References Bacon, C.R. and Druitt, T.H. 1988. Compositional evaluation of the zone calc-alkaline magma chamber of Mount Mazama, Crater Lake, Oregon: Contributions to Mineralogy and Petrology, v.98, p.224-256. Bernier, F., Stevenson, R.K., Gariepy, C., and Franklin, J.M. 1999. Nd isotopic studies in the south Sturgeon Lake Greenstone Belt, northwestern Ontario: a progress report: Fifth Annual Workshop, Lithoprobe Western Superior Transect, Canada, p.117-121. Bouma, A.H.. 1962. Sedimentology of some Flysch Deposits: A Graphic Approach to Facies Interpretation: Elsevier, Amsterdam, 198p. Busby-Spera, C.J. 1984. Large volume rhyolitic ash flow eruptions and submarine caldera collapse in the lower Mesozoic Sierra Nevada, California. J. Geophysical Research, v.89, p.8417-8427. Campbell, I.H., Franklin, J.M., Gorton, M.P., Hart, T.R., and Scott, S.D. 1981. The role of subvolcanic sills in the generation of massive sulphide deposits: Economic Geology, v.76, p.2248-2253. Cas, R.A.F. and Wright, J.V. 1987. Volcanic Successions: Modern and Ancient. Allen and Unwin, London. Davis, D.W. and Trowell, N.F. 1982. U-Pb zircon ages from the eastern Savant Lake-Crow Lake metavolcanicmetasedimentary belt, northwestern Ontario. Canadian Journal of Earth Sciences, v.19, p.868-877. Davis, D.W., Krogh, T.E., Hinzer, J., and Nakamura, E. 1985. Zircon dating of polycyclic volcanic at Sturgeon Lake and implications for base metal mineralization. Economic Geology, v.80, p.1942-1952. Dimroth, E., Cousineau, P., Leduc, M., and Sanschagrin, Y. 1978. Structure and organization of Archean subaqueous lava flows, Rouyn-Noranda area, Quebec, Canada: Canadian Journal of Earth Sciences, v.15, p.902-918. Doyle, M.G. and Allen, R.L. 2003. Subsea-floor replacement in volcanic-hosted massive sulphide deposits: Ore Geology Reviews, v.23, p.183-222. Druitt, T.H. and Francaviglia, V. 1991. Caldera formation on Santorini and the physiography of the islands in the late Bronze Age: Bulletin of Volcanology, v.54, p.484-493. Dube, B., Koopman, E.R., Franklin, J.M., Poulsen, K.H., and Patterson, M.R. 1989. Preliminary study of the stratigraphic and structural controls of the Lyon Lake massive sulphide deposit, Wabigoon Subprovince, northwest Ontario: in Current Research, Part C, Geological Survey of Canada, Paper 89-1C, p.275284. Fiske, R.S., Naka, J., Iizasa, K., Yuasa, M., and Klaus, A. 2001. Submarine silicic caldera at the front of the Izu-Bonin Arc, Japan: voluminous seafloor eruptions of rhyolite pumice: Geological Society of America Bulletin, v.113, p.813-824. Franklin, J.M. 1996. Volcanic-associated massive sulphide base metals; in Eckstrand, O.R., Sinclair, W.D., Thorpe, R.I. (eds.), Geological Survey Canada, Geology of Canada 8, p.158-183. Franklin, J.M., Kasarda, J., and Poulsen, K.H. 1975. Petrology and chemistry of the alteration zone of the Mattabi massive sulphide deposit: Economic Geology, v.70, p.63-79. Galley, A. 2002. Characteristics of composite subvolcanic intrusive complexes associated with Precambrian VMS districts: in Balley, A., Bailes, A., Hannington, M., Holk, G., Katsube, J., Paquette, F., Paradis, S., Santaguida, F, and Taylor, B., Database for Camiro Project 94E07: Interrelationships between subvolcanic intrusions, large-scale alteration zones, and VMS deposits: Geological Survey of Canada Open File Report 4431, p.1-40. Galley, A. 2003. Composite synvolcanic intrusions associated with Precambrian VMS-related hydrothermal systems: Mineralium Deposita, v.38, p.443-473. - 149 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Galley, A., van Breemen, O., and Franklin, J. M. 2000. The relationship between intrusion-hosted Cu-Mo mineralization and the VMS deposits of the Archean Sturgeon Lake mining camp, northwestern Ontario; Economic Geology, v.95, p.1543-1550. Gibson, H.L., Morton, R.L., and Hudak, G.J. 1999. Submarine volcanic processes, deposits and environments favourable for the location of volcanicassociated massive sulphide deposits: Rev. Economic Geology, v.8, p.13-51. Groves, D.A. 1984. Stratigraphy and alteration of the footwall volcanic rocks beneath the Archean Mattabi massive sulphide deposit, Sturgeon Lake, northwestern Ontario: unpublished M. Sc. Thesis, University of Minnesota – Duluth, 115 p. Groves, D.A., Morton, R.L., and Franklin, J.M. 1988. Physical volcanology of the footwall rocks near the Mattabi massive sulphide deposit, Sturgeon Lake, Ontario; Canadian Journal of Earth Sciences, v.25, p.280-291. Harvey, J.D. and Hinzer, J.B. 1981. Geology of the Lyon Lake ore deposits, Noranda Mines Limited, Sturgeon Lake area, Ontario: Canadian Institute of Mining and Metallurgy Bulletin 74, p.77-84. Holk, G., Taylor, B., Galley, A., Hannington, M., and Timbal, A. 2002. Geochemical and alteration studies of the Sturgeon Lake Caldera Complex: in Balley, A., Bailes, A., Hannington, M., Holk, G., Katsube, J., Paquette, F., Paradis, S., Santaguida, F, and Taylor, B., Database for Camiro Project 94E07: Interrelationships between subvolcanic intrusions, large-scale alteration zones, and VMS deposits; Geological Survey of Canada Open File Report 4431, p.333-369. Hudak, G.J. 1989. The physical volcanology and hydrothermal alteration associated with the F-Group Archean massive sulphide deposit, Sturgeon Lake, northwestern Ontario: unpublished M. Sc. Thesis, University of Minnesota – Duluth, Duluth, Minnesota, 172p. Hudak, G.J. 1996. The physical volcanology and hydrothermal alteration associated with late caldera volcanic and volcaniclastic rocks and volcanogenic massive sulphide deposits in the Sturgeon Lake region of northwestern Ontario: unpublished Ph. D. Dissertation, University of Minnesota, Minneapolis, Minnesota, 463p. Hudak, G.J., Heine, J.J., Newkirk, T.T., Odette, J., and Hauck, S. 2002. Comparative geology, stratigraphy, and lithogeochemistry of the Five Mile Lake, Quartz Hill, and Skeleton Lake VMS occurrences, Vermilion District, NE Minnesota: Natural Resources Research Institute Open File Report NRRI/TR-2002/03, 350p. Hudak, G.J., Morton, R.L., Franklin, J.M., and Peterson, D.M. 2003. Morphology, distribution, and estimated eruption volumes for intracaldera tuffs associated with volcanic-hosted massive sulphide deposits in the Archean Sturgeon Lake Caldera Complex, northwestern Ontario: American Geophysical Union Geophysical Monograph 140 (Explosive Subaqueous Volcanism), p.345-360. Hudak, G.J., Morton, R.L., Franklin, J.M., Peterson, D.M., Walker, J.S., Jongewaard, P.K., and Murphy, C. in prep. Lithostratigraphy and volcanic evolution of the Archean Sturgeon Lake Caldera Complex. Hudak, G., Morton, R., and Peterson, D. 2008. Field guide to the volcanology, structure, alteration, and mineralization of Archean greenstone belts in the vicinities of Sturgeon Lake and Rainy River, Ontario, and Lake Vermilion, Minnesota: Short Course and Field Investigation of Physical Volcanology, Structure, and Hydrothermal Alteration associated with VMS and Lode Gold Deposits in Archean Greenstone Belts: Precambrian Research Center Guidebook 08-01, 209p. Jongewaard, P.K. 1989. Physical volcanology and hydrothermal alteration of the footwall rocks to the Archean Sturgeon Lake massive sulphide deposit: unpublished M. Sc. Thesis, University of Minnesota – Duluth, Duluth, Minnesota, 141p. Kennish, M.J. and Lutz, R.A. 1998. Morphology and distribution of lava flows on mid-ocean ridges: a review: Earth-Science Reviews, v.43, p.63-90. Kokelaar, P. and Busby, C. 1992. Subaqueous explosive eruptions and welding of pyroclastic deposits: Science, v.257, p.196-201. King, E.M., Valley, J.W., and Davis, D.W. 2000. Oxygen isotope evolution of volcanic rocks at the Sturgeon Lake volcanic complex, Ontario: Canadian Journal of Earth Sciences, v.37, p.39-50. Koopman, E.R. 1993. Stratigraphy, structural geology, and stratigraphic controls of ore distribution of the Lyon Lake massive sulphide deposit, Sturgeon Lake, Ontario: unpublished M. Sc. Thesis, Carleton University, Ottawa, Ontario, 170 p. Lipman, P.W. 1976. Caldera collapse breccias in the western San Juan Mountains, Colorado. Geol. Soc. Am. Bull. 87, p.1397-1410. Lipman, P.W. 1997. Subsidence of ash-flow calderas: relation to caldera size and magma chamber geometry. Bull. Volcanol. 59, p.198-218. Lipman, P.W. 2000. Calderas. In Sigurdsson, H. (ed.), Encyclopedia of Volcanoes, Academic Press, p.643662. Lowe, D.R. 1982. Sediment gravity flows II: Depositional models with special reference to the deposits of high density turbidity currents: Journal of Sedimentary Petrology, v.52, p.279-297. Monecke, T., Petersen, S., and Hannington, M. 2014. Constraints on water depth of massive sulfide formation: evidence from modern seafloor hydrothermal systems in arc-related settings: Economic Geology, v.109, p.2079-2101. - 150 - �Proceedings of the 61st ILSG Annual Meeting - Part 2 Morgan, L.A. and Schulz, K.J. 2012. Physical volcanology of volcanogenic massive sulfide deposit in volcanogenic massive sulfide occurrence model: U. S. Geological Survey Scientific Investigations Report 2010-5070C, chap. 5, 36p. Morton, R.L. and Franklin, J. M. 1987. Twofold classification of Archean volcanic-associated massive sulphide deposits: Econ. Geol. 82, p.1057-1063. Morton, R.L., Hudak, G.J., and Franklin, J.M. 1999. Geology, south Sturgeon Lake area, Ontario. Geol. Surv. Canada Open File Rpt. 3642. Morton, R.L., Hudak, G.J., and Koopman, E. 1996. Physical volcanology, hydrothermal alteration, and massive sulphide deposits of the Sturgeon Lake Caldera. GAC-MAC Field Trip Guidebook B3, 37p. Morton, R.L., Walker, J.S., Hudak, G.J., and Franklin, J.M. 1991. The early development of an Archean submarine caldera complex with emphasis on the Mattabi ash flow tuff and its relationship to the Mattabi massive sulphide deposit: Econ. Geol. 86, p.1002-1011. Moss, R. 1992. An oxygen isotope study of the Lyon Lake area, northwestern Ontario: Potential application to exploration for volcanogenic massive sulphide deposits in medium grade metamorphic terrains: unpublished B. Sc. Thesis, University of Toronto, Toronto, Ontario, 57p. Mueller, W.U., Stix, J., White, J.D.L., Corcoran, P.L., Lafrance, B., and Daigneault, R. 2008. Characterization of Archean subaqueous calderas in Canada: physical volcanology, carbonate-rich hydrothermal alteration, and a new exploration model: in Gottsman, J., and Marti, J., (editors), Developments in Volcanology, Volume 10, Caldera Volcanism: Analysis, Modeling and Response: Elsevier, Amsterdam, p.181-232. Mueller, W.U., Stix, J., White, J.D.L., and Hudak, G.J. 2004. Chapter 4.6: Archean Calderas: in Eriksson, P. G., Altermann, W., Nelson, D. R., Mueller, W. U., and Catuneanu, O., eds., The Precambrian Earth: Elsevier, p.345-357. Nedimović, M.R. and West, G.F. 2002. Shallow threedimensional structure from two-dimensional crooked line seismic reflection data over the Sturgeon Lake Volcanic Complex: Economic Geology, v.97, p.17791794. Poulsen, K.H. and Franklin, J.M. 1981. Copper and gold mineralization in an Archean trondhjemite intrusion, Sturgeon Lake, Ontario: Geol. Surv. Canada Paper 81-1A, p.9-14. Sanborn-Barrie, M. and Skulski, T. 1999. Tectonic assembly of continental margin and oceanic terranes at 2.7 Ga in the Savant Lake-Sturgeon Lake greenstone belt, Ontario: in Current Research 1999-C, Geological Survey of Canada, p.209-220. Sanborn-Barrie, M., Skulski, T., and Parker, J. 2001. 300 m.y. of tectonic history recorded by the Red Lake greenstone belt, Ontario: in Current Research 2001C, Geological Survey of Canada, p.15-22. Sanborn-Barrie, M., Skulski, T., and Whalen, J.B. 1998. Tectono-stratigraphy of central Sturgeon Lake, Ontario: deposition and deformation of submarine tholeiites and emergent calc-alkaline volcanosedimentary sequences: in Current Research 1998-C, Geological Survey of Canada, p.115-126. Seyfried, W.E., Jr., Ding, K., Berndt, M. E., and Chen, X. 1999. Experimental and theoretical controls on the composition of mid-ocean ridge hydrothermal fluids: Reviews in Economic Geology, 8, 181-200. Severin, P. W. A., 1981, Geology of the Sturgeon Lake Cu0Zn-Pb-Ag-Au deposit: C.I.M. Bulletin, v. 75, no. 846, p. 107-123. Smith, R. L. and Bailey, R. A. 1968. Resurgent Cauldrons: Geol. Soc. America Mem. 116, p.613-662. Trowell, N.F. 1974. Geology of the Bell Lake – Sturgeon Lake area, Districts of Kenora and Thunder Bay. Ontario Div. Mines Rep. 114, 67p. Trowell, N.F. 1983. Geology of the Sturgeon Lake area, Districts of Thunder Bay and Kenora: Ontario Geol. Surv. Rept. 221, 97p. Walker, J.S. 1993. Physical volcanology and hydrothermal alteration of the footwall rocks to the Archean Mattabi massive sulphide deposit, northwestern Ontario: unpublished M. Sc. Thesis, University of Minnesota – Duluth, 174p. Walker, J.S. and Hudak, G.J. 1990. Detailed geological mapping, F-Group reclamation trench “A”: Mattabi Mines Ltd. Unpublished map. White, J.D.L. 2000. Subaqueous eruption-fed density currents and their deposits; Precambrian Res. 101 (24), p.87-109. White, N.C. and Hedenquist, J.W. 1990. Epithermal environments and styles of mineralization: variations and their causes, and guidelines for exploration: Journal of Geochemical Exploration, v.36, p.445474. White, N.C. and Hedenquist, J.W. 1995. Epithermal gold deposits: styles, characteristics, and exploration: Society of Economic Geology Newsletter, v.23, p.113. Whitham, A.G. and Sparks, R.S.J. 1986. Pumice: Bulletin of Volcanology, v.48, p.209-223. Winchester, J.A. and Floyd, P.A. 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements: Chem. Geol. 20, p.325-343. Wright, I C., Gamble, J.A., and Shane, P.A.R. 2003. Submarine silicic volcanism at the Healy Caldera, southern Kermadec Arc (SW Pacific): I – volcanology and eruption mechanisms: Bulletin of Volcanology, v.65, p.15-29. - 151 - � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2015 Description An account of the resource Institute on Lake Superior Geology. Dryden, Ontario. May 20-24, 2015. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2015 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/88b14dde3e7a16aeb5b6b67dc352d17c.pdf f477a80c8302e8a3243aa0d3acca505a PDF Text Text Institute on Lake Superior Geology 62ND ANNUAL MEETING May 4-8, 2016 Duluth, Minnesota Sponsored by PRECAMBRIAN RESEARCH CENTER AND DEPARTMENT OF EARTH AND ENVIRONMENTAL SCIENCES AT THE UNIVERSITY OF MINNESOTA DULUTH Meeting Co-Chairs James Miller, Christian Schardt, and Dean Peterson Proceedings Volume 62 Part 1 – Program and Abstracts Edited by Christian Schardt and Jim Miller \ i �ii �Table of Contents Institutes on Lake Superior Geology, 1955-2016 iv Sam Goldich and the Goldich Medal vi Goldich Medal Guidelines viii Goldich Medalists and Goldich Medal Committee x Citation for Goldich Medal Award to Mark Jirsa xi Memorial to Leon Gladen xiii Eisenbrey Student Travel Awards xiv Joe Mancuso Student Research Awards xv Doug Duskin Student Paper Awards and Award Committee xvi Board of Directors, Local Committee, and Session Chairs xvii Field Trip Leaders xviii Corporate and Individual Sponsors of Student Travel Scholarships xix Report of the Chair of the 61st Annual Meeting xx Duluth Entertainment and Conventions Center Floor Plan xxii Technical Program xxiii Poster Presentations xxx Abstracts 1-160 Reference to abstracts in Part 1 should follow the example below: Authors, 2016, abstract title. 62nd Institute on Lake Superior Geology Proceedings v. 62, Part 1-Program and Abstracts, p. XX. Proceedings Volume 62, Part 1—Program and Abstracts, and Part 2—Field Trip Guidebook are published by the 62nd Institute on Lake Superior Geology and distributed by the Institute Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca Some figures in this volume were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the volume when it is available on-line at http://www.lakesuperiorgeology.org. ISSN 1042-99 iii �Institutes on Lake Superior Geology, 1955-2016 95 o o 85 o Wabigoon subprovince90 o 80 48 o Wawa-Abitibi subprovince 48 o Wawa-Abitibi subprovince o 45 45o Minnesota River Valley subprovince MEETING LOCATIONS Phanerozoic Mesoproterozoic Map by Mark Jirsa 95o Paleoproterozoic o 90 85o Archean Superior Province # Date Place Chairs 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 1955 1956 1957 1958 1959 1960 1961 1962 1963 1964 1965 1966 1967 1968 1969 1970 1971 1972 1973 1974 1975 1976 Minneapolis, Minnesota Houghton, Michigan East Lansing, Michigan Duluth, Minnesota Minneapolis, Minnesota Madison, Wisconsin Port Arthur, Ontario Houghton, Michigan Duluth, Minnesota Ishpeming, Michigan St. Paul, Minnesota Sault Ste. Marie, Michigan East Lansing, Michigan Superior, Wisconsin Oshkosh, Wisconsin Thunder Bay, Ontario Duluth, Minnesota Houghton, Michigan Madison, Wisconsin Sault Ste. Marie, Ontario Marquette, Michigan St. Paul, Minnesota C.E. Dutton A.K. Snelgrove B.T. Sandefur R.W. Marsden G.M. Schwartz & C. Craddock E.N. Cameron E.G. Pye A.K. Snelgrove H. Lepp A.T. Broderick P.K. Sims & R.K. Hogberg R.W. White W.J. Hinze A.B. Dickas G.L. LaBerge M.W. Bartley & E. Mercy D.M. Davidson J. Kalliokoski M.E. Ostrom P.E. Giblin J.D. Hughes M. Walton iv �# Date 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 Place 1977 1978 1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 Chairs Thunder Bay, Ontario Milwaukee, Wisconsin Duluth, Minnesota Eau Claire, Wisconsin East Lansing, Michigan International Falls, Minnesota Houghton, Michigan Wausau, Wisconsin Kenora, Ontario Wisconsin Rapids, Wisconsin Wawa, Ontario Marquette, Michigan Duluth, Minnesota Thunder Bay, Ontario Eau Claire, Wisconsin Hurley, Wisconsin Eveleth, Minnesota Houghton, Michigan Marathon, Ontario Cable, Wisconsin Sudbury, Ontario Minneapolis, Minnesota Marquette, Michigan Thunder Bay, Ontario Madison, Wisconsin Kenora, Ontario Iron Mountain, Michigan Duluth, Minnesota Nipigon, Ontario Sault Ste. Marie, Ontario Lutsen, Minnesota Marquette, Michigan Ely, Minnesota 56 2010 International Falls, Minnesota 57 58 59 60 61 62 2011 2012 2013 2014 2015 2016 Ashland, Wisconsin Thunder Bay, Ontario Houghton, Michigan Hibbing, Minnesota Dryden, Ontario Duluth, Minnesota v M.M. Kehlenbeck G. Mursky D.M. Davidson P.E. Myers W.C. Cambray D.L. Southwick T.J. Bornhorst G.L. LaBerge C.E. Blackburn J.K. Greenberg E.D. Frey & R.P. Sage J. S. Klasner J.C. Green M.M. Kehlenbeck P.E. Myers A.B. Dickas D.L. Southwick T.J. Bornhorst M.C. Smyk L.G. Woodruff R.P. Sage & W. Meyer J.D. Miller & M.A. Jirsa T.J. Bornhorst & R.S. Regis S.A. Kissin & P. Fralick M.G. Mudrey & Jr., B.A. Brown P. Hinz & R.C. Beard L. Woodruff & W.F. Cannon S. Hauck & M. Severson M. Smyk & P. Hollings A. Wilson & R. Sage L. Woodruff & J. Miller T. Bornhorst & J. Klasner J. Miller, G. Hudak, & D. Peterson M. Jirsa, P. Hollings, & T. Boerboom, P. Hinz & M.Smyk T. Fitz P. Hollings T. Bornhorst & A. Blaske J. Miller & M. Jirsa R. Cundari & P. Hinz J. Miller, C. Schardt, & D. Peterson �Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970’s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 vi �INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL vii �Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. viii �Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. ix �Goldich Medalists 1979 Samuel S. Goldich 1997 Ronald P. Sage 1980 not awarded 1998 Zell Peterman 1981 Carl E. Dutton, Jr. 1999 Tsu-Ming Han 1982 Ralph W. Marsden 2000 John C. Green 1983 Burton Boyum 2001 John S. Klasner 1984 Richard W. Ojakangas 2002 Ernest K. Lehmann 1985 Paul K. Sims 2003 Klaus J. Schulz 1986 G.B. Morey 2004 Paul Weiblen 1987 Henry H. Halls 2005 Mark Smyk 1988 Walter S. White 2006 Michael G. Mudrey 1989 Jorma Kalliokoski 2007 Joseph Mancuso 1990 Kenneth C. Card 2008 Theodore J. Bornhorst 1991 William Hinze 2009 L. Gordon Medaris, Jr 1992 William F. Cannon 2010 William D. Addison & Gregory R. Brumpton 1993 Donald W. Davis 2011 Dean M. Rossell 1994 Cedric Iverson 2012 James D. Miller 1995 Gene La Berge 2013 Tom Waggoner 1996 David L. Southwick 2014 Laurel Woodruff 2015 Rodney J. Ikola 2016 GOLDICH MEDAL RECIPIENT Mark Jirsa Goldich Medal Committee Serving through the meeting year shown in parentheses. Mark Smyk (2016) Ontario Geological Survey Hélène Lukey (2017) Cliffs Natural Resources Shannon Zurevinski Lakehead University x �Citation for the Goldich Medal Award to Mark A. Jirsa I am honored to be able to present the 2016 Goldich Medal to Mark Jirsa. I first met Mark in 1983, when I was hired for a six month stint at the Minnesota Geological Survey. The first time I walked into the MGS office I was immediately introduced to both Mark Jirsa and Jim Miller, and the three of us were quickly dispatched to northern Minnesota to conduct field mapping for an unusual project involving the Minnesota Waste Management Board. I soon realized that I was working with a mapping Zen-master, and was very fortunate to have started my career with not one, but two people who are now both at the forefront of their respective areas of expertise. It was also the beginning of a long friendship and comradery. Mark’s career has never strayed far from Lake Superior – he earned a B.S. degree from the University of Wisconsin – Eau Claire in 1976, and an M.S. from the University of Minnesota – Duluth 1980. His Master’s thesis was on the petrology and tectonic significance of the interflow sediments in the Keweenawan North Shore Volcanic Group of Northeastern Minnesota. He worked a short time in northwestern Minnesota for Exxon Minerals, and then joined the Minnesota Geological Survey. As if the ‘normal’ work load of annual stints of field mapping, compilation, and publication of geologic maps and reports isn’t enough, Mark is also the Technical Editor for every map and report that is published at the MGS. The wide range of topics for the things he reviews – including Precambrian, Phanerozoic, and Quaternary geology, as well as hydrostratigraphy, testifies to the breadth of Mark’s geologic knowledge. Mark has been a mainstay of the ILSG for many years – his first paper was in 1978 in Milwaukee, on the topic of his Master’s thesis. He has probably attended every meeting since then. As most of you know, Mark has been one of the most active and involved members of the Institute on Lake Superior Geology for many years. He has co-chaired meetings in Minneapolis, International Falls, and Hibbing. He has led or co-lead 13 field trips starting with Eveleth in 1993 through the most recent 2014 meeting in Hibbing, and after this meeting we can bump that total up to 15. He has submitted over 30 abstracts for oral talks and posters, and has been session chair multiple times. He was the Secretary-Treasurer from 1994-2002, returned as Treasurer in 2005 and still is today. He has also led field trips for GSA meetings. If you have ever been on an ILSG trip that Mark is taking part in, you also know he has his nose right on the outcrop talking to the leaders and other participants about the rocks, not only for that stop, but for how they tie into the rest of the world. If you are a new young member of ILSG you should stick close to Mark and ask any question you want because he will fully engage you, and to him there is no such thing as a ‘stupid question’. However, the Goldich award is for more than just involvement with the ILSG. It is also about one’s contribution to the geology of the Lake Superior region. In the latter, Mark has contributed a great amount. His work has focused mainly on deciphering the complex xi �geology of the Archean rocks in both northern Minnesota and in the Minnesota River Valley, but he has also contributed a great deal to understanding the Paleoproterozoic terranes of Minnesota, including the east-central Minnesota batholith and environs, the Sioux Quartzite, the Biwabik and Gunflint Iron Formations, and the Sudbury ejecta deposits in Minnesota. In all of these cases his work has benefited not only Minnesota, but has applications elsewhere in the Lake Superior region. He has authored or co-authored more than 60 maps and reports published by the MGS, has authored or co-authored numerous publications in refereed journals, and selflessly agrees to give presentations to the public on a wide variety of topics pertaining to his work Mark’s latest focus is on unraveling the Timiskaming-type assemblages in northern Minnesota, which has given him cause to lead nine Precambrian Research Center capstone projects aimed at tracing and deciphering these assemblages. These capstone projects have given dozens of aspiring geologists the opportunity to map with a great mentor. He always found a way to sandwich these capstone projects in between all the other contractual mapping obligations of the MGS. Mark first started at the Minnesota Geological Survey in 1979, as a Junior Geologist, and one of his first projects was making a geologic map of Paleozoic strata of the Twin Cities basin. Fortunately for us hard rock types, he quickly moved on to what he loves, Precambrian rocks. Mark has an uncanny knack for field mapping (especially picking out graded beds!) and accompanying drill core logging. His ability to unravel the structural attributes of everything from a single outcrop to an entire greenstone belt never ceases to amaze me. Equally as amazing and inspirational to me is his tireless work ethic, be it long days in the field or dark winter days in the office. I’ve never known him to knock off a day of field work because of any type of weather conditions – more than once I’ve been back indoors, warm and dry, for several hours due to atrocious field conditions, but I won’t see Mark until after dark when he comes in stomping mud off his boots telling me about some great thing he discovered that day. Subsequent to my being hired full-time at MGS in 1987, I have worked almost continuously with Mark on a wide variety of mapping and drilling projects throughout all of the different Precambrian terranes of Minnesota. Early on I had the job of being his field assistant, which was great fun since my main task was helping peel outcrops – and boy did we peel! This led to other larger mapping projects where we divided up map areas, drilling projects that went through entire winters, and independent mapping projects. During every one of these, and continuing to this day, Mark has always set the bar when it comes to initiating projects, field work, and interpreting the rocks – from the outcrop through map compilation and publication. More than once Mark would have some idea for a grand mapping project for which my initial reaction was “Really? You think we can do that?”, then a couple of years later there it was - a finished project. The year 1983 was 33 years ago. Since I am currently 56 years old, that means I have known Mark for well over half my life. We have spent months living out of the same motel room, driving to and from field work, to ILSG meetings, etc. I don’t know if the feeling is mutual, but I had a great time through all of it. And my conclusion after having lived half my life with the guy is that he is most deserving of this medal. Terry Boerboom Precambrian Geologist Minnesota Geological Survey xii �In Memoriam Leon Wayne Gladen, 82, of Hibbing, died April 23rd, 2016. He was born in Bemidji, MN on April 25th, 1933, to Leonard and Hattie (Holmberg) Gladen. Leon grew up on the family farm in Bemidji and he served in the U.S. Army during the Korean War. He was a graduate of UMD and he held a Master’s Degree in Geology. He worked as a geologist for the Minnesota Department of Natural Resources for ten years and later for Lehman & Associates out of Minneapolis. Leon and Bernice (Grzybowski) were united in marriage on July 1, 2001 in Kenai, Alaska. Leon and Bernice enjoyed traveling the world together for his work assignments. He enjoyed reading, hunting, collecting fly fishing rods, and gardening. xiii �Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xiv �Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive one-half of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. In Fall 2015, the ILSG Board of Governors awarded three awards from the Joe Mancuso Student Research Fund. The winners were: Amanda Van Lankvelt University of Massachusetts-Amherst, Department of Geosciences Current degree program: PhD candidate (Advisor: M.L. Williams) Determining the Deformation Age of the Baraboo Syncline Award: $500 Detaya Johnson University of Wisconsin-Milwaukee, Department of Geosciences Current degree program: Bachelors of Science (Advisor: Dyanna Czeck) Geochemical analysis of deformed metaconglomerates Award: $500 Laura Cuccio Utah State University, Department of Geology Current degree program: MS Candidate (Advisor: James Evans) Evaluating the nature of sedimentary rock-crystalline basement interface and its control on hydrologic processes. Award: $500 xv �Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and longtime friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. 2016 Student Paper Awards Committee Karl Everett – KEA Associates Dyanna Czeck - University of Wisconsin-Milwaukee Tim Kroeger – Bemidji State University (MN) Michael Zieg – Slippery Rock University (PA) Dorothy Campbell – Ontario Geological Survey xvi �Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Christian Schardt (2016-2019) – University of Minnesota Duluth Rob Cundari (2015-2018) – Ontario Geological Survey Jim Miller (2014-2017) – University of Minnesota Duluth Allan Blaske (2013-2016) – AECOM Pete Hollings - Secretary (2013-2016) – Lakehead University Mark Jirsa – Treasurer (2011-2014) – Minnesota Geological Survey Local Committee Jim Miller – General Meeting Chair Department of Earth and Environmental Sciences University of Minnesota Duluth Christian Schardt – Technical Program Chair Department of Earth and Environmental Sciences University of Minnesota Duluth Dean Peterson – Field Trip Chair Peterson Geoscience LLC and Precambrian Research Center University of Minnesota Duluth Session Chairs Robert Lodge – University of Wisconsin – Eau Claire Gerry White – Ontario Geological Survey, Thunder Bay, ON Ben Drenth – U.S. Geological Survey, Denver, CO Jeff Lynott – Foth, Green Bay, WI Latisha Brengman – University of Minnesota Duluth Esther Kingsbury Stewart – Wisconsin Geological and Natural History Survey xvii �Field Trip Leaders Field trips have been the mainstay of the ILSG since its inception 62 years ago. We want to give a special thanks to the field trip leaders who volunteered their time and talent in carrying that tradition forward. 1) GLACIAL GEOLOGY OF THE LAURENTIAN UPLANDS Phil Larson – Vesterheim Geoscience PLC Howard Mooers – Dept. of Earth and Environmental Sciences, UMD 2) NEOARCHEAN GEOLOGY OF THE WESTERN VERMILION DISTRICT Mark Jirsa, Amy Radakovichm Terry Boerboom - Minnesota Geological Survey 3) Cu-Ni-PGE DEPOSITS OF THE DULUTH COMPLEX Mark Severson – Tech American Andrew Ware – PolyMet Mining Kevin Boerst – Twin Metals Minnesota Steve Monson-Geerts – Natural Resources Research Institute, UMD 5) GEOLOGY OF THE ENDION SILL ALONG THE DULUTH LAKEWALK Jim Miller - Dept. of Earth and Environmental Sciences, UMD 6) GEOLOGY AND TROUT FISHING ALONG AMITY CREEK, DULUTH Dean Peterson – Peterson Geoscience LLC George Hudak - Natural Resources Research Institute, UMD 7) ARCHEAN AND PROTEROZOIC GEOLOGY OF THE GUNFLINT TRAIL Mark Jirsa - Minnesota Geological Survey 8) KEWEENAWAN GEOLOGY OF THE HOVLAND AREA Terry Boerboom – Minnesota Geological Survey John Green - Dept. of Earth and Environmental Sciences, UMD 9) DULUTH HARBOR GEOLOGIC HISTORY BOAT CRUISE: QUATERNARY TO ANTHROPOCENE Irv Mossberger, Mehgan Blair, Eric Dott – Barr Engineering Andy Breckenridge – University of Wisconsin - Superior Todd Kremmin - Dept. of Earth and Environmental Sciences, UMD xviii �Sponsors The following organizations and individuals made general contributions to the 62nd Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. All of the funds contributed this year go toward travel awards for student registrants. INDIVIDUAL CONTRIBUTORS TO STUDENT TRAVEL SCHOLARSHIPS WILLIAM EVERETT JOHN BERKLEY HENRY DJERLEV STEVE HOAGLUND ALLAN MACTAVISH RYAN DAYTON MARY ARTHUR DAN COSTELLO HARVEY THORLIEFSON DANIEL ROMANELLI GORDON MEDARIS, JR. ERIC DOTT With an especially generous donation provided by RON SEAVOY xix �Report of the Chair of the 61st Annual Meeting Dryden, Ontario The 61st ILSG was held in Dryden, Ontario on May 19-24, 2015. The meeting was chaired by Robert Cundari (Ontario Geological Survey) and Peter Hinz (Ministry of Northern Development and Mines) with considerable assistance from the organizing committee (Mark Smyk, Al MacTavish and Pete Hollings). The meeting was attended by a total of 123 delegates including 31 students. Special thanks to individuals who provided financial support for the meeting (Mary Arthur, Steve Baumann, Leonard Espinosa, Gordon Medaris Jr., Allan MacTavish, Jim Miller and Paul Weiblen) as well as the Thunder Bay Branch of the Canadian Institute of Mining and Metallurgy (CIM) for its generous donation. The two-day technical session began on Thursday May 21st which focused on Midcontinent Riftrelated geology and special topics including an extended oral presentation summarizing research to date on the Sudbury Impact Event in the Lake Superior Region. Technical talks continued through Friday morning with talks focusing largely on Archean geology. A total of 21 talks were given, 8 of which were presented by students. A total of 24 posters were displayed, 9 of which were presented by student authors. The 2015 Goldich Medal was awarded to Rodney J. Ikola from Esko, Minnesota. Thomas Waggoner presented the award during the annual banquet citing Rodney’s many contributions to the geoscience and mining community of Minnesota. The evening banquet speaker was Steve Beneteau – Senior Diamond Advisor / Chief Gemmologist for the Province of Ontario and the Manager of the Diamond Sector Unit for the Ontario Ministry of Northern Development and Mines. The title of his talk was: “Ontario’s Diamonds: A Journey from Mine to Market”. The meeting offered three multi-day field trips, three one-day field trips and three half-day field trips covering the Archean Geology of northwestern Ontario. Three pre-meeting field trips were offered on Tuesday May 19th and Wednesday May 20th, including Red Lake Geology (2-day) led by Andreas Lichtblau and Carmen Storey (Ontario Geological Survey), a Western Wabigoon Subprovince Transect (Dryden to Meggisi Lake) led by Mark Puumala and Dorothy Campbell (Ontario Geological Survey) and the Geological Setting of the Thunder Lake Gold Deposit led by Treasury Metals Inc. personnel. Three half-day trips were offered on Friday May 22nd, including Classic outcrops of the Dryden Area led by Peter Hinz (Ministry of Northern Development and Mines), Gold Occurrences of Van Horne Township led by Steve Meade (Ontario Geological Survey) and the Unique mineralizing event at the Pidgeon Molybdenum Occurrence led by Craig Ravnaas (Ontario Geological Survey). Three post-meeting field trips were offered starting Friday afternoon running through Sunday May 24th, including the Historic Pickle Lake Camp (1.5-day) led by Mark Smyk (Ontario Geological Survey), Pete Hollings (Lakehead University) and Neil Pettigrew (Fladgate Exploration Consulting Corp.), the Ghost Lake Batholith and Related Pegmatites led by Shannon Zurevinski (Lakehead University) and the Mattabi/Sturgeon Lake Historic VMS Camp (2-day) led by George Hudak (University of Minnesota Duluth). xx �The Institute’s Board of Directors met on Thursday May 21st to discuss the business of the Institute. The meeting was attended by meeting co-chair Robert Cundari, Treasurer Mark Jirsa, Secretary Peter Hollings and board members Jim Miller (2014 chair), Theodore Bornhorst (2013 chair) and Al MacTavish (2012 chair). Secretary Hollings took the minutes of the Board meeting that are as follows: 1. Accepted report of the Chairs for the 60th ILSG, Hibbing, Minnesota; as printed in the Proceeding Volume (Miller), and minutes of last Board meeting, May 15, 2014 (Hollings) 2. Received, discussed, and accepted 2014-2015 ILSG Financial Summary (Jirsa). 3. Received, discussed, and accepted 2014-2015 report of the Secretary (Hollings). 4. Approved Rob Cundari as on-going ILSG Board member 5. Approved Duluth as the site for the 62nd annual ILSG meeting. The meeting will be hosted by Jim Miller and Christian Schardt. 6. Discussed and approved replacing Bernhardt Saini-Eidukat as the “member from academia” on Goldich Committee (end of term 2015) with Shannon Zurevinski 7. Discussed student research grants. It was agreed that the application deadline will be switched to April 1 with results announced by April 30 each year. The next competition will be held in 2016. Jirsa to provide a list of all recipients of these awards so their attendance at the Annual Meeting can be tracked. 8. Jirsa to investigate the possibility of obtaining Director’s insurance including the cost and what would be covered. Also to investigate the implication of turning the Institute into a LLC. 9. Jirsa to develop a budget template for Meeting Chairs to help with the reporting of expenses and revenue when submitting the financial summaries to the Treasure The chairs would like to thank all those who assisted with the meeting including the organizing committee, the field trip leaders, the session chairs, service providers and those who provided support behind the scenes. The chairs would also like to thank those who participated in the meeting including the field trip attendees and the oral and poster presenters for their enthusiastic involvement with the Institute. Respectfully submitted, Robert Cundari and Peter Hinz Co-chairs, 61st Institute on Lake Superior Geology xxi �xxii �TECHNICAL PROGRAM WEDNESDAY MAY 4, 2016 All trips leave from the Harbor Side entrance (G) of the Duluth Entertainment and Convention Center 8:00am - 5:30pm PRE-MEETING FIELD TRIPS 1) GLACIAL GEOLOGY OF THE LAURENTIAN UPLANDS Phil Larson – Vesterheim Geoscience PLC Howard Mooers – Dept. of Earth and Environmental Sciences, UMD 2) NEOARCHEAN GEOLOGY OF THE WESTERN VERMILION DISTRICT Mark Jirsa, Amy Radakovichm Terry Boerboom - Minnesota Geological Survey 3) Cu-Ni-PGE DEPOSITS OF THE DULUTH COMPLEX Mark Severson – Tech American Andrew Ware – PolyMet Mining Kevin Boerst – Twin Metals Minnesota Steve Monson-Geerts – Natural Resources Research Institute, UMD 1:00pm - 5:30pm HALF-DAY PRE-MEETING FIELD TRIP 5) GEOLOGY OF THE ENDION SILL ALONG THE DULUTH LAKEWALK Jim Miller - Dept. of Earth and Environmental Sciences, UMD 4:00 pm - 10:00 pm Registration (Horizon Foyer) 7:00 pm - 10:00 pm Welcoming Reception (Horizon Foyer) Poster Session (Horizon Foyer and Room 202) xxiii �THURSDAY MAY 5, 2016 Asterisk * denotes a student eligible for Best Student Paper Award 7:30 am - noon REGISTRATION 8:00 am OPENING REMARKS Jim Miller and Christian Schardt, Co-Chairs, 2016 ILSG TECHNICAL SESSION I Session Chairs: Robert Lodge – University of Wisconsin – Eau Claire Ann Wilson – Ontario Geological Survey 1A) PETROLOGY AND METALLOGENESIS OF ARCHEAN AND PALEOPROTEROZOIC IGNEOUS RX 8:10 Dave Peck, Lionnel Djon, Cameron McLean, Gary DeSchutter, Jill Maxwell, Kelsey Privett, Denis Decharte, Chris Roney, Michelle Huminicki, & Bob Stewart The Lac Des Iles PGE-Cu-Ni deposit, Canada: an organized mega-breccia unit? 8:25 M. L. Djon*, G.R. Olivo, J.D. Miller, D.C. Peck, and B. Joy PGE Mineralization in the Northern Ultramafic Center of the Lac des Iles Complex, Ontario: Evidence of Magmatic and Hydrothermal processes 8:40 Erik Haroldson*, Brian Beard, Aaron Satkoski, Clark Johnson, and Philip Brown U-Th-Pb isotopes of the Reef Deposit; a Au-Cu occurrence in central Wisconsin 8:55 Ashley Quigley*, Thomas Monecke, Eric Anderson, Nigel Kelly, and Patrick Quigley Setting of volcanogenic massive sulfide deposits of the Paleoproterozoic Penokean volcanic belt 9:10 Robert Lodge, Geoffrey Pignotta, Brigitte Gélinas, Kelly Schwierske, and George Hudak Volcanological, Geochemical, and Geochronological Comparisons of the Gafvert Lake Sequence in Minnesota and Shebandowan Assemblage in Ontario 9:25 9:40 G. Gamelin*, V. Stinson, Yuanming Pan, and M. Nadeau A comparative study of mafic and felsic lithologies from the Borden Belt and adjacent greenstone belts in the Wawa-Abitibi Terrane Brigitte Gélinas* and Peter Hollings The Geology and Geochemistry of the Laird Lake Property, Red Lake Greenstone Belt, Ontario 9:55 COFFEE BREAK AND POSTER SESSION xxiv �1B) PRECAMBRIAN GEOCHRONOLOGY / ARCHEAN SEDIMENTATION AND STRUCTURE 10:15 Michael Mudrey Continued Evaluation of the Dilatancy Model for Discordant Uranium-Lead Age Determination of Zircon 10:30 Ben Frieman*, Yvette Kuiper, Nigel Kelly, and Thomas Monecke Provenance and tectonic evolution recorded by successor basins in the Abitibi-Wawa terrane: Insights from new U-Pb LA-ICP-MS analyses of detrital zircon 10:45 Sophie Kurucz* and Philip Fralick Giant Domes of the Mosher Carbonate, Steep Rock, Ontario 11:00 Matthew Svensson* and Philip Fralick The Badwater gabbro as an analogue for the weathering of Martian basalts 11:15 Victoria Stinson*, Yuanming Pan, Gleceria Gamelin, and Matthew Nadeau A re-examination of the Kapukasing structural zone 11:30 Tracy Carson*, Brittany Deley, and Mary Louise Hill Microstructural comparison of the Hardrock Project at Geraldton, Ontario and the Coffee Gold Project, Yukon 11:45 LUNCH BREAK ILSG BOARD MEETING TECHNICAL SESSION II Session Chairs: Ben Drenth – U.S. Geological Survey, Denver, CO Jeff Lynott – Foth, Green Bay, WI 2A) MIDCONTINENT RIFT GEOLOGY AND MINERALIZATION 1:15 Klaus Schulz and Suzanne Nicholson The Geochemistry of the Siemens Creek Formation and the Nature of Early Midcontinent Rift Basaltic Magmatism in the Western Lake Superior Region 1:30 Mark Smyk, Peter Hollings, and Philip Fralick A Preliminary Investigation of Enigmatic Igneous Rocks on Big Powder Island, Northern Lake Superior: A Possible Mesoproterozoic Magmatic Event 1:45 Sean O’Brien*, Peter Hollings, and Jim Miller Petrology, geochemistry and sulphur isotopes of the Crystal Lake gabbro and Mount Mollie dyke, Northwestern Ontario 2:00 Robert Cundari, Peter Hollings, David Good, and Sarah Davis Geochemistry and petrogenesis of volcanic rocks in the Coldwell Alkaline Complex; new insights from the Wolfcamp Lake volcanic rocks 2:15 David Good, Robert Linnen, and Iain Samson The Cu/Pd diagram and metal/sulfur variation as an exploration tool: Examples from the Coldwell Alkaline Complex, Ontario xxv �2:30 Robert Mahin and Steven Beach The Eagle East Magmatic Nickel-Copper Discovery 2:45 Connor Mulcahy*, Jim Miller, Robert Mahin, Steven Beach, and Bob Nowack Emplacement and Crystallization History of Ni-Cu-(PGE) Sulfide-mineralized Peridotites in the Eagle Intrusion, Upper Michigan 3:00 Christian Schardt Metal isotopic signatures in the Duluth Complex associated with magmatic Cu-NiPGE mineralization 3:15 COFFEE BREAK AND POSTER SESSION 2B) ENVIRONMENTAL GEOLOGY RELATED TO MINING AND EXPLORATION 3:35 Alex Brown Unique characteristics of sediment-hosted stratiform copper mineralization resulting from exceptional latent volcanic heat at White Pine, northern Michigan 3:50 Robert Seal, Perry Jones, Nadine Piatak, and Laurel Woodruff Potential value of pre-mining baseline oxygen, hydrogen, and sulfur isotopic data from surface waters for proposed large mining projects in northern Minnesota 4:05 Andrew Manning, Richard Wanty, and Jean Morrison Preliminary groundwater age and chemistry data from cover overlying Duluth Complex Ni-Cu-PGE deposits, NE Minnesota 4:20 Nadine Piatak, Robert Seal, Perry Jones, and Laurel Woodruff Copper toxicity and dissolved organic matter: Resiliency of mineralized watersheds in northern Minnesota and Michigan 4:35 Andrea Reed, Barry Frey, and Kevin Hanson Expanding the historical exploration document collection at the Minnesota Department of Natural Resources: the Polaris Joint Venture exploration program 4:50 George Hudak, Monson Geerts Stephen, Larry Zanko, Sara Post, and Euan Reavie The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter ‐ 2015 Update 6:00 RECEPTION/POSTER SESSION – CASH BAR (Harbor Side Foyer) 7:00 ANNUAL BANQUET (Harbor Side Ballroom)  Announcement of 63rd Annual Meeting Location  2016 Goldich Award Presentation to Mark Jirsa  Banquet Presentation - Peter Clevenstine, Asst. Director of Minerals, MN DNR “Managing Minnesota's Mineral Resources and the DNR’s Conservation Agenda” xxvi �FRIDAY MAY 6, 2016 Asterisk * denotes a student eligible for Best Student Paper Award 8:00 OPENING REMARKS, UPDATES Jim Miller and Christian Schardt, Co-Chairs, 2016 ILSG TECHNICAL SESSION III Session Chairs: Latisha Brengman –University of Minnesota Duluth Esther Kingsbury Stewart– Wisconsin Geological and Natural History Survey 3A) PROTEROZOIC TECTONICS AND SEDIMENTATION 8:10 Paul Bedrosian Making it and breaking it in the upper Midwest: Constraints on continental assembly and rifting from EarthScope2 8:25. Esther Kingsbury Stewart and Jeffrey Mauk Sequence stratigraphy and basin evolution of the Mesoproterozoic Nonesuch Formation, Ashland syncline, northern Wisconsin 8:40 V.J.S. Grauch, Michael Powers, and Eric Anderson Progress on 3D modeling of the Midcontinent Rift System in the western Lake Superior region and an isopach map of the Oronto Group 8:55 Robyn Jones* and Philip Fralick Sedimentology of a pre-vegetation prograding deltaic assemblage: the Mesoproterozoic Kama Hill and Outan Island Formations, Ontario 9:10 Philip Fralick and Kamil Zaniewski Sedimentology of a Pre-Vegetation Floodplain Assemblage: the Mesoproterozoic Hele Member of the Sibley Group, Ontario 9:25 Julie Bartley, John Berger, Tanner Eischen, Sydney Firmin, and Lindsey Reiners Hypersaline conditions for stromatolite growth in the Rossport Formation (Mesoproterozoic, Ontario) 9:40 Richard Ojakangas What Happened in Northern Minnesota Between 2700 Ma and 1900 Ma? The Answer Is in the Pokegama Formation: A Multicycle Sedimentary History! 9:55 COFFEE BREAK AND POSTER SESSION 3B) GUNFLINT IRON FORMATION AND BARABOO QUARTZITE 10:25 William Cannon, Laurel Woodruff, and Stacy Saari Traces of the Sudbury meteor impact in the western Gogebic Iron Range, northern Wisconsin 10:40 Ruby Reid-Sharp* and Mary Louise Hill Characterizing deformation of Gunflint Formation in contact with Archean basement rocks east of Thunder Bay, Ontario xxvii �10:55 Carli Nap* and Philip Fralick Mesoproterozoic Alteration of the Paleoproterozoic Gunflint Formation: Analogies with Martian Blueberries 11:10 Esther Kingsbury Stewart, Eric Stewart, and Matthew Lamb Discovering hidden folds and faults in the Precambrian: new insights into Baraboointerval stratigraphy and deformation in southern Wisconsin 11:25 Gordon Medaris, Jr. Quantifying Mass Fluxes of Potassium in Weathering and Metasomatism of Paleosols 11:40 LUNCH BREAK TECHNICAL SESSION IV 10 YEAR ANNIVERSARY OF THE PRECAMBRIAN RESEARCH CENTER AT UMD 1:15 Jim Miller, Dean Peterson, and George Hudak Ten Years of Educating the Next Generation of Precambrian Field Geologists 1:30 Dean Peterson, Jim Miller, and George Hudak The PRC's Precambrian Field Camp ‐ A Decade of Training Students Geologic Mapping of the Canadian Shield 1:45 George Hudak and Dean Peterson Future Directions for the Precambrian Research Center 2:00 Mark Jirsa Nine years of capstones: A summary of PRC field camp capstone projects in the Neoarchean Knife Lake Group and associated rocks, central BWCAW, Minnesota 2:15 UPCOMING FIELD OPPORTUNITIES BEST STUDENT PAPER AWARDS STUDENT TRAVEL AWARDS 3:00 END OF TECHNICAL SESSION 4:00 pm POST-MEETING FIELD TRIPS 6) GEOLOGY AND TROUT FISHING ALONG AMITY CREEK, DULUTH Dean Peterson – Peterson Geoscience LLC George Hudak - Natural Resources Research Institute, UMD 7) ARCHEAN AND PROTEROZOIC GEOLOGY OF THE GUNFLINT TRAIL Mark Jirsa - Minnesota Geological Survey 8) KEWEENAWAN GEOLOGY OF THE HOVLAND AREA Terry Boerboom – Minnesota Geological Survey John Green - Dept. of Earth and Environmental Sciences, UMD xxviii �SATURDAY MAY 7, 2016 8:00am – 5:00pm POST-MEETING FIELD TRIP 9) DULUTH HARBOR GEOLOGIC HISTORY BOAT CRUISE: QUATERNARY TO ANTHROPOCENE Irv Mossberger, Mehgan Blair, Eric Dott – Barr Engineering Andy Breckenridge – University of Wisconsin - Superior Todd Kremmin - Dept. of Earth and Environmental Sciences, UMD xxix �POSTER PRESENTATIONS Asterisk * denotes a student eligible for Best Student Paper Award Eric Anderson, V.J.S Grauch, and Michael Powers Reprocessed seismic data image geology and structure near the Douglas fault on the Bayfield Peninsula, Wisconsin Kira Arnold* and S.E. Zurevinski An Investigation of the Ney’s Lookout Lamprophyric Dyke, Marathon, ON Kristofer Asp* and Christian Schardt An Investigation of Ni and Cu Isotopic Fractionation in Basal Duluth Complex Cu-Ni-PGE Mineralization, Northeastern Minnesota Steven Baumann, Alexandra Cory, and Sandra Dylka Lithological sedimentary divisions of the Copper Harbor Formation in Gogebic and Ontonagon Counties, Michigan Thomas Buchholz, Falster, Alexander, and Wm. B. Simmons The occurrence of Li, B, Sn, and W in the Nine Mile Pluton, Wausau Syenite Complex, Marathon County, Wisconsin William Cannon Mobilization of silica by flash heating of silica gel beneath the Sudbury Impact Layer, Baraga Basin, Michigan Val Chandler and Amy Radakovich Utility of the horizontal-to-vertical spectral ratio (HVSR) passive seismic method for determining Quaternary sediment thickness and bedrock elevation in north-central Minnesota: Fun with little control and generally poor data Jonathan Clark, Kristen Eshler, Patrick Groff, Taylor McClendon, Alexander Rode, Emily Salings, Kristen Spinelli, Kyle Vander Wyst, Aiden Walsh, Kristofer Asp, and Phillip Larson Bedrock Geology of the Devilfish Lake Area, Cook County, Minnesota Sarah Davis*, Peter Hollings, and Rob Cundari Mineralogy and Geochemistry of the Wolfcamp Lake Basalts Brittany Deley* and Mary Louise Hill Origin of the gold-hosting porphyry at Geraldton, Ontario Benjamin Drenth and Chad Ailes Re-digitized public aeromagnetic data for parts of the west-central Upper Peninsula, Michigan Benjamin Drenth, Raymond Anderson, Klaus Schulz, Joshua Feinberg, Val Chandler, and William Cannon Progress on Geophysical Mapping of the Northeast Iowa Intrusive Complex Simon Dolega* and Philip Fralick Geochemistry of deep and shallow water Archean banded iron formations, and their postdepositional implications in the western Superior Province, Canada xxx �Espree Essig*, Howard Mooers, and Karen Gran 3D Geological Mapping Using Terrestrial LiDAR at Soudan Underground Mine F. Glass Evidence in the Eastern Canadian Shield of regular fault patterns of crustal origin for the loci of some mineral deposits and late-stage intrusive events. Michael Felzan * and Marcia Bjørnerud Multi-stage development of breccias in the Baraboo Quartzite, Rock Springs, Wisconsin Carol Finn, Michael Zientek, Paul Bedrosian, Benjamin Bloss, Bethany Burton, Dean Petersen, and Heather Parks Geophysical Imaging of Layered Mafic Complexes and Relation to Platinum Group Element Exploration Carlin Green*, Robert Seal, William Cannon, and Nadine Piatak Quantitative abundance and preliminary morphological characterization of amphiboles in the Ironwood Iron-Formation, Gogebic Iron Range, Wisconsin Gregory Guenther* and Esther Kingsbury Stewart Paleocurrent interpretation of the Cambrian Elk Mound group using geophysical optical borehole image (OBI) logs from two new boreholes, Dodge county, Southern Wisconsin Samuel Helmuth* and Robert Lodge A New Rusk County: Producing an new Precambrian geological map from new field observations and compilations of historic geological/geophysical datasets Benjamin Hinks*, Joyashish Thakurta, Robert Mahin, and Steve Beach Geochemical and petrological studies on the origin of Ni-Cu sulfide mineralization at the Eagle and Eagle East intrusions in Marquette County, Michigan Sheree Hinz* and Peter Hollings Preliminary Observations of the Ultramafic Metavolcanic Rocks in the Eastern Portion of the Shebandowan Greenstone Belt, northwestern Ontario Nathaniel Jackson*, Bruno De Moura Merss, and Robert Lodge Lithostratigraphy and Ore Petrology of the Eisenbrey Zn-Cu-Pb Deposit, Rusk County, Wisconsin Shiyun Jin and Huifang Xu Incommensurately modulated structure of plagioclase as an indicator of cooling history of igneous rock Detaya Johnson* and Dyanna Czeck Geochemistry of Seine River metaconglomerates from Mine Centre, Ontario: interpreting fluid flow and volume changes during deformation with implications for strain analysis Alexandra Kozlowski* and S.E. Zurevinski Mineralogy and petrology of the diamondiferous Madonna Dyke, Marathon, ON Timothy Kroeger Preliminary Report on the Palynology of the Gervais Formation (Pleistocene), Red Lake County, Minnesota xxxi �Matthew Lamb* and Esther Kingsbury Stewart A comparison of Baraboo-Interval (Late Paleoproterozoic) Iron-Formation, Southern Wisconsin Crystal Lambert* and John Swenson Millennial-scale shoreline bluff retreat rates in the western arm of Lake Superior Matthew Matko* and Christian Schardt Small scale microanalysis of rock and mineral textures and its relationship to mineral separation S. Metteer* and S.E. Zurevinski Mineralogy and Petrology of the Rabbit Foot Dyke, White River, ON Jim Miller, Aaron Balles, Ellie Brown, Ryan Helms, Greta Penzel, and Luke Smith Geology of the Cherokee Lake area of the Boundary Waters Canoe Area, Cook County, MN - 2015 Precambrian field camp capstone mapping D. Nikkila*, R.H..Mitchell, and S.E. Zurevinski Investigations of the Layered Series Nepheline Syenite within Center II of the Coldwell Complex, Marathon, ON Maile Olson* and Robert Lodge Ore Petrography and Precious Metals of the Primary Flambeau Massive Sulfide Ore Mark Puumala and Seamus Magnus A preliminary evaluation of the structural controls on gold mineralization in the Jackfish Lake area, northwestern Ontario Lindsey Reiners*, Tanner Eischen*, and Julie Bartley The building blocks of stromatolites: Comparisons across time and environment Tyler Sager* and Nigel Wattrus Evaluating H/V analysis of passive seismic data as a means to map sediment thickness in the Duluth-Superior harbor Andrew Sasso* and Joyashish Thakurta Geochemical and Petrological Comparisons of Peridotite Units in Marquette County, Michigan Ruth Schulte, Nadine Piatak, Robert Seal, and Laurel Woodruff Acid-Generating and Acid-Neutralizing Potential of Silicate Rocks from the Basal Mineralized Zone of the Duluth Complex, Minnesota V. Smith* and S. Zurevinski The Mineralogy, Petrography and Geochemistry of the Anderson Lake Pegmatite Occurrence Matthew Svensson* and Stephen Kissin Source of Native Iron in Canadian Arctic Artifacts Margaret Upton, Ryan Puzel, Jaron Christenson, Morgan Kent, Steven Spreitzer, and Mark Jirsa Geologic mapping of Neoarchean and Proterozoic rocks near Kekekabic Lake, northeastern Minnesota, by students of the Precambrian Research Center’s 2015 field camp xxxii �Gerrit VanderWaal* and Christian Schardt Influence of mineral liberation on metal leaching and dissolution rates in ore material and associated host rock Blake Wallrich* and Michael Zieg Small-Scale Petrographic Variations in a Nipigon Diabase Sill Zacharie Zens* and Robert Lodge Geochemistry and Petrography of the Volcanic Strata Hosting the Flambeau Cu-Zn-Au Deposit in Rusk County, WI: A Re-examination of Wisconsin’s Only Past-Producing Volcanogenic Massive Sulfide Mine. Michael Zieg and Blake Wallrich Evidence for Episodic Emplacement History of a Nipigon Diabase Sill Michael Zientek, Klaus Schulz, Laurel Woodruff, William Cannon, Suzanne Nicholson, Lukas Zürcher, Heather Parks, and Connie Dicken Assessment of Undiscovered Nickel-Copper-Platinum Group Element (Ni-Cu-PGE) Resources Related to Conduit-Type Mineralization in the Midcontinent Rift System, Michigan, Minnesota, Ontario, and Wisconsin xxxiii �ABSTRACTS 1 �Reprocessed seismic data image geology and structure near the Douglas fault on the Bayfield Peninsula, Wisconsin ANDERSON, Eric D., GRAUCH, V.J.S., POWERS, Michael H., US Geological Survey, MS 964, PO Box 25046, Denver, CO 80225 USA A prominent gravity low lies over the Bayfield Peninsula in northern Wisconsin. The mapped bedrock geology includes sedimentary rocks of the Oronto and Bayfield Groups that overlie Midcontinent rift-related volcanic and intrusive rocks. The nearly 100 mGal amplitude anomaly has been interpreted to reflect low density Archean granite that is surrounded by higher density basalt (White, 1966; Allen and others, 1997), informally called White’s Ridge (Figure 1). Reprocessed seismic reflection data are helping to understand the geology and structures that are encompassed by the gravity low. Seismic reflection data acquired in 1984 were obtained for portions of several lines on the Bayfield Peninsula (Figure 1). Initially, only stacked time-series sections were available. The raw shot records were obtained for two lines within the western gradient of the gravity low and reprocessed using modern seismic data processing techniques. New stacking velocities were determined and used to create a new migrated time section and convert it to depth. The results show significantly more detail than was evident in the original time sections. These reprocessed data provide new insights about the geology to depths as great as 12 km. The reprocessed seismic lines run north-south and east-west for a total of about 65 line km. The north-south line is perpendicular to the mapped Douglas fault and confirms that it is south-dipping, thrusting Keweenawan volcanic rocks over as much as 3.7 km of younger, southdipping Bayfield and Oronto Group rocks to the north. At approximately 1 km depth, an anticlinal structure is evident south of the mapped Douglas fault. The velocity model indicates that the anticlinal structure is within the sedimentary rock package that overlies the volcanic rocks. North of the Douglas fault, both seismic sections indicate the Oronto Group increases in thickness towards the fault. The reprocessing of the east-west line has revealed several significant reflections in what was previously considered as reflection-poor Archean gneiss. These events are at depths from 9 to 12 km and may indicate the presence of layered strata of unknown origin at the easternmost end of the seismic line. The reprocessed data show in both lines a wavy texture in parts of the volcanic rocks. This texture is not evident in the vintage time sections. In the east-west line, the wavy texture is more pronounced where the volcanic rocks are in unconformable contact with the overlying Oronto Group. In the north-south line, the wavy texture only appears at depth in the volcanic rocks north of the Douglas fault. This texture may be due to heterogeneity in the physical properties and/or layering of the volcanic rocks or be the expression of local intrusive activity. The new seismic observations are being integrated into broader scale gravity and magnetic models that span the gravity low over the Bayfield Peninsula. Preliminary models indicate that relatively low density pre-rift rocks can explain the gravity low. The seismic data add considerable detail to the geologic model and are helping to define stratigraphic and structural relationships between the lithologies at depth. 2 �Figure 1: Generalized geology map of western Lake Superior showing location of reprocessed seismic data within the western gradient of the gravity low that overlies the Bayfield Peninsula. Anomalous gravity lows occur at both White’s and Grand Marais Ridges. The seismic lines provide detailed subsurface images that are being integrated with 2D gravity and magnetic models to help characterize the source of the gravity lows. REFERENCES Allen, D.J., Hinze, W.J., Dickas, A.B., and Mudrey, M.G., 1997. Integrated geophysical modeling of the North American Midcontinent rift system: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota. In Ojakangas, R.W., Dickas, A.B., and Green, J.C. (Eds.), Middle Proterozoic to Cambrian Rifting, central North America: Geological Society of America Special Paper 312: 47-72. White, W.S., 1966. Tectonics of the Keweenawan basin, western Lake Superior region. U.S. Geological Survey Professional Paper 524-E: E1-E23. 3 �An Investigation of the Ney’s Lookout Lamprophyric Dyke, Marathon, ON ARNOLD, Kira1, and ZUREVINSKI, S.E.1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, kaarnold@lakeheadu.ca The Coldwell Complex of the Superior Province is known to host multiple lamprophyric dykes varying from ultramafic to alkaline affinities (Heaman and Machado 1992). The 1108 +/1 Ma age of the Coldwell complex and close spatial proximity supports a strong relationship to the magmatism of the Keweenawan Midcontinent Rift. The Ney’s Lookout dyke is located within the Center II portion of the Coldwell Complex, crosscutting assimilated syenites. This area of Center II is highly brecciated and assimilated, which is interpreted to be representative as evidence of the cauldron collapse (Mitchell and Platt 1978). The lamprophyre is devoid of diamonds, and has not been thoroughly analyzed other than brief field assessments. The objective of this research is to complete a mineralogical assessment of the lamprophyre dyke in order to characterize and classify the lamprophyre according to the IUGS classifications of lamprophyre-clan rocks. The main minerals that comprise the lamprophyre are euhedral zoned pyroxene phenocrysts and flow-aligned anorthoclase laths in a fine-grained groundmass composed of biotite, pyroxene, amphibole and feldspars. The pyroxene phenocrysts are poikilitic with inclusions of biotite, amphibole and relic olivine similar to the groundmass. Relic chloritized olivine phenocrysts are present in the groundmass and as inclusions in pyroxenes. The pyroxenes present in the sample are classified as diopside with minor augite. Mineralogical compositions of the Ney’s Lookout amphibole classify as Ferroan Paragasitic Hornblende. When compared to other lamprophyres within the Coldwell Complex, there were mineralogical and textural similarities between Ney’s Lookout Lamprophyre and local sannaites, such as porphyritic texture and mantling of amphiboles on pyroxene (Mitchell et al. 1991). Aside from minor variances in modal mineralogy, the lamprophyre best resembles a sannaite under the alkaline lamprophyre classification. The chill margin along the dyke is a fine-grained rim with no contact metamorphism present, interpreted as representing a moderate emplacement temperature. Laths of anorthoclase are directionally aligned in the lamprophyre and are interpreted as flow textures. Zonation in the pyroxene is distinct in the phenocrysts, representing increasing Cr and Ti outward toward the rim with increased Fe in the core and decreasing amounts in the rim. Uralitization is the altertion of pyroxene phenocrysts to amphibole around the outter rim. The zonation and uralitization often occur from magma mixing events. The Ney’s Lookout Sannaite Lamprophyre, has intruded into an area of highly brecciated and assimilated country rock. As well, the Little Pic River fault, a North-South trending fault, is located to just to the West of the lamprophyre. It is likely that the Ney’s Lookout lamprophyre was emplaced along planes of weakness similar to other lamprophyres in the area. 4 �Figure 1. A. Pyroxene phenocrysts present in the lamprophyre dyke are poikilitic and euhedral with simple twinning. B. The distinct zonation of pyroxene phenocrysts with a Fe rich core. Cr and Ti wt. %’s increase towards the rim. Al wt. % varies between zones, with no distinctive trend. References: Heaman, L. M., and Machado, N. 1992. Timing and origin of midcontinent rift alkaline magmatism, North America: evidence from the Coldwell Complex. Contributions to Mineralogy and Petrology, v. 110, 289-303. Mitchell, R. H., and Platt, R. G., 1978. Mafic mineralogy of ferroaugite syenite from the Coldwell alkaline complex, Ontario, Canada. Journal of Petrology, v. 19, 627-651. Mitchell, R. H., Platt, R. G., Downey, M., and Laderoute, D. G. 1991. Petrology of alkaline Lamprophyres from the Coldwell alkaline complex, northwestern Ontario. Canadian Journal of Earth Sciences, v. 28, 1653-1663. 5 �An Investigation of Ni and Cu Isotopic Fractionation in Basal Duluth Complex Cu-Ni-PGE Mineralization, Northeastern Minnesota Asp, Kristofer1, Schardt, Christian 1 1 Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 1114 Kirby Dr. Duluth, MN 55812 USA Cu-Ni-PGE magmatic sulfide-style mineralization occurs along the western margin of the Duluth Complex in northeastern Minnesota. Previous studies have demonstrated a notable fractionation of 60Ni and 58Ni in terrestrial materials, including both primary and secondary phases, with a total range of 2.1 ‰ [1 – 5]. Additional work has indicated a fractionation of 65Cu and 63Cu, with pronounced differences between primary copper sulfides and secondary copper phases in a variety of magmatic deposit types [6, 7]. Prior to this research, no δ60/58Ni or δ65/63Cu values have been measured in Duluth Complex rocks. The primary goal of this study is to measure Ni and Cu isotope values in a variety of Duluth Complex samples, and develop a possible model for the δ60/58Ni and δ65/63Cu isotopic systems in this geologic terrane. A secondary goal is to determine whether or not Ni and Cu isotope values in surface material could be used as an exploration tool for identifying Cu-Ni-PGE mineralization at depth. Based on the findings of previous studies, samples were collected to determine the isotopic differences between sulfide-bearing and sulfide-barren material. The potential effects of weathering were also taken into account by collecting samples at the surface in addition to unweathered, primary mineralized drill core. Samples were collected from a variety of locations in the basal Duluth Complex, including glacial till beds and surface outcrops in the vicinity of the Spruce Road, Maturi, Serpentine, Mesaba, and NorthMet deposits. Drill core from the Birch Lake, Wyman Creek, and Wetlegs deposits, along with those listed previously, was provided by several companies, including Duluth Metals/Twin Metals, Teck, PolyMet, and Encampment Minerals. Several pieces of drill core were also obtained from the MN DNR core facility in Hibbing, MN. A detailed characterization of till, weathered surface, and primary drill core samples revealed three main sources of nickel in Duluth Complex material: silicate, sulfide, and secondary oxide. The 24 δ60/58Ni values have an overall range from -0.97 to 0.22 ‰, but are correspondingly distinct in each type of material: silicate (-0.03 ‰ average), sulfide (-0.36 ‰ average), secondary oxide (-0.50 ‰ average). Further geochemical and microprobe work, along with the isotopic values, indicate two main stages of Ni fractionation in basal Duluth Complex rocks: a high temperature stage during crystallization, and a low temperature stage during surficial weathering. High-T fractionation is defined by a preferential incorporation of 58Ni into sulfide, while silicates, especially olivine, are reflective of the Bulk Silicate Earth value [3]. Low-T fractionation results in a preferential incorporation of 58Ni into secondary oxide, while 60 Ni possibly enters solution and leaves the system [1; 5]. The 22 measured Duluth Complex δ65/63Cu values have an overall range from -1.28 ‰ to 0.36 ‰, with an overall average of -0.35 ‰. This range is roughly similar to the magmatic sulfide values measured by [6], but are a significant departure from values measured in other deposit types. Specifically, deposits containing abundant secondary copper phases, including Cucarbonates and hydroxides, are significantly more enriched in 65Cu, with values up to 2.41 ‰ reported [7]. Due to the lack of observed secondary Cu minerals in project samples, it is possible that the low-T fractionation of 65Cu into secondary phases may not occur as readily in the basal Duluth Complex. Accordingly, the only indicated fractionation in the Duluth Complex is a high 6 �temperature fractionation, where 63Cu is preferentially incorporated into Cu sulfides during crystallization. Based on the measured isotopic values, there are distinct differences between sulfidebearing and sulfide-barren material at the surface that may indicate the presence of mineralization at depth. There is potential for these isotopic systems to be used as an exploration tool, but further research is required to fully evaluate the processes associated with these systems. Figure 1: Developed model for Ni isotopic fractionation in the basal Duluth Complex. Primary fractionation occurs at high temperatures during the crystallization of olivine and sulfide, while secondary fractionation leads to a preferential incorporation of 58Ni into secondary oxides. Also during secondary fractionation, 60Ni enters solution, and may be incorporated into other secondary phases including ferromanganese crusts and secondary Ni silicates (garnierite). References 1. Cameron, V. and Vance, D. (2014). Heavy nickel isotopic compositions in rivers and the ocean. Geochemica et Cosmochimica Acta 128: 195-211 2. Gall, L., Williams, H.M., Siebert, C., Halliday, A.N., Herrington, R.J., Hein, J.R. (2013). Nickel isotope compositions of ferromanganese crusts and the constancy of deep ocean inputs and continental weathering effects over the Cenozoic. Earth and Planetary Science Letters 375: 148-155 3. Gueguen B., Rouxel O., Ponzevera E., Bekker A., Fouquet Y. (2013) Ni isotope variations in terrestrial silicate rocks and geological reference materials measured by MC-ICP-MS. Geostandards and Geoanalytical Research 3: 297-317 4. Hiebert RS., Rouxel, O., Houlé, MG., Bekker, A. (2014) Ni isotope fractionation between komatiite and sulfide mineralization at the Neoarchean Hart deposit, Abitibi greenstone belt, Canada. Geological Society of America Abstracts 46: 467 5. Wasylenski, L.E, Howe, Haleigh D., Spivak-Birndorf, L.J., Bish, DL. (2015) Ni isotope fractionation during sorption to ferrihydrite: implications for Ni in banded iron formations. Chemical Geology 400: 56-64 6. Larson, P.B., Maher, K., Ramos, F.C., Chang, Z., Gaspar, M., Meinert, L.D. (2003). Copper isotope ratios in magmatic and hydrothermal ore-forming environments. Chemical Geology 201: 337-350 7. Markl, G., Lahaye, Y., Schwinn, G. (2006) Copper isotopes as monitors of redox processes in hydrothermal mineralization. Geochemica et Cosmochimica Acta 70: 4215-4228 7 �Hypersaline conditions for stromatolite growth in the Rossport Formation (Mesoproterozoic, Ontario) BARTLEY, Julie K., 1BERGER, John, EISCHEN, Tanner, 2FIRMIN, Sydney, and REINERS, Lindsey Department of Geology, Gustavus Adolphus College, St. Peter, Minnesota 56082 1 Present Address: Bay West LLC, St. Paul, Minnesota 55103 2 Present Address: Department of Geography and the Environment, University of Denver, Denver, CO 80208 The Mesoproterozoic Rossport Formation of Ontario, Canada is approximately 1.4 billion years old (Franklin et al., 1980) and is generally interpreted to have been deposited in an intracratonic basin, most likely a rift-related lake Figure 1 (Rogala et al., 2005). While the Rossport consists dominantly of sandstone and shale, the Middlebrun Bay Member, in in the middle of the formation, is a carbonate unit. The Middlebrun Bay Member, in exposures on the Channel Islands and along the North Shore of Lake Superior, consists of massive limestone and cherty, stromatolitic carbonate. Key features of these carbonates suggest an interval of low lake level, low clastic influx, high salinity, and local conditions suitable for microbialite development. They provide an important Figure 2 environmental constraint on conditions in this lake, and additionally suggest environments in which complex, carbonate-precipitating microbial communities could thrive. The presence of a massive, recrystallized limestone on Copper Island suggests of hypersaline, evaporite-producing conditions. This unit is devoid of stromatolites or microbial laminae and has an unusual bright white color lacking internal structure (Fig. 1) with a coarsely recrystallized texture. These features suggest dissolution and replacement of a primary, soluble phase such as an evaporite mineral. The presence of large sandstone clasts let down from the overlying bed is reminiscent of collapse breccia associated with dissolution of that primary mineralogy (Fig. 2). Additionally, geochemical data suggest broad similarity with other Proterozoic carbonates interpreted as calcitized evaporites (Manning-Berg and Kah, 2013) Based on these data, we interpret the massive carbonate exposed on Copper Island as a calcitized evaporite, probably deposited originally as gypsum and replaced by calcite during diagenesis (Firmin and Bartley, 2014). Previous work on the Rossport Formation suggests that the Middlebrun Bay Member formed when lake levels and clastic influx were low (Rogala et al., 2007). Combined with an evaporite interpretation of the massive carbonate unit, we suggest the Middlebrun Bay interval may have been deposited during a period of increased aridity in the region. In this model, stromatolites would have formed in a hypersaline lake environment during intervals of low clastic influx. 8 �Figure 3 Middlebrun Bay stromatolites have low-relief stratiform to columnar morphology (Figure 3). Early diagenetic chert is locally abundant, and the meso- to micro-scale texture suggests in situ precipitation of microbial laminae and rapid cementation of stromatolite form. Stromatolites occur in lacustrine environments throughout geologic history, commonly under conditions of high alkalinity or high salinity (e.g., Gomez et al., 2013). During the Mesoproterozoic, though, stromatolites are more often associated with marine environmental conditions (Kah et al., 2009) and only occasionally occur in association with indicators of elevated salinity (e.g., Neudert and Russell, 1981). The Middlebrun Bay stromatolites, therefore, provide an important datum in space and time, allowing comparison of their morphology, texture, and fabric in relation to the features of younger, more commonly hypersaline stromatolites and contemporaneous, normal-marine forms. REFERENCES Firmin, S., and Bartley, J.K., 2014, An unusual Mesoproterozoic carbonate unit: Relic of a saline lake? Institute on Lake Superior Geology, v. 60, p. 45-46. Franklin, J. M., McIlwaine, W.H., Poulsen, K.H., and Wanless, R.K., 1980, Stratigraphy and depositional setting of the Sibley Group, Thunder Bay district, Ontario, Canada: Canadian Journal of Earth Sciences, v. 17, p. 633-651. Gomez, F.J., Kah, L.C., Bartley, J.K., and Astini, R.A., 2014, Microbialites in a high-altitude Andean lake: Multiple controls on carbonate precipitation and lamina accretion: PALAIOS, v. 29, p. 233-249. Kah, L.C., Bartley, J.K., and Stagner, A.F., 2009, Reinterpreting a Proterozoic enigma: ConophytonJacutophyton stromatolites of the Mesoproterozoic Atar Group, Mauritania: International Association of Sedimentologists Special Publication 41, p. 277-295. Manning-Berg, A.R., and Kah, L.C., 2013, Calcitized Evaporites and the Evolution of Earth’s Early Biosphere: Geological Society of America Abstracts with Programs, v. 45(7), p. 628. Neudert, M.K., and Russell, R.E., 1981, Shallow water and hypersaline features from the Middle Proterozoic Mt. Isa Sequence: Nature, v. 293, p. 284-286. Rogala, B., and Fralick, P.W., 2005, Stratigraphy and sedimentology of the Mesoproterozoic Sibley Group and related igneous intrusions, northwestern Ontario: Ontario Geological Survey Open File Report 6174, 128 pp. Rogala, B., Fralick, P.W., Heaman, L.M., and Metsaranta, R., 2007, Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada: Canadian Journal of Earth Sciences, v. 44, p. 1131-1149. 9 �Lithological sedimentary divisions of the Copper Harbor Formation in Gogebic and Ontonagon Counties, Michigan BAUMANN, Steven D.J.1, CORY, Alexandra B.1, DYLKA, Sandra K.1 1 Geology Section, Midwest Institute of Geosciences and Engineering, 1321 W. Touhy Ave. 2S, Chicago, IL 60626 The sedimentary assemblage of the Copper Harbor Conglomerate shows four to five distinct sedimentary units (with interbedded volcanics) exposed for a length of approximately 35 miles along Lake Superior in Gogebic and Ontonagon Counties. 7.5’ quadrangle scale mapping in 2014 to 2015 led to the recognition of four distinct sedimentary facies extending from the North Ironwood east along Lake Superior to at least the White Pines 7.5’ quadrangles (see figure 1). A fifth sedimentary unit at the top of the formation exists in the Black River Harbor 7.5’ quadrangle east to the Carp River East 7.5 minute quadrangle. The five sedimentary units can be lumped into two main facies. The lower two units are conglomerate dominated (Units 1 and 2), while the upper three units are dominantly sandstones (Units 3, 4, and 5). These two basic facies appear to be separated by an unconformity, where the contact is exposed along the Black River. The basal unit of the lower facies (Unit 1) is reddish brown to brown, massive, clast supported, polyometic orthoconglomerate about 1700 to 3000 feet thick. The clasts are well rounded and range in size from pebbles to boulders (<14” in diameter). Lithic arkose makes up the majority of the matrix at <15% of the rock. Along the Black River this unit contains about 480 feet of volcanics at about the halfway point within the unit, informally called the “Black River Flows”. These flows are dominantly andesite with some beds of sedimentary rock. The flows are stratigraphically lower in section than the base of the Lake Shore Traps. Unit 2 overlies Unit 1 in a gradational to intertonguing relationship over about 200 feet, except where the Lake Shore Traps are present. The Lake Shore Traps separate the two units with sharp contacts. Unit 2 is a deep red to reddish brown, thick bedded, clast supported, polyometic sandy orthoconglomerate. The clasts are all well rounded and generally larger than in Unit 1 (<22” in diameter). Lenses of deep red, medium to coarse grained, lithic arkosic arenite exist within the unit and are <4’ thick by <50 feet wide. Unit 2 is about 700 to 2800 feet thick. Unit 2 is separated from the upper facies (Unit 3) by a sharp contact, which is a probable unconformity. This unconformity separates the two main facies. Unit 2 thickens at the expense of Unit 3 in the Carp River West and East quadrangles. Unit 3 is deep purplish gray to brown, thin to medium cross bedded, fine to coarse grained, arkosic arenite. It contains beds of polyometic sandy diconglomerate with well rounded clasts <8” in diameter. Ripple marks are very common on bedding planes. Unit 3 is 800 to 3000 feet thick. Unit 3 grades up into Unit 4 over about 30 feet. The lithology of Unit 4 closely resembles the Freda Formation. It is deep red mottled pale yellow brown becoming purplish near the top, medium bedded to cross bedded, fine to coarse grained, arkosic arenite. The unite fines up section. It contains isolated thick beds of polyometic sandy paraconglomerate, with clasts <6” in diameter. Ripple marks are common. Unit 4 has the most consistent thickness at 1000 to 1700 feet. Unit 5 shows an intertonguing and gradational contact with Unit 4. Gradation is over about 20 feet. Unlike the lower four units, it is not continuous throughout the area. It exists as a large lens extending about 21 miles from the North Ironwood to Carp River East quadrangle. It is a deep purple to dark gray, laminated to massive, sandy shale, with some polyometic sandy paraconglomerate beds. Clasts are <3” in diameter. The sandy parts of the unit are mostly coarse lithic arenites. It is the thinnest of the units at 0 to 400 feet thick. The top of the unit is locally covered, except in the Carp River East quadrangle, where the top is gradational with the overlying Nonesuch Formation. The top of the Copper Harbor Formation at Nonesuch Falls along the Little Iron River (White Pine quadrangle), resembles Unit 5 and may be equivalent. However, complex local folding and faulting makes a definite correlation difficult. Units 1 and 2 thin dramatically on the east side of the Porcupine Mountains and pinch out altogether just south of the mountains (at about 46.76o by -89.62o). However, lithologies resembling Units 1 and 2 appear to be traceable on the north shore of the Keweenaw Peninsula, where Unit 1 can be traced along with a facies change in Unit 2. At Horseshoe Harbor (47.473o by -87.809o) Unit 2 includes more continuous beds of sandstone with shale, along with 10 �stromatolites. The thinning of Units 1 and 2 observed around the Porcupine Mountains may just be a local phenomenon caused by the extremely thick Porcupine Volcanics. The four to five fold division of the Copper Harbor Formation is persistent and traceable from about -90.19o (North Ironwood quadrangle) to -89.57o longitude (White Pine quadrangle) along Lake Superior. The four to five fold division west of longitude -90.19o, begins to break down. West of longitude -90.28o (in the Little Girls Point quadrangle), the Copper Harbor Formation appears to be undivided. The Copper Harbor pinches out entirely in Copper Falls State Park, north of Mellen, Wisconsin. References: Bornhorst, T.J., Rose, W.I., 1994. Self Guided Geologic Field Trip to the Keweenaw Peninsula, Michigan. Institute on Lake Superior Geology, volume 40, pp. 161-164 Cannon, W.F., 1995, Geologic Map of the Ontonagon and Part of the Wakefield 30’ x 60’ Quadrangles, Michigan, United States Geological Survey, Miscellaneous Investigation Series, Map I-2499 Dickas, A.B., Mudrey Jr., M.G., 1992. Keweenaw Sedimentary Rock of the South Shore, Lake Superior. Institute on Lake Superior Geology, volume 38, pp. 43-102 White, W.S., Wright, J.C., Lithofacies of the Copper Harbor Conglomerate, Northern Michigan, United States Geological Survey, Professional Paper 400-B, pp. B5-B8 Figure 1: Mapped Quadrangles in Gogebic and Ontonagon Counties 11 �Making it and breaking it in the upper Midwest: Constraints on continental assembly and rifting from EarthScope Paul A. Bedrosian U.S. Geological Survey, Denver Federal Center, MS 964, Denver, CO, 80225 The North American mid-continent presents a window into craton growth and stabilization as well as the 1.1 Ga rifting event that nearly tore Laurentia apart. Unique to this region is the preservation of this tectonic collage, largely unmodified by subsequent tectonic events, which permits examination of if and how such events are preserved in the continental lithosphere. Focusing on the upper Midwest, I will discuss the implications of a three-dimensional resistivity model derived from EarthScope magnetotelluric data [Bedrosian, 2016]. The resistivity model reveals the distribution of highly conductive Penokean age metasedimentary rocks in Minnesota, Michigan, and Wisconsin. These rocks are correlated with metagraywackes of the Michigamme Formation in MI and WI, and with graywackes of the Virginia and Rove Formations in MN. The electrical signature of these rocks is unique throughout the entire midcontinent region. Their high conductivity is attributed to metallic sulfides and in some cases graphite. The former is considered a potential source of sulfur for certain types of mineral deposits found in the region. A more detailed magnetotelluric survey, in addition to a reconnaissance airborne electromagnetic survey, is being carried out to map these rocks in greater detail. An isolated sliver of similarly conductive rocks is mapped in the subsurface in northwest Iowa (sub-parallel to the Spirit Lake Tectonic Zone); I speculate that Penokean-age rocks may be preserved further west than currently assumed. The Paleoproterozoic structural collage was interrupted by the 1.1 Ga Mid-continent Rift System (MRS). The type electrical signature of the MRS is found in Iowa, where a resistive medial horst is imaged, flanked by deep sedimentary basins filled with Bayfield and Oronto Group equivalent sediments. A distinction is observed between the moderately conductive Oronto Group and the highly conductive Bayfield Group. Translating this basic picture to the Lake Superior graben, a pronounced west to east asymmetry is seen, with the Thiel fault being the most obvious division. The asymmetry is taken to reflect the different geometric response of each half of the basin to compression during the Grenville orogeny [Cannon, 1994]. The orientation of the rift-bounding faults in relation to the stress field at the time is speculated to have resulted in a greater degree of compression in the western half of the Superior graben than in the eastern half. Additionally, there appears to be a horst and graben geometry across each half of the Superior graben. At a much larger scale, the resistivity of the mantle lithospheric beneath the region is surprisingly heterogeneous. The spatial pattern of these variations bears little resemblance to the crustal imprint of past tectonic events or to the direction of North American absolute plate motion. I argue that these resistivity variations reflect differing degrees of hydration (metasomatism) preserved within the lithosphere. Lithospheric hydration in the upper Midwest is speculated to have occurred during MRS magmatism. This interpretation is consistent with geochemical and isotopic analyses of MRS basalts and their inferred mantle sources [Nicholson et al., 1997]. REFERENCES Bedrosian, P.A., 2016. Making it and breaking it in the Midwest: Continental assembly and rifting from modeling of EarthScope magnetotelluric data, Precambrian Research, doi:10.1016/j.precamres.2016.03.009 Cannon, W.F., 1994. Closing of the Midcontinent rift-A far-field effect of Grenvillian compression. Geology 22, 155–158. Nicholson, S.W., Schulz, K.J., Shirey, S.B., Green, J.C., 1997. Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development. Canadian Journal of Earth Sciences 34, 504–520. 12 �Unique characteristics of sediment-hosted stratiform copper mineralization resulting from exceptional latent volcanic heat at White Pine, northern Michigan BROWN, Alex C., 13250 rue Acadie, Pierrefonds, QC, H9A 1K9, acbrown@polymtl.ca Most studies of copper mineralization hosted by basal greybeds of the Nonesuch Formation at White Pine, Michigan, have concluded that the main-stage sediment-hosted stratiform copper (SSC) mineralization resulted from an upward influx of cupriferous brine from a coarse-grained footwall aquifer, the Copper Harbor Conglomerate (White and Wright, 1966; Brown, 1971). An early diagenetic timing is interpreted mainly from textural evidence that copper deposited largely by replacement of in situ fine-grained syndiagenetic pyrite (Fig. a), that infiltrations of oreforming brine would have been relatively rapid before advanced compaction and lithification of the basal fine-grained Nonesuch strata, and that main-stage mineralization pre-dated later structurally controlled mineralization (Mauk, 1993). Curiously, the replacement of fine-grained disseminations of euhedral pyrite by cupriferous sulfides at White Pine is clearly visible in a transition zone at the top of the cupriferous zone (cms in width), but not within the cupriferous zone proper (meters in width) where disseminated chalcocite (Fig. b) occurs with a grain-size considerably greater than that of disseminated pre-ore pyrite found above the cupriferous zone. The only strong visual evidence for chalcocite replacement of pyrite within the cupriferous zone is the occurrence of chalcocite nodules which presumably were initially syndiagenetic pyrite nodules; commonly, the chalcocite nodules are surrounded by halos of hematite which probably represent iron released during pyrite replacement. Curiously too, the White Pine deposit is unique in that most other SSC deposits world-wide exhibit fine-grained euhedral pyrite replacements within their cupriferous zones. Also, the White Pine copper mineralization is composed virtually only of disseminated chalcocite with a remarkably narrow transition (Py→Cp→Bn→Cc) into overlying pyritic strata (Fig. c), whereas most SSCs world-wide exhibit gradual copper-iron sulfide transitions over the breadth of their cupriferous zones. The recent proposal (Brown, 2013, 2014) that the upwardly infiltrating cupriferous brine at White Pine was warmed to ~100oC by latent heat from the underlying buried Porcupine Volcanics dome (Fig. d) could offer explanations for the unique character of the White Pine SSC deposit: a) the anomalously warm ore-stage environment could have resulted in a more rapid and complete reaction through the Py→Cp→Bn→Cc sequence to form uniquely chalcocitic mineralization rapidly within the cupriferous zone proper while the copper front advanced upward through the basal Nonesuch graybeds; b) under warm conditions, fine-grained chalcocite could have recrystallized to coarser-grained chalcocite; and c) the preservation of fine-grained zoned replacement textures and pseudomorphs after pyrite euhedral along the top of the cupriferous zone could represent replacements of pyrite at low temperatures prevailing toward the end of the main-stage copper mineralization event. References Brown, A.C., 1971, Zoning in the White Pine Copper Deposit, Ontonagon County, Michigan: Economic Geology, v. 66, p. 543-573. Brown, A.C., 2013, Brine viscosity vs. temperature: A key to explaining copper mineralization in the finest-grained basal Nonesuch Formation in the White Pine-Presque Isle district, northern Michigan: Institute on Lake Superior Geology, Proc. of Annual Meeting, Houghton, Michigan, v. 59, p. 9-10. Brown, A.C., 2014. Latent thermal effects from Porcupine Volcanics calderas underlying the White PinePresque Isle stratiform copper mineralization, northern Michigan: Economic. Geology, v. 109, p. 2035-2050. 13 �Mauk, J.L., 1993, Geological and geochemical investigations of the White Pine sediment-hosted stratiform copper deposit, Ontonagon County, Michigan: Ph.D. thesis, Univ. of Mich., Ann Arbor, 194 p. White, W.S. and Wright, J.C., 1966, Sulfide-mineral zoning in the basal Nonesuch Shale, northern Michigan: Economic Geology, v. 61, p. 1171-1190. a) Disseminated very fine-grained pyrite (some framboidal) in Nonesuch graybeds above the cupriferous zone, White Pine deposit. (Relatively coarse grain of pre-ore stage chalcopyrite (Cp) in center of view). c) Narrow pyrite to chalcocite transition (Py→Cp→Bn→Cc) at the top of the cupriferous zone, White PinePresque Isle district. b) Fine-grained chalcocite (white) of the cupriferous zone, basal Nonesuch graybeds. Chalcocite grain-size decidedly greater than that of pyrite in a). Note also the absence of pseudomorphs after euhedral pyrite. d) Schematic illustration of Porcupine Volcanics dome contributing latent volcanic heat during early diagenetic infiltrations of Cu-bearing brine into basal Nonesuch graybeds, White Pine-. Presque Isle district. See Brown (2014) for further explanations. 14 �The occurrence of Li, B, Sn, and W in the Nine Mile Pluton, Wausau Syenite Complex, Marathon County, Wisconsin. BUCHHOLZ, Thomas W.1, FALSTER, Alexander. U. 2, and SIMMONS, Wm. B. 2 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494; 2Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217. The Nine Mile Pluton is the youngest (≈1505 Ma, Dewane & Van Schmus, 2007) and most silicic of the four intrusions comprising the Wausau Syenite Complex, and is primarily composed of granite and quartz monzonite. Over the last several decades we have developed considerable data on the behavior of some minor elements during the formation of this complex, notably Li, B, Sn and W, which are addressed here. Identification was via electron microprobe (EMP), X-ray diffraction (XRD), energy dispersive spectromentry (EDS) and direct coupled plasma spectroscopy (DCPS) as indicated. Although overall sparse, Li mineralization has been identified in several areas of the pluton, associated with pegmatites and greisens exposed in quarrying operations exploiting easily excavated “rotten granite” or grus. In the former Wimmer pits and adjacent operations, the Limica zinnwaldite, KLiFe2+2Al(Al2Si3O10 (an intermediate composition between siderophyllite and polylithionite), has been identified (EMP, DCPS) from several fractionated pegmatites, associated with columbite- pyrochlore- and euxenite-group minerals. Additionally, approximately 2 gms of elbaite tourmaline (XRD, EMP), Na(Li1.5Al1.5)Al6(Si6O18)(BO3)3 (OH)3(OH), were found in a large (2.4x2.2x0.4 m) miarole in a pegmatite located in the former Thurber pit (later Wimmer #3). Zinnwaldite (EMP, DCPS) has also been found in several pegmatites in the former Koss pit in the central portion of the pluton (associated with cassiterite, columbite-, euxenite- and pyrochlore-group minerals); in greisenized pegmatites and aplites in the former Maguire Pit in the south-central portion of the pluton (associated with cassiterite, ferberite/huebnerite, topaz, cryolite, prosopite and other minerals), and in a greisen-related episyenite and vein assemblage in the Ladick operation (associated with cassiterite, pseudobrookite, monazite, molybdenite and other minerals) in the southwest portion of the pluton. Overall, Li contents in Nine Mile Pluton pegmatite wall zones reach 10-20 ppm, while in the granite itself Li contents range from below detection limit to 4 ppm (DCPS). Boron minerals are extremely rare in the Wausau Complex, having only been found as the aforementioned elbaite tourmaline, and a few grams of schorl tourmaline (EMP) found in a small miarole (5 cm diameter) in the former Wimmer pit #2. This scarcity is probably a consequence of very low boron levels in the Complex. In general, DCPS analysis of Nine Mile Pluton pegmatite wall zones indicate B contents ranging from below detection limit to 3 ppm, and of the granite itself from below detection limit to 1 ppm, although DCPS analysis of the wall zone of the elbaite-bearing pegmatite indicated anomalous B contents of 12-15 ppm. Sn mineralization is not abundant in the Nine Mile Pluton, but is relatively widespread. Cassiterite, SnO2, has been found as apparent epitaxial overgrowths on ilmenite in a small miarolitic granite body on the western side of the former Wimmer Pit #3 in the NE portion of the pluton (EDS). Tantalian cassiterite (EMP) was found in a small fractionated pegmatite nearby where it was closely associated with columbite-group minerals, U-rich pyrochlore, a U-niobate mineral and hafnian zircon, and was also identified (EDS, EMP) from a fractionated pegmatite exposed by workings in the former Koss Pit in the central portion of the pluton Sparse cassiterite has also been found as red-brown grains in chloritized biotite from an euxenite and fluorite-rich pegmatite exposed in the Kafka operation in the western portion of the pluton (EDS), and in a 15 �thin quartz-pyrite vein briefly exposed in the floor of the Red Rock Southwest Pit in the southwest portion of the pluton (EDS). Cassiterite was abundant in greisenized pegmatites and aplites exposed by workings in the former Maguire Pit in the southern portion of the Nine Mile Granite (EDS, EMP), and tantalian cassiterite (EMP) was common in the greisen-related episyenite and vein assemblage in the Ladick quarry briefly described above. It is interesting that several analyzed cassiterites have elevated tantalum contents of 0.3 to 0.4 wt% Ta2O5, while their Nb2O5 contents are much lower, suggesting that Ta is more compatible in cassiterite than Nb. Tungsten mineralization is very limited within the Nine Mile Pluton, and has only been found in the greisenized pegmatites and aplites that were exposed in the former Maguire Pit briefly described above. “Wolframite” (sensu lato) was common as crystals to about 1.5 mm, and varied from Fe-rich ferberite, FeWO4, to Mn-rich huebnerite, MnWO4, often within the same crystal (EMP). Small grains of a Pb-tungstate mineral, probably either raspite or stolzite (EDS, EMP, both Pb(WO4)), were present as inclusions in ferberite. Closely associated columbite-(Fe) in some cases contained up to 12 wt.% WO3 (EMP), resulting in tungstenian columbite-(Fe). Such high-W columbites are likely disordered, and may be “wolframoixiolite”, (Nb,W,Ta,Fe,Mn)2O4 (MINDAT, accessed 03/2016), but due to limited amounts of material this hypothesis has not been confirmed. Finally, scheelite (EMP), CaWO4, was noted as small grains in heavy mineral separates from several dikes, and ferberite pseudomorphs of octahedral to cubic morphology were found in one dike (No. 5), and are likely ferberite replacements of scheelite. Reference: Dewane, T. J., Van Schmus, W. R. (2007): U-Pb geochronology of the Wolf River batholith, north-central Wisconsin: Evidence for successive magmatism between 1484 Ma and 1468 Ma. Precambrian Research, V. 157, pp. 215-234. 16 �Mobilization of silica by flash heating of silica gel beneath the Sudbury Impact Layer, Baraga Basin, Michigan CANNON, W.F. U.S. Geological Survey, Reston, VA 20192 The Baraga Basin in northern Michigan is a structural depression in which sedimentary rocks of the Paleoproterozoic Baraga Group are preserved. The section includes ejecta-bearing rocks of the Sudbury Impact Layer (SIL) deposited 1,850 million years ago (Cannon and others, 2010). At the time of the Sudbury impact, the area of the Baraga Basin was covered by a shallow, perhaps intertidal, sea in which chemical sediments, now interlayered chert, carbonate minerals, and minor iron minerals were accumulating. The ejecta layer in this area, about 500 kilometers west of the impact site, is typically a few meters thick and consists mostly of impact glass in the form of spherules and shards and some interlayers rich in accretionary lapilli. The SIL was deposited across the area from a fast-moving curtain of ballistic ejecta that, in some of the area, was transformed into a base surge that incorporated parts of the underlying rock units into the ejecta-bearing layer. A distinguishing feature of the SIL in the Baraga Basin is the prevalence of a very siliceous matrix surrounding ejecta particles. The SIL also contains many particles that are themselves very siliceous, many of which are highly amygdular. Both features suggest an unusual involvement of silica in the deposition of the ejecta layer. Some of the uppermost beds of the sedimentary sequence at the time of the impact may have been silica gel. Gels can contain more than 90 wt. % water. When these gels came into contact with newly deposited high temperature ejecta, or were fragmented and incorporated into it, flash heating by the hot ejecta resulted in explosive release of steam and generation of silica-rich fluids that infused much of the ejecta layer. Most ejecta particles appear to have been solid during deposition, but flattened and stretched forms suggest that they were hot enough to have been soft and easily deformed. The presence of soft gels is shown microtexturally by siliceous (now cherty) inclusions in the ejecta that are distorted as indicated by both their external shape and internal features (Fig. 1). Especially compelling textures are shown by slivers of hardened silica that sharply penetrated into gelatinous clasts (Fig. 2). Many chert clasts have an unusual internal structure consisting of closely spaced amygdules whose outlines are shown by a film of sericite grains, but are otherwise entirely a recrystallized mosaic of fine quartz grains (Fig. 3). The term “popcorn chert” seems appropriate for them in that they appear to have formed by flash vaporization of their original high water content to produce these once highly vesicular silica masses. Commonly the amygdules are moderately to strongly flattened parallel to bedding (Fig. 4). Some samples show arcuate films of sericite grains that may be fragments of broken bubble walls (Fig. 5 ) suggesting that boiling also produced independent hollow structures that were later collapsed and broken. A well preserved contact of the base of the ejecta with underlying chert, seen in one drill hole, appears to have preserved this flash boiling in progress (Fig 6). The cuspate and discordant upper contact of chert containing thin magnetite laminae is overlain by a selvage of popcorn chert that may have formed in situ as chert gel was converted to a silicasteam mixture upon heating from the overlying ejecta. These features, together, suggest that a substrate of hydrous silica played an important role in determining the character of the SIL in the Baraga Basin. Cannon, W.F, Schulz, K.J., Horton, J. Wright, Jr., and. Kring, David A., (2010) The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA: Geological Society of America Bulletin, v. 122, p. 50-75. 17 �1. Irregular mass of highly vesicular chert surrounded by glassy ejecta. 2. Sliver of chert that penetrated and deformed a clast of chert-magnetite gel. 3. Fragment of popcorn chert in glassy ejecta. Amygdules are outlined by a coasting of sericite. 4. Fragment of popcorn chert whose upper half is compressed and highly flattened. 5. Films composed mostly of sericite in a siliceous matrix. These curved films may be the walls of broken bubbles. 6. Upper contact of chert with magnetite lamellae and overlying glassy ejecta. Popcorn chert at contact. 18 �Traces of the Sudbury meteor impact in the western Gogebic Iron Range, northern Wisconsin CANNON, W.F.1, WOODRUFF, Laurel G.2, and SAARI, Stacy3 1 U.S. Geological Survey, MS 954, Reston, VA 20192 2 U.S. Geological Survey, 2280 Woodale Drive, Mounds View, MN 55112 3 Global Minerals Engineering, LLC, Hibbing MN 55746 Rock layers containing ejecta particles and other indications of the 1850 Ma Sudbury meteor impact have been observed in core from 12 drill holes along 40 km of the Gogebic Iron Range in northern Wisconsin. This ejecta-bearing material, found about 650-700 km west of the impact site near Sudbury, Ontario, is distinct from other occurrences of the Sudbury Impact Layer around the Lake Superior region. It occurs in multiple beds of reworked ejecta that commonly have ultra-potassic compositions. Ejecta-bearing material is found in two types of deposits: 1thick (up to 20 m) debris flows of coarse breccia, and 2- multiple thinner (0.02- 2 m) turbidite beds throughout the lower 25 m of the Tyler Formation (Fig.1). Figure 1. Stratigraphic relationships of ejecta-bearing beds to the regional stratigraphy of the Gogebic Iron Range. The Pence Member of the Ironwood Iron Formation is an even-bedded, generally moderately magnetic, iron formation containing minor units of granular ironformation. The Tyler Formation, in its lower part, is laminated argillite and shale containing variable amounts of carbon and pyrite. It contains thin units of chert and magnetic chert in some drill holes. Debris flows- Three drill cores contain coarse breccia from 10-20 m thick at the contact of the Pence Member of the Ironwood Iron Formation and Tyler Formation. They appear to be three separate units because breccias are absent in intervening drill holes. Where breccia is present, the thickness of the underlying Pence Member is less than in drill holes without breccia, suggesting that the debris flows incised their own channels into the Pence. The flows consist of a mixture of apparently local rocks and ejecta, including accretionary lapilli, and altered impact glass. They are massive to crudely bedded. A sparse set of quartz grains with shock metamorphic features (Fig. 2A) is a definitive indication of a relationship to the Sudbury impact. Turbidite beds- Ejecta-bearing beds that appear to be turbidites are interbedded with laminated shale and siltstone of the lower 25 m of the Tyler Formation. The number of beds varies from 0 to 8 in various drill holes and the beds vary in thickness from a few cm to 2 m. Many beds have prominent normal size grading. These beds cannot be correlated either in number or thickness between closely spaced drill holes showing that they are individual tongues of material with limited lateral extent, at least along the current strike of the iron range. The occurrence of multiple ejecta-bearing beds leads to the obvious conclusion that they are 19 �reworked material derived from the original ejecta blanket. These beds were commonly identified as tuff on drill logs, but a suite of features characteristic of ejecta, including a sparse but well developed suite of impact shocked quartz grains (Fig. 2B) indicate a link to the Sudbury event. Figure 2. Examples of quartz grains with relict planar deformation features indicative of intense shock. A-from a debris flow; B-from a turbidite bed. Figure 3. A-microcline phenocrysts in a vesicular glass fragment, B-microcline phenocrysts in a quench-textured microcline-rich fragment, C-amygdular glass fragment composed almost entirely of fine-grained mosaic-textured microcline. Both types of beds have ultra-potassic compositions. K2O ranges from 8.4 to 13.0 wt. % in five samples analyzed. The K content results both from abundant small phenocrysts and fragments of microcline, and from a finer mosaic of microcline that overprints original textures of clasts. The phenocrysts are a very early component of the ejecta, being enclosed both in accretionary lapilli and glass fragments (Fig. 3A, B). The finer microcline may be secondary-a result of seafloor K-metasomatism akin to formation of K-bentonites in younger volcanic ash. Individual clasts of K-rich ejecta indicate that they were altered before being incorporated into the present beds. The ejecta-bearing beds record a set of processes including: 1) deposition of an original ejecta blanket in a shallow marine setting probably not far from the current study area, 2) alteration of the ejecta to K-rich partly consolidated sediments, and 3) remobilization of the metasomatized ejecta to produce the debris flows and multiple ejecta-bearing turbidite beds. 20 �Microstructural comparison of the Hardrock Project at Geraldton, Ontario and the Coffee Gold Project, Yukon CARSON, Tracy, DELEY, Brittany, and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON, P7B 5E1, Canada The Superior Province of the Canadian Shield and the North American Cordillera both host extensive gold mineralization. These regions are distant from each other and very different in age, but the similarities between microstructures within gold-hosting units are striking. Separate microstructural studies were done on gold-hosting units in the Coffee Gold Project south of Dawson, Yukon and the Hardrock gold project at Geraldton, Ontario and the results are compared. The focus in each project is to examine the microstructures within gold-hosting rock units to determine the protolith of the host and assess structural control on gold mineralization. Based on mineralogy, each study concludes that a felsic igneous rock is the likely protolith to the gold-hosting lithology. Each study finds similar styles of deformation resulting in microstructures typical of a mylonite. The microstructures include a foliation defined by parallel alignment of muscovite and rigid porphyroclasts of feldspar in a fine-grained groundmass. Together, these observations lead to the conclusion that the host for gold mineralization in each study is a deformed felsic plutonic rock. Both gold-hosting lithologies have similar mineralogy consisting dominantly of feldspar, quartz and muscovite. Rigid porphyroclasts of plagioclase are situated within a groundmass composed of <0.1mm quartz, feldspar and muscovite. Plagioclase is the dominant feldspar in both studies. Feldspar is present both in the groundmass and as rigid porphyroclasts. The plagioclase porphyroclasts commonly have asymmetrical tails, indicating non-coaxial strain, and commonly contain deformation twins. Sericite alteration on porphyroclasts is evidence for the presence of hydrous metamorphic fluids. In the Geraldton study these rigid porphyroclasts are fractured. Quartz is common within the groundmass in both studies, and late quartz veins are abundant within the Yukon study. Undulose extinction and subgrain boundaries indicate dislocation creep in quartz. Dimensional preferred orientation of muscovite defines the foliation and muscovite bends around the rigid porphyroclasts. The Coffee Gold Project is hosted within the Yukon–Tanana Terrane bounded between the Tintina and Denali Faults and along strike of the Teslin Fault, this area is located within the Tintina Gold Province. The study examines three zones, Kona, Supremo and Latte. In each of the zones a felsic plutonic rock with varying degrees of deformation was interpreted to be the protolith. The Kona zone is mapped as granite and is the least deformed of the three zones. The Kona zone’s mineralogy consists of feldspar, quartz, sericite and biotite with an average grainsize of approximately 2.5mm. Deformation in the Kona zone includes undulose extinction, subgrain boundaries and serrated grain boundaries in the quartz as well as sericite alteration of feldspars. Supremo and Latte zones show increased deformation with the development of a foliation bending around rigid porphyroclasts and deformation features in quartz and feldspar. Quartz microstructures such as subgrain boundaries, serrated grain boundaries and undulose extinction are common and indicate deformation by dislocation creep. Rigid porphyroclasts of feldspar are common providing a minimal original grainsize of ~3mm. Deformation twins observed within feldspars provide further evidence of deformation. The Latte zone is mapped as biotite schist, however biotite is present in only small amounts. The mineralogy in the Latte zone 21 �consists dominantly of quartz and feldspar suggesting a felsic igneous protolith. A relationship between deformation and gold mineralization is observed; as deformation increases the gold mineralization increases. Gold mineralization based on published assay values was lowest in the Kona zone and increased in the more deformed Latte and Supremo zones. The protolith of all three zones is likely a deformed felsic plutonic rock with varying degrees of metamorphism and deformation. The Hardrock Project is located within the Beardmore-Geraldton belt. The BeardmoreGeraldton belt lies between the Wabigoon and Quetico subprovinces and it is comprised of multiple fault-bounded metavolcanic and metasedimentary belts. The Hardrock Project is one of several past-producing gold mines situated within the Beardmore-Geraldton gold camp. The Beardmore-Geraldton gold camp is bounded by major steeply dipping structures that are comparable to structures in the Tintina Gold Province. Mineralogy and microstructure of the “porphyry” within the Hardrock Project are consistent with a deformed felsic plutonic rock similar to the Latte and Supremo zones in the Coffee Gold Project. The Latte and Supremo zones in the Coffee Gold Project have mylonitic features that are comparable to the “porphyry” at the Hardrock Project. The “porphyry” is known to have significantly more mineralization than its protolith, Croll Lake stock, suggesting that gold mineralization is related to deformation. This is similar to deposits within the Tintina Gold Province such as the Coffee Gold Project. The Beardmore-Geraldton belt may be an older Cordilleran-type gold province with structural controls on gold mineralization similar to the Tintina Gold Province. This comparison gives insight into the similarities of shear-zone-hosted gold deposits. In both examples, mylonitized felsic plutonic rock hosts gold mineralization. 22 �Utility of the horizontal-to-vertical spectral ratio (HVSR) passive seismic method for determining Quaternary sediment thickness and bedrock elevation in north-central Minnesota: Fun with little control and generally poor data Val W. Chandler and Amy L. Radakovich Minnesota Geological Survey, 2609 Territorial R., St. Paul, MN 55114 Work continues in Minnesota on using the horizontal-to-vertical-spectral ratio (HVSR) passive seismic method to help map the thickness of Quaternary glacial sediments and the topography of the bedrock surface. A total of 570 passive seismic stations have been acquired in and around Becker, Cass, Hubbard, and Wadena Counties in north-central Minnesota as part of the County Geologic Atlas (CGA) program of the Minnesota Geological Survey (MGS). This area is characterized by thick and complex glacial sequences, widely scattered drillhole control, and, unfortunately, generally poor HVSR data. Ideally a HVSR spectrum yields a largeamplitude, singular peak that represents the bedrock surface (Chandler and Lively, 2014; Lane and others, 2008). In north-central Minnesota, the observed HVSR spectra instead commonly consist of low-amplitude and multiple peaks, where recognition of bedrock signatures is difficult when observing spectra individually. The cause of the poor signal quality is unknown, but could be related to a highly irregular bedrock surface or to the complexity of the overlying glacial sequences. In areas with these types of issues we have made increased use of multi-station spectral cross-sections that have been converted to approximate elevation sections (Fig. 1). This conversion uses a power-law calibration that is based on observed HVSR results at control points, usually consisting of water wells or seismic refraction soundings. Due to the lack of a sufficient number of control points in the study area, a generalized calibration is used, which is based on 303 widely distributed control points in Minnesota. More refined sections may eventually be possible with locally derived calibrations, but our generalized approach produces useful sections that readily allow lateral correlation of HVSR features within a framework of available geologic control. Figure 1. Color HVSR section across central Becker County. West end of profile located at UTM 252686E/5203087N; East end of profile located at UTM 335129E/5210337N) White X’s designate HVSR-based estimates of the bedrock surface. White diamonds indicate bedrock horizons at control points, either drill-holes with well number (6 digit label), or seismic refraction sites (4 digit label). Bedrock horizons are: K, the top of Cretaceous sediments; Sap the top of saprolite; and Fbr the top of fresh bedrock. Individual stations are labeled in black. AM approximates the western margin of the Alexandria Moraine. Vertical scale is elevation and horizontal scale is cumulative distance between successive stations. 50X vertical exaggeration. 23 �Using the cross-sectional approach as a guide, viable maps have been produced that estimate the elevation and depth of the Precambrian bedrock surface that underlies the study area. Trough- like depressions in the Precambrian surface are indicated beneath northeastern and central Becker County and beneath northeastern Wadena and eastern Hubbard Counties, where total depths are 500-1000 ft. (150-300 m). These depressions are known to locally contain Cretaceous strata, but the HVSR method cannot discriminate the soft Cretaceous rocks from the glacial sediments. Broad rises in the Precambrian surface are indicated near the junction of Becker, Hubbard, and Wadena Counties, and much of southern and eastern Cass Counties. These preliminary interpretations have been largely corroborated by follow-up seismic refraction profiling and test drilling. The results of this study indicate that when used with the proper precautions, the HVSR method can be useful for bedrock depth and elevation studies, even under less-than-ideal conditions. REFERENCES Chandler, V. W., and Lively, R. S., 2014, Evaluation of the horizontal-to-vertical spectral ratio (HVSR) passive seismic method for estimating the thickness of Quaternary deposits in Minnesota and adjacent parts of Wisconsin: Minnesota Geological Survey Open File Report 14-01, 52 p. Lane, J. W., Jr., White, E. A., Steele, G. V., and Cannia, J. C., 2008, Estimation of bedrock depth using the horizontal-to-vertical (H/V) ambient-noise seismic method: in Symposium on the Application of Geophysics to Engineering and Environmental Problems, Proceedings of the Environmental and Engineering Geophysical Society, 13 p. 24 �Bedrock Geology of the Devilfish Lake Area, Cook County, Minnesota CLARK, Jonathan1, ESHLER, Kristen2, GROFF, Patrick3, MCCLENDON, Taylor4, RODE, Alexander5, SALINGS, Emily6, SPINELLI, Kristen7, VANDER WYST, Kyle8, WALSH, Aiden9, ASP, Kristofer10, and LARSON, Phillip11 1 Northwest Missouri State University, Maryville, MO, 2Temple University, Philadelphia, PA,3California State University, East Bay, Hayward, CA, 4Huffington Department of Earth Science, Southern Methodist University, Dallas, TX, 5Wayne State University, Detroit, MI, 6Missouri State University, Springfield, MO, 7Department of Geological Studies, Binghamton University, Binghamton, NY, 8University of Wisconsin-Milwaukee, Milwaukee, WI, 9 School of the Environment, Washington State University, Pullman, WA, 10Department of Earth and Environmental Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth, MN 55812, 11Vesterheim Geoscience PLC, 120 Greenwood Ln, Duluth, MN 55803 and Department of Earth and Environmental Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth, MN 55812   Grout and others (1959) has heretofore been the primary published work investigating the contact between the reversed polarity lower northeast sequence of the North Shore Volcanic Group (NSVG) and the early magmatic stage of the Duluth Complex (DC) in the vicinity of Devilfish Lake, and by extension the eastern limit of the Mesoproterozoic Duluth Complex, Minnesota, USA. Geologic mapping and gravity surveying by the Precambrian Research Center Field Camp in the summers of 2014 and 2015 provides additional detail to better contextualize Grout’s pioneering work with more recent understanding of the geological and geophysical nature of the DC, NSVG, and Midcontinent Rift as a whole (e.g. Miller et al., 2012). Targeted mapping addressed 1) understanding the contact between the DC and hangingwall lower northeast sequence NSVG volcanic rocks, 2) identifying small-scale mafic intrusive dikes and sills cross-cutting the NSVG and/or DC, and 3) infilling state-wide gravity station coverage along roads and trails constructed since prior gravity surveying in the late 1960s. In this area, volcanic rocks of the ~1108 Ma lower northeast series of the NSVG form an ~9 km thick, shallow, southwest-dipping sequence overlying the Paleoproterozoic Rove Formation. The base of the sequence is composed of relatively primitive tholeiitic basalts of the Grand Portage lavas, and is capped by the relatively evolved basaltic andesites, andesites, icelandites, and rhyolites of the Hovland lavas. The Hovland lavas are characterized by abundant feldspar phenocrysts. This study recognizes a sequence of thick-bedded, relatively evolved, aphyric basalts and basaltic andesites intermediate to the Grand Portage and Hovland lavas, herein named the Esther lavas. The Grand Portage and Esther lavas successively truncate against the DC toward the west. The base of lower northeast sequence was subsequently intruded by the Crocodile Lake gabbro (CLG) and Cucumber Lake granophyre (CL); these intrusions have been correlated with the ~1109 Ma Poplar Lake intrusion. A general lack of volcanic or sedimentary xenoliths in the CLG and CL suggest they intruded as sheet-like bodies along the basal contact of the NVSG with the underlying Rove Formation, with little incorporation or removal of wall rock, suggesting the pinch-out of the Grand Portage and Esther lavas may be a primary characteristic reflecting original volcanic basin geometry. The contact between the DC (CL) and overlying NSVG is much more irregular than indicated by previous mapping, consistent with a relatively shallow, southward-dipping contact. Hornfelsed NSVG volcanic rocks immediately overlying this contact form a prominent linear trend of hills immediately south of the contact, and apophyses of granophyre are common in the hornfels. Overall, the contact between the DC and NSVG appears to mirror the dip of flows within the NSVG. 25 �Mapping identifies at least four mafic intrusive phases with distinct mineralogy and geochemistry cross-cutting the NSVG and DC. Sill-like(?) gabbroic intrusions cross-cutting the Esther lavas are the only pre- or syn-DC intrusions recognized by this mapping. Multiple mafic intrusive phases cross-cutting the DC occur as dikes in both NW-SE and SW-NE trending structures. E-W oriented, normal polarity dikes (Chester dikes) contains abundant anorthosite xenoliths, suggesting correlation with the ~1096 Ma Beaver River diabase. Mafic dikes cross-cutting the CLG display both coarse-grained and chilled, aphanitic textures, suggesting multiple intrusive events during the cooling of the CLG. Previous to this study, state-wide gravity data indicated an ~45 mgal gradient in the Bouguer anomaly over ~5 km between points collected over the Rove Formation (north) and the contact between the CLG and CL intrusions and hangingwall NSVG volcanics (south). Infill surveying of this gap suggests that the CLG and CL intrusions are southward-dipping sheet-like bodies. Truncation of this E-W oriented gradient coincident with the eastern extent of the DC suggests intrusion of the CLG/CL was controlled in part by subsidence along a rift-perpendicular tear fault. Mapping results support the hypothesis that the early, eastern extent of the DC was emplaced as sheet-like intrusions at the base of the lower northeast sequence NSVG volcanic pile. The NSVG pile may preserve pinch-out of individual volcanic units reflecting original basin geometry. Multiple geochemically, texturally, and mineralogically distinct mafic dikes cross-cutting both the early DC and NSVG attest to a heretofore unrecognized long and complex history of post-early magmatic stage igneous activity in this sector of the Midcontinent Rift. References Grout, F.F., Sharp, R.P., and Schwartz, G.M., 1959, The geology of Cook County Minnesota: Minnesota Geological Survey Bulletin 39, 163 p., 16 pls. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., 2002, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations 58, 207 p. 26 �Geochemistry and petrogenesis of volcanic rocks in the Coldwell Alkaline Complex; new insights from the Wolfcamp Lake volcanic rocks CUNDARI, Robert1, HOLLINGS, Peter2, GOOD, David3, and DAVIS, Sarah2 1 Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, 435 James St. S., Suite B002, Thunder Bay, ON, P7E 6S7 Canada 2 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada 3 Department of Earth Sciences, University of Western Ontario, London, ON N6A 5B7 Recent work carried out on volcanic rocks present within the Coldwell Alkaline Complex has unveiled new insights into the genesis of Midcontinent Rift-related volcanic rocks along the North Shore of Lake Superior, Ontario. Coldwell Complex volcanic rocks display broadly similar trace element patterns to basalt type I and II (cf. Nicholson et al. 1993) suggesting eruption in the early magmatic stage of the Midcontinent Rift (Good et al. 2015; Cundari et al. 2012). This is supported by the intrusion of the Eastern Gabbro suite into Upper and Lower metabasalt units at 1108±1 Ma and the intrusion of centre II syenites into the Coubran Lake and Wolfcamp Lake basalts at 1109+8-4 Ma (Heaman and Machado, 1992). Good et al. (2015) proposed three distinct basaltic packages present within the Coldwell Complex, listed here as progressively more evolved units based on Mg number and Ni abundance (Figure 1a); the Lower meta-basalts (LMB), the Upper meta-basalts (UMB) and the Coubran Lake basalts (CLB). The Wolfcamp Lake basalts (WLB) represent a fourth unit in this sequence more evolved than previously recognized volcanic units in the Coldwell Complex. WLB display an average TiO2 content of 1.86 compared to 0.86 for the CLB, 1.15 for the UMB, and 1 for the LMB suggesting the Wolfcamp rocks are distinctly different and more evolved than those suites (Figure 1b). The WLB also exhibit the highest total alkali abundance of all four volcanic units present within the Coldwell complex. WLB display greater LREE abundances compared to the CLB but lack a negative niobium anomaly suggesting the LREEenrichment is not due to crustal contamination. Additionally, SiO2 content for the WLB is too low to be explained by contamination of continental crust. The CLB were interpreted to be the volcanic equivalent of the Two Duck Lake gabbro based on trace element abundances by Cundari (2012). Trace element abundances for the Geordie Lake gabbro trend towards the WLB suggesting a co-genetic relationship whereby the WLB may represent the volcanic equivalent of the Geordie Lake gabbro. Gd/Ybn ratios for the WLB are broadly similar to those for the CLB suggesting a similar depth of partial melting. The extreme LREE enrichment exhibited by the WLB compared to the CLB may be due to lower degrees of partial melting although it should be noted that field evidence such as a reddish staining of alkali material suggests that WLB have been altered by syenite intrusions which may have contributed to the LREE content of these rocks. A number of mechanisms may be responsible for the LREE enrichment characteristics observed in the Coldwell Complex volcanic rocks, notably the CLB and the WLB. Extreme LREE enrichment may be controlled by apatite accumulation, suggested by a positive correlation between Th and Gd with P2O5 although LREE abundances may be too high to be accounted for entirely by apatite. An alternative mechanism for generation of LREE enriched volcanic rocks may be partial melting of an enriched mantle source suggesting a possible relationship with Coldwell syenitic intrusions. Further work is required to develop the relationships between the volcanic units present within the Coldwell Complex, most notably, the relationship between the Coubran Lake and Wolfcamp Lake volcanic units. Modelling of fractional crystallization of average Coubran basalt composition trending toward Geordie Lake gabbro and Wolfcamp basalt may explain the petrogenesis of these units and establish the magmatic relationship between them. 27 �a b c d Figure 1: Major and trace element abundances of Coldwell Complex volcanic units compared to Coldwell Complex intrusive units. Figure 2: La/Sm vs. Gd/Yb plot of Coldwell Complex volcanic units compared to Coldwell Complex intrusive units. References Cundari R., 2012, Geology and geochemistry of Midcontinent rift-related igneous rocks: M.Sc. thesis, Thunder Bay, ON, Lakehead University, 122 p. Good, D.J., Hollings, P., Cundari, R. and Ames, D. Significance of LREE-enriched mantle source to genesis of basalt in the Coldwell Alkaline Complex, Midcontinent Rift, Ontario. 61st Institute on Lake Superior Geology, Dryden, ON, May 20-24, 2015, Proceedings Volume 61, Part 1, p.70-71. Heaman, L.M. and Machado, N., 1992. Timing and origin of midcontinent rift alkaline magmatism, North America: evidence from the Coldwell Complex; Contr. to Mineralogy and Petrology, 110, p.289-303. Nicholson, S.W., Shirey, S., Schulz, K., and Green, J., 1997. Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development. Canadian Journal of Earth Sciences 34: 504-520. 28 �Mineralogy and Geochemistry of the Wolfcamp Lake Basalts DAVIS1, Sarah, HOLLINGS1, Pete and CUNDARI2, Rob 1. Department of Geology, Lakehead University, Thunder Bay, ON, Canada 2. Ontario Geological Survey, Ministry of Northern Development, Mines and Forestry, Suite B002, 435 James St. South Thunder Bay, ON P7E 6S7 Canada The Wolfcamp Lake Basalts are Midcontinent Rift-related volcanic rocks (Walker et al., 1993) that have been linked with other MCR volcanic rocks such as the Coubran Lake, Penn Lake, Bamoos Lake and Foster Island volcanic units along the north shore of Lake Superior which have all been interpreted to be roof pendants within the Coldwell Complex (Puskas, 1967; Walker et al., 1993). Basalts of the Wolfcamp Lake volcanic unit are exposed in two main areas bisected by the TransCanada highway northwest of Marathon, Ontario. The entire unit is approximately 2 km east-west by 4 km north-south with one large exposure northeast of the highway east of Wolfcamp Lake and a second exposure south of the TransCanada highway cropping out as large cliffs, railway cuts and outcrops on the north shore of Lake Superior near Port Munroe. Flow thicknesses are generally 2-4 m with variations present locally. The main lithology comprises basalt flows which are dominantly ophitic or sub-ophitic. The mineralogy is dominated by feldspars (primarily plagioclase), olivine and pyroxenes. Alteration minerals including hornblende, sillimanite, chlorite, biotite, opaques (primarily magnetite) are present in all samples at concentrations from 3% up to 40%. All of the samples of the Wolfcamp Lake basalts show very consistent trace element geochemistry with OIB-like characteristics on primitive mantle normalised plots. They are characterised by negative zirconium, hafnium and titanium anomalies with no negative niobium anomalies observed. The Wolfcamp Lake basalts were compared to the closest volcanic rocks which are the Coubran Lake Basalts, but differences were observed when compared to other volcanic rocks in the area. When compared to the three types of basalts identified in the Coubran Lake unit the Wolfcamp Lake basalts lack the strong negative niobium anomaly displayed by the Coubran Types A and B which was interpreted to be the result of contamination (Fig. 1). The Wolfcamp Lake Basalts also show lower Mg numbers than the Coubran Lake Basalts indicating that the former are more evolved. This suggests that the Wolfcamp basalts are uncontaminated MCR magmas despite their more evolved compositions. The Wolfcamp Lake basalts are broadly similar to those of the Lower Osler volcanic rocks, with similar LREE patterns and no negative niobium anomaly (Fig. 2). However, the Osler volcanics display a lower La/Smn ratio on a Gd/Ybn vs La/Smn diagram (Fig. 3) indicating a more primitive composition compared to the Wolfcamp Lake Basalts. 29 �Figure 1: Spider plots comparing geochemistry of Average Wolfcamp Lake Basalts to Coubran Lake Basalt Types A, B and C. Comparative data from Cundari (2012). Figure 2: Spider Plot comparing Wolfcamp Lake Basalts to the Lower Osler Volcanics. Comparative data from Hollings et al. (2007). Figure 3: Plot of La/Sm vs Gd/Yb showing Midcontinent Rift volcanics. Normalizing values from Sun and McDonough (1989). Comparative data from Cundari (2012). REFERENCES Cundari, R. 2012. Geology and Geochemistry of Midcontinent Rift-related igneous rocks. Masters Thesis, Lakehead University, Thunder Bay, Ontario. Good, D.J. 1993. Genesis of Copper-Precious Metal Sulfide Deposits in the Port Coldwell Alkalic Complex, Ontario. Ph.D. Thesis, McMaster University, Hamilton. Ontario Geological Survey, Open File Report 5839. Hollings, P., Fralick., P. and Cousens, B., 2007c. Early history of the Midcontinent Rift inferred from geochemistry and sedimentology of the Mesoproterozoic Osler Group, northwestern Ontario. Canadian Journal of Earth Sciences 44: 389-412. Puskas, F.P. 1967. Geology of Port Coldwell Area, District of Thunder Bay, Ontario. Ontario Department of Mines. Open File Report 5014. Sun, S.S. and McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes, In: Summary of field work. Ontario Geological Survey, Miscellaneous Paper 100, P. 26-29. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T., Penczak, R.S., 1993. Precambrian Geology of the Coldwell Alkalic Complex. Ontario Geological Survey. Open File Report 5868. 30 �Origin of the gold-hosting porphyry at Geraldton, Ontario DELEY, Brittany and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON, P7B 5E1, Canada The lithologic unit known as the “porphyry” at Geraldton, Ontario is host to widespread gold mineralization, but its origin is poorly understood. The porphyry occurs within the BeardmoreGeraldton greenstone belt in the Superior Province of the Canadian Shield. Gold was first discovered in the Beardmore-Geraldton belt in 1917, and production occurred between 1930 and 1970 at 20 mines throughout the belt, producing 4.1 million ounces of gold (www.premiergoldmines.com, 2016). Exploration is currently active again in the BeardmoreGeraldton belt. Microstructural analysis was done on samples from the “porphyry” to better understand its origin. Samples from the porphyry contain approximately 1mm porphyroclasts of plagioclase within a groundmass composed of approximately 0.1mm grains of quartz, plagioclase, and muscovite. The fine-grained groundmass is formed as a result of grainsize reduction by dislocation creep. The plagioclase porphyroclasts commonly have asymmetrical tails, and muscovite crystals tend to bend around the porphyroclasts. Asymmetrical tails are an indication of non-coaxial strain. Plagioclase porphyroclasts are commonly fractured with calcite infilling the fractures. The presence of calcite indicates a CO2-bearing fluid phase. Sericite alteration occurs on porphyroclasts indicates the presence of H2O-bearing metamorphic fluids. Plagioclase commonly contains deformation twins. Undulose extinction and subgrain boundaries indicate dislocation creep in quartz. The porphyry also has a well-developed foliation defined by parallel alignment of muscovite. Based on microstructural analysis, it appears that the most likely protolith for the porphyry is a felsic plutonic rock. The microstructures are typical of a mylonite. Characteristic features of mylonitic rocks include a strong foliation produced by the parallel alignment of minerals, the presence of a fine grained matrix produced by grainsize reduction mechanisms with porphyroclasts, and the presence of a certain asymmetry. Approximately 30km east of Geraldton, near Longlac, Ontario, is a 150 km2 elliptical, granodiorite intrusion, the Croll Lake stock. This intrusion is the nearest felsic and plutonic rock to the porphyry at Geraldton. The Croll Lake stock is also deformed. Deformational features in quartz include undulose extinction, subgrains and serrated grain boundaries. Plagioclase commonly contains deformation twins. Evidence for deformation of the stock increases westward toward its margin where it resembles the porphyry. Microstructural analysis of the porphyry and the Croll Lake stock suggest that the “porphyry” is a mylonitized fragment of the Croll Lake stock. 31 �PGE Mineralization in the Northern Ultramafic Center of the Lac des Iles Complex, Ontario: Evidence of Magmatic and Hydrothermal processes Djon, M. L.1,3, Olivo, G.R.1 Miller, J.D.2, Peck, D.C.3and Joy, B1 1 Queen's University, Department of Geological Sci. and Geological Engineering, Kingston, Ontario K7L 3N6 Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812 3 North American Palladium Ltd, 10th Avenue, Thunder Bay, Ontario, P7B 2R2. 2 The Archean Lac des Iles Complex in northwest Ontario consists of a mafic complex (South LDI complex) and a predominantly ultramafic complex (North LDI complex). The South LDI complex is ~6 km long and 1.5-2.5km wide and comprises the Mine block intrusion, South LDI intrusion, and Camp Lake Intrusion. The Mine Block Intrusion is the best known of the individual intrusions within the LDIC due to the presence of significant Pd-rich platinum group element (PGE) mineralization that has been mined since 1993 (Lavigne and Michaud 2001). The North LDI complex is a composite, tadpole-shaped, intrusive body with dimensions of 6 km by 4.5 km and featuring a preponderance of ultramafic over mafic cumulate rocks that are distributed in two centres: the Southern Ultramafic Centre and the Northern Ultramafic Centre (Sutcliffe and Sweeney 1985). Various exploration programs conducted over a 50 year span have identified PGE occurrences with >1 g/t Pd+Pt. The best documented of these ultramafic-hosted PGE occurrences are in the Sutcliffe Zone, which consists of four subparallel, stratiform PGE-enriched horizons exposed within the Northern Ultramafic Centre (NUC). Detailed field mapping, petrographic studies, and lithogeochemistry of recently acquired drill cores indicate that the 4 km diameter, >1900m-thick, lopolith-shaped NUC can be subdivided into an Eastern Marginal Zone and a Layered Series. Furthermore, the Layered Series can be divided into fourteen cyclic units defined by cumulus mineral paragenesis of two general types. Cyclic unit type A (CUA) displays a cumulate stratigraphy of Ol+Sp (dunite)  Ol+Opx+Cpx±Sp (olivine-websterite) Cpx+Opx (websterite)  Pl+Opx±Cpx (gabbronorite). Cyclic unit type B (CUB) displays a simpler cumulate sequence of Ol+ Sp (dunite)  Ol+Cpx (olivine clinopyroxenite). Cyclic unit boundaries are sharp between similar types but also where CUA overlies CUB. In contrast, where CUB overlies CUA, the contact is gradational and is marked by the complex hybrid unit. Within individual cyclic units, cyclical cryptic variation in mg# of olivine (Fo) and pyroxenes (En) shows a progressive upward decrease followed by an abrupt to gradual increase at or just below cyclic unit boundaries. Olivines of CUB rocks have lower Ni concentrations and higher Fo contents than olivines in CUA rocks. CUA and CUB samples are further distinguished by having different abundance ranges in incompatible elements, and Zr/Y and Ce/Yb ratios. Djon et al. (submitted) interpreted CUA and CUB type cyclic units to have formed from two compositionally distinct parental magmas. CUA parental magmas would have been enriched in Si, Ni and incompatible trace elements and depleted in Ce/Yb ratio relative to CUB parental magmas. The cyclical cryptic variation within the NUC cyclic units is suggested to have formed by crystal fractionation within individual magma inputs randomly interrupted by alternating recharge of CUA or CUB magmas. The PGE-enriched horizons occur exclusively in four of the CUA-type cyclic units and show very different enrichment patterns depending on whether it is sharply overlain by another CUA unit or is capped by hybrid zone below a CUB unit (Fig. 1). The upper three PGE- enriched zones occur in CUA6, -8 and -11 (which are overlain by the hybrid units and CUB) and are hosted in websterite and/or gabbronorite units. They form 15-20m-thick disseminated zones, where the PGE grades gradually increase toward the top of the cycle and rapidly decrease into the hybrid unit. The lowermost PGEenriched zone occurs halfway through the CUA-5, which is overlain by CUA-6. It is hosted in olivine websterite and websterite cumulates, and forms 10-15m-thick zone in which PGE grades slowly decreases toward the top of the cycle (Fig. 1). 32 �The host rocks show weak to moderate hydrothermal alteration with variable proportion of mainly actinolite-tremolite, serpentinite, and chlorite. The primary PGE mineralization is commonly found in association with disseminated and blebby sulfide (0.2-2 vol%; mainly pyrrhotite, chalcopyrite, and pentlandite with minor cubanite, troilite, and cobaltite) and locally with primary hydrous phase (fluorine- bearing magnesiohornblende). In altered rocks, the primary sulfides have been partially replaced by chalcopyrite, pentlandite, sphalerite, heazlewoodite, and chalcocite. Palladium occurs either in solid solution with primary pentlandite and cobaltite or as Pd-telluride; however, Pd-bearing minerals containing mainly Pb and to lesser extent Ge, Sn, Bi, Rh, and Ag (e.g. zvyagintsevite; Pbbearing palladium telluride) occur at the margin of secondary sulfides (e.g. heazlewodite, pentlandite) and altered chrome-spinel, along pyroxene fractures, or included in serpentine and amphibole. Ptbismuthotellurides and sperrylite commonly occur associated with primary sulfides at sulfide–silicate contacts and, minor laurite is found enclosed in chrome-spinel. In general, PGE enrichments are related to increases in total S, Cu and Zr contents and a decrease in Mg:Fe ratios of pyroxenes. The mineralized zones averaging 0.358 ppm Pd+Pt ( 643 ppm Pd+Pt in 100% sulfide), Pd/Pt and Pd/Ir ratios ranging from 0.9 to 3.5 and 35 to 537, respectively, and a wide range of S/Se ratios (500-6000). The highest PGE (Pd+Pt) grades up to 11 ppm (4377 ppm in 100% sulfide) are found in serpentinized olivine websterite, which yield an average Pd/Pt ratio of 3.5 and a S/Se ratio of approximately 2.000. Our data suggest that magmatic crystal fractionation leading to early sulfur saturation and late concentration of base and precious metals in a volatile-rich fluid and possibly magma mixing can be accounted as dominant process for the different PGE occurrences throughout the CUA-type cyclic units. The PGE contents could have been enhanced during a post-magmatic stage by hydrothermal fluids which have interacted with early-formed assemblages resulting in dissolution and redistribution of PGE. Figure 1: (A) Detailed crosssection of the Sutcliffe Zone within the Layered Series showing the distribution of the PGE enriched horizons with their stratabound Pd and Pt concentrations. (B) illustrates the stratigraphic location of the PGE-rich horizons throughout the cyclic unit types from drill hole NL12101, in relation with the variation of modal rock type, the whole rock concentrations of selected PGE and trace elements (from North American Palladium Ltd database) and the average enstatite (En') content of clinopyroxene (Cpx) from (Djon et al., submitted) REFERENCES Lavigne, M.J., Michaud, M.J. 2001 Geology of North American Palladium Ltd.'s Roby Zone Deposit, Lac des Iles. Exploration and Mining Geology 10: 1-17 Djon, M. L., Miller, J.D., Olivo, G.R., and Peck, D.C. (submitted) Petrogenesis of Cyclic Units in the Northern Ultramafic Centre of the Lac des Iles Complex, Ontario, Canada: Evidence of Two Distinct Parental Magmas. Cont to Min & Petro. Sutcliffe, R.H. and Sweeny, J.M., 1986. Precambrian Geology of the Lac des Iles Complex, District of Thunder Bay, Ontario. Ontario Geological Survey, Map 3047, Geological Series-Preliminary Map, scale 1:15840. 33 �GEOCHEMISTRY OF DEEP AND SHALLOW WATER ARCHEAN BANDED IRON FORMATIONS, AND THEIR POST DEPOSITIONAL IMPLICATIONS IN THE WESTERN SUPERIOR PROVINCE, CANADA Dolega, S., and Fralick, P. Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario P7B 5E1 Canada (sdolega@lakeheadu.ca) Geochemical studies have been conducted on banded iron formations for over 40 years to attempt to better understand their synthesis, depositional environment and ocean chemistry. Previous studies have made assumptions that the geochemistry of the iron formation reflects the exact geochemistry of seawater during time of formation (Konhauser et al. 2011; Planavsky et al. 2014). These studies have not taken to account the possibilities of a post depositional mechanism, such as diagenesis and metamorphism, which can affect the mobility of elements and can alter interpretations inferred from geochemical analyses. Several locations containing Archean banded iron formations have been chosen for this study. They include the banded iron formations of the Beardmore-Geraldton Belt, Wabigoon Subprovince, Lac-St Joseph of the Uchi Subprovince, Shebandowan Greenstone Belt of the Wawa Subprovince, Weagamow – North Caribou Greenstone Belt of the Sachigo Subprovince, Red Lake Greenstone Belt of the Uchi Subprovince and the Steep Rock Group of the Wabigoon Subprovince. These locations were chosen based on their differences in age, depositional environment and differences in metamorphic grade. Geochemical analyses were conducted using an Inductively Coupled Plasma Optical Emission Spectrometer and Inductively Coupled Plasma Mass Spectrometer at Lakehead University. Preliminary results show that the Na/K ratio is greater in the magnetite-rich samples and lower in the hematite-rich samples (Figure 1). This is consistent for all the iron formations chosen for this study. The differences in the Na/K ratio cannot reflect the original geochemistry of seawater because the magnetite and hematite layers form by the same mechanism, the precipitation of ferric hydroxides. During diagenesis, magnetite will form where organic carbon reduces some of the iron (Drever, 1974). Changes in the Na/K ratio within the iron formation indicate that there has been element mobility after deposition either by diagenesis or regional metamorphism. Therefore, any geochemical conclusions inferred from banded iron formations without taking into account element mobility must be questioned. 34 �Na vs K ratios in Banded Iron Formations 140 Beardmore, Gerladton and Lac St. Joseph Magnetite 120 Fe2O3/FeO 100 Beardmore, Geraldton and Lac St. Joseph Hematite 80 60 40 Musselwhite Magnetite 20 0 0 20 40 60 80 100 120 Na2O/K2O Figure 1: Ratio scatter plot showing how Na and K concentrations vary with the mineralogy of the iron bearing phase in the iron formations studied. References Drever, J. I. (1974). Geochemical Model for the Origin of Precambrian Banded Iron Formations. Geological Society of America Bulletin, 85, pp. 1099-1106. Konhauser, K. O., Lalonde, S. V., Planavsky, N. J., Pecoits, E., Lyons, T. W., Mojzsis, S. J., Rouxel, O. J., Barley, M. E., Rosìere, C., Fralick, P. W., Kump, L. R., Bekker, A. (2011). Aerobic bacterial pyrite oxidation and acid rock drainage during the Great Oxidation Event. Nature, 478(7369), 369-373. doi:10.1038/nature10511 Planavsky, N. J., Asael, D., Hofmann, A., Reinhard, C. T., Lalonde, S. V., Knudsen, A., Wang, X., Wang, X., Ossa Ossa, F., Pecoits, E., Smith, A. J., Beukes, N. J., Bekker, A., Johnson, T. M., Konhauser, K. O., Lyons, Rouxel, O. J. (2014). Evidence for oxygenic photosynthesis half a billion years before the Great Oxidation Event. Nature Geoscience Nature Geosci, 7(4), 283286. doi:10.1038/ngeo2122 35 �Progress on Geophysical Mapping of the Northeast Iowa Intrusive Complex DRENTH, Benjamin1, ANDERSON, Raymond2, SCHULZ, Klaus3, FEINBERG, Joshua M.4, CHANDLER, Val5, and CANNON, William3 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 2 Dept. Earth and Environmental Sciences, Univ. Iowa, Iowa City, IA, 52242 3 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA, 20192-6320 4 Dept. Earth Sciences, Univ. Minnesota, 310 Pillsbury Dr. SE, Minneapolis, MN, 55455-0219 5 Minnesota Geological Survey, 2642 University Avenue W., St. Paul, MN, 55114-1032 Large amplitude gravity and magnetic highs over northeast Iowa are interpreted to reflect a buried intrusive complex composed of mafic/ultramafic rocks, the northeast Iowa Intrusive Complex (NEIIC), intruding Yavapai Province (1.8-1.72 Ga) rocks. The age of the complex is unproven, although it has been considered to be Keweenawan (~1.1 Ga). Because only four boreholes reach the complex, which is thought to be covered by 200-700 m of Paleozoic sedimentary rocks, geophysical methods are critical to developing a better understanding of the nature and mineral resource potential of the NEIIC. An initial airborne data collection campaign in the region of Decorah, IA, included high-resolution magnetic and gravity gradient data. Geologic interpretation of those data (Drenth et al., 2015) highlighted circular magnetic and gravity gradient highs thought to represent an alkaline ring complex or a mafic anorogenic intrusive complex, concepts that may be applicable to other similar anomalies over the broader NEIIC. Interpreted northeast trending dikes, thought to be Keweenawan, are also critical to understanding the age of the NEIIC and country rocks. A large (~22,700 line km) follow-up aeromagnetic-only survey was flown during the fall of 2015 over the southern half of the NEIIC, in the region of Cedar Rapids, Waterloo, and Manchester, IA. Those data are here presented publicly for the first time, in the context of other older geophysical datasets in the region. No rigorous interpretive effort has yet been made, and much ground gravity data remains to be acquired in the region, but the aeromagnetic data show many interesting features that are similar to those observed in and interpreted from the Decorah geophysical data: Large circular aeromagnetic anomalies are consistent with significant individual intrusive complexes, and numerous northeast trending linear anomalies are consistent with dikes of possible Keweenawan age. Follow-up geochronologic and paleomagnetic work may also help to clarify the geologic relations. Reference Drenth, B.J., Anderson, R.R., Schulz, K.J., Feinberg, J. M., Chandler, V.W., and Cannon, W.F., 2015, What lies beneath: Geophysical mapping of a concealed Precambrian intrusive complex along the Iowa-Minnesota border: Canadian Journal of Earth Sciences, v. 52, p. 279-293. doi: 10.1139/cjes2014-0178. 36 �Re-digitized public aeromagnetic data for parts of the west-central Upper Peninsula, Michigan DRENTH, Benjamin1, and AILES, Chad 1 1 Crustal Geophysics and Geochemistry Science Center, U.S. Geological Survey, PO Box 25046 MS 964, Denver, CO, 80225 USA The public aeromagnetic database (Daniels et al., 2009) for Michigan’s west-central Upper Peninsula (UP) is widely regarded as unsuitable for intermediate- and detailed-scale geologic mapping and mineral exploration applications. There are several limitations of the data, including being available only in a native analog format, being acquired with too wide of a line spacing and too high of a terrain clearance, and being digitized at a much lower level of detail than shown on original contour maps. This abstract describes an ongoing effort to re-digitize aeromagnetic data from original contour maps in the greatest detail possible. To date, the Michigan F (Case and Gair, 1965), E (U.S. Geological Survey 1967a, 1967b, and 1967d), and C (U.S. Geological Survey 1967c and 1967d) surveys in the west-central UP have been redigitized. Each survey was a fixed-wing total-field aeromagnetic survey flown at a nominal terrain clearance of 150 meters (Daniels et al., 2009). The Michigan C survey was flown in 1948 with 533 m line spacing along N-S lines, the F survey was flown in 1950 with 400 m line spacing along N-S lines, and the E survey was flown in 1949 with 400 m line spacing along E-W lines. After removal of an unspecified base level, the acquired data were interpolated onto contour maps with a minimum contour interval of 50 nT. A subsequent digitization effort from the contour maps followed at an unknown time, and the resulting digitized data are those publically available today from the USGS (e.g., Daniels et al., 2009). However, that digitization effort inexplicably sampled the contour maps along only every other flightline, effectively simulating surveys with double the actual line spacing. This resulted in poor geologic resolution. The same problem plagues several other vintage aeromagnetic datasets acquired in the central and western UP. We have re-digitized each of these aeromagnetic datasets in their entirety, sampling each contour from the original contour map and effectively capturing all of the available detail. The results are a much more effective representation of the region’s geology. Several known geologic features in the region are well represented, including the Marquette and Menominee iron ranges, dike swarms, blocks of Archean rocks, and the plutons hosting the Eagle and Eagle East deposits. However, the recovered data still have several major and minor limitations that must be considered by interpreters. First, the surveys were flown at too wide a line spacing and too far above the ground for detailed-scale geologic mapping and mineral exploration. Second, the minimum contour interval of 50 nT shown on the original contour maps means that more subtle anomalies and geologic details undoubtedly present in the original flightline data will never be recovered. Third, the exact terrain clearance of the magnetometer was not recorded, and in several localities may have varied significantly from the nominal 150 meter clearance. Finally, the base level removed from the magnetic data wasn’t recorded, meaning that the total-field intensity and formal total-field anomalies cannot be calculated. 37 �REFERENCES Case, J.E., and Gair, J.E., 1965. Aeromagnetic map of parts of Marquette, Dickinson, Baraga, Alger and Schoolcraft Counties, Michigan, and its geologic interpretation: U.S. Geological Survey Geophysical Investigations Map GP-467. Daniels, D.L., Kucks, R.P., Hill, P.L., and Snyder, S.L., 2009, Michigan magnetic and gravity maps and data: a website for the distribution of data: U.S. Geological Survey Data Series 411 available online at http://pubs.usgs.gov.ds/ds411. U.S. Geological Survey, 1967a, Aeromagnetic map of the Crystal Falls quadrangle and part of the Florence quadrangle, Iron County, Michigan, U.S. Geological Survey Geophysical Investigations Map GP-607. U.S. Geological Survey, 1967b, Aeromagnetic map of the Ned Lake quadrangle and part of the Witch Lake quadrangle, Iron, Baraga, and Marquette Counties, Michigan, U.S. Geological Survey Geophysical Investigations Map GP-609. U.S. Geological Survey, 1967c, Aeromagnetic map of parts of the Ralph and Norway quadrangles, Dickinson County, Michigan, U.S. Geological Survey Geophysical Investigations Map GP-610. U.S. Geological Survey, 1967d, Aeromagnetic map of Sagola quadrangle and part of the Iron Mountain quadrangle, Dickinson, Iron, and Marquette Counties, Michigan, U.S. Geological Survey Geophysical Investigations Map GP-611. 38 �3D Geological Mapping Using Terrestrial LiDAR at Soudan Underground Mine ESSIG, Espree, MOOERS, Howard, and GRAN, Karen Department of Earth and Environmental Sciences, University of Minnesota, Duluth, MN essig008@d.umn.edu Light Detection and Ranging, or LiDAR, data technology, has the potential to revolutionize the way we visualize and interpret 3 dimensional relationships in space. Conceptualization of underground geology and mine workings can be particularly challenging. This study evaluates the use of terrestrial LiDAR to visualize the physical setting and geology at Soudan Underground Mine, which today serves as a historical State Park, physics laboratory and point of significant geological interest. Using a Faro Focus 3D laser scanner, courtesy of Karen Gran and the University of Minnesota Duluth’s Department of Earth and Environmental Sciences, 3D point clouds and associated photographs have been collected on at the 27th level at Soudan. These point clouds provide submillimeter resolution and spatially accurate representations of anything in range of the laser beam. These individual scans are then stitched together in the Faro Scene® software using a combination of spherical target, plane, and point recognition approaches with an error of less than 2 mm. Data collection, stitching and rendering of these point clouds is ongoing. The goal of this project is to create a working 3 dimensional model of the underground drift workings at Soudan using LiDAR that can serve as a platform for various applications. Geology of the Soudan Mine mapped by Thompson (2016), Peterson and Patelke (2003), and Vallowe, Thalhamer, Rhoades, and Peterson (2010) are overlain on the 3D imagery. The scope of this project uses modelling as a tool to clearly visualize the spatial relationships among mapped geologic units in 3D space. The Soudan Iron Formation is an Algoma-type banded iron formation (BIF) on the southern end of the Ely Greenstone Belt. Originally chemically deposited as magnetite with chert in association with mafic volcanics, the Soudan Iron Formation was later hydrothermally altered to rich hematite ore through metasomatism and syntectonic deformation. Interpretations and details of its genesis have caused ongoing geological debates, many of which depend on structural and resulting geochemical controls. We evaluate the use of 3D mapping as a geological interpretive tool at Soudan Mine. Pictured to the left are 3D rendered point cloud representations of the 27th level at Soudan Mine 39 �References Oliver Iron Mining Company., 1953-1954. Soudan Mine Area Worksheet. Peterson, D. M. and Patelke, R.L., 2003. Bedrock Geologic Map of the Soudan Mine Area, St. Louis County, Northeast ern Minnesota. Natural Resources Research Institute Technical Report NRRI/TR-2003/29. Scale 1:5,000. Peterson, D.M. and Patelke, R.L., 2003. Geologic Map of the 27th Level East Drift, Soudan Mine. Peterson, D.M. and Patelke, R.L., 2003. National Underground Science and Engineering Laboratory (NUSEL) Geological Site Investigation for the Soudan Mine, Northeastern Minnesota. Economic Geology Group, Natural Resources Research Institute, University of Minnesota Duluth. Peterson D.M., Rhoades, D.L., Thalhamer, E.J. and Vallowe, A.M., 2010. Surface and Subsurface Geologic Maps of the Soudan Underground Mine State Park, St. Louis County, Northeastern Minnesota. Precambrian Research Center Map Series. PRC/MAP2010-01. Scale 1:2,500. Thompson, Adam, 2015, M.S. Thesis (Advisor: Mooers). A hydrothermal model for metasomatism of neoarchean Algoma-Type banded iron formation to massive hematite ore at the Soudan Mine, NE Minnesota. 40 �Multi-stage development of breccias in the Baraboo Quartzite, Rock Springs, Wisconsin FELZAN, Michael and BJØRNERUD, Marcia Geology Department, Lawrence University, 711 E Boldt Way, Appleton, Wisconsin 54911 USA. The ca. 1650 Ma Baraboo Quartzite contains localized zones of breccia in which angular fragments of the red to purple quartzite are engulfed in a stockwork of coarse white vein quartz and subsidiary kaolinite. The time of formation of these breccias relative to the formation of the Baraboo Syncline is not well constrained because the highly competent quartzite experienced little internal deformation, and thus there are no clear cross-cutting relationships between the breccias and folding-related mesoscale structures. At two breccia localities, muscovite that appears to be in textural equilibrium with the vein quartz and kaolinite has yielded 40Ar/39Ar plateau ages of 1405 + 5 Ma and 1459 + 3 Ma (Medaris et al., 2002), broadly consistent with 40Ar/39Ar ages of 1460 to 1484 Ma obtained from muscovite that lies parallel to cleavage planes in the Seeley Slate, which overlies the Baraboo Quartzite but is only known from drill core samples (Medaris et al., 2009). Fluid inclusion analyses of the white vein quartz, combined with the equilibrium coexistence of quartz + kaolin + muscovite, constrain the conditions of breccia formation to 200-280° C and 0.5-2 kbar, or about 2-8 km depth (Medaris et al., 2011). The most extensive brecciated zone in the Baraboo Range occurs on the north limb of the syncline near the town of Rock Springs, WI, extending from the west side of Ablemans Gorge to the Rock Springs Quarry east of the Baraboo River. At this location, the brecciated zone is at least 50 m thick, >3 km long, and approximately parallel with the vertical, ENE-striking bedding. We recently obtained permission to collect samples at the Rock Springs Quarry, where the brecciated zone is well exposed on the south wall of the main pit. Here it is clear that the white vein quartz breccia in fact overprints two other distinctive types of broken and veined rock. The older of these are cm-scale bands of fragmented host quartzite and greyish vein quartz, the latter recognizable in thin section because of its significantly larger grain size and elongate grain shapes. Grain boundary textures suggest that the host rock had already been metamorphosed to quartzite at the time of this early fragmentation. Broken zircon grains occur in these early breccias, and in one thin section, two halves of what seems to have been a single original grain can be seen on opposite walls of the brecciated zone. This suggests that the initial failure along these zones occurred under transient stresses high enough to fracture zircon, a famously tough mineral. The vein quartz in these oldest breccia bands exhibits kink-like deformation lamellae, indicative of quasi-ductile deformation and temperatures just below the onset of quartz plasticity (ca. 350°C). This is close to the peak metamorphic temperature inferred from the absence of biotite in the Seeley Slate (Medaris et al. 2011). These deformed early breccias are in turn cut by distinctive black, hematite-rich veins that occur in three different geometries. The most common form is a mesh-like framework around cm-size fragments of host quartzite. In these occurrences, the hematite-rich zones coincide with areas of very fine, apparently cataclasized, quartz grains much smaller (0.1 mm) than the grain size typical of the host quartzite (>1 mm). The hematite occurs as tiny (<0.1 mm) needle-like crystals that ‘decorate’ the edges of the quartz grains. In many instances, white (late-stage) vein quartz clearly cuts across this black, hematitic cataclasite. The second mode of occurrence of the hematite-rich material is as planar zones – apparently bounded by tensile fractures – that show no contrast in grain size with the adjacent, unmineralized host quartzite. As in the cataclastic zones, hematite occurs as very fine fibrous crystals that fill interstitial spaces around and between quartz grains. The third geometric mode of the hematitic material is as selvages around quartzite fragments that are enclosed in a matrix of white (late) vein quartz. Here the black material seems to have adhered to the edges of clasts when they were separated by the emplacement of the coarse vein quartz. The textural similarities of the hematite in these three types of occurrence suggest that it formed under the same conditions, involving a fluid of distinct composition, over a limited period of time. Finally, the white vein quartz was emplaced in complex three-dimensional networks that cut through not only the two earlier types of broken/mineralized rock but also previously unfractured host quartzite. This stage of brecciation and fluid flow affected a much larger mass of rock than the earlier two and involved large dilatant strains. In representative hand specimens, the white vein quartz constitutes as much as 40% of the rock volume. The fragments of host quartzite have a wide range of sizes (mm to m scales) and morphologies, ranging from tabular to equant to highly irregular three-dimensional shapes. Crescentic and splinter-like clasts suggest energetic spallation from the walls of the veins. Although the separated clasts have a ‘jigsaw’-like appearance, it is difficult in most cases to fit the pieces back together by matching their shapes or truncated internal features. 41 �This could indicate significant translation and/or rotation of the fragments within the zone prior to the formation of the vein quartz. Fluid inclusions are very abundant in the vein quartz and typically occur in parallel planar arrays within a given crystal, pointing to episodic growth over many cycles of fluid infiltration. In some cases, the vein quartz occurs as zoned, phantom, and terminated prismatic crystals up to 5 cm long, growing into open space or into masses of kaolin. In thin section, the vein quartz shows sweeping undulose extinction, indicating that it too experienced significant deviatoric stress after its crystallization (van Lankvelt and Bjørnerud, 2010). Collectively, these observations suggest that brecciation and veining at Rock Springs occurred in several distinct stages, involving different modes of deformation and fluids of varying composition, over a protracted period of time. First, an early period of brittle failure, apparently at very high – possibly co-seismic -- stresses, fractured the host quartzite and even cracked detrital zircon grains. An early generation of quartz veins were precipitated into the fractures, and these veins then became the locus of quasi-ductile deformation at temperatures close to the inferred metamorphic peak of 350°. Next, a period a cataclasis along anastomosing shear fractures created cm-scale zones of finely comminuted quartz grains. These cataclasized zones acted as conduits for an iron-rich fluid that infiltrated the rock and was overpressured enough to create additional pathways for itself through hydrofracturing. Next, a large volume of rock, including but extending beyond the previously fractured area, was pervasively broken into fragments whose varied and complex shapes suggest violent, implosive failure. Silicic fluids at temperatures of 200-280°C then flowed repeatedly through this brecciated rock, eventually enclosing the clasts in a matrix of white, fluid inclusion-rich quartz and lesser kaolinite. Finally, continuing stresses caused slight deformation of the vein quartz. The temporal relationships between these inferred events and the regional folding is not clear. The parallelism of the Rock Springs breccia zone with bedding in the Baraboo Quartzite could indicate that flexural slip during the main folding event played a role in the formation of the breccia. The Rock Springs breccia occurs within an especially massive stratigraphic interval of the quartzite, with little interbedded phyllite (Van Hise Rock being the notable exception). Perhaps at some point in the folding process, bending stresses became so great within the stiff quartzite beam that it failed along a bedding-parallel zone. Failure then allowed overpressured fluids to rush into the fragmented rock, leading to further fracturing and successive stages of veining. The source of the fluids, and in particular the iron-rich fluid that left the unusual hematite veins, is not known. The Baraboo Quartzite as a whole – one of Earth’s earliest ‘red bed’ sequences -- owes its color to oxidized iron, so the iron could have been scavenged from the host rock itself. Alternatively, a little-known iron formation, the Freedom Formation, lies stratigraphically above the Baraboo Quartzite and Seeley Slate and occurs (in the subsurface) in the core of the Baraboo syncline (Roe and Bjørnerud, 2012). If the formation of the hematite-rich areas post-dated the folding and rotation of bedding in the north limb to its present vertical orientation, fluid flow from the Freedom Formation might have been lateral rather than downward. Another fundamental question is whether all the breccias in the Baraboo district (and other Baraboo-interval quartzites) are of the same age and origin or whether they reflect only local stresses and fluid flow. The 14001460 Ma ages of muscovite from other breccia localities as well as the Seeley Slate suggest that deformation and metamorphism at Baraboo were linked to Wolf River Batholith magmatism. In any case, it seems the Baraboo breccias have a more complex story to tell than previously thought and may provide a new window into the processes that contributed to the formation of the still-enigmatic Baraboo Mountains. References cited Medaris L.G., Jr., Singer B., Dott, R.H., Jr., Naymark, A., Johnson, C., and Schott, R., and 2003. Late Paleoproterozoic climate, tectonics, and metamorphism in the southern Lake Superior region and Proto-North America: Evidence from Baraboo-Interval quartzites. Journal of Geology 111, p. 243-257. Medaris, L.G., Jr., Jicha, B., Dott, R.H., Jr. and Singer B., 2009. A 1465 Ma 40Ar/39Ar age for the Seeley Slate: Implications for metamorphism in the Baraboo Range, Wisconsin. Institute on Lake Superior Geology Program with Abstracts 55, p. 59-60. Medaris, L.G., Jr., Dott, R.H., Jr., Craddock, J., and Marshak, S., 2011. The Baraboo District: A North American Classic. Geological Society of America Field Guide 24, p. 63-82. Roe, C. and Bjørnerud, M., 2012. The ca. 1650 Ma Freedom Formation: A late iron formation in the Lake Superior region. Institute on Lake Superior Geology Program with Abstracts 58, p. 73-74. Trepmann, C.A., Stöckhert, B., 2002. Cataclastic deformation of garnet: a record of synseismic loading and postseismic creep. Journal of Structural Geology 24, p. 1845–1856. van Lankvelt, A. and Bjørnerud, M., 2010. Revisiting the Baraboo breccias, Institute on Lake Superior Geology Program with Abstracts 56, p. 67-68. 42 �Geophysical Imaging of Layered Mafic Complexes and Relation to Platinum Group Element Exploration FINN, Carol A., ZIENTEK, Michael, BEDROSIAN, Paul, BLOSS, Benjamin, BURTON, Bethany, PETERSEN, Dean and PARKS, Heather Layered ultramafic to mafic intrusions such as the Duluth Complex are economically important because they can host magmatic ore deposits containing economic concentrations of nickel, copper, titanium, vanadium and platinum-group elements (PGE’s). Because most new discoveries lie under cover, modern exploration for PGE’s relies heavily on understanding the geophysical signature of the entire magmatic system in which buried deposits form. Geophysical surveys and methods of analysis provide higher resolution views of the environments in which PGE’s develop than were previously available. Combined analysis of potential field and electromagnetic data can provide constraints on the volume of the intrusion, its extent under cover, possible locations of sulfide mineralization. These data can also effectively map smallscale (~3 meters) structures within layered intrusions that host PGE-bearing magmatic ore deposits. Previous interpretations of aeromagnetic and gravity data helped map the geology of the Duluth Complex under cover and confirmed that the complex consists of multiple intrusions (Chandler, 1990; Chandler et al., 1998; Miller et al., 2001). The advent of new filtering, imaging, and modeling techniques will allow us to refine these earlier interpretations, and new high resolution magnetic and electromagnetic (EM) data provide detailed views of the Kawishiwi intrusion. For example, the tilt derivative of the aeromagnetic data significantly enhances small scale features, such that linear aeromagnetic anomalies in the Poplar Lake and Misquah Hills intrusions, Duluth Complex, possibly related to layering, can be observed. In other intrusions, like the Stillwater Complex, the linear aeromagnetic anomalies primarily relate to boundaries between major stratigraphic units and olivine-bearing rock layers altered to a mixture of serpentine and magnetite. Gravity highs characterize the exposed and interpreted buried extent of several intrusive components of the Duluth Complex (Chandler, 1990; Chandler and Ferderer, 1989; Chandler and Lively, 1998). Two dimensional modeling of the Duluth Complex indicate that the Complex thickens to the southeast (Chandler and Ferderer, 1989; Chandler and Lively, 1998). Similar to the Bushveld (Cole et al., 2013; Finn et al., 2015a) and Stillwater Complex layered mafic intrusions (Finn et al., 2015b), modeling of the gravity data constrained by other geophysical data, may help constrain the 3D extent of the Duluth Complex. Existing gravity data over the Kawishiwi intrusion indicate that the Cu-Ni-PGE deposits along its base align along a gravity gradient indicating a density contrast between basement and mafic intrusive rocks (Condor Consulting Company, 2011, internal report). Stochastic inversion of the EM data aids in enhancing low resistivity features that could indicate undiscovered deposits. Existing MT data images the configuration of low resistivity sulfides within the Animikie basin adjacent to the Duluth Complex (Bedrosian, 2016), important because the degree of reaction and assimilation of the Animikie rocks with the mafic magmas resulted in a variety of mineralization zones. In addition, the Animikie basin may have influenced the extent of the Duluth Complex as the Transvaal basin did for the Bushveld intrusion (Finn et al., 2015a). 43 �REFERENCES Bedrosian, P.A., 2016., Making it and breaking it in the Midwest: Continental assembly and rifting from modeling of EarthScope magnetotelluric data. Precambrian Research, doi:10.1016/j.precamres.2016.03.009. Chandler, V.W. and Ferderer, R.J., 1989. Copper-nickel mineralization of the Duluth Complex, Minnesota; a gravity and magnetic perspective. Economic Geology, 84(6), pp.1690-1696. Chandler, V.W., 1990. Geologic interpretation of gravity and magnetic data over the central part of the Duluth Complex, northeastern Minnesota. Economic Geology, 85(4), pp.816-829. Chandler, V. W. and R. S. Lively, 1998, Gravity and magnetic modeling of the Duluth Complex in the Allen 7.5minute quadrangle, St. Louis County, Minnesota, University of Minnesota, Minnesota Geological Survey Miscellaneous Map Series, Map M-90. Cole, Janine, Susan J. Webb and Carol A. Finn, 2014, Reassessing geophysical models of the Bushveld Complex - have we come full circle?: Journal of South African Earth Sciences, v. 92, p. 97-118 http://dx.doi.org/10.1016/j.jafrearsci.2014.01.012 1464-343X. Finn, Carol A., Paul Bedrosian, Janine Cole, Tshepo David Khoza and Susan J. Webb, 2015a, Mapping the extent of the Northern Lobe of the Bushveld layered mafic intrusion from geophysical data: Precambrian Research, v. 268, 279-294, doi:10.1016/j.precamres.2015.07.003. Finn, CA, Michael Zientek, Paul Bedrosian, Janine Cole, Susan Webb and Heather Parks, 2015b, Geophysical exploration for PGE deposits in layered mafic intrusions, Abstract NS42A-01 presented at 2015 Fall Meeting, AGU, San Francisco, CA. 44 �Sedimentology of a Pre-Vegetation Floodplain Assemblage: the Mesoproterozoic Hele Member of the Sibley Group, Ontario FRALICK, Philip and ZANIEWSKI1, Kamil Department of Geology, Lakehead University, Thunder Bay, ON philip.fralick@lakeheadu.ca Descriptions of fluvial systems operating prior to significant terrestrial macrophyte vegetation have concentrated on assessing the impact of plant evolution on channel style; this is probably in part due to the scarcity of well developed floodplain successions in fluvial assemblages of these ages. This study describes wet, pre-vegetation floodplain deposits and processes observed and inferred form a continuous succession of drill core through an extensive, 1.4 Ga, delta-top channelfloodplain assemblage forming a portion of the Sibley Group, Ontario, Canada. Sub-aerial deposits are dominated by flaser, wavy and lenticular bedded, find-grained sandstones, siltstones, and mudstones, with abundant small mudstone rip-up clasts. Soft-sediment deformation of these units is ubiquitous, with loading and injection features being the most prominent. Thicker medium-grained sandstone beds, representing crevasse splays, commonly have poorly developed protosols in their upper portions. Well-laminated sediments with wave ripples, and only rarely containing rip-up clasts and soft sediment deformation, were deposited in floodplain ponds. These deposits differ from post-vegetation floodplain sediments in having: (i) better preservation of layering without rootlet bioturbation; (ii) dominance of rippled sand on the floodplain, probably due to lack of vegetation-induced baffling and thus higher velocities of overbank flows away from the levees; (iii) large-scale generation of small, intraformational clasts caused by intense drying of the thin upper layer of sediment due to the lack of shade; (iv) desiccation crack fills consisting of peds and locally derived intraclasts, probably transported by overland flow during rainfall events that did not result in sediment delivery from the main channels; (v) ubiquitous soft-sediment deformation features in sub-aerial deposits denoting a groundwater table very close to the surface; and (vi) well-laminated, commonly oxidized sediment that accumulated in floodplain ponds attributable to low levels of organic loading, though some green pond sediments indicate that microbial microflora, and probably microfauna, did exist in these ponds. These attributes are the direct result of the lack of macrophyte vegetation, and produce floodplain assemblages that are distinctly different from those currently forming in similar climatic settings. 45 �Provenance and tectonic evolution recorded by successor basins in the Abitibi-Wawa terrane: Insights from new U-Pb LA-ICP-MS analyses of detrital zircon FRIEMAN, Ben M.1, KUIPER, Yvette D. 1, KELLY, Nigel M.2, MONECKE, Thomas1 Dept. of Geology & Geological Engineering, Colorado School of Mines, 1516 Illinois St., Golden, CO, 80401; 2Dept. of Geological Sciences, University of Colorado at Boulder, 2200 Colorado Ave., Boulder, CO, 80309 Understanding the detrital zircon provenance of syntectonic successor basins can provide insight into the processes related to amalgamation of the Superior Province. The aim of this study was to investigate the provenance of successor basins in the southern Abitibi subprovince (SAS), which comprises the southeastern most extent of the Superior Province. To establish the provenance of sedimentary rocks in successor basins of the SAS we have utilized statistically robust U-Pb dates of detrital zircon grains obtained by laser ablation-inductively coupled plasma-mass spectrometry (LA-ICP-MS). The results were compared to detrital zircon data from successor basins in the Wawa subprovince from Lodge et al. (2013) to establish the characteristics of provenance throughout the Abitibi-Wawa terrane. The Abitibi-Wawa terrane (Stott et al., 2010) is one of the largest, best-exposed, well-studied, and economically well-endowed Archean greenstone belts in the world. The Abitibi and Wawa subprovinces occur to the east and west, respectively, of the Paleoproterozoic Kapuskasing uplift in Ontario. In recent years, extensive high-resolution, U-Pb zircon geochronology coupled with new and existing mapping has revealed a characteristic lithotectonic progression for the development of the Abitibi-Wawa terrane (e.g., Ayer et al., 2002). The research, primarily focused on the SAS, indicates that construction of the terrane was dominated by bimodal volcanism in a subaqueous environment from ~2750 to ~2695 Ma. The termination of submarine volcanism at ~2695 Ma was marked by the onset of regional tectonism, which progressively localized along major regional deformation zones and resulted in the formation of orogenic gold deposits. Synchronous with tectonism was the development of thick sedimentary successions in successor basins. In the SAS, two distinct types of successor basins are recognized based on sedimentological, structural, and geochronological constraints: the ~2695-2685 Ma, submarine, turbidite-dominated Porcupine assemblage and the ~2680-2670 Ma, largely subaerial, coarse clastic-dominated Timiskaming assemblage. The Timiskaming assemblage is typically fault- and/or unconformitybounded, represents a critical temporal and structural marker unit, and commonly hosts orogenic gold deposits. To establish the provenance of successor basins in the SAS a representative set of sixteen samples was collected from across the subprovince. Sampling targeted well-constrained localities with a variety of stratigraphic relationships, spanning many of the major structural zones and mining camps in the SAS. To evaluate the detrital zircon date distributions of samples from these successor basins, probability density (PDF) and cumulative distribution (CDF) functions were calculated. These data provide a robust characterization of the provenance of Porcupine and Timiskaming assemblage sedimentary rocks. The PDFs for all of the samples display similar principal populations of detrital zircon grain dates that are dominated by ~28002650 Ma results. However, primary differences between Porcupine and Timiskaming assemblage detrital zircon populations were indicated by differing proportions of older, pre-2800 Ma dates. The CDF calculations indicate that zircon grains with pre-2800 Ma dates comprise ≤5% of the Porcupine assemblage data and ~10-15% of the Timiskaming assemblage data. Similar to SAS 1 46 �samples, Wawa detrital zircon samples (Lodge et al., 2013) are predominately composed of grains with 2800-2650 Ma dates. Furthermore, the Wawa data displays a similar proportion of pre-2800 Ma dates to Porcupine assemblage samples (i.e., <5%), but also contain a relatively high proportion of ~2650 Ma and younger zircon grains. Lodge et al. (2013) interpreted the young dates as a result of metamorphism, and therefore they may not reflect primary depositional differences in provenance between the southern Abitibi and Wawa subprovinces. The Wawa successor basins display sedimentological characteristics that are ‘Timiskaming like’ (i.e., dominated by coarse clastic sedimentary rocks), although the maximum age of deposition is locally well-constrained by detrital zircon geochronology to be ~2690 Ma (Lodge et al., 2013). Consequently, despite broad similarities to the Timiskaming assemblage in the SAS, the Wawa successor basins appear to be the same age as, and contain detrital zircon provenance that is most comparable to, the Porcupine assemblage in the SAS. In general, the provenance of SAS detrital zircon samples indicates that the successor basins are dominantly composed of local detritus, consistent with geochemical and sedimentological studies (e.g., Feng and Kerrich, 1990; Cocoran and Mueller, 2007). However, the occurrence of pre-2800 Ma detrital zircon grains may be indicative of provenance from adjacent subprovinces, as the SAS does not contain rocks in this age range (e.g., Ayer et al., 2002). Therefore, we interpret these data as an indicator of exhumation and erosion of an emergent hinterland to the north (i.e., the Winnipeg River, Marmion, and Opatica subprovinces). Potential sources for pre2800 Ma grains also occur to the south in the Minnesota River Valley terrane. However, these rocks also contain zircon grains with ~3500-3300 Ma dates (Bickford et al., 2006), which are not observed in the SAS data. In the SAS, it is well-established that the Porcupine and Timiskaming assemblages were deposited at different times based, in part, on the ages of the youngest detrital zircon grains they contain (e.g., Ayer et al., 2002). Our data additionally indicates that Timiskaming assemblage rocks contain a higher abundance of zircon grains with pre-2800 Ma dates relative to Porcupine assemblage rocks. This suggests that detritus from the hinterland was more prevalent during the later stages of tectonism at ~2680-2670 Ma, supporting models that invoke north to south propagation of the tectonic front. REFERENCES Ayer, J., Amelin, Y, Corfu, F., Kamo, S., Ketchum, J., Kwok, K., and Trowell, N., 2002. Evolution of the southern Abitibi greenstone belt based on U-Pb geochronology: Autochthonous volcanic construction followed by plutonism, regional deformation and sedimentation. Precambrian Research, v. 115, p. 63–95. Bickford, M.E., Wooden, J.L., and Bauer, R.L., 2006. SHRIMP study of zircons from Early Archean rocks in the Minnesota River Valley: Implications for the tectonic history of the Superior Province. GSA Bulletin, v. 118, p. 94–108. Corcoran, P.L., and Mueller, W.U., 2007. Time-transgressive Archean unconformities underlying molasse basin-fill successions of dissected oceanic arcs, Superior Province, Canada. Journal of Geology, v. 115, p. 655–674. Feng, R., and Kerrich, R., 1990. Geochemistry of fine-grained clastic sediments in the Archean Abitibi greenstone belt, Canada: Implications for provenance and tectonic setting. Geochimica et Cosmochimica Acta, v. 54, p. 1061–1081. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., Jirsa, M.A., and Hamilton, M.A., 2013. New U-Pb geochronology from the Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa Subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province. Precambrian Research, v. 235, p. 264–277. Stott, G.M., Corkery, M.T., Percival, J.A., Simard, M., and Goutier, J., 2010. A revised terrane subdivision of the Superior Province. Ontario Geological Survey Open File Report 6260, p. 20-21–20-10. 47 �A comparative study of mafic and felsic lithologies from the Borden Belt and adjacent greenstone belts in the Wawa-Abitibi Terrane GAMELIN, G., Stinson, V.R., Pan, Yuanming, and Nadeau, M. Department of Geology, University of Saskatchewan, 114 Science Place, Saskatoon, SK, Canada S7N 5E2, glecy.gamelin@usask.ca The Borden Belt in the Kapuskasing Structural Zone of the Superior Province of Canada contains schists and gneisses that reached upper amphibolite to granulite facies of metamorphism. Whole rock geochemistry and trace element analyses indicate igneous protoliths of sub-alkaline affinity, ranging from ultramafic to felsic compositions. Igneous tectonic geochemical signatures are preserved which indicate shallow mantle melting environment and crustal input, characteristic to arc-subduction zone magmatism. The Wawa-Abitibi Terrane can be defined as a tectonostratigraphic terrane, by comparing trace element geochemistry and field relationships of the Borden belt, Michipicoten Greenstone Belt, and the southern Abitibi Greenstone Belt in Ontario. All three belts share similar geochemical signatures which indicate lithologies associated with the formation and progression of a mature volcanic arc in a shallow marine setting within the Wawa-Abitibi Terrane. 48 �The Geology and Geochemistry of the Laird Lake Property, Red Lake Greenstone Belt, Ontario GÉLINAS, Brigitte1 and HOLLINGS, Pete1 1 Department of Geology, Lakehead University, Thunder Bay, Ontario The Laird Lake property is situated on what has been previously interpreted as the angular unconformity between the Balmer (2.99 to 2.96 Ga) and the Confederation (2.74 to 2.73 Ga) assemblages on the south-western end of the Red Lake greenstone belt, northwestern Ontario (Corfu and Wallace, 1986; Corfu and Andrews, 1987; Stott and Corfu, 1991). The Balmer assemblage is characterized by mafic tholeiites, ultramafic flows, local banded-iron formations and minor felsic volcanic rocks, whereas the Confederation comprises mafic to felsic volcanic rocks with a calc-alkalic affinity (Parker, 2000; Sanborn-Barrie et al., 2001, 2004). Contrary to previous interpretations, the Laird Lake property displays metamorphic assemblages up to amphibolite grade facies which are attributed to regional metamorphism, rather than being related to the metamorphic aureole of the Killala-Baird batholith (2704 ± 1.5 Ga) north of the field area (Sanborn-Barrie et al., 2004; Dubé et al. 2000). Multiple gold occurrences are found on the property and generally occur within 200 m of the interpreted unconformity. The occurrences are hosted by multiple rock types with some of the highest recorded values from grab samples of quartz veins (101.90 g/t), banded-iron formation (35.34 g/t), altered mafic volcanic rocks (38.57 g/t) and quartz-feldspar porphyritic crystal-tuff (>0.1 g/t; LeBlanc, 2015). This study will develop an integrated model for the geology and gold mineralization of the Laird Lake area by 1) characterizing the primary host rocks and 2) investigating the nature, timing and origin of the Au-mineralization. Mapping and sampling of the various rock types in the field area at surface and within five drill cores that intersect the gold bearing zone will provide the geological constraints for this study. Petrographic and SEM analysis will characterize the textures and mineralogy of the rocks, whereas whole rock geochemistry will be used to characterize their composition in order to differentiate the Balmer and Confederation assemblages. Quartz vein samples (barren and mineralized) will be analyzed for oxygen isotopes and fluid inclusion work will be conducted if possible. Additionally, Sm-Nd isotope analyses will be conducted on carefully selected samples in order to acquire information on the source of the host rocks to help establish the tectonic history of the area, and last, a geochronology sample of a felsic volcanic unit within the Balmer assemblage will be dated by U-Pb zircon analysis. Previous models suggest that the Balmer and Confederation assemblages cannot be distinguished in the field, however, observations made over the field season indicate that the two assemblages show clear differences. The Balmer comprises a fine-grained, aphyric mafic metavolcanic, locally pillowed, with various amounts of biotite and carbonate alteration and banded-iron formation. The Confederation was observed to have phenocrystic (feldspar and/or amphiboles) and rare aphyric mafic metavolcanic rocks intercalated with intermediate and felsic metavolcanic rocks. The mafic to felsic metavolcanic rocks were not observed to have pillows and display much weaker alteration than the Balmer. Both Balmer and Confederation rocks show a dominant east-trending fabric with increasing foliation observed when approaching the unconformity. Detailed mapping at the “Gold Bearing Zone” trench and whole-rock geochemistry indicates that the outcrop is part of the Balmer assemblage with aphyric mafic metavolcanic tholeiites and banded iron formations. Known gold occurrences at the trench are localized within a banded iron formation cut by a gold-bearing shear. This relationship would suggest a 49 �sulphidation reaction in order to precipitate the gold, samples of which have yielded values up to 35.34 g/t over 2m (LeBlanc, 2015). Primitive mantle-normalized trace element profiles for Balmer and Confederation volcanic units show distinct trends that support the subdivision of assemblages within the field. As expected, the Balmer mafic volcanic rocks show a tholeiitic signature, with two distinct trends; trend 1 has a low Th/Nb ratio and a weak negative Ti anomaly whereas trend 2 has a more LREE depleted profile with a distinctly higher Th/Nb ratio and a stronger negative Ti anomaly. The Balmer ultramafic volcanic units equally show two trends; trend 1 is enriched in LREE while trend 2 is depleted in LREE. Both show high Th/Nb ratios with flat HREE. The Confederation assemblage volcanic rocks show calc-alkaline signatures with enriched LREE, flat HREE, high Th/Nb ratios and strong negative Ti anomalies, all indicative of rocks formed above a subduction zone. Figure 1: Primitive mantle-normalized trace element profiles for Balmer and Confederation assemblage metavolcanic rocks (normalization factors are after McDonough et al. 1992). (A) Balmer assemblage. (B) Confederation assemblage. REFERENCES Corfu, F. and Andrews, A.J. 1987. Geochronological constraints on the timing of magmatism, deformation, and gold mineralization in the Red Lake greenstone belt, northwestern Ontario; Canadian Journal of Earth Sciences, v.24, p.13021320. Corfu, F. and Wallace, H. 1986. U–Pb zircon ages for magmatism in the Red Lake greenstone belt, northwestern Ontario; Canadian Journal of Earth Sciences, v.23, p.27-42. Dubé, B., Balmer, W., Sanborn-Barrie, M., Skulski, T. and Parker, J. 2000. A preliminary report on amphibolite-facies, disseminated-replacement-style mineralization at the Madsen gold mine, Red Lake, Ontario; Geological Survey of Canada, Current Research 2000-C17, 12p. LeBlanc, J. 2015. Project Summary Report for the Laird Lake Gold Project, Red Lake, Ontario, Canada. Unpublished Company Report for Bounty Gold Corp. 50p. McDonough, W. F., S. Sun, A. E. Ringwood, E. Jagoutz, and A. W. Hofmann (1992), K, Rb, and Cs in the earth and moon and the evolution of the Earth mantle, Geochim. Cosmochim Acta, S. R. Taylor Symposium volume, 1001-1012 Parker, J.R. 2000. Gold mineralization and wall rock in the Red Lake greenstone belt: A regional perspective; in Summary of Field Work and Other Activities, 2000, Ontario Geological Survey, Open File Report 6032, p.22-1 to 22-28. Sanborn-Barrie, M., Rogers, N., Skulski, T., Parker, J., McNicoll, V. and Devaney, J. 2004. Geology and tectonostratigraphic assemblages, east Uchi Subprovince, Red Lake and Birch–Uchi belts, Ontario; Geological Survey of Canada, Open File 4256; Ontario Geological Survey, Preliminary Map P.3460, scale 1:250 000. Sanborn-Barrie, M., Skulski, T. and Parker, J. 2001. Three hundred million years of tectonic history recorded by the Red Lake greenstone belt, Ontario; Geological Survey of Canada, Open File 4594, 30p. Stott, G.M. and Corfu, F. 1991. Uchi Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.145-238. 50 �Evidence in the Eastern Canadian Shield of regular fault patterns of crustal origin for the loci of some mineral deposits and late-stage intrusive events Glass, F., Quebec Qualitative analysis of aeromagnetic data and examination of other data sets (geomorphology, mapped outcrops ) have been used to determine the presence of interpreted lineaments at known mineral occurrences or late-stage intrusive events ( mapped or aeromagnetic observation ). These lineaments are thought to represent ductile or brittle fault traces and to have origin within the crust. The prevalence of particular fault-trace orientations is consistently regular over a large area of the Canadian Shield. The age of the dated mineral deposits or intrusive events associated with particular fault traces or fault loci vary from 2,7 Ga to 1,1 Ga. Within a 1 000 km map block area or greater, the inter-relationship of certain mineral deposits and late-stage intrusive events appears to consistently correlate to specific geographic orientations. Within the boundary of a smaller block ( 100 km or smaller in size ) anywhere inside the larger block, the inter-relationship between local deposits or nearby late-stage intrusives, coincides with some of those lineaments evident at the regional scale. In addition, there are also different orientations, not necessarily observed with extensive continuity, unequivocally present. These features of local extent are, however, oddly present, with near exact correlation to specific regional orientations, over a similarly large area of the Canadian Shield. In order to explain the origin of the above observations requires likely two different depths for the far and local stress fields or, more plausibly, a single cause for the stress field possibly acting on two crustal layers, the near-surface one delaminated from the deeper crust below. Prior to at least 1,1 Ga, the Canadian Shield appears to have behaved as a single rigid crustal block. ( The Mount Polley alkaline complex has a N27W faulted boundary to the west; the Labrador Trough to the M7 earthquake ( 1929 ) in the St-Lawrence ( Quebec Embayment Structure ) is at N27W ). 10 slides have been prepared showing repetitions of a few common lineament orientations ( all flat-map representations ) with a final slide illustrating the necessity to use polar projection ( spherical maps ) over large regional distances. All 11 slides are presented on one poster. 51 �The Cu/Pd diagram and metal/sulfur variation as an exploration tool: Examples from the Coldwell Alkaline Complex, Ontario GOOD, David J.1, LINNEN, Robert L.1 and SAMSON, Iain M.2 1 2 Department of Earth Sciences, Western University, London, ON, Canada, N5A 5B7, dgood3@uwo.ca Department of Earth & Environmental Sciences, University of Windsor, Windsor, ON, Canada N9B 3P4 Mapping out variations in Cu/Pd, Cu/S and Pd/S through a Ni-Cu-PGE sulphide deposit provides important insight to the mechanisms acting during deposit formation. For instance, the Cu/Pd diagram has been used: to illustrate PGE depletion by previous sulphide segregation (Barnes et al., 1993); to estimate R-factors required to form deposits in the Duluth Complex (Thériault and Barnes, 1998); and as an indicator of extreme PGE enrichment (Barnes and Ripley, 2016). We propose here that patterns and features shown by a large data set on the Cu/Pd diagram can provide insight into the cumulative history of the deposit, particularly in a magmaconduit setting where intrusive bodies formed by build up of multiple sills. This study compares spatial and compositional relationships for 15,000 assays from 11 mineralized zones within the Marathon, Geordie Lake, and Area 41 deposits and the Four Dams and Redstone occurrences located within the 1.1 Ga Midcontinent Rift related Coldwell Alkaline Complex. The deposits are proposed to be co-genetic and, except for Geordie Lake, exhibit common petrologic features indicating formation in a magma conduit setting at or just above the contact between mafic meta-volcanic rocks and the Archean basement (Good et al., 2015). The Geordie Lake deposit is located closer to the middle of the complex and was cut by syenitic rocks during Cycle I of the Coldwell intrusive event. The host rocks for the deposits consist of some combination of ophitic gabbro and pegmatite; apatitic clinopyroxenite; or augite troctolite. Mineralization consists predominantly of disseminated assemblages of chalcopyrite ± pyrrhotite ± bornite. Trends for metal abundances across mineralized intervals in all zones, except for the W Horizon at Marathon, show a correlation between Cu, Pd, Pt and S. But metal tenor (metal/S) and/or Cu/Pd in these sections commonly exhibit saw-tooth patterns that change gradually up through intrusions with sharp steps occurring within and between individual zones. Mineralized intervals that exhibit increasing Cu, S and Pd with decreasing Cu/Pd are consistent with models for accumulation of sulphides from magma with similar R-factor. Step-like changes in Cu/Pd or Cu/S within a zone are interpreted to represent contacts between individual pulses of sulphidebearing magma with an inherent R-factor attribute. Taken together, the range of Cu/Pd for all samples in a mineralized zone represents the cumulative history of individual magma pulses. Contouring the data for each zone by point density defines a characteristic shape and trend line for the zone, the slope of which is dependent on the total range of R-factors for the accumulated magma pulses. Shallow-dipping trend lines, such as that observed for the Geordie Lake deposit and the Marathon footwall zone, indicate a simple intrusive history with magma pulses having a limited range of R-factors. Steeply negative trend lines extending from the Pd-depleted to the Pd-enriched fields represent magma pulses with a very broad range of R-factors and indicate a very dynamic intrusive history, as is the case for data from the W horizon and the top unit at Area 41. The latter relationship shows there is a spatial, and possibly a temporal, relationship between the most copper-rich and Pd-rich zones at Marathon and Area 41. 52 �REFERENCES Barnes, S-J, Couture, J-F, Sawyer, EW, & Bouchaib, C., 1993. Nickel-copper sulfide occurrences in BelleterreAngliers belt of the Pontiac sub-province and the use of Cu/Pd in interpreting platinum-group element distributions. Econ Geol 88: 1402–1418. Barnes, S.-J. & Ripley, E.M., 2016. Highly Siderophile and Strongly Chalcophile Elements in Magmatic Ore Deposits, Reviews in Mineralogy & Geochemistry 81: 725-774. Good, D.J., Epstein, R., McLean, K., Linnen, R.L. & Samson, I.M., 2015, Evolution of the Main Zone at the Marathon Cu-PGE sulfide deposit, Midcontinent Rift, Canada: spatial relationships in a magma conduit setting. Economic Geology 110, p. 983-1008. Thériault, R. D., Barnes S.-J., and Severson M. J., 1997, The influence of country-rock assimilation and silicate to sulfide ratios (R factor) on the genesis of the Dunka Road Cu - Ni - platinum-group element deposit, Duluth Complex, Minnesota, Can. J. Earth Sci. V. 34, p. 375-389. 53 �Progress on 3D modeling of the Midcontinent Rift System in the western Lake Superior region and an isopach map of the Oronto Group GRAUCH, V.J.S., POWERS, Michael H., and ANDERSON, Eric D. U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 Efforts are underway to develop a three-dimensional (3D) model of the structure and configuration of sedimentary and volcanic basins related to the Midcontinent Rift System (MRS) in the western Lake Superior region (Fig. 1). The current efforts build on a 3D model previously developed in the 1990s (Allen et al., 1997) and concepts developed by many other previous workers. Technical advances in modeling capabilities, expanded and improved geophysical data coverage, and renewed interest in the mineral resources of the MRS provide the motivation for new attempts at 3D digital modeling of the MRS. An improved 3D model of the area helps visualize geologic relations, provides mechanisms to test hypotheses about tectonic history and the spatial distribution of mineralization, and helps identify areas where more detailed analysis is required. Our 3D model is regional and intended to show broad variations in geology. Only three generalized, MRS-related geologic packages are represented, following Allen et al. (1997). The packages, from oldest to youngest, are 1) undivided Keweenawan plutonic and volcanic rocks, 2) Oronto Group sedimentary rocks, and 3) Bayfield Group and equivalent sedimentary rocks. A fourth package represents undivided pre-rift rocks (basement). In addition, three major fault systems are modeled: the Douglas, Lake Owen, and Keweenaw fault systems (Fig. 1). The modeling strategy involves digitizing the bases of the geologic packages, fault locations, and general orientation data along 2D sections and from geologic maps. The 3D modeling software then connects the digitized points into surfaces and volumes in 3D space, using simple geologic rules for stratigraphic and onlap relations and for the lateral extents of the influences of faults. Digitized points from the previous 3D model (Allen et al., 1997) serve as a guide, but we are re-evaluating geophysical data and making sure that updated geologic concepts have been incorporated. In particular, we are revisiting the analyses of available seismic-reflection sections with modern software and developing new gravity and magnetic models to substantiate the interpretations. For example, we have applied modern reprocessing techniques to proprietary seismic-reflection data in the Bayfield Peninsula to document previously unpublished results and provide a better image of the geology. We are also developing rootmean-square velocity maps for analog seismic-reflection sections to establish depths to strong reflections. Some of these new depth interpretations are revealing discrepancies between interpreted seismic sections where they cross each other, which merit further work. An isopach map of Oronto Group thickness constructed from the 3D model (Fig. 2) illustrates the tremendous thickness of sediments that were deposited within present-day, western Lake Superior during post-magmatic subsidence and before Bayfield basin formation. The thickest parts of the basin are generally located in between White’s and Grand Marais ridges, respectively. A significant observation is the broad asymmetry of the Oronto basin in western Lake Superior, with the thickest part of the basin (>12 km) near the south shore adjacent to the Porcupine Mountains. The asymmetry is present in similar, previous maps (e.g., Allen et al., 1997), but is more pronounced in our map. The broad, southeastdeepening asymmetry next to the Porcupine Mountains contrasts with isopach thicknesses and previous workers' interpretations of GLIMPCE seismic line C (Fig. 2). Along line C, the thickest Oronto Group is concentrated within a narrow (~25-km) zone adjacent to the south shore. The local thickening next to the Douglas fault at the southeast side of White's Ridge is present in both our and Allen's model, derived independently from informal observations of proprietary seismic data. Reference Allen, D. A., Hinze, W. J., Dickas, A. B., and Mudrey, M. G., Jr., 1997, Integrated geophysical modeling of the North American Midcontinent Rift System: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota, in Ojakangas, R. W., Dickas, A. B., and Green, J.C., ed., Middle Proterozoic to Cambrian Rifting, Central North America: Geological Society of America Special Paper 312, p. 47-72. 54 �Figure 1: Generalized rock units of the Midcontinent rift system in the western Lake Superior region and area covered by the 3D model. Geographic boundaries are shown by dashed lines. Figure 2: Isopach map of the Oronto Group from the 3D model, shown as a color-shaded relief image with illumination from the west. Gravity lows interpreted as basement highs (Allen et al., 1997): White's Ridge (WR) and Grand Marais Ridge (GM). Location of GLIMPCE seismic line C is indicated by red and white line. 55 �Quantitative abundance and preliminary morphological characterization of amphiboles in the Ironwood Iron-Formation, Gogebic Iron Range, Wisconsin GREEN, Carlin J., SEAL, Robert, R., II, CANNON, William F., and PIATAK, Nadine U.S. Geological Survey, MS 954, Reston, VA 20192 The western portion of the Gogebic iron range in northern Wisconsin constitutes one of the largest undeveloped iron resources of the Lake Superior region (Cannon et al., 2007). Knowledge of mineralogical changes related to metamorphism in the Gogebic iron range will be essential to planning for potential future resource development, especially solid mine waste management practices. The purpose of this study is to document the distribution and morphological character of amphibole minerals in the Ironwood Iron-Formation related to the thermal impact of the Mellen Intrusive Complex (MIC). The Paleoproterozoic Ironwood Iron-Formation is the principal iron-bearing unit in the Gogebic iron range. The Ironwood has five named members (from the base upward): Plymouth, Yale, Norrie, Pence, and Anvil (Anvil Member is absent in the western Gogebic iron range) (Cannon et al., 2007). The Ironwood contains ferruginous chert, Fe-oxides, Fe-silicates, and Fecarbonates in granular or oolitic wavy and finely laminated beds. Contact metamorphism related to the Mesoproterozoic MIC had the combined effects of increasing the magnetite content and progressively changing the silicate mineralogy of the Ironwood Iron-Formation (Cannon et al., 2007). Three drill holes, selected for study based on the completeness of their stratigraphic section and location relative to the MIC contact, were examined to determine variability in mineral assemblages with changes in metamorphic grade. The drill holes are oriented east-west, increasing in distance from the MIC eastward. Drill hole A is approximately 2 km from the contact, B is approximately 3.2 km from the contact, and C is approximately 4.7 km from the contact. Samples were selected from lithologic sub-units within the 4 different members to create a representative sample suite for each member. A separate topical set of samples was also chosen from areas of particular interest to elucidate amphibole paragenesis. A few samples were also collected from the underlying Palms Formation. All samples were examined by optical microscopy and scanning electron microscopy (SEM), and analyzed by powder X-ray diffraction (XRD) using Co Kα radiation. Quantitative estimates of mineral abundances were calculated from XRD data using the Rietveld method. The three drill cores show a well-defined progression from low-grade metamorphic conditions in the easternmost drill hole (A) to high-grade metamorphic conditions in the westernmost drill hole (C). Figure 1 displays the mineral distribution in the representative suite of samples from each drill hole. Samples are shown at the drill hole depth from which they were taken. Drill hole A primarily exhibits low-grade metamorphism characteristics (as described in Klein, 1983). Quartz is commonly present in major amounts. Magnetite typically ranges from approximately 10 to 40 weight percent (wt. %) but is absent in a small number of samples. Siderite and members of the dolomite-ankerite series are common and make up significant portions of the total rock volume. Chlorite-group minerals are commonly present in trace to minor amounts. The presence of Fe- and Mg-carbonates and trace amounts of Fe-silicates such as stilpnomelane are indicative of the low-grade metamorphic conditions. Amphiboles were not detected within the representative samples, but trace to minor amounts were found in several of the topical samples from the Norrie and Plymouth members. 56 �Central drill hole B commonly contains amphiboles of the grunerite-cummingtonite series, diagnostic of medium-grade metamorphic conditions (see Klein, 1983). These amphiboles are present in each member of the Ironwood, ranging in quantities from 3 to 65 wt. %. Magnetite and quartz remain the most common minerals. The near-complete absence of Fe- and Mg-carbonates is due to metamorphic reactions of these minerals with quartz, producing members of the grunerite-cummingtonite and actinolite-tremolite series along with calcite. Chlorite-group minerals remain common but at diminished amounts. Drill hole C has indications of high-grade metamorphism, such as the appearance of pyroxene, fayalite, and garnet. Quartz and magnetite continue to be common. Members of the gruneritecummingtonite and actinolite-tremolite series remain common and are present from trace amounts up to approximately 40 wt. %. Chlorite-group minerals become rare at these conditions. Research on amphibole morphology and mineral chemistry is ongoing. Examination of thin sections using petrographic microscopy and SEM reveals a range of amphibole crystal habits and grain sizes. Distinctions in primary morphological classifications of individual amphibole crystals, based on degree of elongation of one or more dimensions and aspect ratio, is underway. Very fined-grained crystalline aggregates composed of grains less than 10 µm in size, as well as large single crystals up to 5.5 mm also have been observed. The distribution of a variety of amphiboles in the Ironwood Iron-Formation has been documented by powder XRD and SEM-EDS (energy dispersive spectroscopy). Quantitative mineral abundance estimates show that the presence of amphiboles is dependent on the proximity to the MIC. Preliminary microscopic investigation reveals a range of crystal habits and geometric variability. These will be characterized in more detail in continuing studies, which will include electron probe micro-analysis to determine quantitative mineral chemistry. Figure 1. Mineral distribution in weight percent in the Ironwood Iron-Formation. REFERENCES Cannon, W.F., LaBerge, G.L., Klasner, J.S., and Schulz, K.J. 2007, The Gogebic iron range—a sample of the northern margin of the Penokean fold and thrust belt: U.S. Geological Survey Professional Paper 1730, 44 p. Klein, C., 1983, Diagenesis and metamorphism of Precambrian banded iron-formations: In, Trendall, A.F. and Morris, R.C. (Eds.), Iron-formation: Facts and Problems, Elsevier, Amsterdam, p. 417-465. 57 �PALEOCURRENT INTERPRETATION OF THE CAMBRIAN ELK MOUND GROUP USING GEOPHYSICAL OPTICAL BOREHOLE IMAGE (OBI) LOGS FROM TWO NEW BOREHOLES, DODGE COUNTY, SOUTHERN WISCONSIN GUENTHER, Gregory1, KINGSBURY STEWART, Esther2 1Department 2Wisconsin of Geology and Geography, University of Wisconsin - Whitewater, 800 W Main St, Whitewater, WI 53190 USA Geological and Natural History Survey, 3817 Mineral Point Rd, Madison, WI 53705 Paleocurrents measured from cross-beds in sandstone indicate sediment transport direction or the flow direction of water at the time of deposition. Characterizing variation in paleocurrent direction through a stratigraphic succession is one tool for investigating evolution of depositional environment through time. In southern Wisconsin the Cambrian Elk Mound Group is comprised of, from oldest to youngest, the Mount Simon Formation sandstone, Eau Claire Formation shale, and Wonewoc Formation sandstone. In south-central Wisconsin the Eau Claire Formation undergoes a facies change from shale to sandstone. Here the Elk Mound Group is present mostly in the subsurface so it is understood primarily through observation of drill core, drill cuttings, and downhole geophysical logs. The purpose of this study is (1) to test the validity of using measurements from optical borehole image (OBI) logs to measure paleocurrents in sandstones present in the subsurface and (2) investigate how paleocurrent direction varies through two thick, sandstone-dominated intervals of the Elk Mound Group in Dodge County, Wisconsin. OBI logs are oriented, down-hole images taken from drill site boreholes. The images return as 360 degree unwrapped digital representations of the borehole, and these images record sedimentary structures like cross-beds, laminations, and bedding planes. We use Wellcad software to measure the strike and dip of cross-beds and laminations from the Elk Mound Group that were imaged with OBI logs from two boreholes that the Wisconsin Geological and Natural History Survey drilled in Dodge County in 2015. The Slinger site encountered 243 feet (74 m) of Elk Mound Group sandstone before hitting Precambrian iron-formation. The Westphal 2 site encountered 422 feet (129 m) of Elk Mound Group sandstone and did not hit Precambrian basement. We compare OBI logs to drill core collected at each site. We compare our paleocurrent measurements to existing published data. Finally, we use elemental data collected using a handheld x-ray fluorescence (XRF) instrument at one foot intervals along each core to subdivide the Elk Mound Group. We observe how paleocurrent measurements vary between each of these preliminary subdivisions of the Elk Mound Group. The paleocurrent directions we measured from OBI logs are consistent with published data for the Elk Mound Group (Michelson et al., 1973; Driese et al., 1981; Hagadorn et al., 2002). Published data report primarily south-directed paleocurrent directions with a secondary bimodal direction for the Elk Mound Group. At the Slinger and Westphal 2 study locations, we subdivide the Elk Mound Group into four informal units based on changes in elemental concentrations of Al, Ti, Zr, and K. The lower two units are characterized by southwest-directed paleocurrents. Paleocurrents become bimodal in the upper two units. Based on the similarity with published paleocurrent data, we conclude that paleocurrents may be measured accurately from OBI logs. Furthermore, the change in paleocurrent direction we observe from southwest in the lower Elk Mound Group to bi-modal in the upper Elk Mound Group likely reflects a change in environmental conditions that control sediment transport and deposition. Measuring paleocurrent direction from OBI logs is therefore a useful tool to aid subdivision of seemingly monotonous, thick packages of sandstone that are present in the subsurface. 58 �We acknowledge Dr. Prajukti Bhattacharyya, University of Wisconsin-Whitewater, for her assistance with this project. Driese, S. G., Byers, C. W. (1981). Tidal Deposition in the Basal Upper Cambrian Mt. Simon Formation in Wisconsin. Journal of Sedimentary Research, vol. 51, no. 2, p. 367-381. Hagadorn, J. W., Dott, R. H. Jr., Damrow, D. (2002). Stranded on a Late Cambrian shoreline: Medusae from central Wisconsin. Geology, vol. 30, no. 2, p. 147-150. Michelson, P. C., Dott, R. H. Jr. (1973). Orientation Analysis of Trough Cross Stratification in Upper Cambrian Sandstones of Western Wisconsin. Journal of Sedimentary Research, vol. 43, no. 3, p.784-794. NOTE: Journal of Sedimentary Research formerly known as Journal of Sedimentary Petrology. 59 �U-Th-Pb isotopes of the Reef Deposit; a Au-Cu occurrence in central Wisconsin HAROLDSON, Erik1 BEARD, Brian1 SATKOSKI, Aaron1 JOHNSON, Clark1 BROWN, Philip1 1 Department of Geoscience, University of Wisconsin-Madison, 1215 W. Dayton, Madison, WI, 53706 USA The Reef deposit, located within the Wausau Volcanic complex (WVC), is a vein hosted Au-Cu occurrence historically calculated to contain ~454,600 tons grading 0.262 opt gold, 0.25 opt silver and ~0.28% copper located approximately 15 miles east of Wausau, Wisconsin. The Ladysmith-Rhinelander volcanic complex (LRVC), adjacent to the North of the WVC within the Pembine-Wausau volcanic subterrane, is known to contain many volcanogenic massive sulfide (VMS) deposits, several of which are considered to be economically viable to mining (DeMatties 1996). The Reef deposit has been described in the past as a lode gold or shear-zone hosted gold deposit (DeMatties 1996) separating it genetically from the VMS deposits to the north, along with a physical separation of the hosting volcanic complex (WVC vs. LRVC). The Reef deposit has also been described as the root zone of a Cu-Zn or Zn-Cu VMS deposit (Scott 1988) and although there are no known VMS deposits of economic interest, the WVC is known to host minor VMS mineralization (DeMatties 1996). An orogenic lode gold deposit could potentially be associated with various tectonic events in the history of central North America including; ~1700 Ma Yavapai, ~1600 Mazatzal or even anorogenic magmatic activity of the ~1510 Ma Stettin syenite complex or the ~1470 Ma Wolf River Batholith. The Reef deposit consists of seven mineralized zones which consist of Au-Cu bearing quartz-sulfide veins hosted in primarily basaltic material (mafic metavolcanics and gabbroic intrusions) (Kennedy and Harding, 1990). Whilst ore zones of quartz-sulfide veins are also crosscut by similar gabbroic intrusives it is unclear when the Au mineralization was deposited. The zones trend northeast, dip to the northwest and are closely associated with felsic intrusions (Kennedy and Harding 1990). Felsic dikes and sills are intermingled with the host gabbro intrusives as a swarm of granophyric to porphyritic, locally aplitic units. The deposit area is flanked to the west and northwest by dominantly basalt of massive and pillowed flows and mafic tuff. Felsic intrusions flanking the deposit are concentrated in the western and southeastern areas adjacent to the deposit. U-Th-Pb isotopes analyses were made of sulfide phases which were micro-drilled from drill core selections and aliquots of coarse reject from split drill core assay material, previously split and crushed by Aquila Resources. The coarse reject samples are referred to as ‘whole rock’ (WR) samples. WR and sulfide (chalcopyrite, pyrite and pyrrhotite) 206Pb/204Pb, 207Pb/204Pb and 208 Pb/204Pb values range from 15.732 to 22.897, 15.222 to 16.058 and 35.139 to 41.794 respectively (Fig 1). Pb isotope values indicate a source region for sulfides at Reef similar to that of the Flambeau VMS deposit. The age of volcanogenic massive sulfide deposits in the LRVC have been estimated at ~1860 Ma (DeMatties 1996) and a Pb-Pb isochron for Reef of 1925 ± 92 Ma does overlap the estimate. Measured 238U/204Pb (μ) values range from 0.185 to 4.45 for sulfides and 1.016 to 14.202 for WR samples. 232Th/204Pb (ω) values range from 0.0436 to 3.084 for sulfides and 0.623 to 19.051 for WR samples. Only some μ and ω values support highly radiogenic Pb values to have been derived from in-situ decay (Fig 2). Samples are either subject to later coincident uranium and thorium loss, or lead addition from a high μ source. 60 �A genetic model for the Reef deposit remains elusive; but a primary association with the Penokean orogen has been established. Figure 1 – 207Pb/204Pb vs 206Pb/204Pb plot of results from the Reef deposit and the Flambeau and Lynne VMS deposits. Smaller plot is inset with same data plotted. WR = Whole Rock; WRHM = Whole Rock Hand Magnetic fraction. Dark filled squares are previously published values for VMS deposits (Afifi et al 1984). Curve plotted is lead evolution from Stacey and Kramers (1975). Figure 2 - plots investigating the in-situ decay of U-Th in Reef deposit samples. Symbols are same as figure 1. Reference isochron of 1925 Ma is shown in each plot. REFERENCES Afifi, A., Doe, B.R., Sims, P.K., Delevaux, M.H., 1984, U-Th-Pb isotope chronology of sulfide ores and rocks in the early Proterozoic metavolcanic belt of northern Wisconsin, Economic Geology, v. 79, pp. 338-353 Dematties, T.A., 1996, A geologic framework for early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: an exploration model, Institute on Lake Superior Geology, Volcanogenic massive sulfide deposits of northern Wisconsin: a commemorative volume, pp. 31-65. Kennedy, L.P., and Harding, T.A.,1990, Summary report of the Reef joint venture Marathon County, Wisconsin, Noranda Exploration Inc., 63 pp. Scott, W.P., 1988, A volcanic hosted gold occurrence in Marathon County, Wisconsin, A thesis submitted in partial fulfillment of the requirements for the degree of master of science geology, University of WisconsinMadison, 99 pp Stacey, J.S., and Kramers, J.D., 1975, Approximation of terrestrial lead isotope evolution by a two-stage model, Earth and Planetary Science Letters, vol. 26, pp 207-221. 61 �A New Rusk County: Producing an new Precambrian geological map from new field observations and compilations of historic geological/geophysical datasets HELMUTH, Samuel L.1, LODGE, Robert W.D.1 1 Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA Precambrian rocks of the Penokean Orogeny within the Wisconsin Magmatic Terrane host world-class volcanogenic massive sulfide deposits (VMS) near the southern limit of the Canadian Shield (DeMatties, 1994). This is a Paleoproterozoic juvenile and continental arc sequence consisting of mafic to felsic volcanic assemblages, sedimentary sequences, and associated plutonic rocks that are about 1.8-1.9 billion years old (Schulz & Cannon, 2007). These deposits are significant sources of metals such as zinc, copper, lead, silver and gold in the form of sulfide minerals. The size and concentration of VMS deposits in Wisconsin gives the region potential to become a world-class mining district. However, a thick cover of glacial deposits and Paleozoic sedimentary strata limit the Precambrian exposure in the region and make regional tectonic and metallogenic interpretations of the Precambrian geology difficult. The primary objective of this project was to create a new geologic map of the Precambrian geology in Rusk County, Wisconsin, describing the distribution and petrologic characteristics of the Precambrian bedrock beneath this cover by compiling historic geological and geophysical datasets with new field observations and combining these observations with modern geochemical data and interpretations. There were two phases of mapping related to this project. The first phase included sampling rocks from drill core, which was collected during mineral exploration and mining activities, in regions in the county where no bedrock exposure exists. These uppermost parts of these drill cores were described and sampled. The second phase involved field mapping and sampling of Precambrian rock outcrops in the rivers, road cuts, and quarries throughout Rusk County. All samples were processed and analyzed via XRF to determine their major and minor element geochemistry and thin sections from select samples were used for petrographic analysis. The volcanic and tectonic insights gained through petrographic and geochemical interpretations were compiled and geospatially integrated with historic geologic maps published through the Wisconsin Geological and Natural History Survey (Mudrey et al. 1987) and aeromagnetic geophysical surveys available through the USGS (Daniels & Snyder, 2002) to better constrain the Precambrian lithostratigraphic units beneath Cambrian and Quaternary cover. This integration of field mapping, drill core sampling, and data compilation of geological and geophysical datasets in ArcGIS created a new and improved geological map of the Precambrian rocks of Rusk County (Figure 1). In addition, petrographic and geochemical interpretations of these rocks will add a new layer of understanding that will significantly improve our tectonic, volcanological, and metallogenic framework of the Precambrian bedrock in Wisconsin. Based on aeromagnetic geophysical patterns, distribution of various volcanic, plutonic, and sedimentary rocks, and geochemical domains throughout Rusk County, we subdivided the Precambrian geology into four supracrustal units: (1) Blue Hills Metasedimentary Unit, (2) Thornapple Metavolcanic Unit, (3) Flambeau-Jump River Metavolcanic Unit, and (4) Weyerhaeuser-Forks Metavolcanic Unit. The Blue Hills Metasedimentary Unit underlies the northwestern portion of the county and has abundant outcropping quartzite and phyllite. The Thornapple Metavolcanic Unit is located throughout the north-central portion of the map forming a moderately magnetic geophysical domain and hosts the Zn-Cu-Pb Eisenbrey VMS 62 �deposit. Bedrock exposures and core samples observed included primarily intermediate volcanic assemblages with lesser basalt, rhyolites, and iron formation forming. The lower magnetic portion of the southeastern part of the county is predominantly mafic to intermediate volcanic assemblages that hosts the Flambeau Cu-Au VMS deposit. Lastly, the Weyerhaeuser-Forks Unit underlies the highly-magnetic areas in the southwest and northeast parts of the map. Here, alkali basalts and associated intrusions have been discovered. The aeromagnetic geophysical surveys also allowed us to better identify felsic plutonic bodies in the mapping area. Figure 1: Preliminary Precambrian geology of Rusk County, Wisconsin, showing major faults (bolded lines) and contacts (thin lines) shown in relation to the airborne magnetic geophysical map for the state (Daniels & Snyder, 2002). Circles in mapping area show where core and field samples have been collected and analyzed. Numbers 1 through 4 correlate to approximate location of supracrustal units described in text. REFERENCES Daniels, D.L. & Snyder, S.L., 2002. Wisconsin Aeromagnetic and Gravity Maps and Data: A Web Site for Distribution of Data. USGS Open File Report 02-493. DeMatties, T.A., 1994. Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Economic Geology, 89: 1122-1151. Mudrey, M.G., LaBerge, G.L., Myers, P.E., and Cordua, W.S., 1987. Bedrock geology of Wisconsin: Northwest sheet. Wisconsin Geological and Natural History Survey Regional Map Series. Scale: 1:250,000. Schulz, K.J. & Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research, 157: 4-25. 63 �Geochemical and petrological studies on the origin of Ni-Cu sulfide mineralization at the Eagle and Eagle East intrusions in Marquette County, Michigan HINKS, Benjamin1, THAKURTA, Joyashish1, MAHIN, Robert2, and BEACH, Steve2 1 Department of Geosciences, Western Michigan University, 1903 W. Michigan Ave. Kalamazoo, MI 49008, benjamin.d.hinks@wmich.edu 2 Eagle Mine, Lundin Mining Corporation, 4547 County Road, Champion, MI 49814 The Eagle deposit is a small, high-grade Ni-Cu-bearing sulfide deposit located in the north central portion of Michigan’s Upper Peninsula in Marquette County. The Eagle and the Eagle East intrusions, separated by 600m, are parallel to the east-west trending Marquette-Baraga dike swarm that is associated with ~1.1 Ga Midcontinent Rift System magmatism (Ding et al. 2010). Recently, in 2014 a discovery was made within the Eagle East intrusion. Drilling by Lundin Mining Corporation intersected a small section of high grade sulfide minerals near the base of the intrusion. The purpose of this study is to characterize the magmatic sulfide deposit at Eagle and its relationship to the surrounding country rocks. The relationship between Eagle sulfide ores and the newly discovered Eagle East sulfide ores have also been analyzed. Petrographic analysis has shown that Eagle and Eagle East sulfide mineralization consists of three minerals: pyrrhotite, pentlandite and chalcopyrite. Olivine crystals tend to host sulfide minerals, indicating that an immiscible sulfide liquid must have formed prior to olivine formation. Rare earth element (REE) analysis from peridotite samples at Eagle and Eagle East confirms that the intrusions are very similar. The REE diagram depicts similar concentrations and patterns of REEs with shallow slopes for the two intrusions, indicating low degrees of crystal fractionation and high degrees of partial melting that formed the magmas at Eagle and Eagle East. Sulfur isotope analysis was also performed on sulfide ores from Eagle, Eagle East, Michigamme Formation rocks and Archean basement rocks. The three intrusive sulfide zones (disseminated, semi-massive, and massive) for both intrusions are within the range of 0‰ to 5‰. Michigamme Formation rocks display δ34S values from 6‰ to 20‰. Archean basement rocks display a wide range of sulfur isotope values from -11‰ to 7‰. The δ34S values reported from the main sulfide ore bodies of Eagle and Eagle East are indicative of slight enrichment of 34S, most likely caused by interaction and mixing of mantle-sourced sulfur (0‰) with surrounding crustal-sourced sulfur. Country rock contamination δ34S signatures found within the immiscible sulfide liquid may have been averaged or totally erased through interaction of mantle magmas with the crustal sulfide liquid. δ34S signatures could have additionally been altered by the mixing of sulfur from the two crustal sources, Archean and Proterozoic rocks. Archean rocks typically have a range of high and low δ34S values from -11‰ to 7‰, while Proterozoic rocks have high δ34S values of 6‰ to 20‰. Mixing of sulfur sourced from Archean rocks with sulfur sourced from Proterozoic Michigamme Formation rocks could have averaged the δ34S signatures seen in the ore bodies to ~3.5‰. Previous researchers Ding et al. (2012) determined δ34S values at the Eagle deposit for disseminated and massive sulfides ranging from 0.3‰ to 4.6‰, while semi-massive sulfides were characterized by δ34S values ranging from 2.2‰ to 5.3‰. The δ34S values of sulfide ores hosted in the intrusion are much lower than would be expected if assimilation of crustal-sourced sulfur had occurred. Ding et al. (2012) attributed the low δ34S values within the Eagle intrusion 64 �to mantle-sourced sulfur (0‰) coming into contact with sulfide liquid that had higher δ34S values attributed to crustal rock assimilation and depleting the δ34S values closer to 0‰. However, since the δ34S values within intrusive sulfides at both Eagle and Eagle East in this study are so similar, it is proposed that mixing of crustal-sourced sulfur form the Michigamme Formation rocks and Archean basement rocks could have additionally averaged the δ34S values seen within the intrusions to a range between 0‰ to 5‰. Mixing of sulfur derived from crustal rocks and the interaction of mantle-sourced sulfur could have averaged the δ34S values seen within the ore bodies from that of crustal contamination signatures. Figure 1: Hypothetical diagram illustrating a model for the mixing of sulfur from Archean and Proterozoic sources, resulting in δ34S values between 0‰ to 5‰ for sulfides hosted within the Eagle and Eagle East intrusions. REFERENCES Ding, X., C. Li, E. M. Ripley, D. Rossell, and S. Kamo (2010), The Eagle and East Eagle sulfide ore‐bearing mafic‐ ultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution, Geochem. Geophys. Geosyst., 11, Q03003, doi:10.1029/2009GC002546. Ding, X., E. M. Ripley, S. B. Shirey, C. Li (2012), Os, Nd, O and S isotope constraints on country rock contamination in the conduit-related Eagle Cu-Ni-(PGE) deposit, Midcontinent Rift System, Upper Michigan: Geochemica et Cosmochimica Acta, 89, pp. 10-30. 65 �Preliminary Observations of the Ultramafic Metavolcanic Rocks in the Eastern Portion of the Shebandowan Greenstone Belt, northwestern Ontario HINZ, Sheree and HOLLINGS, Pete Department of Geology, Lakehead University, Thunder Bay, Ontario P7B 5E1 The ultramafic metavolcanic rocks in the eastern portion of the Shebandowan greenstone belt are located within Conmee Township (Fig. 3), which is part of the larger Wawa-Abitibi terrane (Stott et al., 2010). The Shebandowan greenstone belt has been divided into three main assemblages based on geochronological studies by Stott and Corfu (1998) and Lodge (2012), namely; the Greenwater assemblage (circa 2720 Ma), the Kashabowie assemblage (circa 2695 Ma) and the Shebandowan assemblage (circa 2690 to 2680 Ma). This study focuses on the ultramafic metavolcanic rocks of the Greenwater assemblage that are best exposed in the Bateman trenches dug by Linear Metals in 2008 and the Freewest and Dawson trenches (Fig. 3). These trenches display near continuous stratigraphic sections including spinifex-textured flows, massive flows, flow breccia and quenched autobreccia (Lodge, 2014). Other rock types in this assemblage include mafic, intermediate and felsic metavolcanic rocks, as well as metasedimentary rocks. Preliminary field observations include the identification of preserved macroscopic primary volcanic textures, despite the greenstone belt being metamorphosed to greenschist facies. The macroscopic textures and associations are important because on a microscopic scale, all primary mineralogy has been altered and metamorphosed. Spinifex texture is widely observed throughout the map area; this texture is formed through undercooling of a very hot magma and is a critical feature of komatiites (Fig. 2). Associated with the spinifex-textured komatiite in the Linear Metals trenches is a unit with polyhedral jointing analogous to columnar jointing in basaltic flows (Arndt 2008). In several localities an ultramafic volcanic breccia with spinifex-textured angular fragments in a glassy groundmass is present (Fig. 1). These macroscopic textures are difficult to put into context with the surrounding units because of the gaps in the trenches where contacts should exist. Preliminary geochemical analyses show that these rocks have lower than normal MgO for a true komatiitic rock, with an average of 11.2 wt% but some samples showing up to 27wt% MgO. The SiO2 content is also higher than normal for a typical komatiite, with an average of 50%, with some samples as low as 40wt%. The Ni content fits well with the typical komatiite composition with an average of 607ppm and average Cr content of 1133ppm. This suggests that the ultramafic rocks in the study area are predominantly komatiitic basalts with rare komatiites. Both komatiites and komatiitic basalt display spinifex texture. Further mapping will be completed in the summer of 2016 in order to correlate the units over a wider area. Petrographic studies are underway and geochemical data will be analyzed once obtained. By combining detailed petrographic studies with whole-rock geochemical data this study will investigate fractionation trends, mixing and crustal contamination signatures to understand the evolution of the komatiites. This study will produce a detailed map of the komatiites and develop a model for the formation of these rocks. 66 �Figures 1) Spinifex texture in ultramafic rock, from Linear Metals trench 2) Ultramafic breccia from Freewest trench Figure 3) Regional geology map showing the location of study areas within Conmee Township in relation to the Shebandowan greenstone belt. Bedrock geology modified from Santaguida (2001a, 2001b) and Lodge et al. (2015). REFERENCES Arndt, N.T., Lesher, C.M, Barnes, S.J. 2008. Komatiite. Cambridge University Press. 467p. Corfu, F. and Stott, G.M. 1998. Shebandowan greenstone belt, western Superior Province: U/Pb ages, tectonic implications, and correlations; Geological Society of America Bulletin, v.110, p.1467-1484. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Franklin, J.M. and Hudak, G.J. 2015. Geodynamic setting, crustal architecture, and VMS metallogeny of ca. 2720 Ma greenstone belt assemblages of the northern Wawa subprovince, Superior Province; Canadian Journal of Earth Sciences, v.52, p.196-214. Lodge, R.W.D. 2012. Preliminary results of uranium–lead geochronology from the Shebandowan greenstone belt, Wawa Subprovince; in Summary of Field Work and Other Activities 2012, Open File Report 6280, p.10-1 to 10-10. Lodge, R.W.D., Ratcliffe, L.M., and Walker, J.A. 2014. Geology and Mineral Potential of Sackville and Conmee Townships, Wawa Subprovince; in Summary of Field Work and Other Activities 2014, Open File Rpt 6300, p 9-1 to 9-17. Santaguida, F. 2001a. Precambrian geology compilation series—Quetico sheet; Ontario Geological Survey, Map 2663, scale 1:250 000. ——— 2001b. Precambrian geology compilation series—Thunder Bay sheet; Ontario Geological Survey, Map 2664, scale 1:250 000. Stott, G.M., Corkery, M.T., Percival, J.A., Simard, M and Goutier, J. 2010. A revised terrane subdivision of the Superior Province; in Summary of Field Work and Other Activities 2010, Ontario Geological Survey, Open File Report 6260, p.20-1 to 20-10. 67 �The Minnesota Taconite Workers Health Study: Environmental Study of Airborne Particulate Matter - 2015 Update HUDAK, George1, MONSON GEERTS, Stephen1, ZANKO, Larry1, POST, Sara1, and REAVIE, Euan1, 1 Natural Resources Research Institute, 5013 Miller Trunk Highway, Duluth, MN, 55811 The Natural Resources Research Institute (NRRI) conducted a detailed characterization of mineral dust in northeastern Minnesota. The purpose of this research was to evaluate the effects of present emissions from taconite mining and processing on air quality throughout the Mesabi Iron Range (MIR) (Figure 1) by characterizing airborne mineral particulate matter (PM) within currently operating taconite processing plants, in MIR communities surrounding taconite mining/processing operations, and in population centers in Minnesota not associated with taconite mining. Characterization studies of age-dated lake sediments were also conducted to determine the composition of past PM deposition. NRRI’s sampling and characterization work represents the community/environmental component of the Minnesota Taconite Workers Health Study, a broad University of Minnesota (UM) research effort involving both the NRRI and the School of Public Health. Figure 1. Locations of taconite processing plants on the Mesabi Iron Range being sampled during this study (after Oreskovich and Patelke, 2006) Air sampling was performed within taconite operations, MIR communities, and non-MIR communities by NRRI scientists during both winter and summer seasons from 2009-2012. Sampling was conducted at four process locations within taconite operations, including: 1) secondary crushers; 2) magnetic separators/concentrators; 3) agglomerators/ball drums; and 4) kiln/pellet discharge areas. Community sampling took place on centrally-located rooftops of public buildings, or in the case of the northern most background site, in a remote sampling location to evaluate the air quality away from the MIR. Sampling and analytical techniques for all project work are described in detail in several reports that are in preparation. NRRI’s research methods do not produce exposure data, and are not meant to provide data for regulatory purposes. 68 �NRRI has evaluated the physical (gravimetric, morphology, concentration), mineralogical, and chemical characteristics of the PM obtained from sampling at the taconite operations and MIR/non-MIR communities. This included analysis of 55 taconite plants; 73 northeastern Minnesota community and 6 Minneapolis samples. Age-dated lake sediment cores were collected from Silver Lake in Virginia, MN, on the central MIR, and “North-of-Snort” Lake on MIR’s eastern end, near Babbitt, MN, and 36 analyses were completed. The results provide historical data regarding potential mineralogical inputs from iron mining and processing from ~1840 (which pre-dates iron mining on the MIR) to the present, including the natural ore mining to taconite mining transition period. Community results are as follows:  measured particulate matter concentrations for PM2.5 in all MIR communities have been below 12 µg/m3, and for total PM have been below 16µg/m3;  particulate matter concentrations on the MIR are similar to those in the two NE Minnesota background sites (Duluth NRRI, Ely Fernberg site), and are lower than those obtained in Minneapolis (UM Mechanical Engineering Building rooftop);  mineral particulate matter in community air samples reflects the mineralogy of the Biwabik Iron Formation and other Minnesota rock types and geological materials;  elongate mineral particles (EMP) are present in MIR community ambient air samples; however, asbestiform amphiboles were rarely observed (1 asbestiform amphibole EMP in ~22,800m3 of air). Taconite plant results are as follows:  plant environments can be dusty, with the most dusty environments associated with the agglomerator and kiln discharge areas;  particulate matter levels (PM1, PM2.5, PM10, and total PM) show a slight increase in the five MIR communities during plant/mine activity, but this increase is not statistically significant compared to when the plants were not operating.  significantly higher concentrations of EMPs, including amphiboles, were detected in the eastern most plant compared with the other five plants, but the morphology of these structures more closely resembles cleavage fragments rather than asbestiform morphologies. Lake sediment results are as follows:  a water elutriation method developed by Webber et al. (2008) was effective for isolating PM2.5 particles from age-dated sediment intervals.  the mineralogy of isolated PM2.5 EMPs reflects bedrock and glacial geology in the vicinity of both lakes.  a portion of the insoluble PM within the sediment of the two MIR lakes is most likely attributable to atmospheric inputs (fugitive dust) generated by historic iron ore/taconite mining activity References MDH. Method 852 (1999) T.E.M. analysis for mineral fibers in air – 852. Minnesota Department of Health, Microparticulate Unit, St. Paul, MN. 42 pp. Oreskovich, J. A., and Patelke, M. M., 2006, Historical use of taconite byproducts as construction aggregate materials in Minnesota: A Progress Report: Natural Resources Research Institute Report of Investigation NRRI-RI-200602, 10 p. Webber, J. A., Blake, D. J., Ward, T. J., and Pfau, J. C., 2008, Separation and characterization of respirable amphibole fibers from Libby, Montana, Inhalation Toxicology, 20:8, p 733-740. 69 �Lithostratigraphy and Ore Petrology of the Eisenbrey Zn-Cu-Pb Deposit, Rusk County, Wisconsin JACKSON, Nathaniel, DE MOURA MERSS, Bruno, and LODGE, Robert W.D. Department of Earth Science, University of Wisconsin-Eau Claire, Eau Claire, WI The primary objective of this research is to study the geological characteristics of the poorly understood Eisenbrey Zn-Cu-Pb deposit in Rusk County, Northwestern Wisconsin. Volcanogenic massive sulfide (VMS) deposits are significant sources of metals such as zinc, copper, lead, silver and gold in the form of sulfide minerals (Franklin et al. 2005). Despite the proximity of the Eisenbrey deposit to the better known, past-producing Flambeau Cu-Au VMS deposit, there has been essentially no research completed on the rocks hosting the Eisenbrey nor has there been any volcanic and tectonic linkages made to the strata hosting the Flambeau. Understanding the tectonic and metallogenic framework of the Eisenbrey and any potential genetic relationship to the Flambeau deposit will significantly improve our understanding of the Precambrian geology of northwestern Wisconsin (e.g. DeMatties 1996). The samples of the metalliferous ores and their host rocks are currently being analyzed to determine their petrographic and geochemical characteristics and to re-interpret the economic geology of this region. Initial phases of research included visiting the outcrop exposures of the Paleoproterozoic rocks that host the deposit that are present along the Thornapple River approximately 4 miles northwest of the Flambeau deposit. The ore horizon is present as interlayered sulfide-bearing magnetite-chert iron formation and strongly foliated chloritic mafic units exposed with a total stratigraphic thickness of ~10 m. The ore horizon is hosted in variably silicified intermediate to felsic volcaniclastic rocks. These Paleoproterozoic rocks are intruded to the north and west by gabbro and pyroxenite dykes associated with the midcontinent rift. The second phase of this research involved re-logging and sampling of historic drill core collected during mineral exploration (Figure 1). The ores and altered rocks are hosted within a thick pile of moderately foliated dacitic volcaniclastic unit consisting of tuffs and lapilli tuffs that have feldspar phenocrysts up to 3mm in size and composes 1-2% of the rock. These rocks are altered to an anthophyllite-magnetite±sericite schist that is generally fine grained and has a moderate foliation. This altered unit is variably mineralized and contains thin pyrite stringers and disseminations. Based on lithostratigraphic associations, the Eisenbrey deposit is a bimodal felsic-type deposit. In addition, the presence of an ore horizon dominated by sphalerite, chalcopyrite, pyrite and pyrrhotite supports this classification (Franklin et al., 2005). However, massive to semimassive ore are generally low grade and are composed of mostly pyrrhotite and pyrite with local abundances of chalcopyrite and sphalerite composing a combined 10% over 1-2 meters. Reflected light and scanning electron microscopy are currently being carried out on these ores. In addition to confirming the major ore minerals, trace amounts of silver telluride have been found. 70 �Figure 1: Representative cross section of the Eisenbrey VMS deposit showing location and stratigraphic context of holes logged for this study (bold lines and labels). Figure modified from LeBerge (1996). REFERENCES DeMatties, T.A., 1994. Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Economic Geology, 89: 1122-1151. Franklin, J.M., Gibson, H.L., Jonasson, I.R., Galley, A.G., 2005. Volcanogenic massive sulphide deposits, in: Hedenquist, J.F.H., Goldfarb, R.J., Richards, J.P. (Eds.), Economic Geology, 100th Anniversary Volume, pp. 523-560. LeBerge, G.L. (ed), 1996. Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume. Institute on Lake Superior Geology, Proceedings, 42nd Annual Meeting, Cable, WI, vol. 42, part 2, 179 p. 71 �Incommensurately modulated structure of plagioclase as an indicator of cooling history of igneous rock JIN, Shiyun, and XU, Huifang Department of Geoscience, University of Wisconsin-Madison, 1215 W. Dayton Street, Madison, WI 53706 Plagioclase feldspar is the most abundant group of mineral in the earth’s crust. Intermediate plagioclase formed at low temperature display a highly ordered aperiodic structure, resulting in satellite diffractions with irrational indices (e-reflections), which has been an enigma for more than half a century since its first discovery in 1940(Chao and Taylor 1940). Three plagioclase crystals of similar chemical composition from Duluth igneous complex were analyzed with single crystal X-Ray diffraction. The incommensurately modulated structure were solved in (3+1)D hyperspace. Detailed differences of the crystal structures were revealed, while the structures solved with only main reflections are basically the same as volcanic labradorite. The plagioclase in troctolite (An57) and gabbro (An54) from Duluth Igneous complex has a modulation period of 42Å and 45Å respectively, which are significantly larger than the modulation period of a metamorphic plagioclase with similar composition (~30Å). The structures display density modulation with an amplitude of less than 10 mole%, which is hard to characterize accurately due to the limited second order satellite reflections (f-reflections). The plagioclase in the anorthosite with composition of about An60, on the other hand, shows a modulation period of 40Å. The structure is obviously more ordered than the other two crystals, displaying a density modulation of 18 mole% in amplitude. The number of f-reflections collected is about triple of the troctolite sample and more than 10 times of the gabbro sample. The fitness of the refinement result and the experimental data is significantly better. The degree of orderliness characterized by the XRay diffraction data indicates a different origin of the anorthosite from the Duluth Layered Serious. Thus the modulation period and amplitude of e-plagioclase may be used as an indicator of the cooling history of the host rock. Chao, S.H. and Taylor, W.H. (1940) Isomorphous replacement and superlattice structures in the plagioclase feldspars, pp. 76-87. 72 �Nine years of capstones: A summary of Precambrian Research Center field camp capstone projects in the Neoarchean Knife Lake Group and associated rocks, central Boundary Waters Canoe Area Wilderness, Minnesota JIRSA, Mark A. Minnesota Geological Survey, University of Minnesota; jirsa001@umn.edu An important mission of the University of Minnesota-Duluth’s Precambrian Research Center (PRC) is training students to map the geology of Precambrian terranes. The arrowhead region of Minnesota generally—and the Boundary Waters Canoe Area Wilderness specifically—contains some of the best localities for such mapping. Since its inception in 2006, the PRC has conducted 9 seasons of geologic field camps in the region (2007-2015), with an emphasis on the rock types, alterations, and structures that host much of the world’s metallic mineral endowment. The training involves several weeks of mapping exercises, followed by a “capstone” project conducted in small groups with practicing geologists. These projects test student skills by creating new geologic maps in areas of poorly known geology, which benefits both students and mentor organizations. Students conduct field work for 7-8 days in remote locations, then remarkably turn their observations into geologic maps during 4 days and present their results before an audience of academic, government, and industry professionals. This presentation describes one set of capstone projects in the central part of the Boundary Waters Canoe Area Wilderness (BWCAW). The geology was mapped to varied levels of detail by the author and 41 students during 9 individual capstone projects. Although these capstones focused on lithologic and structural complexities of the Neoarchean Knife Lake Group (Fig. 1), the resulting maps provide details about other parts of the Wawa subprovince of the Archean Superior Province, rare diabasic dikes, unconformitybounded Paleoproterozoic iron-formation, and the basal Mesoproterozoic Duluth Complex. Figure 1. Generalized bedrock geologic map of NE Minnesota showing the area of capstone mapping projects (outline labeled “Knife Lake Capstones;” ~ area of Fig. 2). The Neoarchean unit labeled “Supracrustal Rocks” encloses both older volcanic sequences and younger, largely sedimentary ones. Outline of BWCAW is dashed. Volcanic, sedimentary, and intrusive rocks of the Knife Lake Group comprise a Timiskaming-type extensional basin and its apparent wall- and floor-rocks. The geologic units are parceled into structural lozenges separated by anastomosing shear and fault zones (Fig. 2). Although rock types are comparatively pristine within each lozenge, correlation of units from one fault-bounded block to another is challenging. Nevertheless, these projects attempt to “unstrain” the rocks within each lozenge to reveal stratigraphic variations that may reflect fluctuations in original basin geometry and progressive erosional dissection of basin wall rocks. Understanding the lithologic details and the apparent post-depositional tilt of individual lozenges of rock is essential to this objective. An integrated depositional and tectonic model of the basin is evolving using Gruner’s (1941) work, unpublished theses, and capstone mapping. Basin evolution appears to have involved early subaerial calc-alkalic volcanism synchronous with sediment deposition, uplift, erosion, and surface weathering that contributed alluvium to localized fault basins. This was followed by development of braided streams, and transitioned to subaqueous deltaic and turbiditic deposition. Strata containing clasts of the ca. 2690 Ma Saganaga Tonalite were deformed and metamorphosed at ca. 2680 Ma, which brackets basin development within a 10 Ma period. 73 �Figure 2 Historic geologic map of central BWCAW from Gruner (1941) showing fault zones that bound lozenges or segments of internally coherent strata. Bounding faults are modified from Gruner’s work based on subsequent mapping. Mapping funded in part by the Precambrian Research Center contributed to the following publications: Abstracts—Institute on Lake Superior Geology, 2008-2016. Lead authors/meeting years: Jirsa 2008, 2009, 2012; Fahrenkrog 2010; Birkmeier 2011; Korman 2013; Mulcahy 2014; Krogmeier 2015; Christenson 2016. Driese, S.G., Jirsa, M.A., Ren, M., Sheldon, N.D., Brantley, S.L., Parker, D., and Schmitz, M., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, v. 189, p. 1-17. Jirsa, M.A., 2011, Bedrock geology of the western Gunflint Trail area, northern Minnesota: Minnesota Geological Survey Miscellaneous Map M-191, scale 1:24,000. Jirsa, M.A., Boerboom, T.J., Chandler, V.W., Mossler, J.H., Runkel, A.C., and Setterholm, D.R., 2011, Geologic map of Minnesota-Bedrock geology: Minnesota Geological Survey State Map Series S-21, scale 1:500,000. Jirsa, M.A., Leu, A., and Miller, J.D., Jr., 2013, Preliminary bedrock geologic map of the Pagami Creek fire area, northeastern Minnesota: Minnesota Geological Survey Open-File Report OFR-13-01, scale 1:24,000. Jirsa, M.A., Starns, E.C., and Schmitz, M.D., in press, Bedrock geologic map of the 2006 Cavity Lake forest fire area, Boundary Waters Canoe Area Wilderness, northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map M-193, scale 1:24,000 (currently MGS Open-File Report OFR-2008-05). Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., Jirsa, M.A., and Hamilton, M.A., 2013, New U-Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research v. 235, p. 264-277. REFERENCE CITED Gruner, J.W., 1941, Structural geology of the Knife Lake area, northeastern Minnesota: Geological Society of America Bulletin 52:1577-1642, and map at scale 1:42,240. 74 �Geochemistry of Seine River metaconglomerates from Mine Centre, Ontario: interpreting fluid flow and volume changes during deformation with implications for strain analysis JOHNSON, Detaya1, and CZECK, Dyanna M.1 1 Department of Geosciences, University of Wisconsin-Milwaukee, P.O. Box 413, Milwaukee, WI 53201 USA Ductile shear zones are formed by tectonic plate interactions and are often thought to be deep-seated equivalents to faults. They can act as barriers or conduits for fluid flow, with many acting as both conduits for zone-parallel flow and barriers to cross-zone flow (e.g. Selverstone et al., 1991; Yonkee et al., 2013). Fluids have major impacts on deformation; they can alter deformation mechanisms, metamorphic reactions, and strain accumulation. Even though fluid is such an important factor in deformation, it is difficult to study due to its transient nature. Geochemical analyses may be used to interpret fluid conditions during deformation. In particular, major element geochemical analyses may be conducted to determine whether fluids assisted depletion of soluble elements resulting in volume changes (e.g. Newman and Mitra, 1993; Srivastava et al., 1995; Yonkee et al., 2003). This method has been used to demonstrate significant volume loss in some shear zones (e.g. up to 70% reported in Newman and Mitra, 1993) and none in others (Srivastava et al., 1995). The Seine River metaconglomerates in the Rainy Lake region of northwestern Ontario, Canada are located in a wedge of deformed rocks between the Quetico Fault and the Seine RiverRainy Lake Fault. The metaconglomerates underwent ductile deformation and greenschist facies metamorphism. They have clasts with varied lithology including felsic to ultramafic volcanic and granitoid clasts. Selected samples were previously analyzed to determine strain magnitudes by Czeck et al. (2009); they demonstrated that different clast types had differing rheologies that resulted in a range of strain magnitudes. Strain magnitude also varied across the region (Fig. 1). In the most highly strained outcrops, significant carbonate alteration suggests that fluid flow was an important factor during deformation, and there may have been a symbiotic relationship between fluid localization and enhanced deformation. Geochemical analysis and evaluation for fluid-rock interaction had not previously been performed on the Seine River metaconglomerates and was the focus of this study. Compositions of several clasts types (granitoids, mafic volcanics, and felsic volcanics) were determined for a range of strain magnitudes using major and minor element X-Ray Fluorescence (XRF). Results from 36 analyses were reviewed to see if there were significant changes in composition from low to high strain sites. These preliminary results show a high degree of variability amongst major constituents within the same clast type at individual sites. For a given clast type, the mean concentrations of all major elements are indistinguishable at low and high strains (Table 1). For example, in granitoid clasts, the mean concentration of SiO2 is 70.43% (st. dev. of 4.17%) at low strain and 69.13% (st. dev. of 6.06%) at high strain. The high variability and overlap of data for all major elements indicates that either A) there was no significant clast volume change between low and high strain regions or B) the geochemical signal of any volume change is masked by a large variability in initial compositions. So while there is clear evidence from alteration patterns that fluid flow differed between low and high strain outcrops, this did not result in measurable volume change with increasing strain. The inherent heterogeneity in metaconglomerates allows them to be extremely useful for strain analysis, but complicates geochemical characterization. In order to capture the variability within the population of clasts and determine the extent of fluid-assisted alteration in 75 �deformation, we may require a larger sample size similar to the study by Yonkee et al. (2013) in deformed diamictites. Figure 1. Seine metaconglomerates at a variety of strain magnitudes. A) Low strain ($1 CAD for scale). B) Moderate strain (lens cap for scale). C) High strain with carbonate alteration ($1 CAD for scale) Table 1. Results of XRF analysis. The compositional signatures of each clast type are highly variable. In all cases, the mean concentrations of elements are indistinguishable between low and high strain when the standard deviation is taken into account. References Czeck, D. M., Fissler, D. A., Horsman, E., and Tikoff, B., 2009. Strain analysis and rheology contrasts in polymictic conglomerates: an example from the Seine metaconglomerates, Superior Province, Canada. Journal of Structural Geology 31, 1365-1376. Newman, J., Mitra, G., 1993. Lateral variations in mylonite zone thickness as influenced by fluid-rock interactions, Linville Falls fault, North Carolina. Journal of Structural Geology 7, 849-863. Selverstone, J., Morteani, G., Staude, J.M., 1991. Fluid channelling during ductile shearing; transformation of granodiorite into aluminous schist in the Tauern Window, Eastern Alps. Journal of Metamorphic Geology 9, 419-431. Yonkee, W. A., Czeck, D. M., Nachbor, A., Barszewski, C. B., Pantone, S., Balgord, E., and Johnson, K. R., 2013. Strain accumulation and fluid-rock interaction in a naturally deformed diamictite, Willard thrust system, Utah (USA): Implications for crustal rheology and strain softening. Journal of Structural Geology 50, 91-118. Yonkee, W.A., Parry, W.T., Bruhn, R.L., 2003. Relations between progressive deformation and fluid-rock interaction during shear-zone growth in a basement cored thrust sheet, Sevier Orogenic Belt, Utah. American Journal of Science 303, 1-59. 76 �SEDIMENTOLOGY OF A PRE-VEGATATION PROGRADING DELTAIC ASSEMBLAGE: THE MESOPROTEROZOIC KAMA HILL AND OUTAN ISLAND FORMATIONS, ONTARIO JONES, Robyn and FRALICK, Philip Water Resource Science, Department of Geology, Lakehead University, ON, Canada, rjones2@lakeheadu.ca This study focused on the Kama Hill and Outan Island Formations of the Sibley Group located approximately 30 to 130 km east and northeast of Thunder Bay, Ontario. A lack of associated extrusive igneous rocks has resulted in an imprecise age for the unit of between approximately 1460 to 1339 Ma (Rogala, 2003; Franklin, 1978). A paleomagnetic pole position for the Sibley plots at 1400 Ma on the apparent polar wander path (Robertson, 1973) and further defines its age. The principle objective of this study was to more precisely understand the depositional environment of this section of the Sibley Group. Cheadle (1986) interpreted the Kama Hill Formation as a mud-flat environment, but Rogala (2003) and Rogala et al. (2007) believed it to be a prograding deltaic assemblage. Previous studies have found evidence of desiccation cracks in outcrop, as reported by Cheadle (1986) in the lower Kama Hill Formation. While other studies have interpreted the depositional environments as deltaic to fluvial; with fining upwards sequences as channel-fill deposits and coarsening upwards sequences as sub-aqueous distributary mouth bars (Rogala, 2003; and similar to Haszzeldaine, 1984). The lithofacies associations in this study were constructed based on four cored drill-holes within the Nipigon Plate and currently reside in the MNDM Conmee core library. Along with drill-core logging of the 340m thick interval, multiple thin sections were prepared to analyze the microstructures and grain arrangements. The drill cores were cross-correlated and divided into different lithofacies associations (LA) based on grain-size and primary sedimentary structures: fine-grained, silt-rich, rippled sandstone LA; cross-stratified LA; massive medium-grained sandstone LA; fine-grained sandstones LA; silt-rich, very fine-grained sandstone LA, and upper fine-grained LA. The lower two thirds of the succession studied consists of two coarsening- and thickening-upward sequences, with the lower one terminating in distributary mouth bar sediments and the upper one overlain by a fluvial floodplain-channel assemblage. Finally, there are desiccation cracks within the interval studied; however, they are located within areas in the upper third of the core interpreted as subaerial floodplain, where a sediment moisture deficit could easily occur. Also, the lower contact with the Rossport Formation was difficult to determine in some core and as the Rossport has mud cracks it may have led to confusion in the past. The most striking difference between this delta and modern examples is that the upper distributary mouth bar is mostly composed of massive, thick sandstones that appear to have been deposited by slurry flows down the distributary channels, and may be related to the lack of vegetation. This study gives unique insight into the primary structures of deltas, the role that flora and fauna play within deltaic systems, and has determined that the Kama Hill and Outan Island Formations were formed in deltaic environments. 77 �Figure 3: He-02-02 depth; 166.3 to 170.8 m; nice ripple lamination with numerous mud drapes present, part of the fine grained lithofacies association. Figure 1: NB-97-04 depth: 489.75-495.64 m; shows possible mud crack located within the floodplain assemblage. Figure 2: NB-97-04: 692.00 to 697.87m; mud-chip conglomerate between rippled sandstones located in the silt-rich lithofacies association. References Cheadle B. A. 1986. Stratigraphic and sedimentation of the Middle Proterozoic Sibley Group, Thunder Bay District, Ontario. Unpublished PhD. thesis, the University of Western Ontario, London, 434p. Franklin, J.M., 1978. The Sibley Group, Ontario. In, Ed. By R.K. Wanless and W.D. Loveridge. Geological Survey of Canada, Paper 77-14, 31-34. Haszeldine, R.S. 1984. Muddy deltas in freshwater lakes, and tectonism in the Upper Carboniferous Coalfield of NE England. Sedimentology, 31, 811-822. Robertson, W.A., 1973. Pole position from thermally cleaned Sibley Group sediments and its relevance to Proterozoic magnetic stratigraphy. Canadian Journal of Earth Sciences. 10, 180-193. Rogala, B. 2003. The Sibley Group: A lithostratigraphic, Geochemical and Paleomagnetic Study. Master’s Thesis, Geology Department, Lakehead University, Thunder Bay. Rogala, B., Fralick, P.W., Heaman, L.M., and Metsaranta, R., 2007. Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada. Canadian Journal of Earth Sciences, 44, 1131-1149. 78 �Mineralogy and petrology of the diamondiferous Madonna Dyke, Marathon, ON KOZLOWSKI, Alexandra1, and ZUREVINSKI, S.E.1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, akozlows@lakeheadu.ca 1 Diamondiferous intrusive lamprophyric dykes have been identified occurring just North of Lake Superior, near Marathon, Ontario. The Madonna Dyke is a diamondiferous occurrence that lies within the Superior Province within a region host to multiple alkalic and carbonatitic complexes, many of which are linked to the intrusive activity resulting from the Midcontinent rifting event at 1.1 Ga. The Madonna Dyke is a hypabyssal rock with medium- to fine-grained phenocrysts of pseudomorphed olivine, spinel, pyroxene, and amphibole set in a dark green to black altered groundmass of mainly calcite after melilite, REE-poor apatite, phlogopite and spinel. Pseudomorphed olivine occurs as microphenocrysts, phenocrysts and rare macrocrysts replaced by serpentine, magnetite and calcite. A few fresh olivine macrocrysts show mantle compositions ranging from Fo91 to Fo92. Clinopyroxenes are aluminous diopside with Al2O3 ranging from 3.11 to 14.47 wt.%. Groundmass mica show kinoshitalite – phlogopite compositions with up to 4 wt.% BaO and 20.9 wt.% Al2O3. Spinel-group mineral compositions follow Magnetic Trend #2 – the Titanomagnetite Trend, where spinels range in composition from aluminous magnesian chromites to titanian magnesian chromites to titanian chromites to members of the ulvöspinel-magnetite series. Spinel-group minerals occur as red chromium spinel phenocrysts to macrocrysts with magnesium-rich cores and iron-rich rims, often associated with olivine phenocrysts and macrocrysts. They also occur as fine-grained opaque groundmass titanomagnetites with altered cores, and as reaction products forming a necklace texture around olivine. Atoll spinels are present. Although the Madonna Dyke shows some textural and petrogenetic features of kimberlites, the mineralogy, including the presence of calcite after melilite and amphibole, are analogous with an ultramafic lamprophyre of Alnöitic affinity. Figure 1: Regional map of the area north of Lake Superior between Terrace Bay and Marathon showing the alkaline complexes and structures of the area (after Smyk et al., 1993). The location of the Madonna Dyke is denoted by the red star. 79 �Figure 2: Microphotographs of the Madonna Dyke. (A) SEM-BSE image with red arrow pointing to an atoll spinel with a resorbed spinel core, calcite lagoon and magnetite rim. (B) Transmitted light PPL image of a pseudomorphed olivine macrocryst and reddish-brown chromium spinels. 80 �Preliminary Report on the Palynology of the Gervais Formation (Pleistocene), Red Lake County, Minnesota Timothy J. Kroeger Center for Environmental, Economic, Earth and Space Studies, Bemidji State University, 1500 Birchmont Dr. NE, Bemidji MN 56601. tkroeger@bemidjistate.edu The Gervais Formation (Harris and others, 1974) is the oldest glacial till stratigraphic unit exposed in northwestern Minnesota. The Gervais is exposed just above river level in a cutbank of the Red Lake River about 2 km northwest of the city of Red Lake Falls. The Gervais is described as very dark gray, unbedded, silty, clay loam that contains few pebbles and cobbles. The outcrop sampled is unusual because there are abundant wood fragments, small logs, insects, and mollusk fragments—concentrated in the lower portion of the formation. Reported radiocarbon dates for the organic materials are: >39,000, >39,900 (wood), >46,900 (coleopterous material) (Shotten and others, 1975). An Early Wisconsinan or pre-Wisconsinan age is suggested for the formation (Harris and others, 1974). Most of the pebbles within the Gervais Formation are carbonate rock although a few pebbles of coal were found. Harris and others (1974) suggested that the till of the Gervais Formation was derived largely from glacially modified lacustrine and fluvial sediment and the silts and clays were probably locally derived. Five samples for palynological analysis were collected from the Gervais Formation. The samples were macerated using standard palynological extraction techniques (Doher, 1980). Palynomorph bearing residues were mounted on microscope slides using glycerine jelly as a mounting medium. All five of the samples from the Gervais Formation proved to be palyniferous. Palynomorph preservation ranges from excellent to moderately good. Algal palynomorphs are common, including Pediastrum, multiple forms of peridinioid dinoflagellate cysts, chorate dinoflagellate cysts, and several additional palynomorph morphologies that are most likely produced by algae. Some of the algal forms are more typical of marine or marginal marine environments. Very few spores of ferns and mosses are present; they include spores that were probably produced by Sphagnum. Conifer grains are relatively common and include grains morphologically similar to Pinus, Picea, Abies, and Thuja. Angiosperm pollen is also common. In addition to pollen forms that would be typical of Pleistocene environments, angiosperm pollen forms that are stratigraphically restricted to the Upper Cretaceous and Paleogene are also present. Included are the form taxa, Aquilapollenites sp., Tricolpites microreticulatus, and Retitrescolpites anguloluminosus. Overall, there is a mixture of palynomorphs that were most likely locally produced and palynomorphs that are reworked from Upper Cretaceous and Paleogene rocks. There is little obvious difference in preservation between grains that may represent local palynomorph production versus those that are reworked. The presence of pollen indicative of Upper Cretaceous and Paleogene rocks and the abundance of peridinioid and chorate dinoflagellate cysts strongly suggests that Upper Cretaceous and Paleogene rock materials were entrained by the glacier; such rocks are common to the west and northwest of the Red Lake Falls area in North Dakota and Saskatchewan where Upper Cretaceous and Paleocene aged marine and terrestrial rock units are exposed or subcrop beneath glacial materials. The clay-rich lithology of the Gervais Formation is also consistent with glacial transport of the shales and mudrocks from a northwesterly source area. 81 �References Cited Doher, L.I., 1980, Palynomorph preparation procedures currently used in the paleontology and stratigraphy laboratories, U.S. Geological Survey: U.S. Geological Survey Circular 830, 29 p. Harris, K.L., Moran, S.R., and Clayton, L., 1974, Late Quaternary stratigraphic nomenclature Red River Valley, North Dakota and Minnesota: Miscellaneous Series 52, North Dakota Geological Survey, 47 p. Shotton, F.W., Williams, R.E.G., and Johnson, A.S., 1975, Radiocarbon 1975, Birmingham University Radiocarbon Dates IX: Radiocarbon, v. 17, no. 3, p. 255-275. 82 �Giant Domes of the Mosher Carbonate, Steep Rock, Ontario Kurucz, Sophie and Fralick, Philip Department of Geology, Lakehead University, Thunder Bay, ON, skurucz@lakeheadu.ca The Giant Dome Lithofacies of the Mosher Carbonate is located within the Steep Rock Group, 5 km north of Atikokan in northwestern Ontario. The Mosher carbonate has been studied for over a century and is one of the most well preserved Archean carbonate sequences in the world (Grotzinger, 1989). The giant domes that constitute the uppermost 70m of the 500m thick carbonate sequence are referred to as the Elbow Point Member. The giant domes are meter-sized, elongate in shape, and are internally composed of alternating crystal fan fabric and cuspate and net-like microbialite fabric. Crystal fan fabric consists of centimeter to decimeter tall radiating fans that have been argued to be originally aragonite that precipitated directly on the seafloor. Microbialites have been described by Sumner (1997) as being composed of draping, mat-like laminae, vertically oriented microbial support structures, and cement-filled voids. While both microbialite and crystal fan fabrics are common in the late Neoarchean to early Proterozoic, their occurrence with one another, and even as alternating lithologies, may be isolated to the Neoarchean. Therefore, the environmental factors that controlled the development of the giant domes, and their unusual internal composition, is not well understood. The image on the left shows the giant domes of the Mosher Carbonate with a hat for  scale. The image on the right shows the lithologies seen within the giant domes and their  interbedded nature. Polished slab is 10cm wide.  The giant domes have been interpreted to have formed in a rimmed platform environment, where a fluctuating redox boundary resulted in the alternation in precipitation of aragonite and calcite (Fralick and Riding, 2015). Major element geochemical data suggests that the crystal fan fabric contains lower concentrations of Mn and Fe than adjacent fenestrate microbialite fabric, while Mg concentrations do not show any change in concentration. This trend may be a good indication that the primary mineralogy of the crystal fan fabric and fenestrate microbialite fabric is different. Similarly, cements that are interstitial and mantle the crystal fans contain relatively higher concentrations of Mg than the crystal fans themselves, indicating that the deposition of the fans and void-filling cements was not synchronous. 83 �Mn Sample S1 1.2 Mn net‐like fabric Mn crystal fan fabric Fe Sample S1 0.8 Fe net‐like fabric Fe crystal fan fabric 0.7 1 0.6 0.5 Fe % Mn % 0.8 0.6 0.4 0.3 0.4 0.2 0.2 B.D. 0 47.5 48 48.5 49 49.5 50 0.1 B.D. 0 50.5 47.5 Ca % 48 48.5 49 Ca % 49.5 50 50.5 Atomic % of Mn and Fe vs. Ca for adjacent net‐like microbialite and crystal fans. The net‐like  fabric displays higher concentrations of both Fe and Mn relative to the associated crystal fan  fabric. B.D.=below detection. There are also features that suggest periodic subaerial exposure of the giant domes. A wellpreserved desiccation surface with a typical polygonal crack pattern can be seen on a sample that in cross-section displays cuspate fenestrate fabric. The association of the support structures with the cracks that are expressed on the surface, leads to consideration of a desiccation related process in their formation. While there is an otherwise complete lack of preserved desiccation surfaces, Fe-rich red-brown surfaces that intervene between the interbedded layers of the giant domes at periodic intervals may represent desiccation or subaerial exposure surfaces that were later destroyed. Lastly, Fe- and carbon-rich dissolution surfaces separate some adjacent lithologies. These surfaces are continuous and can be seen to mark the boundary between crystal fan fabric and fenestrate microbialite fabric within certain samples. They are composed of microcrystalline quartz and zoned dolomite. The dissolution surfaces may represent hardgrounds that formed as a result of a hiatus in sedimentation and/or as result of early cementation and provided preferential fluid pathways for later silicification. References Fralick, P., and Riding, R., 2015. Steep Rock Lake: Sedimentology and geochemistry of an Archean carbonate platform. Earth Science Reviews, v. 151, p. 132-175. Grotzinger, John P., 1989. Facies and evolution of Precambrian carbonate depositional systems: Emergence of the modern platform archetype. SEPM Special Publication No. 44, p. 79106. Sumner, Dawn Y., 1997. Late-Archean calcite-microbe interactions: Two morphologically distinct microbial communities that affected calcite nucleation differently. PALAIOS, vol. 12, no. 4, p. 302-318. 84 �A COMPARISON OF BARABOO-INTERVAL (LATE PALEOPROTEROZOIC) IRON-FORMATION, SOUTHERN WISCONSIN LAMB, Matthew T1. and KINGSBURY STEWART, Esther2 1 Department of Geography, Geology, and Environmental Science, UW-Whitewater, 120 Upham Hall, 800 Main St., Whitewater, WI, 53190. 2 Wisconsin Geologic and Natural History Survey, 3817 Mineral Point Road, Madison, Wi 53705 The Freedom Formation is part of the Baraboo-interval sediments that were deposited after ca. 1.71 Ga and later deformed by the 1.6 Ga Mazatzal Orogeny (Holm et al., 1998; Medaris et al., 2003). In the Baraboo Range of Sauk County, Wisconsin, the Freedom Formation conformably overlies the Seeley Slate, which conformably overlies the Baraboo Quartzite (Weidman, 1904). The Freedom Formation is only present in the subsurface and is known from 100-year-old drill cores that were recovered along the hinge of the Baraboo syncline (Sauk County). In Sauk County, the Freedom Formation is divided into two members: a lower iron-rich unit made up of finely-interbedded to interlaminated, hematite-rich micrite, chert, and silt- to clay-sized siliciclastic material and an upper member comprised of dolomitic micrite. Drill cuttings from three wells between 5 and 50 miles (8 to 80 km) east of the Baraboo Range recovered hematite- and magnetite-rich fine-grained sediments and quartzite, suggesting that the Freedom Formation is present, along with Baraboo-interval quartzite, in the subsurface across a broad area in southern Wisconsin. Additionally, a drill core recovered in 2015 by the Wisconsin Geologic and Natural History Survey encountered iron-formation beneath the Cambrian Elk Mound Group in the subsurface of Dodge County, some 50 miles (80 km) east of Baraboo. The purpose of this research is to compare the Freedom Formation recovered from the Baraboo Range, Sauk County to the iron-formation recovered from drill core in Dodge County in order to better constrain correlation of Baraboo interval sediments. To accomplish this, we identify and log primary lithofacies from the Sauk and Dodge County drill cores. We also analyze thin sections from each lithofacies for mineralogical similarities. Lastly we collect elemental concentrations using a handheld Thermo Fisher Niton XL3t GOLDD+ X-ray florescence (XRF) analyzer on each core at 1 foot intervals to compare geochemistry with lithofacies observed through core logging. We identify three main lithofacies in core from both location: 1) gray to white, apparently massive carbonate, 2) laminated carbonate and gray siltstone, 3) laminated red (hematite-rich) siltstone, gray siltstone, and white to gray carbonate. In thin sections we observe a primary mineral assemblage of fine-grained carbonate, quartz, hematite, chlorite, and magnetite. The relative abundance of these minerals varies by lithofacies. Recrystallization is evident in thin section and core from both locations, but is more pronounced in the Dodge County location. Elemental XRF analyses broadly correlate with lithofacies for the Sauk County location, but the eight feet of iron-formation recovered from Dodge County did not provide enough material to meaningfully compare XRF analyses from both locations. From observation of core and thin section, we show that the Precambrian iron-formation recovered from drill core in Dodge County has similar lithofacies and mineralogy to the Freedom Formation of Sauk County. Based on these similarities, we conclude that the iron-formation from Dodge County is correlative with the Freedom Formation of Sauk County. Therefore, the Baraboo-interval sediments include an ironformation that is laterally continuous over a more than 1,000 km2 area in southern Wisconsin. 85 �Figure 1. A. Location map. Dots show location of cores that recovered iron-formation in Sauk and Dodge counties. B. Core photographs showing examples of iron-formation recovered from each location. Holm, D., Schneider, D.A., Coath, C., 1998. Age and deformation of Early Proterozoic quartzites in the southern Lake Superior region: implications for extent of foreland deformation during final assembly of Laurentia. Geology 26, 907–910. Medaris, L.G., Singer, B.S., Dott, R.H., Naymark, A., Johnson, C.M., and Schott, R.C., 2003. Late Paleoproterozoic Climate, Tectonics, and Metamorphism in the Southern Lake Superior Region and Proto-North America: Evidence from Baraboo Interval Quartzites. The Journal of Geology. 111, 243-257. Weidman, 1904. The Baraboo Iron-Bearing District of Wisconsin. Wisconsin Geological and Natural History Survey Bulleting No. 13, Economic Series No. 8, 190 pp. 86 �Millennial-scale shoreline bluff retreat rates in the western arm of Lake Superior LAMBERT, Crystal A., and SWENSON, John B. Department of Earth and Environmental Sciences, University of Minnesota Duluth, Duluth, MN 55812 In the far western arm of Lake Superior, shoreline bluff retreat is a significant problem for landuse planners and policymakers, who require knowledge of the bluff retreat rate. Previous studies of shoreline bluff retreat employ careful differencing of aerial photos, in combination with direct field observations (e.g., Johnson and Johnston, 1995; Swenson et al., 2006; NRPC; 2012). The large variability in retreat rate reported in these decadal-scale studies reflects the intrinsically stochastic nature of hillslope and coastal processes. By extension, these short-term studies downplay the importance of the well-documented, long-term rise in relative lake level (RLL) driven by differential uplift of the Lake Superior basin (e.g., Mainville and Craymer, 2005). On longer (geologic) timescales, this RLL rise is the fundamental driver for bluff retreat. We developed a simple geometric-kinematic model of bluff retreat rate that addresses explicitly the role of RLL rise. Our model is motivated by the observation that the nearshore lake bottom is a surface of non-deposition and erosion (Kemp et al., 1978; Thomas and Dell, 1978). In the sequence-stratigraphic nomenclature, the lake bottom is a wave-cut transgressive ravinement surface, sensu Posamentier and Allen (1999). On the north shore, where the bluff material is Precambrian bedrock, this transgressive surface shows a very sparse cover lag of cobbles and boulders atop a clean (eroded) bedrock surface; on the south shore, where the bluffs are composed of unconsolidated glacial till, the scoured surface is located lakeward of a zone of littoral drift. With knowledge of the rate of RLL rise (Vz), the dip () of this transgressive surface constrains the long-term bluff retreat rate (Vx) (Fig. 1). Model predictions of bluff retreat rate hold for multi-century to millennial timescales and, as such, remove much of the variability that plagues short-term measurements. One limitation of our model is that it is tied explicitly to knowledge of the long-term rate of RLL rise, which is only moderately well constrained in the western arm of Lake Superior (Mainville and Craymer, 2005; Yu et al., 2013). We applied our model to the extreme western arm of Lake Superior. With Duluth as the western boundary, our study area extended eastward to the mouths of the Knife and Brule rivers on the north and south shores, respectively. Our study area is well suited to analyze the importance of lithology in determining long-term bluff retreat rate: Bluffs on the north shore are composed of Keweenawan-aged igneous rocks of various lithologies, whereas those on the south shore are composed of relatively homogeneous Holocene-aged glacial till. Common to both coastlines was the history of RLL rise. Recent LiDAR data (NOAA, 2010; Fig. 1) provide highresolution bathymetry for the north shore; on the south shore, we digitized the low-resolution NOAA navigation charts. Preliminary results of our modeling efforts suggest that millennial-scale bluff retreat rates on the south shore are as much as an order of magnitude greater than those of the north shore; this result is not surprising, given the significant difference in bulk lithology between the coastlines. Using a widely accepted RLL rise rate of 2.5 mm·a-1 in the Duluth area (Mainville and Craymer, 2005), bluff retreat rates on the north shore average approximately 5 cm·a-1, with considerable variability that correlates broadly with lithology. In map view (Fig 1), this variability manifests itself as resistant headlands and erodible bays with tall escarpments. Thick felsic volcanic units show the greatest bluff retreat rates (>7 cm·a-1), whereas mafic intrusives, e.g. the Endion Sill, are most resistant to erosion, e.g. the Endion sill. The correlation between retreat rate and 87 �lithology is reduced by ‘buttressing’ effects, i.e. a strong mafic intrusive unit protecting or supporting an adjacent felsic unit. Figure 1. (Left) Cartoon illustrating relationship between transgressive surface, rate of RLL rise, and bluff retreat rate. (Right) Map of eastern Duluth shoreline, showing nearshore LiDAR-derived bathymetry (NOAA, 2010) and mapped bedrock units (Green and Miller, 2008). Relatively low quality bathymetric data hinder somewhat the application of our model to the south shore, which consists of unconsolidated till. The relatively smooth coastline—lacking bays and headlands—and the generally planar offshore bathymetry together suggest little variability in bluff retreat rate with position along the coast. The primary difficulty in applying our model here is identifying the closure depth of the littoral zone—the region of active longshore sand transport—and, by extension, the nearshore extent of the transgressive surface. Our preliminary best estimate for the slope of the transgressive surface yields a millennial-scale bluff retreat rate of between 50 and 100 cm·a-1. Our model predictions for long-term bluff retreat rates on both coastlines are within the ranges of values reported from decadal-scale studies (Johnson and Johnston, 1995; Swenson et al., 2006). REFERENCES Green, J.C., and Miller, J.D., Jr. (2008) Bedrock geology of the Duluth quadrangle, St. Louis County, Minnesota. Minnesota Geological Survey Miscellaneous Map M-182, scale 1:24,000. Johnson, B. L., and Johnston, C. A. (1995) Relationship of lithology and geomorphology to erosion of the western Lake Superior coast. Journal of Great Lakes Research, 21(1), 3-16. Kemp, A.L.W., Dell, C.I., and Harper, N.S. (1978) Sedimentation rates and a sediment budget for Lake Superior. Journal of Great Lakes Research, 4, 276-287. Mainville, A., and Craymer, M. R. (2005) Present-day tilting of the Great Lakes region based on water level gauges. Geological Society of America Bulletin, 117(7-8), 1070-1080. National Ocean and Atmospheric Administration (NOAA) (2010) Data verification report Lake Superior bathymetric LiDAR. Northwest Regional Planning Commission (NRPC) (2012) Lake Superior South Shore Bluff Recession Rate Study. Posamentier, H.W., Allen, G.P. (1999) Siliciclastic sequence stratigraphy: concepts and applications. SEPM Concepts in Sedimentology and Paleontology, no. 7, 210 p. Swenson, M. J., Wu, C. H., Edil, T. B., & Mickelson, D. M. (2006) Bluff recession rates and wave impact along the Wisconsin coast of Lake Superior. Journal of Great Lakes Research, 32(3), 512-530. Thomas, R.L., Dell, C.I. (1978) Sediments of Lake Superior. Journal of Great Lakes Research, 4(3-4), 264-275. Yu, SY., Colman, S.M., and Milne, G.A. (2013) Radiocarbon Dating of Basal Peats Supports Separation of Lake Superior from Lakes Michigan-Huron about 1250 years ago. Earth and Planetary Science Letters, 375, 319-325. 88 �Volcanological, Geochemical, and Geochronological Comparisons of the Gafvert Lake Sequence in Minnesota and Shebandowan Assemblage in Ontario LODGE, Robert W.D. 1, PIGNOTTA, Geoffrey S.1, GÉLINAS, Brigitte, R.2, SCHWIERSKE, Kelly L.1, and HUDAK, George J.3 1 Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI 2 Department of Geology, Lakehead University, Thunder Bay, ON 3 Precambrian Research Center, Minnesota Natural Resources Research Institute, University of Minnesota-Duluth The magmatic and tectonic evolution of Archean greenstone belts is complex and rocks are formed and modified by episodes of pre-deformation volcanic and plutonic events, syndeformation orogenic magmatism and sedimentation, and post-orogenic plutonism. In addition to contributing to the understanding Archean magmatic and tectonic processes, unravelling the different episodes of the formation of these greenstone belts have important metallogenic implications. Many of the greenstone belts in the Superior province are host to a variety of base and precious metal deposits in a variety of deposit types. The well exposed Neoarchean greenstone belts in the westernmost part of the Wawa-Abitibi terrane present an excellent opportunity to study each of these phases of formation. This study examines and compares the volcanic and plutonic rocks that are part of the syndeformational Timiskaming-type assemblages in the Vermilion and Shebandowan greenstone belts in Minnesota and Ontario (Figure 1). The Gafvert Lake sequence of the Vermilion greenstone belt and the Shebandowan assemblage of the Shebandowan greenstone belt contain coeval volcanic assemblages that were deposited in D2 transtensional basins formed during the accretion of the Wawa-Abitibi terrane to the rest of the Superior Province around 2690 Ma (Lodge et al, 2013). The interpreted geodynamic setting for these assemblages is similar to that of the Timiskaming group in the Kirkland Lake area that hosts several world-class lode gold deposits. The parallel settings between these areas have metallogenic implications and therefore it is critical to fully understand the characteristics of the Timiskaming-like assemblages in the western Wawa-Abitibi terrane to understand their potential precious metal endowments. Newly acquired geochemical and geochronological data from the Gafvert Lake sequence and published geochemical data from the volcanic and plutonic assemblages in the Shebandowan greenstone belt (Gélinas et al. 2016) have allowed a more detailed comparison of the two assemblages. The most notable difference is that these two volcanic-plutonic deposits are compositionally distinct despite being coeval in formed in similar geodynamic settings. The Gafvert Lake sequence is notably more felsic with calc-alkalic, quartz-phyric volcanic and plutonic rocks dominating the strata. In contrast, the volcanic and plutonic rock in the Shebandowan assemblages are calc-alkalic to alkalic in composition with predominately andesitic volcanic compositions and monzonite to syenite plutonic suites. Geochemically, the Shebandowan assemblage are notably more enriched in their LREE and other incompatible elements relative to the Gafvert Lake Sequence, but are still significantly less alkalic than the volcanic and plutonic rocks that are found in the Kirkland Lake region. These compositional variations in the magmatic rocks are likely contributing to the relative differences in gold prospectivity in these greenstone belts. 89 �Figure 1: Regional geologic map of the Vermilion and Shebandowan greenstone belts in Minnesota and Ontario. Inset map shows location of these greenstone belts in relation to the western part of the Wawa-Abitibi terrane if the Superior Province. VGB, SGB, WGB, and MGB are referring to the Vermilion, Shebandowan, Winston Lake, and Manitouwadge greenstone belts, respectively. Figure is modified from Lodge et al. (2013). References Gélinas, B.R., Lodge, R.W.D., Gibson, H.L., 2016. Characterization of the Mineralization and Alteration at Tower Mountain, Conmee Township, Shebandowan Greenstone Belt, Ontario Ontario Geological Survey, Miscellaneous Release - Data 330. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., Jirsa, M.A. and Hamilton, M.A., 2013, New U–Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince,Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province. Precambrian Research (235), 264-277. 90 �The Eagle East Magmatic Nickel-Copper Discovery MAHIN, Robert A. and BEACH, Steven T. Eagle Mine LLC/Lundin Mining Corp., 200 Echelon Dr., Negaunee, MI 49866 In June 2015, Lundin Mining Corporation announced the discovery of high grade magmatic nickel-copper mineralization similar in style to its 100%-owned Eagle deposit (5.2 MT reserves @ 3.12% Ni, 2.56% Cu), located in Michigan’s Upper Peninsula. The new zone, Eagle East, is approximately two kilometers east of the Eagle mine and 950 meters deep. Assay highlights include:  30.85 meters of massive (MSU) and semi-massive sulphide (SMSU) with 5.23% Ni, 8.74% Cu, and 9.49 g/T combined Pt-Pd-Au. Included in this interval was 16.38 meters of MSU with 6.70% Ni, 13.59% Cu, and 16.43 g/T combined Pt-Pd-Au (DDH 14EA331I; 1,139.85m to 1,170.70m).  23.85 meters of MSU and SMSU with 5.34% Ni, 4.41% Cu, and 4.37 g/T combined PtPd-Au, including 9.83 meters of MSU with 8.16% Ni, 7.10% Cu, and 5.14 g/T combined Pt-Pd-Au (DDH 14EA331H; 1,142.18m to 1,166.03m). Eagle and the new Eagle East zone are hosted by separate 1Ga Midcontinent Rift-related ultramafic intrusions that intrude the Baraga Basin, a gently dipping syncline of Paleoproterozoic slates, shales, and greywackes. The intrusions are interpreted to be part of the plumbing system for the Keweenaw flood basalts that host the renowned native copper deposits. Massive sulphides at the Eagle East discovery consist of generally greater than 90% pentlandite, chalcopyrite, and pyrrhotite. Semi-massive sulphides consist of net-textured pentlandite, chalcopyrite, pyrrhotite, and silicates. Also present are magmatic breccias containing inclusions of basal cherts/quartzites and granitic Archean basement. The mineralisation is filling a subhorizontal chonolith or conduit. Similar to the Eagle deposit, the new zone also contains horizontal MSU sills that intrude from the conduit laterally out into the surrounding metasediments. The discovery was made by applying a dynamic conduit model of magmatic sulfide mineralisation in which multiple deposits should exist along a main feeder dike. The nickel and copper tenors of the uneconomic mineralisation of Eagle East (aka Yellow Dog Peridotite) were known to be higher than that of Eagle. If mineralisation was found, the tenors predicted it could be higher grade than Eagle. The drilling program involved attempting to trace the feeder dyke of from the lower portion of Eagle East into undrilled territory using directional drilling. Critical to the success of the program was the ability to use Devico directional drilling to steer each hole to a specific target and using of a single parent drill hole as a platform for multiple directional kick-off daughter holes. Semi-massive sulphides and magmatic breccias occupying the core of a 15 to 40 meter wide peridotite dyke were intersected early in the program and spurred further drilling along strike. Semi-massive sulphides continued to be intersected for over a 250 meter strike length to where thicker SMSU and high-grade MSU were encountered. The discovery drilling was made by using just two parent holes and a total of 15 kick-offs. The high grade zone has been traced for over a 60 meter strike length and is open in all directions. The project is still in an exploration and drill-out phase and the full extent of Eagle East has not been determined. 91 �Preliminary groundwater age and chemistry data from cover overlying Duluth Complex Ni-Cu-PGE deposits, NE Minnesota MANNING, Andrew H.1, WANTY, Richard B.2, and MORRISON, Jean M.2 1 Central Mineral and Environmental Resources Science Center, U.S. Geological Survey, Denver, CO 2 Crustal Geophysics and Geochemistry Science Center, U.S. Geological Survey, Denver, CO The U.S. Geological Survey initiated a project in 2015 aimed at evaluating geochemical exploration methods for covered deposits in the northern Midcontintent Rift. A first round of soil and groundwater samples were collected in September from unconsolidated material overlying the Spruce Road, Wyman Creek, and Skibo deposits with the objective of determining effective sampling methodologies and characterizing the general geochemical signature of these deposits within the shallow cover. Twenty-seven water samples were collected, including 21 groundwater samples and 6 surface water samples. Groundwater samples were collected from minipiezometers (plus one spring) having depths of mainly 0.5 to 1.5 m installed mostly in areas of suspected local groundwater discharge to wetlands. Five piezometers were installed along a flow-line-parallel transect ending at Filson Creek overlying the Spruce Road deposit. All samples were analyzed for major and trace element chemistry and stable isotopes of water (2H and 18O). Ten samples were also analyzed for groundwater age tracers, including dissolved noble gases (He, Ne, Ar, Kr, and Xe), 3He/4He ratio, tritium, and/or chlorofluorocarbons (CFC11, -12, and -113). Age data were collected along with chemistry to survey the range of groundwater residence times in the cover and to investigate chemical evolution and metal transport processes along flow paths. Groundwater conditions in the site area presented several challenges to obtaining wellconstrained age determinations. For example, high DOC concentrations led to oxygen depletion in CFC sample bottles prior to analysis, resulting in CFC degradation and erroneously old computed CFC ages. However, difficulties with individual dating methods were overcome by utilizing multiple age tracers together, allowing samples to be sorted into the following age categories with reasonable confidence: <0.5 yr old; 0.5 to 2 yr old; 2 to 10 yr old, and 15 to 30 yr old. Water <0.5 yr old was distinguishable mainly based on heavy 18O values (-8 to -10.5‰ compared to -11 to -12‰; Levy et al., 2014) and warm noble gas recharge temperatures (9 to 15°C compared to 2 to 5°C), along with apparent 3H/3He ages of 0 yr, indicating 2015 spring or summer recharge. Water >2 yr old was distinguishable mainly based on apparent 3H/3He ages, with 15-30 yr old water containing substantial terrigenic 4He (produced from U-Th decay) and resulting low 3He/4He ratios 30 to 60% below atmospheric values. Tritium concentrations range from 7 to 10 TU, suggesting samples contained little pre-modern water >60 yr old. Groundwater Cu concentrations range from <0.5 to 150 g/L (median of 5.2 g/L) and Ni from <1 to 348 g/L (median of 5.5 g/L), similar to concentrations reported for the site area in previous studies (Siegel and Ericson, 1980; Miller et al., 1992). Concentrations are generally greater at Spruce Road, where Cu and Ni medians are 21 and 19 g/L, respectively. Groundwater Cu and Ni are roughly correlated with soil Cu and Ni concentrations at the piezometer location/depth, as expected. However, a relatively well defined negative correlation is apparent between Cu concentration and pH, as well as Ni and pH, suggesting that pH is another important control on Cu and Ni mobility in the groundwater system (Figs. 1A, B). Measured pH ranges from 5.7 to 8.6 (median of 7.2), and dissolved Cu and Ni concentrations are commonly observed to decrease as pH increases above the acidic range in mineralized settings 92 �due to adsorption on negatively charged mineral surfaces. Figure 1C suggests that pH generally increases with age along flow paths, probably due to weathering of abundant mafic minerals (Fig. 1D). This increasing pH is the likely cause of an apparent decrease in Cu and Ni concentrations with greater age (Figs. 1E, F). The observed negative correlation between Cu/Ni concentrations and age is significant because it may provide a limit on Cu/Ni concentrations and mobility within the deeper groundwater system in the site area, and must be taken into account in geochemical exploration approaches. It also corroborates the finding of Walton-Day et al. (1990) that discharging deeper, older groundwater in the cover may not be a primary contributor to enriched Cu and Ni concentrations in surface water in streams and wetlands in the Spruce Road area. Figure 1. (A and B) Dissolved Cu and Ni concentration vs. pH. WC = Wyman Creek, SK = Skibo, SR = Spruce Road. (C) pH vs. interpreted groundwater age. Interpreted age is approximate midpoint of assigned age range referred to in text. Power law trend line is for Spruce Road samples. (D) Dissolved Ca (solid symbols) and Mg (open symbols) concentration vs. interpreted groundwater age. (E and F) Dissolved Cu and Ni concentration vs. interpreted groundwater age. Power law trend line is for Spruce Road samples. REFERENCES Levy, Z.F., Siegel, D.I., Dasgupta, S.S., Glaser, P.H., and Welker, J.M., 2014. Stable isotopes of water show deep seasonal recharge in northern bogs and fens. Hydrological Processes, v. 228, p. 4938–4952. Miller, W.R., Ficklin, W.H., and McHugh, J.B., 1992. Geochemical exploration for copper-nickel deposits in the cool-humid climate of northeastern Minnesota. Journal of Geochemical Exploration, v. 42, p. 327-344. Siegel, D.I., and Ericson, D.W., 1980. Hydrology and water quality of the copper-nickel study region, northeastern Minnesota. U.S. Geological Survey Water-Resources Investigations Open-File Report 80-739, 87 p. Walton-Day, K., Filipek, L.H., and Papp, C.S.E., 1990. Mechanisms controlling Cu, Fe, Mn, and Co profiles in peat of the Filson Creek Fen, northeastern Minnesota. Geochimica et Cosmochimica Acta, v. 54, p. 2933-2946. 93 �Small scale microanalysis of rock and mineral textures and its relationship to mineral separation MATKO, Matthew W.1 and SCHARDT, Christian1 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 Ore deposits form in a wide variety of geologic settings, and the processes involved in the formation of these deposits are well understood. Previous research along these lines has typically focused on more macro-scale processes such as fluid migration, chemical and metal transfer, and the associated physical and chemical changes that result (Cathles, 1981; Zientek, 2012). The ideas behind these processes have since been applied to large-scale features based on field observations, laboratory experiments, and theoretical assumptions to create models for the formation of these ore deposits. However, our understanding of how small-scale properties can influence the mineralization and alteration processes in these deposits is currently insufficient. Small-scale in-situ properties such as mineral grain size and shape variation, fracture dispersal, and the interconnectivity of pore space, along with the spatial distribution of these properties, may well play an important role in how ore deposits develop. Because mineralization styles such as disseminated, net-texture, or vein-style ores are all influenced by these properties it is important that they are investigated to see what influence they have on larger-scale processes (Prince et al., 1995; Wennberg et al., 2009; Holzheid, et al., 2000; and Liu, et al., 2014). The primary goal of this project is to examine selected small-scale physical rock properties such as pore space, pore connectivity, fracture dispersal, as well as mineral grain size and orientation. Data collected from the examination of these properties will be used to create preliminary models for the physico-chemical formation of different mineralization and alteration styles. The investigation of these small-scale features will result in improved ore deposit models that take into account a greater set of formation variables. Materials were selected from a range of ore deposit types exhibiting different mineralization styles and/or alteration. Samples will be studied using a combination of analytical techniques such as X-ray Computed Tomography (xCT), Electro Pulse Disaggregation (EPD), and Mineral Liberation Analysis (MLA). xCT will be used to examine small cores taken from sample material to examine in-situ properties such as pore-space distribution, mineral grain morphology or fracture dispersal. EPD is being utilized as an alternative to traditional methods of material separation as this method is a touch-free way to liberate individual mineral grains. Electric pulses traveling through the material along zones of weakness, such as mineral grain boundaries or fractures, allow material disaggregation preferentially along these boundaries and preserve the original mineral grain morphology (Cabri et al., 2008). MLA will be performed using a scanning electron microscope (SEM) on portions of the samples disaggregated by the electro pulse technique. The MLA will be employed to characterize the degree of ore mineral separation from gangue material and other ore minerals, grain shape patterns, and mineral grain size variation. Preliminary qualitative analysis of disaggregated material under a binocular microscope indicates that the EPD technology excels at separating ore minerals from silicate material and gangue, while preserving the original crystal morphology. The separation of individual ore minerals from each other in the case of massive ore will need to be assessed using the MLA technique because visual observation through the binocular microscope is not adequate. Samples of the disaggregated material will be scanned using an optical particle analyzer to assess the preserved crystal morphology. 94 �Figure 1) A sample of chromite ore disaggregated using EPD technology (top left). Sample image of a porosity analysis from a sample run through an xCT (bottom left) A sample grain mount after an MLA has been performed (right). References Cabri, Louis J., et al. "Electric-pulse disaggregation (Epd), hydroseparation (Hs) and their use in combination for mineral processing and advanced characterization of ores." Canadian Mineral Processors, 40th Annual Meeting, Ottawa, Proceedings. Vol. 211. 2008. Cathles, L.M. (1981) Fluid flow and ore genesis of hydrothermal ore deposits: Economic Geology 75th Anniversary Volume, p. 424 – 457 Zientek, M.L. (2012) Magmatic Ore Deposits in Layered Intrusions - Descriptive Model for Reef-Type PGE and Contact-Type Cu-Ni-PGE Deposits: U.S. Geological Survey Open File Report 2012-1010, p. 48 Prince, C. M., Ehrlich, R., Anguy, Y. (1995) Analysis of spatial order in sandstones; II, Grain clusters, packing flaws, and the small-scale structure of sandstones: Journal of Sedimentary Research 65, p. 13 - 28 Wennberg, O.P., Rennan, L., Basquet, R. (2009) Computed tomography scan imaging of natural open fractures in a porous rock; geometry and fluid flow: Geophysical Prospecting 57, p. 239 – 249 Holzheid, A., Schmitz, M.D., and Timothy L. Grove, T.L. (2000) Textural equilibria of iron sulfide liquids in partly molten silicate aggregates and their relevance to core formation scenarios: Journal of Geophysical Research, 105, p. 13,555 - 13,567 Liu, P.P., Zhou, M.F., Chen, W.T., Boone, M., and Cnudde, V. (2014) Using Multiphase Solid Inclusions to Constrain the Origin of the Baima Fe–Ti–(V) Oxide Deposit, SW China: Journal of Petrology, v. 55, p. 951 - 976 95 �Quantifying Mass Fluxes of Potassium in Weathering and Metasomatism of Paleosols MEDARIS, L. Gordon Jr. Department of Geoscience, University of Wisconsin-Madison, Madison, WI 53706 medaris@geology.wisc.edu Paleosols provide significant insights into ancient climates, and weathering mass fluxes can be used to evaluate various climofunctions such as mean annual precipitation (MAP), mean annual temperature (MAT), and atmospheric pCO2. However, the chemical compositions of many paleosols have been modified by potassium metasomatism, and such modification must be taken into account when calculating mass fluxes associated with weathering. Method - The Chemical Index of Alteration (CIA) is a good measure of degree of weathering, attaining a value of 100 upon complete removal of labile oxides (Nesbitt & Young, 1982; CIA = 100 × molar Al2O3/[Al2O3+CaO*+Na2O+K2O], where CaO* is CaO in silicate minerals). It has been demonstrated that modern weathering produces chemical trends subparallel to the A-C*N side of an A-C*N-K plot (Nesbitt & Young, 1984; Fig. 1). The compositions of many paleosols are displaced from weathering trends towards the K apex, as illustrated by the saprolite in Fig. 1, which is a result of potassium metasomatism subsequent to weathering. The original CIA value (CIACALC) for the metasomatized saprolite can be determined by constructing a line from the K apex through the measured saprolite composition to intersect the weathering trend. A second line is then constructed horizontally from the point of intersection to the vertical CIA scale on the left, giving the original CIA value (Fedo et al., 1995). In an analogous manner, the pre-metasomatic K2O content of the saprolite is determined by constructing a line from the C*N apex through the CIA intersection point on the weathering trend to the A-K side of the figure, where the molar K2O:Al2O3 ratio is given (Fig. 1). The pre-metasomatic K2O content of the saprolite is then calculated from this ratio and the measured Al2O3 content of the saprolite, converting molar values to wt. % (Medaris et al., 2015). However, K2O values calculated in this manner are only minima, because after complete removal of plagioclase, further weathering with removal of K feldspar may have proceeded along the A-K side of the figure toward the A apex. Note that the difference between minimum calculated K2O contents and possible actual contents is a function of lithology (Fig. 2), where the difference between the two decreases with a decrease in original K feldspar content, e.g. progressing from granite to granodiorite to tonalite. 96 �Application - The Baraboo paleosol is a mature, 800 cm-thick weathering profile that developed by intense weathering of Baxter Hollow granite at ~1700 Ma, which resulted in complete removal of both plagioclase and K feldspar from the paleosol and corresponding absence of detrital K feldspar in the overlying Baraboo Quartzite (Driese & Medaris, 2008). With one exception, all saprolite and regolith samples are displaced from the weathering trend due to ~1460 Ma potassium metasomatism (Fig. 3). Relative to the Al2O3 content of mean granite protolith, ~30% K2O was added to the paleosol (Fig. 4), which integrated over the entire profile, amounts to 0.13 mols/cm2 K2O. The calculated amount of K2O removed by weathering, following the method described above, was ~25%, corresponding to 0.16 mols/cm2. Because both plagioclase and K feldspar were completely removed by weathering, the actual % change of K2O may be taken as equal to that of Na2O, yielding an estimate of 0.59 mols/cm2 K2O removal over the profile (Fig. 4). Atmospheric pCO2 - Calculated atmospheric pCO2 is a function, among other things, of the total mass flux of weathering. Despite the large relative difference between calculated and estimated weathering fluxes of K2O, this difference has little effect on calculated pCO2, because the flux of K2O is small compared to the combined flux of SiO2, MgO, CaO, and Na2O ( = -5.03 mols/cm2). Thus, pCO2 for calculated and estimated K2O fluxes is 18.1 vs. 19.0 × PIAL, respectively. Calculated pCO2 is much more sensitive to duration of weathering and paleosol removal by erosion than to the uncertainty in K2O flux, e.g. an increase in weathering duration from 105 to 106 years corresponds to a reduction in pCO2 from 19.0 to 1.9 × PIAL, and pCO2 prior to 10% erosion was 26.4 × PIAL, rather than a post-erosion value of 19.0 × PIAL. References Driese SG & Medaris LG Jr (2008) Journal of Sedimentary Research, v. 78, 443-457 Fedo CM et al. (1995) Geology, v. 23, 921-924 Medaris LG Jr et al. (2015) Precambrian Research, v. 257, 83-93 Nesbitt HW & Young GM (1982) Nature, v. 299, 715-717 Nesbitt HW & Young GM (1984) Geochimica et Cosmochimica Acta, v. 48, 1523-1534. 97 �Mineralogy and Petrology of the Rabbit Foot Dyke, White River, ON Metteer, S1, and Zurevinski, S.E.1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1 Diamond-bearing macrocrystic, and xenolithic rocks have been identified near the town of White River, Northwestern Ontario. Initial assessment of these rocks has led to their being classified “melnoite”, a term used to describe potentially diamondiferous ultramafic lamprophyres. Olivine occurs as two distinct phases within the rocks of the Rabbit Foot Dyke; macrocrystal olivine and groundmass olivine. Macrocrystal olivine ranges in Mg/(Mg+Fe) from Fo to Fo , while phenocrystal and groundmass olivine ranges from Fo to Fo . Phlogopite compositions range from Ba-phlogopite to tetraferriphlogopite. Spinel-group minerals commonly have chromite cores with titanomagnetite rims. Spinel compositions follow the “Magmatic trend 2”, the titanomagnetite trend. Spinel-group minerals can be observed displaying atoll textures. Spinel and perovskite are commonly spatially related, and are observed surrounding larger macrocrysts such as olivine in a necklace texture. Perovskite compositions are relatively pure CaTiO , lacking any REE concentrations. In at least one outcrop of xenolithic rock, spherical magma clasts occur up to 10cm in diameter, with cores of fragmented olivine macrocrysts. The rocks of the Rabbit Foot Dyke are in many ways analogous to kimberlite in texture and mineralogy, however, significant petrogenetic overlap with melnoites, or ultramafic lamprophyres, is evident. The macrocrystic and xenolithic rocks of the Rabbit Foot Dike are characterized here as kimberlite with melnoitic (ultramafic lamprophyre) affinity. 98 �GEOLOGY OF THE CHEROKEE LAKE AREA OF THE BOUNDARY WATERS CANOE AREA, COOK COUNTY, MN - 2015 PRECAMBRIAN FIELD CAMP CAPSTONE MAPPING Jim MILLER, Aaron Balles, Ellie Brown, Ryan Helms, Greta Penzel, Luke Smith Precambrian Research Center, University of Minnesota Duluth, Duluth, MN 55812 As a capstone mapping project for the 2015 Precambrian field camp, a crew of five students under the supervision of instructor Jim Miller conducted four days of field mapping bedrock geology in the Cherokee Lake area. This area is located in the Boundary Waters Canoe Area west of Brule Lake in Cook County, Minnesota. The area is accessible from the Caribou Trail, which heads north from Tofte to a canoe landing on Brule Lake. The 8-mile paddle and 4 portages totaling over 300 rods between the landing and Cherokee Lake started out with attempting a 5-mile paddle across the western half of Brule Lake into a strong wind and high waves. After a harrowing episode of one of our canoes being blown off course and “temporarily separated from the group - long story”, we took refugee on an island and resumed our trip to Cherokee Lake under much calmer conditions the next day. The main objective of this capstone project was to conduct bedrock geologic mapping of intrusive igneous rocks to the northwest of the 2014 capstone projects on North and South Temperance Lakes (Beaver et al., 2014; Miller et al., 2015). That capstone mapping revealed that the footwall to the well differentiated Sawbill Lake intrusion, previously defined by capstone mapping in 2007 (Frost et al., 2007), 2009 (Blakely et al., 2009), 2010 (Brooker et al., 2010), and 2011 (Asp et al., 2011) and additional mapping by Ben Brooker for his MS thesis (Brooker and Miller, 2013), is composed of yet another as yet unnamed, well differentiated mafic layered intrusion. It was hoped that this footwall layered mafic intrusion defined in the Temperance Lakes area would project into the southern part of Cherokee Lake. This turned out not to be the case, as mostly Anorthositic Series and Felsic Series rocks of the Duluth Complex were found in the Cherokee Lake area. Previous studies of the Cherokee Lake area include reconnaissance mapping by Grout et al. (1959) and Davidson (1977). Grout’s 1:100,000-scale map of Township 63 North, Range 4 West (Fig. XXIII, Grout et al., 1959) shows the area around Cherokee Lake to contain mostly gabbro with minor granophyric granite, intermediate rocks, and anorthositic gabbro. Davidson’s (1:24,000-scale) reconnaissance map of the Cherokee Lake 7.5’ quadrangle (Davidson, 1977) shows the dominant rock type to be anorthositic gabbro with minor areas of granophyric granite, intergranular granite, intermediate intrusives, metavolcanics, and olivine gabbro. The discovery that the gabbroic anorthosite mapped by Davidson in North and South Temperance Lakes (Davidson, 1977) and in Homer Lake (Davidson and Burnell, 1977) to the southeast was incorrect led to the suspicion that Davidson’s identification of gabbroic anorthosite in the Cherokee Lake area was also erroneous. Turns out Davidson got it right. The 2015 capstone mapping project focused mapping shoreline exposures around the perimeter and along the many islands in Cherokee Lake. Overall, about 350 outcrops were mapped. Four general rock types were encountered in the area. In order of abundance, these were gabbroic anorthosite (90% of exposure), granophyric granite, diabase dikes, and mixed diabase/granophyre dikes. As is typical of Anorthositic Series rocks found elsewhere in the Duluth Complex, anorthositic rocks in the Cherokee Lake area are comprised of various modal types including gabbroic anorthosite, olivine leucogabbro, olivine gabbroic anorthosite and anorthosite. Most exposures show subophitic to ophtic textures with Cpx oikocrysts up to 5cm 99 �diameter. Olivine, when present, ranges from small anhedral grains to poikilitic oikocrysts up to 3cm diameter. Fe-Ti oxides are typically subpoikilitic clots up to 0.5 cm diameter. Most rocks show some degree of foliation that is variably oriented on an outcrop scale. Granophyric granite is salmon pink, fine- to medium fine-grained, intergranular to micrographic leucogranite, quartz monzonite, and quartz ferromonzodiorite. It contains 3-10% mafics (amphibole, oxide, and Fe-pyroxene). The large mass of granophyre east of Cherokee Lake is the western extension of the Misquah Hill granophyre of the Felsic Series of the Duluth Complex, which predates the Anorthositic Series (1108 vs. 1099Ma, Vervoort et al., 2007). Although smaller granite bodies north and west of Cherokee Lake were initially interpreted as Felsic Series equivalents (Miller et al., 2001), contact relationship observed in this study indicate that these granites cut the Anorthositic Series and are thus not part of the Felsic Series. Small-scale intrusions cutting the Anorthositic Series rocks throughout the area include 1) thin (5-30 cm) granophyre dikes, 2) narrow (5cm-5m), very fine-grained, columnar-jointed diabase dikes, and 3) hybrid dikes of mixed diabase and granophyre. The hybrid dikes contain a mixture of medium fine-grained granophyre and aphanitic to fine-grained massive diabase, as well as intermediate hybrid lithologies. The felsic and mafic components commonly occur in sharp lobate contacts suggesting two-magma mixing. Locally, the diabase is cut by more angular apophyses of granophyre. This rock type is similar to late hybrid intrusions cutting the Sawbill Lake intrusion to the south (Brooker and Miller, 2013). Plans for a future capstone mapping project in the area would be to map between Sawbill Lake and Cherokee Lake in order to better define the Sawbill Lake “footwall” mafic layered intrusion exposed in the Temperance Lakes. References Asp, K., Leu, A., Parisi, A., Sletten, D., Brooker, B., Miller, J., 2011, Bedrock geology of the Sawbill Lake area: University of Minnesota Duluth, Precambrian Research Center, PRC/MAP-2011-04, 1: 12,000. Blakely, S., Brown, A., Foley, D., Rowland, A., Stifter, E., and Miller, J., 2009, Bedrock geology map of Homer Lake and adjacent areas; Cook County, Northeastern Minnesota: University of Minnesota Duluth, Precambrian Research Center, PRC/MAP-2009-01, 1: 12,000. Brooker, B.P., and Miller, J.D., 2013, Bedrock geologic map of the Sawbill Lake Intrusion, Cook County, MN. Precambrian Research Center Map Series PRC/Map-2013-01, scale 1:24,000. Brooker, B.P.,Hadley, M.L., Markwood, L.W., Olson, J., Tomlinson, A.P., and Miller, J.D.,2010, Bedrock geologic map of the Jack Lake and Weird Lake areas, Cook County, northeastern Minnesota: University of Minnesota Duluth, Precambrian Research Center, PRC/Map-2010-05, 1: 12,000. Chandler, Val W, 1983, Aeromagnetic map of Minnesota, Cook and Lake counties: Minnesota Geological Survey,Aeromagnetic Map Series, Map A-1, scale 1:250,000 Davidson, D.M., 1977, Reconnaissance geologic map of the Cherokee Lake quadrangle, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series, M-30, scale 1:24,000 Davidson, D.M, Jr., and Burnell, J.R., Jr., 1977, Reconnaissance geologic map of the Brule Lake quadrangle, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series, M-29, scale 1:24,000 Frost, S.J., Juda, N.A., and Miller, J., 2007, Bedrock Geology Map of Homer Lake and Adjacent Areas; Cook County, Northeastern Minnesota: University of Minnesota Duluth, Precambrian Research Center, PRC/MAP2007-02, 1: 12,000 Grout, F.F., Sharp, R.P., and Schwartz, G.M., 1959, The geology of Cook County, Minnesota: Minnesota Geological Survey Bulletin 39, 163 p. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.E., 2001, Geologic map of the Duluth Complex and related rocks, northeastern Minnesota. Miscellaneous Map Series, M-119, scale 1:200,000 Vervoort, J.D., Wirth K., Kennedy, B., Sandland , T., Harpp, K.S., 2007, The magmatic evolution of the Midcontinent rift: New geochronologic and geochemical evidence from felsic magmatism: Precambrian Research 157, p. 235–268. 100 �Continued Evaluation of the Dilatancy Model for Discordant Uranium-Lead Age Determination of Zircon MUDREY, M.G., Jr. 106 Ravine Road, Mount Horeb, WI 53572, USA - mgmudrey@mhtc.net The dilatancy model for lead loss in zircon explains the apparent linearity of data for suites of zircon plotted on a concordia diagram (Goldich and Mudrey, 1972; 1975). The lower intercept age of the discordia dates the approximate time of loss of lead through the escape of fluids which entered the metamict zircon. Time is required for the cumulative effects of U and Th decay to produce radiation damage . Uplift and erosion bring the zircon-bearing rock close to the surface and the resultant relief of pressure causes micro-fracturing and a concomitant increase in the volume or dilation of the rock. This, in turn, permits the entry of water and reactions between the water and the solid phase Under these conditions some of the solution carrying radiogenic lead escapes. The zircon U- Pb ages become discordant. The decay of uranium to lead yields decay equations which can be coupled by the construction of a concordia diagram, permitting the calculation of 206Pb/238U, 207Pb/235 U, and 207 Pb/206Pb ages. Substitution of uranium for zirconium leads to radiation damage in zircon that produces a complex of mixture of crystallites, and amorphous compounds. The extent of the damage depends on the original uranium and thorium content and the age of the mineral. Upon subsequent events, zircon can lose or gain both uranium and lead leading to the plot of data not in concordance, but commonly along a line from the original age to a lower intercept that has been variously interpreted. The figure 1 is a concordia diagram for the Rainy Lake district, Ontario and includes discordant heritage data (Hart and Davis, 1969), and modern single crystal concordant isotopic data (Lodge and others, 2013). The discordia line with an apparent Paleozoic lower intercept was rejected by Hart and Davis on the basis that its interpretation in the episodic lead-loss model requires an event at approximately 500 Ma for which there is no evidence. These same data, however, are subject to different interpretation in terms of the dilatancy model and may represent the exhumation of paleoplanes. Goldich and Mudrey submit that the excellent fit of the data to the 2,720-60 Ma chord clearly favors the dilatancy model in explaining the age discordance in the Rainy Lake district. The non-analytical, or 'geologic,' reasons for discordance are (1) mixing of two or more populations of zircon, (2) Pb loss or U gain, (3) intermediate progeny product disequilibrium, and (4) initial Pb isotopic composition. Other factors such as the U isotopic composition of the sample may also be germane. Peterman and others (1986) illustrate the difficulty in interpreting of poly-episodic events in the Lake Superior region. Episodic lead loss may be by subsequent high temperature metamorphic or thermal events, or by low-temperature dilatancy and uplift. Water, either as H2O or as hydroxyl (OH)- occurs both in the primary crystal structure, and in micro capillary channels and pores which result from radiation damage and stress release. The release of pressure during uplift and erosion results in expansion of the rock mass from micro fractures, rifting and jointing, particularly in quartz-rich rocks. This dilatancy permits 101 �some of the water with dissolved radiogenic lead to escape. This low temperature lead loss model, occurs relatively late in the history of the zircon, and, therefore, the 207Pb/206Pb age commonly approaches the true age, provided the area is not reburied under a significant stratigraphic sequence. Time since crystallization and uranium concen- tration determine the extent of dilantancy. As shown in figure 2, U/Pb analyses on abraded and HF leached single zircons of AS3 of the Duluth Complex containing ap- proximately 100 ppm U (Min and others, 2000) yield a 207Pb/206Pb age of 1097 Ma. Also is plotted an unabraded zircon from a Logan Sill near Lake Nipigon (Davis and Sutcliffe, 1985) with approximately 3000 ppm U and an 207Pb/206Pb age of 1107 Ma. The unabraded zircon has not had sufficient time for dilatancy, although the uranium content is high. The Illinois-Deep-Hole-Project, drilled three deep core holes along the Illinois Wisconsin state line (Coates and others, 1983). The zircon discordia upper intercept and the whole rock Rb/Sr are in agreement, 1,430 Ma (Peterman and others, 1986). The lower intercept of the discordia gives an apparent age of 180 Ma. Zimmerman (1986) reports an apatite fission track age of 140 Ma at the Precambrian surface at 614 m, suggesting very strongly that the lower discordia intercept is a geologically interpretable event of uplift and exposure of the Mesoproterozoic granite. I estimate that based on in situ stress and strain measurements (Engelder, 1993, p.267), the depth at which dilantancy is occurs is greater than 200 m but less than 1000 m. SELECTED REFERENCES Coates, M.S. and others, 1983, Introduction to the Illinois Deep Hole Project: Journal of Geophysical Research. v. 88, p. 7267-7285. Davis, D.W. and Sutcliffe, R.H., 1985, U-Pb ages from the Nipigon plate and northern Lake Superior Geological Society of America Bulletin, v. 96, p. 1572-1579. Engelder, T., 1993, Stress Regimes in the Lithosphere: Princeton Press, Princeton, New Jersey, 451 p Goldich, S.S., and Mudrey, Jr., 1972, Dilatancy model for discordant U-Pb zircon ages, in Contributions to recent geochemistry and Analytical chemistry (Vinogradov Volume), A.I. Tugarinov, ed., pp. 415-18 Moscow: Nauka Publ. Office 1972 (in Russian). Goldich, S.S., and Mudrey, M.G., Jr., 1975, Dilatancy model for discordant U-Pb zircon ages, in Tugarinov, A.I., ed., Recent contributions to geochemistry and analytical chemistry: New York, John Wiley & Sons, p. 466470 (English version of 1972 paper). Hart, S.R. and Davis, G.L., 1969, Zircon U-Pb and Whole-Rock Rb-Sr Ages and Early Crustal Development near Rainy Lake, Ontario: Geological Society of America Bulletin, v. 80, p. 595-616. Lodge, R.W.D. and others, 2013, New U–Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince,Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research, v 235, p. 264-277. Min, A. and others, 2000, , A test for systematic errors in 40Ar/39Ar geochronology through comparison with U/Pb analysis of a 1.1-Ga rhyolite: Geochimica et Cosmochimica Acta, v. 64, p. 73–98, 2000 Peterman, Z.E., and others, 1986, Geochronology of Basement Granite, Stephenson County, Illinois: US Geological Survey Bulletin 1622, p. 41-50. Peterman, Z.E., and others, 1986 , A protracted Archean history in the Watersmeet gneiss dome, northern Michigan: US Geological Survey Bulletin 1622, p. 51-64 Zimmermann, R.A. 1986, Fission-track dating of the Illinois drill-hole core: US Geological Survey Bulletin 1622, p. 99-108. 102 �Emplacement and Crystallization History of Ni-Cu-(PGE) Sulfide-mineralized Peridotites in Eagle Intrusion, Upper Michigan MULCAHY, Connor1, MILLER, Jim1, MAHIN, Robert2, BEACH, Steven2, and NOWACK, Robert2 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Duluth, MN 55812; 2 Eagle Mine, 4547 County Road 601, Champion, MI 49814. The Eagle deposit and associated Eagle East prospect near Marquette, MI, are small, high-grade orthomagmatic Ni-Cu-(PGE) sulfide deposits hosted in ultramafic intrusive rocks associated with the Midcontinent Rift (Ding et al., 2010). The Eagle deposit is currently being developed by Lundin Mining Corporation. This study, which is the focus of the lead author’s MS thesis at UMD, has four main goals: 1) to determine the number of magmas involved in Eagle’s genesis, 2) to test an emplacement model for Eagle proposed by Lundin geologists, 3) to determine the petrogenesis of enigmatic “pyroxenite” inclusions (actually melagabbronorite) in parts of the Eagle intrusion, and 4) to determine the emplacement and crystallization history of the Eagle East intrusion. This talk will report on data that address the first and second goals. With the exception of pyroxenite inclusions, the host rock of the Eagle intrusion is generally characterized by Lundin geologists as peridotite, which they subdivide into four grades of sulfide mineralization (mineralization unit types): < 5 vol% poorly mineralized (Per); 5-25% - disseminated (MPer), 25-75% - semimassive (SMSU), and > 75% - massive sulfide (MSU). Petrographic work on silicate mineralogy conducted by Ding et al. (2010) defined four main modal rock types in Eagle and Eagle East – dunite, feldspathic peridotite, melatroctolite and melagabbro. However, they used a non-traditional classification scheme that did not discriminate between pyroxene types and did not factor in textures. In this study, petrographic observations of 121 samples from three Eagle drill cores and three Eagle East drill cores reveal that most samples representing all four mineralization types modally classify as a feldspathic lherzolite to feldspathic olivine websterite. Interpreting cumulate nomenclature from mineral habits and modes, the samples uniformly classify as olivineclinopyroxene-orthopyroxene cumulates with variable amounts of intercumulus plagioclase, hornblende, Fe-Ti oxide, and apatite (OCpH to OCpH [α, f, a] in the cumulate code nomenclature of Miller et al., 2002). In evaluating the number of magmas involved in the formation of the Eagle intrusion, Ding et al. (2010) interpreted variations in incompatible trace element ratios to indicate the involvement of at least two to as many as four compositionally distinct parental magmas. However, reevaluating the lithogeochemistry of the peridotitic rocks in terms of mixtures of cumulus minerals and postcumulus minerals (=trapped liquid), the range of incompatible trace element ratios can be attributed more simply to a single parental magma type. Orthocumulates, which contain elevated amounts of trapped liquid (now represented by intercumulus plagioclase, amphibole, Fe-Ti oxide, and apatite), should have incompatible trace element ratios closer to their parental magmas. In contrast, adcumulates, with little to no trapped liquid component, would be expected to have incompatible trace element ratios more reflective of the partition coefficients of their cumulate mineralogy. As shown in Figure 1, the variation of incompatible trace element ratios such as Zr/Y correlates well with the amount of Eu, Ti, and P, which are all elements associated with intercumulus minerals. Using the mineral-liquid partition coefficients reported by Bedard (1993) and taking the Zr/Y and La/Yb ratios of 6.5 and 5.0, respectively, for orthocumulates as approximating parental magma ratios (green ovals, Fig. 1), Zr/Y and La/Yb were calculated for adcumulates ranging in Ol:Cpx:Opx mode from 80:15:5 to 10:60:30 (black ovals, Fig. 1). It is noteworthy that samples with more abundant sulfide are more likely to be hosted in adcumulates. This may indicate that low density intercumulus silicate melt was displaced by the infiltration of high density sulfide liquid. Lundin geologists have developed a two-stage emplacement/mineralization model based on spatial relationships of the three styles of mineralization. They speculate that initial pulse was intruded into a narrow Y-shaped conduit as a sulfide-oversaturated ultramafic magma that then experienced density driven settling of sulfide liquid to create the massive sulfide unit (MSU) in the neck of the intrusion and the disseminated mineralization in the upper parts of the intrusion. This was followed by emplacement of an olivine porphyritic magma along the margin of the still molten massive sulfide resulting in the creation of the semi-massive body adjacent to the massive sulfide and projecting upward into the disseminated lherzolite. While the trace element data suggest that a single parental magma composition was likely involved in the creation of the Eagle intrusion, it is still possible that this magma was emplaced more than once. 103 �Figure 1) Zr/Y and La/Yb vs. TiO2 + P2O5 (A & B) and vs. Eu (C & D) for Eagle samples. Ti, P and Eu are taken as proxies of intercumulus minerals of Fe-Ti oxide, apatite, and plagioclase, respectively, and give a qualitative measure for the amount of trapped liquid in Ol-Cpx-Opx cumulates. The green ovals represent estimated trace element ratios for the parent magmas approximated by orthocumulates. The black ovals represent the compositional range of trace element ratios of adcumulates in equilibrium with such parental magmas based on experimentally determined mineral-liquid partitions coefficients (Bedard, 1994). Pink areas indicate the range of trace element ratios that might be expected in orthocumulates to adcumulates generated from a single magma composition. Another possible test off whether two major magma pulses were involved in the creation of the Per-MPerMSU sequence and the SMSU is to evaluate whether the Ni content of olivine is different between the two mineralized suites. While the presence of euhedral olivine phenocrysts in the SMSU suggest that olivine crystallized prior to sulfide liquation in the second pulse, the timing of olivine crystallization and sulfide liquation in the initial pulse is not clear. If sulfide liquation occurs before olivine begins to crystallize, Ni should be strongly depleted in olivine. Electron microprobe analyses of olivine from all four mineralization types show undepleted Ni abundances (Fig. 2). This implies that, if there were two pulses of magma, olivine crystallized prior to sulfide liquation in both. As shown in Figure 2, almost all analyses of olivine within a particular sample show a positive correlation between Ni abundance and Fo content, which is consistent with Ni depletion due to fractional crystallization of olivine. With the exception of one sample, the positive correlation Ni and Fo, especially from SMSU and MSU samples, indicates that olivine did not re-equilibrate with enclosing sulfides. This contrasts with Li et al’s (2007) analyses of olivine in mineralized rocks from Voisey’s Bay which show evidence of re-equilibration by their negative Ni-Fo trends. Figure 2) Ni-Fo trendlines by sample for all samples analyzed in the study. Orange points denote semi-massive sulfide, blue points denote disseminated sulfide. Red trendline shows the only negative Ni-Fo trend indicative of olivine-sulfide reequilibration. References Bedard, Jean H. (1994). A procedure for calculating the equilibrium distribution of trace elements among the minerals of cumulate rocks, and the concentration of trace elements in the coexisting liquids. Chemical Geology 118, 143-153. Ding X., Li C., Ripley E. M., Rossell D. and Kamo S. (2010). The Eagle and Eagle East sulfide ore-bearing mafic–ultramafic intrusions in the Midcontinent Rift System, Upper Michigan: geochronology and petrologic evolution. Geochemistry Geophysics Geosystems. Li, C, Naldrett, A.J., and Ripley E. (2007). Controls on the Fo and Ni Contents of Olivine in Sulfide-bearing Mafic/Ultramafic Intrusions: Principles, Modeling, and Examples from Voisey’s Bay. Earth Science Frontiers Vol. 14, Issue 5. Miller, J. D., J. C. Green, M. J. Severson, V. W. Chandler, S. A. Hauck, D. M. Peterson, and T. E. Wahl (2002), Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota, Minn. Geol. Surv. Rep. Invest., 58, 207 pp., Minn. Geol. Surv., Saint Paul.  104 �Mesoproterozoic Alteration of the Paleoproterozoic Gunflint Formation: Analogies with Martian Blueberries NAP, Carli1, FRALICK, Philip2 1 Department of Geology, Lakehead University, Thunder Bay, ON, cnap@lakeheadu.ca 2 Department of Geology, Lakehead University, Thunder Bay, ON, pfralick@lakeheadu.ca NASA’s Opportunity rover landed on Meridiani Planum in summer of 2004 with the intention of studying a rich concentration of hematite in much finer detail than what the preliminary images from the orbiting Mars Global Surveyor could possibly allow. Small, diagenetic, 4mm spherules composed primarily of hematite were discovered embedded in sand blasted bedrock, arming the scientific community with further evidence for a past presence of water on Mars (NASA, 2012). This thesis is an attempt at providing a terrestrial analogue for the formation of Martian spherules by using hematite-rich concretions observed in the minimally metamorphosed, 1.8Ga Gunflint formation as proxy. Approximately 40 minutes eastbound of Thunder Bay at intersections with Mirror Lake (ML) and West Loon (WL) road lie two recently exposed, iron oxidized grainstone outcrops standing ~2-3m vertically. Original deposition occurred in the Paleoproterozoic at 1,876Ma in a storm-dominated shallow shelf. The grainstone has a typical grey-green ankeritic to white chert colour with spherical to rhombic hematite-rich concretions averaging <2mm to 2cm in diameter present in thin layers oriented parallel to the shallowly dipping, lensy bedding or randomly distributed within lenses as independent concretions or hematitic masses. Three fundamental questions occur in response to these outcroppings: Why are these rocks so iron stained when compared with the upper cherty and shaley Gunflint at outcrops of nearby Pass Lake (PL) road?; How do these concretions form?; Is the mechanism responsible for the formation of these concretions a reasonable analogue for those observed on Mars? A scaling down approach to these queries consisted of site mapping and stratigraphic sectioning, qualitative and quantitative petrographic and SEM analysis, and quantitative geochemical analysis using data acquired from ICP-MS, ICP-AES and XRD techniques. It is well known that the Gunflint is high in iron that precipitated out of seawater solution as an insoluble chemical precipitate. In the studied upper cherty member, individual grainstone grains have fine, iron-rich laminae that had precipitated onto the grain surface itself or were accumulated by rolling over an Fe-rich substrate. An implicit, unconformable upper contact with the Sibley group, present above and just tens of meters away from the Gunflint outcrop, which would have allowed for iron rich and potentially oxidative fluid migration into the underlying Gunflint. Large, centimeter scale, iron-rich fracture sets as well as very fine, micrometer scale capillary networks provide evidence for fluid migration. Variability in the red colouration, from blood red to maroon, can be differentiated on concentric layers of individual concretions, as well as overprinting masses, and is suggestive of multiple phases of redox fluid front migrations. It is by some combination of intrinsic and extrinsic iron combined with oxidation that gave these rocks their ultimate red colouration. Hematite concentrations within the ML and WL outcrops are always associated with carbonates that are at varying stages of decay. These hematite-bearing carbonates have been identified through geochemical, XRD and SEM analysis to be of ferroan dolomite to ankerite in composition. They are often found nucleating on or within siliceous and hematite altered grains as rhombs, and commonly mimic the entire grain. Spherical concretions occur when several carbonate altered grains are enclosed by the growth of successive poikilitic carbonate and rarely 105 �display a distinguishable nuclei: the appearance of framboidal pyrite central to a select group of concretions and scattered within the groundmass strongly suggests the influence of bacterial sulfate reduction (BSR) on carbonate growth. The variation in size, morphology and distribution of iron-bearing carbonates within individual grainstone lenses as concretionary spheres or masses, as well as the presence of framboidal pyrites, suggests that primary, hematite-poor carbonate concretion formation occurred in a shallow phreatic, anaerobic environment at some time after cementation and prior to compaction as the grainstone layers accumulated. Gunflint hematite concretions differ in several ways to those observed on Mars. Negative weathering, original iron-carbonate composition, the promotion of growth by BSR and a lesser random to common bedding parallel stratigraphic distribution define the concretions observed at the terrestrial site whereas positive weathering, jarosite-hematite-alunite composition, spherulitic growth by supersaturation of Fe-rich fluids, and non-conformable growth over all stratigraphic units define the concretions at the Martian sites (Morris et al, 2010). Concretions preserve fluid chemistry and are blueprints for flow regimes and are as such important to the evolution of water on Mars and early Earth. References Morris, R.V., Golden, D.C., Ming, D.W., 2010, Spherulitic Growth of Hematite Under Hydrothermal Conditions: Insights into the Growth Mechanism of Hematite Spherules at Meridiani Planum Mars [abstract]: NASA, 41st Lunar and Planetary Science Conference (2010), PDF 2541. NASA, 2012, Opportunity on Mars - Eight Years and Counting - January 24, 2012 Animation: [http://science.nasa.gov/missions/mars-rovers/]. 106 �Investigations of the Layered Series Nepheline Syenite within Center II of the Coldwell Complex, Marathon, ON Nikkila, D1., Mitchell, R.H1., and Zurevinski, S.E1. 1 Department of Geology, Lakehead University, ON The Coldwell Alkaline Complex is the largest alkaline intrusion in North America, and was emplaced during initial magmatism of the Keweenawan Midcontinent Rift (MCR) at 1.1 Ga. Located on the north shore of Lake Superior, the Coldwell Complex was emplaced along the Thiel fault (the northern component of the Trans-Superior Tectonic Zone), and is host to rare earth elements (REE), Cu, Ni, PGE, and other high field element mineralization. Emplaced from east to west, the oldest – termed “Center I” is host to gabbro and Fe-rich augite syenite; “Center II” hosts biotite-gabbro and nepheline syenite; and “Center III” is host to a variety of syenite. The focus of this study is to understand the magma systematics involved in the emplacement and crystallization of the different intrusive centers, specifically Center II. This involves field mapping, extensive sampling and mineralogical study of the complex syenites. There is an emphasis on identifying the REE minerals occurring in the nepheline syenites. This detailed assessment will help to understand the complex systematics of alkaline rocks in the Superior Province, and will assist with mineral exploration within the complex, which is currently underexplored for REEs. Layered series nepheline syenite rocks display a cumulus texture perthitic feldspar, with post cumulus amphibole and associated pyroxene, biotite, and zeolites. The feldspar exhibits secondary albitization of an earlier alkali feldspar which has produced lamellae of albite (An% 04.43). Amphiboles dominantly plot in the hastingsite range; with zoning compositions displaying a trend to Fe and Mn enrichment with Mg, Ca, Ti depletion. Biotite are classified as annite with rare siderophyllite, and occur as alteration products associated with amphibole. REE analysis of the layered series has displayed elevated levels of LREE in apatite (La, Ce, Nd) along with Th and Y, and minor abundances of britholite-(Ce, La) and wohlerite. Massive syenites associated with the layered series at higher structural levels contain abundant wohlerite and britholite-(Ce, La) with apatite containing minor amounts of LREE. Pegmatites representing altered massive syenite; contain the most diverse suite of accessory minerals including: wohlerite, britholite, NbAl-Fe bearing titanite, U-bearing pyrochlore, zoned apatite, zircon, and minor sulphides. Apatite has proven to be a viable mineral for future tracer isotope analysis; dominantly forming 20100µm (rarely up to 300µm) subhedral elongate to hexagonal crystals. Commonly displaying irregular to concentric zoning with an increased LREE rim composition, apatite is found disseminated in the feldspar groundmass or along fractures and grain boundaries of amphibole. 107 �Figure 1: Geological Map of the Coldwell Complex (Note: Star represents sample area). 108 �Petrology, geochemistry and sulphur isotopes of the Crystal Lake gabbro and Mount Mollie dyke, Northwestern Ontario O’BRIEN, Sean1, HOLLINGS, Peter1, and MILLER, Jim2 Department of Geology, Lakehead University, Thunder Bay, ON, P7B 5E1 Canada 2 Department of Geological Sciences, University of Minnesota Duluth, Duluth, MN 55812, United States 1 The Midcontinent Rift (MCR) formed from ~1150 to 1087 Ma, with the majority of the igneous activity occurring in two pulses between ~1108-1105 Ma and ~1100-1094 Ma (Heaman et al., 2007). A magnetic polarity reversal is recognized by Davis and Green (1997) in MCR related volcanic and intrusive rocks during the gap in time between the main pulses. In this study two mineralized MCR-related intrusions, the Crystal Lake gabbro (CLG) and the Mount Mollie dyke (MMD), have been investigated using petrography, geochemistry and sulphur isotopes. The CLG is a “Y” shaped intrusion with a 5 km long north arm and a 3.75 km long southern arm (Fig. 1). The ~ 35 km long MMD extends east from the CLG (Fig. 1). The spatial relationship and similar rock types led to the belief that the two were co-genetic and/or or contemporaneous, however, recent age dating has revealed that the CLG formed at 1099.6 ± 1.2 Ma and the MMD formed at 1109 ± 6.3 Ma (Heaman et al., 2007; Hollings et al., 2010). There are still unresolved issues with this age gap, most notably that both have the same paleomagnetic N-polarity, where one would expect a R-polarity for MCR related rocks that are older than 1105 Ma. This study will use petrography, whole rock geochemistry and sulphur isotopes in an attempt to resolve the conflicting evidence regarding the relationship between the CLG and MMD. Samples of the CLG were collected from an 828 m diamond drill core from the relatively little known about southern arm, whereas MMD samples were collected from a 1087 m diamond drill core. Figure 1. Generalized map and magnetic polarity of the Crystal Lake gabbro, Mount Mollie dyke and surrounding rocks within the Logan Basin (Cundari, 2013). Petrographic studies have revealed that the CLG and MMD are mineralogically and texturally similar, with troctolite and sub-ophitic to ophitic olivine gabbro being the most abundant rock types. The bottom of the MMD hole begins with a sequence of troctolite to olivine gabbro with variable amounts of sub-ophitic to ophitic clinopyroxene. Around -600 m a ~15 m sequence of alternating very coarse- and medium-grained olivine gabbro. The next unit is a clinopyroxene bearing troctolite to gabbro sequence, with sub-ophitic to ophitic clinopyroxene at the bottom and intergranular clinopyroxene in the top 100 m of the sequence. The top 150 m of the drill core consists of granophyre, hornfels, and diabase units. The bottom of the CLG drill core consists of 109 �shale of the Rove Formation and a thin 20 m fine-grained diabase unit. Next is a 100 m unit of medium- to coarse-grained troctolite and olivine gabbro with thin (mm) layers of fine-grained chrome spinel chadacrysts present 10 m above the base of this unit. A 75 m unit of fine-grained and plagioclase phyric diabase, petrographically similar to the 20 m diabase unit, separates two CLG units. The CLG units have chilled contacts with the diabase unit above and below, suggesting that it is significantly younger. The main CLG consists of a sequence of troctolite to gabbro with ophitic to sub-ophitic clinopyroxene with thin layers of chrome spinel chadacrysts near the base. The top 40 m of the hole consists of strongly altered sulphidic olivine gabbro. The MMD and main gabbro unit in the CLG drill core show relatively smooth fractionation trends, with increasing concentrations of Fe2O3, SiO2, Na2O, TiO2, Ba and V and decreasing concentrations of Al2O3, CaO, MgO and Ni from the bottom to the top of the hole; however, CaO values are constant in the CLG. In and around the very coarse-grained unit of the MMD, there are deviations from the smooth trends observed in the other units. This unit is currently being investigated using mineral chemistry to help explain these variations. The 20 m and 75 m thick diabases in the CLG drill hole are geochemically distinct from the CLG. This suggests that these units are likely from a separate source, consistent with the petrographic observation of chilled margins. To determine the role of crustal assimilation in the CLG, δ34S was analyzed from the visible sulphides in various horizons of the drill core. The CLG had visible sulphides occurring at the top and bottom 10 to 50 m of both units. Overall the δ34S ranged from 4.1 to 21.0 ‰, with higher values generally found in the lower unit. Due to the high variability of the δ34S ~ -1 to 33‰ (Johnston et al., 2006) of the surrounding Rove Formation, which the CLG is assumed to have assimilated, it is difficult to determine the degree of assimilation. This is because a high degree of assimilation of a moderate δ34S would yield similar results to a lower degree of a high δ34S value material. Regardless, there is enough data to suggest that crustal assimilation played a role in the sulphur saturation history of the CLG. References Cundari, R.S., Campbell, D., and Puumala, M., 2013. Geology, Geochemistry and Cu-Ni-PGE Mineralization of the Crystal Lake Gabbro, 6th Annual PRC Professional Workshop Cu-Ni-PGE Deposits of the Lake Superior Region, Duluth, Minnesota. Davis, D. W. and Green, J. C., 1997. Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences, 34(4): 476-488. Heaman, L., Easton, R., Hart, T., Hollings, P., MacDonald, C. and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences, 44(8): 1055-1086. Hollings, P., Smyk, M., Heaman, L.M. and Halls, H., 2010. The geochemistry, geochronology and paleomagnetism of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada. Precambrian Research, 183(3): 553-571. Johnston, D. T., Poulton, S. W., Fralick, P. W., Wing, B. A., Canfield, D. E., & Farquhar, J., 2006. Evolution of the oceanic sulfur cycle at the end of the Paleoproterozoic. Geochimica et Cosmochimica Acta, 70(23): 5723-5739. 110 �What Happened in Northern Minnesota Between 2700 Ma and 1900 Ma? The Answer Is in the Pokegama Formation: A Multicycle Sedimentary History! OJAKANGAS, Richard W. Department of Earth and Environmental Sciences, University of Minnesota Duluth, Duluth, Minnesota 55812 rojakang@d.umn.edu The Giants Range Batholith is the core of the Algoman Mountains in the area just to the north of the Mesabi Iron Range. This composite batholith, the Vermilion Complex, and the Saganaga Batholith were emplaced beneath volcanic arcs at about 2700 Ma by subduction processes. Except for 2100 Ma mafic dikes, there is no preserved geologic record in northern Minnesota until the sedimentary rocks of the Paleoproterozoic Animikie Basin, a foreland basin, at ~ 1850 to 1900 Ma. Certainly erosion was going on, for the Algoman Mountains were essentially eroded away as topographic prominences by the time the Animikie Basin developed to the south of the mountains. This is a time span of about 800 million years. If we can assume that the Algoman Mountains were indeed respectable mountains, uplift and erosion must have generated thousands of feet of sediment. We also can reasonably assume that a significant portion of the detritus consisted of resistant quartz eroded from the granitic bodies that intruded into what was likely a volcanic roof, as documented by amphibolitic inclusions in the granitic outcrops. Based on regional relationships and cross bedding, we can assume that the paleoslope was to the south, toward and into the Animikie Basin. The lowest stratigraphic unit in the basin, the Pokegama Formation, less than 50 m thick in total, contains abundant detrital quartz only in its uppermost member, which was a quartz sandstone and is now a quartzite. Therefore the question is: “Where is the expected large quantity of quartz, the likely resistant detritus of 800 million years of erosion?” At Blueberry Hill within the Hibbing Taconite complex of open pits, there is a conglomerate at the base of the Pokegama Formation, which overlies the Archean tonalite to granodiorite basement and underlies the Biwabik Iron Formation. The conglomerate is especially interesting in that it contains numerous rounded pebbles of quartz arenite, which consist of well-sorted and spherical unit quartz grains. At least a partial answer to the above question is that older quartz sands (pre-Pokegama) were transported, abraded, deposited, lithified, re-eroded, and redeposited during those 800 million years. The pebbles of quartz sandstone are therefore proof of a multicycle history of at least two cycles, and three counting the Pokegama. How many other undetected quartz sandstones may have been present but have been totally removed by erosion? A cross bedding plot for the Biwabik Iron Formation is bimodal-bipolar, indicating a tidally influenced environment of deposition that is probably applicable to the Pokagama as well. Although the excellent rounding of resistant quartz grains also fits a tidal model, the total lack of vegetation on land during both Late Archean time and Paleoproterozoic time would have resulted in constant wind abrasion of sand-sized quartz grains by wind, the “best rounder” of sand grains. Therefore, the quartz grains (plus rare feldspar grains) were very likely wellrounded before they reached their final depositional site in the Pokegama Formation. 111 �Rounded pebbles of quartz arenite in a basal conglomerate of the Pokegama Formation 112 �Ore Petrography and Precious Metals of the Primary Flambeau Massive Sulfide Ore OLSON, Maile J. and LODGE, Robert W.D. Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI The Flambeau volcanogenic massive sulfide (VMS) deposit, near Ladysmith, WI, was the only one of several potentially economic deposits to be extracted in the Penokean Orogeny of Wisconsin. Northern Wisconsin is home to at least thirteen VMS deposits hosted within the Wisconsin Magmatic Terrane (DeMatties, 1994), a Precambrian juvenile arc sequence consisting of volcanic rocks, sedimentary sequences, and associated plutonic rocks that are about 1.8-1.9 billion years old (Schulz & Cannon, 2007). The Flambeau mine was considered as a copper-gold deposit because of the supergene-enriched zone that was the only portion of the ore body to be extracted. Interestingly, the enriched zone of the Flambeau deposit that was extracted accounted for ¼ of the total volume of ore and the primary ore was left in the ground. Following an initial surge of academic publications following the discovery and exploration of the Flambeau (e.g. DeMatties 1994), research largely ceased along with the mining and exploration activities. What research was done was predominantly carried out on the secondary enriched zone (e.g. Ross, 1997). Preliminary research indicates the primary ore body being more enriched in zinc and lead (Zens et al. 2015). This project is designed to examine the mineralogy, texture, and composition of the primary ore body, determine the base and precious metal phases that are present in the ore. Petrographic and geochemical analyses of ore-forming minerals have helped constrain the nature and evolution of economic mineralization. The current phase of this research project has involved mineral analyses of polished thin sections of Flambeau ore samples. Using a Scanning Electron Microscope (SEM) in the Materials Science Center at UW – Eau Claire, the habit of gold, silver, empressite, electrum, and Sb-bearing alloy grains have been documented (Fig. 1). As expected with metamorphosed VMS ores, the main ore phases were euhedral pyrite, sphalerite, and lesser galena and chalcopyrite. Most of the thin section samples there are abundant sub-microscopic inclusions of galena, frequently found in pyrite, but also found in sphalerite and chalcopyrite. Bi-Ag-alloy has been observed with these galena inclusions. Various trace minerals found in the primary ore include monazite, barite, cinnabar, and cassiterite within the host mineral pyrite and to lesser extent sphalerite. Within sphalerite and pyrite, native silver, native gold, electrum, empressite, and Bi-Te-Sb alloys were identified. The primary gangue mineral phases are quartz and anthophyllite. Research utilizing reflected light petrography and the SEM have improved the identification of precious metal ore minerals, their chemistry, and characterization of the precious metal mineral-hosts. The data gathered here is contributing to a larger study that is seeking a complete a geochemical and petrographic understanding of Wisconsin’s VMS deposits and to fully characterize the Precambrian and economic geology of rocks hosting the Flambeau mine. 113 �Figure 1: Various images showing mineral hosts of precious metals in the primary Flambeau ore body. Photos A-C are backscatter images from Scanning Electron Microscope. Photo D is photomicrograph from petrographic microscope. Mineral abbreviations: Sp – Sphalerite; Py – Pyrite; Ga – Galena; Au – Native Gold; El – Electrum ± Antimony alloy; Ar – Arsenopyrite; An – Anthophyllite; Em – Empressite (AgTe). References DeMatties, T.A., 1994. Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Economic Geology, 89: 1122-1151. LeBerge, G.L. (ed), 1996. Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume. Institute on Lake Superior Geology, Proceedings, 42nd Annual Meeting, Cable, WI, vol. 42, part 2, 179 p. Schulz, K.J. & Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research, 157: 4-25. Zens, Z.A., Helmuth, S.L., & Lodge, R.W.D. (2015). Geochemistry and petrography of the strata hosting the Flambeau Cu-Zn-Au Deposit: Revisiting Wisconsin’s only past-producing volcanogenic massive sulfide mine. Geological Society of America, Abstracts with Programs, 47(5). Ross, Andrew M. Supergene Gold Enrichment of the Precambrian Aged Flambeau Gossan, Flambeau Mine, Rusk County, Wisconsin. 1997. Print 114 �THE LAC DES ILES PGE-CU-NI DEPOSIT, CANADA: AN ORGANIZED MEGA-BRECCIA UNIT? Dave PECK1, Lionnel DJON1, Cameron McLEAN1, Gary DeSCHUTTER1, Jill MAXWELL1, Kelsey PRIVETT1, Denis DECHARTE1, Chris RONEY1, Michelle HUMINICKI2 and Robert STEWART3 1 North American Palladium Ltd., Exploration Department, 556 Tenth Avenue, Thunder Bay, ON, Canada P7B 2R2 Micro Analytical Facility, Brandon University, 270 18th Street Brandon, Manitoba Canada R7A 6A9 3 330 Ridgevale Drive, Bedford, Nova Scotia, Canada B4A 3M1 2 The Lac des Iles Intrusive Complex (LDI-IC), located in the Thunder Bay Mining District of northwestern Ontario, is one of several ~2.68 Ga mafic +/- ultramafic intrusive bodies that were emplaced into both the Eastern Wabigoon Terrane and the eastern part of the older Marmion Terrane of the western Superior Province. The LDI-IC consists of one mafic complex (South LDI complex) and one predominantly ultramafic complex (North LDI complex). All known mineral resources in the LDI-IC are hosted within the South LDI complex, with the vast majority occurring in the western part of the ~3 km long x 1.5km wide Mine Block intrusion (MBI). Historical PGE-Cu-Ni resources in the western MBI, including mined out and existing resources, now exceed 200 million tonnes at an average grade of >2 g/t Pd. This historical global resource includes approximately 50 million tonnes of higher-grade resources featuring average Pd grades in excess of 4 g/t over true widths of <5 to ~80m that commonly includes narrow widths (1-10m) of ‘bonanza-grade’ Pd mineralization (e.g., >10 g/t to 2-3 oz/t). Most of these resources occur within a small surface area having approximate, known dimensions of ~1 km N-S by 500m E-W by >2 km vertically. Recent insights stemming from 3D assay, wholerock geochemistry, applied mineralogy, geophysical and structural logging data imply a dynamic, multi-stage mineralizing process. Previous research addressing LDI ore forming processes have typically cited one or both of the following: 1) magmatic sulfide collection of PGE and base metals; and, 2) magmatic volatile-related PGE localization and upgrading. Over 50 years of exploration and twenty years of mining at LDI have generated a tremendous geoscientific database including over 500,000 assays and approximately 750,000 metres of exploration and definition drill core and logging. This database underpins several immutable facts relating to the geology, mineralogy, geochemistry and morphology of the PGE-Cu-Ni resources at LDI. These include: 1. Palladium tenors at LDI are amongst the highest documented from magmatic PGE-Cu-Ni deposits, typically exceeding several hundred ppm Pd and locally exceeding several thousand ppm Pd in 100% sulfide; 2. Pd:Pt and Pd:Au ratios increase with increasing Pd grade that, although documented in a small number of other intrusive bodies (e.g. Skaergaard intrusion, Keays and Tegner, 2016) remains an atypical feature of magmatic PGE-Cu-Ni deposits; 3. The highest Pd grades, although principally located along an inferred pre-magmatic regional feeder fault, locally follow other major faults as well as lithological contacts - especially those between vari-textured noritic rocks and earlier-formed intrusive units; 4. Semi-continuous brecciation of pre-existing rock units, including both early phases of the MBI and basement orthogneiss, occurred during the emplacement of fluid-rich noritic magmas and is interpreted as having been coeval with the formation of most of the highest grade PGE-Cu-Ni mineralization; 5. Palladium mineralization is commonly, but not universally, associated with disseminated Fe-CuNi sulfide mineralization comprising a characteristic assemblage of pyrite + pyrrhotite + pentlandite + chalcopyrite +/- millerite with total sulfide mineral abundances generally falling in 115 �6. 7. 8. 9. the range of 0.5 to 2% but with local, narrow occurrences of net-textured and semi-massive sulfide mineralization; Palladium principally occurs in Te-, Bi- and S-rich ore minerals displaying a wide range in texture, locking characteristics and mineral associations but having strong similarities to Pdbearing minerals documented from other PGE deposits; Recent studies of remanent magnetism in the LDI area indicate that the bulk of the PGE-Cu-Ni mineralization was formed after a reversal of the Earth’s magnetic field such that both normallypolarized and reversely-polarized rocks are present in the South LDI-IC; The Offset Fault separates the Roby and Offset structural blocks and is currently interpreted to be an oblique thrust fault – however, the timing of oblique slip movement along the Offset fault remains equivocal as does the relative timing of the currently defined major faults on the mine property and their relative importance in focusing and displacing mineralization; and, Hydrous alteration associated with magmatic volatiles derived from fluid-rich noritic magma produced chlorite-amphibole +/- talc +/- epidote alteration along the contact between a preexisting, largely unmineralized gabbroic unit (EGAB) and mineralized vari-textured noritic rocks. Although many altered structures and contacts are present in the MBI, only a few of these appear to have acted as “sinks” for PGE. Taken together, these observations support an interpretation involving multiple injections of noritic to gabbronoritic magma into the proto South LDI-IC, with different pulses having distinctive geochemical and mineralogical characteristics. The earliest period of magmatism produced massive to layered, iron-enriched and locally oxide-saturated leuco- to ferrogabbronorite and gabbronorite. The second magmatic episode produced massive to weakly layered norite and melanocratic norite featuring local PGE- and base metal sulfide enrichment. The final and principal ore-generating magmatic episode involved the emplacement of water-rich and sulfur-saturated noritic magma that produced a deeply-rooted and laterally- and verticallyextensive magmatic mega-breccia unit. The latter comprises a varitextured leuconorite to melanorite matrix and both cognate and basement-derived fragments of variable size and exhibiting a wide range in alteration intensity and degrees of resorption and partial melting. Although the mega-breccia unit is focused along the main north-south feeder structure it locally tracks along other major faults and intersects most of the pre-existing major rock units in the MBI. With deeper drilling currently confined to the western MBI, the extent (and mineral potential) of the mega-breccia in the central and eastern parts of the intrusion remains unclear. Although having a chaotic appearance on the local scale, recent exploration findings suggest that the mega-breccia is well-organized on the deposit scale. Accordingly, vectoring to higher-grade subzones is becoming possible using routine exploration data including logging information, geochemistry, 3D structural interpretations and geophysical properties. The favourable 3D continuity, thickness, PGE grades and sub-vertical orientation of the LDI PGE-Cu-Ni deposit provides strong motivation for future exploration in the region. The general category of contact-type PGE-Cu-Ni sulfide deposits should now be extended to include deeplyrooted, structurally-controlled magmatic mega-breccia systems like LDI. The higher-grade parts of these breccia systems are expected to occur in areas of maximum magma throughput, such as primary feeder conduits and structural intersections along these feeders. Narrow zones of high-grade PGE mineralization having low total sulfide content represent difficult exploration targets but should carry most of the value of a given mega-breccia system. Reference KEAYS, R.R. and TEGNER, C., 2016: Magma chamber processes in the formation of the low-sulphide magmatic Au–PGE mineralization of the Platinova reef in the Skaergaard intrusion, East Greenland. J. Petrology, in press. 116 �Copper toxicity and dissolved organic matter: Resiliency of mineralized watersheds in northern Minnesota and Michigan. PIATAK, Nadine M.1, SEAL, Robert R. II1, JONES, Perry M.2, WOODRUFF, Laurel G.2, 1 2 U.S. Geological Survey, Reston, VA 20192, npiatak@usgs.gov, rseal@usgs.gov U.S. Geological Survey, Mounds View, MN 55112, pmjones@usgs.gov, woodruff@usgs.gov Copper (Cu) toxicity in surface waters was estimated in watersheds containing contrasting mineral-deposit types in the Duluth Complex (northern Minnesota) and in the Porcupine Mountains Cu district (western Upper Peninsula of Michigan); these waters contain high dissolved organic carbon (DOC) and low to moderate hardness, both of which mitigate metal toxicity. Mineral deposits in these areas are related to rocks of the Midcontinent Rift System. In the Duluth Complex, deposits include magmatic Cu-Ni (nickel)-PGM (platinum group metal) sulfide deposits in mafic rocks and iron (Fe) and titanium (Ti) oxide ultramafic intrusions. Mineral deposits in the Porcupine Mountains district are stratiform Cu deposits hosted by the Nonesuch Formation, which consists of gray to black shales and siltstones. This study compared aquatic life water-quality criteria for Cu calculated on hardness alone, which had been the standard regulatory procedure for the past few decades, and compared it to criteria based on the Biotic Ligand Model (BLM), a newer approach that has supplanted the hardness-based approach for Cu and incorporates major element chemistry, metal speciation, and organic carbon complexation. Surface-water samples were collected along streams in three geologically distinct watersheds in the Duluth Complex: 1. Filson Creek where Cu-Ni-PGM mineralization occurs at the bedrock surface along the basal Duluth Complex; 2. Keeley Creek where Cu-Ni-PGM mineralization occurs only at depth; and 3. the upper St. Louis River in the vicinity of Fe-Ti oxide ultramafic intrusions, which occur at the subcrop beneath glacial cover. Samples were collected during low- and high-flow conditions between September 2012 and April 2014. In addition, surface-water samples were collected in September 2014 from several watersheds in the Porcupine Mountains; a short temporal variation in flow was examined due to a significant storm that occurred between sampling events. Sample locations included 1. upstream of, or not influenced by, the Nonesuch Formation, or 2. downstream of the Nonesuch (influenced) or downstream of mining impacted areas (impacted). The geochemistry of the surface waters reflects underlying rock types, glacially transported unconsolidated materials, mineralization style within each watershed, and geochemical processes occurring in the streams. Waters from the Duluth Complex and the Porcupine Mountains are similar in composition in that they are generally oxic, near neutral to slightly acidic (pH 5.0 to 7.6), and are characterized by moderate carbonate species concentrations (4 – 65 mg/L CaCO3 as bicarbonate) and low sulfate concentrations (< 0.8 – 8 mg/L). Calcium (Ca), sodium (Na), magnesium (Mg), and silica (SiO2) are the main dissolved major cations and the predominant dissolved trace elements include aluminium (Al), Cu, and Fe. Total dissolved solids, Ca, and Cu are generally lower in waters collected from the Duluth Complex compared to the Porcupine Mountains, whereas DOC, Fe, Al, and Ni are generally higher. Despite some variations in their chemistries, these watersheds display an atypical chemical signature when compared to most surface waters in the United States; the surface waters are rich in DOC (18 - 47 mg/L for Duluth; 2 - 22 mg/L for Porcupine Mountains) and have generally either low (10-53 mg/L CaCO3 for Duluth Complex) or low-moderate hardness (21-135 mg/L CaCO3 for Porcupine Mountains). The concentrations of major and trace elements vary seasonally and after large storm events with lower concentrations generally being found during higher flow conditions, consistent with dilution by rain or snowmelt. However, in the Duluth Complex watersheds, dissolved loads for major and trace elements are greater during higher flow conditions, likely from elements accumulating in wetlands and groundwater during dry and winter conditions and then being flushed downstream during higher flow. The aquatic toxicity of most metals (i.e., Ag, Cd, Cr, Cu, Ni, Pb, Zn) is routinely assessed on the basis of hardness-based criteria that adjust for the protective effects of Ca and Mg ions. In 2007, a new guideline was adopted by United States Environmental Protection Agency for determining aquatic life criteria for Cu that relies on the BLM. The BLM evaluates the biological availability of metals in aquatic systems for several organisms (i.e., fishes, water fleas) and incorporates major element chemistry as well as additional water-quality characteristics including metal speciation and organic carbon complexation (Paquin et al., 2002); BLMs for metals other than Cu, including Ag, Cd, Ni, Pb, and Zn, are also being developed. Concentrations of metals in surface waters can be evaluated with regard to the hardness-based and BLM-based criteria using a hazard quotient (HQ), which is the ratio of the dissolved concentration of the metal in the sample to the criterion. Values above 1 imply toxic conditions, whereas those below imply non-toxic conditions. Hazard Quotients based on the hardness-based criteria for Cu are greater than 1 for most sites in Filson Creek and in most waters collected after the storm and, in a few, before the storm in the Porcupine Mountains (Figure 1A), which is a reflection of the low hardness of these waters. However, as shown in Figure 1A, HQs for Cu calculated based on the BLM model are significantly less than 1 for some of 117 �these same sites, in particular for the Filson Creek watershed during low flow; for these sites, the hardness-based approach would predict toxic conditions, whereas the BLM model does not predict toxic conditions. The low HQs based on the BLM reflect the ability of DOC to complex Cu, rendering it unavailable biologically. In contrast, hardness-based HQs for a few sites in the Porcupine Mountains are lower than 1, whereas BLM-based HQs are greater than 1. The different results from the hardness-based and BLM-based approaches suggest that the former may be inadequate to describe metal toxicity especially in watersheds with high DOC and low to moderate hardness. The formation of Cu-DOC complexes significantly reduces the amount of dissolved Cu available to interact with the biotic ligand (the gill) of aquatic organisms. The protective effects of cations, such as Ca and Mg, competing with Cu to complex with the biotic ligand are likely not as important as DOC in many of these waters. The composition of DOC also influences its ability to mitigate metal toxicity and is an input parameter for the BLM model. The humic acid (HA) fraction of DOC is assumed to be the reactive fraction available to complex with dissolved metals. As shown in Figure 1B for a site in the Porcupine Mountains, varying the DOC concentrations and HA fraction in the BLM model significantly changes the predicted chronic water-quality criteria for Cu; the fraction of HA has a greater influence on the criteria as total DOC concentrations are increased (i.e., steeper slope). Based on the measured DOC (16 mg/L) and Cu concentration (38 µg/L) at this site, if the HA fraction of the DOC is less than approximately 40%, the Cu concentration in the water exceeds the criterion, predicting toxic conditions, whereas if HA is greater than approximately 40%, the Cu concentration is less than the criterion. It is also noteworthy that there is a limited range of DOC and HA values for which the BLM is calibrated (Figure 1B); waters collected in the Porcupine Mountains fall within the DOC calibrated range (0.05- 29.6 mg/L), whereas numerous samples from the Duluth Complex exceed the upper limit. Naturally-occurring concentrations of Cu at some sites currently exceed criteria prior to mining. The BLM approach to predicting aquatic water-quality criteria for Cu is likely needed for these waters in order to evaluate the effects of the high DOC and better predict resiliency to increased dissolved Cu concentrations. Additional investigations are needed to examine the composition of the DOC, which influences its ability to mitigate toxicity, as well as some of the limitations of the calibrated range for DOC concentrations in the BLM model. REFERENCES Paquin, P.R., Gorsuch, J.W., Apte, Simon, Batley, G.E., Bowles, K.C., Campbell, P.G.C., Delos, C.G., Di Toro, D.M., Dwyer, R.L., Galvez, Fernando, Gensemer, R.W., Goss, G.G., Hogstrand, Christer, Janssen, C.R., McGeer, J.C., Naddy, R.B., Playle, R.C., Santore, R.C., Schneider, Uwe, Stubblefield, W.A., Wood, C.M., and Wu, K.B., 2002, The biotic ligand model: A historical overview: Comparative Biochemistry and Physiology, v. 133, no. 1-2, p. 3–35. 118 �A preliminary evaluation of the structural controls on gold mineralization in the Jackfish Lake area, northwestern Ontario PUUMALA, Mark1, MAGNUS, Seamus2 1 Ontario Geological Survey, Ministry of Northern Development and Mines, Resident Geologist Program, Suite B002, 435 James St. South Thunder Bay, ON, P7E 6S7, Canada 2 Ontario Geological Survey, Ministry of Northern Development and Mines, Earth Resources and Geoscience Mapping Section, 933 Ramsey Lake Road, Sudbury, ON, P3E 6B5, Canada. Jackfish Lake lies within the Wawa Subprovince, in the western portion of the Schreiber-Hemlo greenstone belt. The supracrustal rocks in this area have been assigned to the Schreiber assemblage (Williams et al. 1991), and include approximately equal proportions of metavolcanic and metasedimentary rocks (Walker 1967). These supracrustal rocks have been intruded by a number of lateto post-tectonic plutons, including the Terrace Bay batholith and the Santoy Lake pluton. No age determinations are available for the supracrustal or intrusive rocks in this portion of the greenstone belt. However, they are presumed to be Neoarchean, based on their stratigraphic relationship to rocks that have been dated elsewhere in the belt (Magnus and Walker 2015). Several late- to post-tectonic faults and shear zones occur in the Jackfish Lake area (Walker 1967). These structures include a major northweststriking shear zone that was recently described by Magnus and Walker (2015) and occurs in the supracrustal rocks between the Terrace Bay batholith and Santoy Lake pluton (see Figure 1). Magnus and Walker (2015) also noted evidence of a late north-south-directed deformation event. As discussed below, this event may have played a role in the localization of gold mineralization at Jackfish Lake. The Jackfish Lake area has a long history of gold exploration that dates back to 1873, when Donald McKellar reported the discovery of gold at Victoria Cape and Jackfish (Schnieders et al. 1996). Since that time, numerous gold occurrences have been found in the Terrace Bay batholith, and in the supracrustal rocks that surround the intrusion. These occurrences include the Empress Mine, which produced 112 ounces of gold during 1896-97 (Schnieders et al. 1996). During 2015, work began on a study to examine the structural controls on gold mineralization in the vicinity of the Terrace Bay batholith. This paper presents the preliminary results of this work, which will be refined and expanded upon during a two year Ontario Geological Survey (OGS) – Lakehead University study set to commence in the summer of 2016. To date, field data have been collected from 11 gold occurrences in the Jackfish Lake area, near the eastern end of the Terrace Bay batholith. This field work has been complemented by the review of information that is on file for 20 other nearby gold and sulphide occurrences at the OGS Thunder Bay South Regional Resident Geologist’s office. Based on a review of the information collected to date, gold has been found to occur in the following structural settings.  Mineralized shear/fault zones that occur in supracrustal rocks near the margins of the Terrace Bay batholith.  Quartz-carbonate vein systems that parallel the batholith-supracrustal rock contact and are located at or near the contact.  Quartz-carbonate veins that occupy late brittle fracture systems within the batholith. The most notable gold mineralized shear zone in the supracrustal rocks is known as the Empress structure (shown on Figure 1). It is a 15 to 25 m wide shear zone that strikes 70˚ with a moderate to steep dip toward the south. This structure hosts a number of gold occurrences, including the Empress Mine. The most significant gold values are obtained where the shear zone is at its widest and exhibits intense folding and quartz vein development. Fold axes in this structure plunge moderately toward the east. The gold-mineralized contact- and batholith-hosted vein systems tend to have the following general characteristics.  One grouping consists of narrow (generally less than 0.2 m wide), shallow-dipping (often northstriking) to near-horizontal veins. Many of these veins occur in en-echelon sets. 119 � The second grouping of veins is approximately north-striking with near-vertical dip. Vein widths are highly variable, ranging from a few millimetres to more than a metre. All of the observations listed above suggest that gold mineralization is most likely to have been associated with a late north-south-directed compression event, such as that proposed by Magnus and Walker (2015). Gold generally appears to be associated with structural traps, both along the Empress structure (i.e., in zones of folding where dilatant structures have developed), and in the Terrace Bay batholith (brittle fractures). In the Empress structure, structural traps are most likely to occur at the nose of significant folds, where quartz veining has occurred in dilatant zones. In the Terrace Bay batholith, the mineralized fracture-fill veins are most likely to have been “dead-end” fractures that were connected on one end to structures that acted as the primary gold bearing fluid conduits. Figure 1. Gold and sulphide occurrences of the Jackfish Lake area (geology from Walker 1967 and Magnus and Walker 2015). All co-ordinates provided in NAD83 Zone 16. REFERENCES Magnus, S.J. and Walker, J. 2015. Geology and mineral potential of Walsh, Tuuri and Syine Townships, SchreiberHemlo greenstone belt; in Summary of Field Work and Other Activities 2015, Ontario Geological Survey, Open File Report 6313, p.14-1 to 14-12. Schnieders, B.R., Smyk, M.C., Speed, A.A. and McKay, D.B. 1996. Mineral occurrences in the Nipigon–Marathon area; Ontario Geological Survey, Open File Report 5951, 912p. Walker, J.W.R. 1967. Geology of the Jackfish-Middleton area, District of Thunder Bay; Ontario Geological Survey, Report 50, 41p. Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L. and Sage, R.P. 1991. Wawa Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.485-539. 120 �Setting of volcanogenic massive sulfide deposits of the Paleoproterozoic Penokean volcanic belt QUIGLEY, Ashley1, MONECKE, Thomas1, ANDERSON, Eric2, KELLY, Nigel3, and QUIGLEY, Patrick1 1 Department of Geology and Geological Engineering, Colorado School of Mines, 1516 Illinois Street, Golden, CO 80401 2 U.S. Geological Survey, MS 964, PO Box 25046, Denver, CO 80225 3 Department of Geological Sciences, University of Colorado-Boulder, UCB 399, Boulder, CO 80309 The Paleoproterozoic (ca. 1875 Ma) Penokean volcanic belt represents one of the world’s principal orogens hosting volcanogenic massive sulfide (VMS) deposits (Franklin et al., 2005). Sporadic exploration from 1970-1995 has identified a large number of VMS deposits and prospects throughout the belt, including the world-class Crandon deposit that comprises an estimated 61 million tonnes of polymetallic massive sulfide ore (Lambe and Rowe, 1987). Despite successful exploration and the significant economic potential in the Penokean volcanic belt, only limited academic research has been conducted focusing on constraining the tectonic, structural, and volcanic setting of the VMS deposits. Many key aspects of the regional geology are not well understood, which is mostly due to extensive glacial cover of the Paleoproterozoic bedrock. The present study aimed to use whole rock major and trace element geochemistry and U-Pb geochronology to investigate the geologic framework and tectonic setting of the belt and to constrain the timing of massive sulfide formation. Regional geochemical samples were taken from geophysically defined geologic domains. The high-precision chemical abrasion ID-TIMS U-Pb dating technique was used on zircon grains from seven felsic volcanic samples from the host rock successions of some of the major VMS deposits within the Penokean volcanic belt. The results of the geochemical analyses reveal subtle differences between the geophysical domains. The majority of the volcanic rocks sampled have a tholeiitic affinity with fewer calc alkaline and transitional rocks. Geochemical evidence points to an island arc subduction environment with either intra-arc or back-arc extension. Results of the ID-TIMS U-Pb geochronological analyses revealed that four of the deposits, namely Bend, Horseshoe, Lynne, and Pelican River were all formed at about 1874 Ma. This is evidence of a major period of volcanism and related VMS deposition that likely occurred in an extensional setting at this time. The Back Forty massive sulfide deposit occurs on the east end of the Penokean volcanic belt. High-precision ID-TIMS U-Pb zircon geochronology of the host rhyolite yielded an age of about 1833 Ma, which is approximately 50 million years younger than the host rock successions of the other deposits of the Penokean volcanic belt. There are two possible explanations for this apparent age. The first explanation is that this represents a crystallization age and the host rocks to the Back Forty are part of a distinctly younger volcanic succession. Alternatively, a thermal event at 1833 Ma may have reset the U-Pb age. The rhyolite sampled from the host rock successions of the Lynne and Back Forty deposits contain rare Archean-aged zircon grains. These zircon grains yielded U-Pb ages of approximately 2700 Ma and are presumably inherited from an Archean basement. These data 121 �support the model that the Superior craton extends beneath the Pembine-Wausau terrane south of the Niagara Fault. REFERENCES Franklin, J.M., Gibson, H.L., Jonasson, I.R., and Galley, A.G., 2005, Volcanogenic massive sulfide deposits: Economic Geology 100th Anniversary Volume, p. 523-560. Lambe, R.N., and Rowe, R.G., 1987, Volcanic history, mineralization, and alteration of the Crandon massive sulfide deposit, Wisconsin: Economic Geology, v. 82, p. 1204-1238. 122 �Expanding the historical exploration document collection at the Minnesota Department of Natural Resources: the Polaris Joint Venture exploration program REED, Andrea, FREY, Barry, and HANSON, Kevin Mineral Potential Evaluation Section, Lands and Minerals, Minnesota Department of Natural Resources The Polaris Joint Venture was a large-scale five year-long exploration program targeting copper-zinc prospects in the greenstone belts of northern Minnesota during the early 1980s. Ernest K. Lehmann & Associates, in partnership with Getty and Billiton, completed airborne and ground geophysical surveys, geologic mapping, geochemical sampling, and 50 drill holes within a 13,000 square mile study area that stretched roughly from Bemidji to Ely. In December 2014, the Lehmann Family Foundation donated the Venture’s results and $10,000 to the Minnesota Department of Natural Resources (DNR). Due to the large areal extent of the donation and quantity of work conducted in previously unexplored areas, the DNR decided to make the data digitally available to the public as soon as possible. Due to the manageable size of this donation, it also gave the DNR the perfect opportunity to experiment with new ways of managing its historical exploration document collection. The full historical exploration document collection can be found online at http://minarchive.dnr.state.mn.us/information.html (Fig. 1). Resulting from the experimentation process with the Polaris Joint Venture data is the website: http://www.dnr.state.mn.us/lands_minerals/polaris/index.html (Fig. 2). From this site, the user should be able to smoothly navigate through and find the wanted data using the DNR’s downloadable Microsoft Excel catalog, downloadable GIS shapefile and PDF links to the data, as well as being able to spatially search online through the web map application. Updates to the Polaris Joint Venture collection are planned and the DNR will be accepting input from outside users. Once the Polaris Joint Venture collection portal is finalized, the DNR’s next steps will be to adapt the new historical document management system to accommodate the other collections. Fig. 1. The old historical document collection system. Fig. 2. A new historical document collection and organization system. 123 �Characterizing deformation of Gunflint Formation in contact with Archean basement rocks east of Thunder Bay, Ontario REID-SHARP, Ruby and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada New expansion along Highway 11/17 east of Thunder Bay, Ontario exposes faults and damage zones that cut through and displace the Gunflint Formation of the Animikie Group, the underlying Archean basement rock, and the basal unconformity between them. Observations and stereographic projections were used to characterize outcrop-scale deformation along a curved road-cut on a hill approximately 1km long and covering an increase in elevation from west to east (where the unconformity is exposed) of about 65m. The Highway 11/17 road cut exposes a fault zone network that has accommodated a minimum of 60m of vertical displacement and unknown horizontal displacement. Two sets of well-preserved slickenlines record normal oblique sense movement on the fault zone. A void-filling calcite vein system is exposed in conjunction with the fault systems and is associated with brecciation. Calcite crystals reach up to 10cm in diameter where larger voids are filled. At least two major fault zones with damage zones at least 1m wide are exposed at the study site and are seen at the far east and far west ends of the road-cut. The Eastern Fault Zone (EFZ) has an average strike and dip of 071°/72° cutting through both Archean basement rock and Gunflint Formation. The Western Fault Zone (WFZ) average strike and dip is 067°/65° and cuts through Gunflint Formation rocks. Two relatively minor reverse faults striking 345° and dipping 48° to 58° cut through the Gunflint Formation about 80m southwest of where the WFZ first appears. The relationship between the EFZ and WFZ is obscured by an oblique fault cutting through Archean rock between them with an average strike and dip of 040° and 60°. The site is a potentially important study location for understanding fault/fluid interactions and how extensional faulting can accommodate fluid transportation. All major extensional features in the study area have a roughly northeast strike which aligns well with the orientation of the Mid-Continent Rift to the southeast suggesting that this fault zone is associated with extension ca. 1Ga. 124 �The building blocks of stromatolites: Comparisons across time and environment REINERS, Lindsey, EISCHEN, Tanner, and BARTLEY, Julie K. Geology Department, Gustavus Adolphus College, St. Peter, MN 56082 Approximately one billion years separate the stromatolites of the Shakopee (Ordovician, Minnesota) and Rossport (Mesoproterozoic Sibley Group, Ontario) formations, yet they may share similar carbonate building blocks, despite stark differences in age, large-scale morphology, and depositional environment. The Shakopee Formation (Ordovician) was deposited in a shallow epeiric sea that covered much of the upper Midwest. Stromatolites in the Shakopee are diagnostic of the carbonate-rich, nearshore facies of the Shakopee (Davis, 1966; Johnson and Simo, 2002), perhaps marking the presence of restricted, warm, shallow marine environments. Stromatolites range in macroscopic form from lowrelief, nearly stratiform morphotypes through low-relief columnar forms to internally complex domal structures. In contrast, the Rossport Formation (Mesoproterozoic) accumulated in an inland lacustrine environment that was at least intermittently hypersaline (Rogala et al., 2007). Robust stromatolite development is restricted to the Middlebrun Bay Member of the Rossport and represents an interval dominated by carbonates within an otherwise mainly siliciclastic succession. In macroscopic form, the Rossport stromatolites are stratiform to columnar, with low inheritance and low synoptic relief. Despite these differences in environment and macroscopic outcrop expression, the two stromatolite occurrences share similar mesoscale to microscale textures – the structural building blocks of stromatolites. Although post-depositional processes, including recrystallization, silicification, and dolomitization, have affected the present-day preservation of these stromatolites, primary textures can be inferred in both units, permitting interpretation of the original carbonate building blocks of these ancient structures. At the mesoscale (visible in hand sample) level, stromatolites show isopachous lamination, oversteepened lamina, and complex internal textures that are characteristic of typical Proterozoic stromatolites formed by in situ precipitation of carbonate during stromatolite growth. At the microscale (visible under a microscope), stromatolites from both localities preserve (1) evidence of a primary micritic texture, with peloidal or clotted form, commonly associated with microbially-influenced carbonate precipitation, and (2) syndepositional cement growth coating stromatolites and forming the mesoscale isopachous laminae. Both stromatolite occurrences lack evidence of trapping-and-binding as a mechanism for stromatolite growth. when. Similarities in stromatolite structural components across time and space suggest that the mechanisms of microbialite construction are broadly comparable in many settings, including ancient marine as well as both ancient and modern lacustrine environments. Modern marine stromatolites, such as those in the Bahamas and building blocks and are likely the outliers in microbialite construction across time and space. REFERENCES Davis, R.A. Jr. 1966. Willow River Dolomite: Ordovician analogue of modern algal stromatolite environments: Journal of Geology v. 74, p. 908-923. Johnson, C.L., and Simo, J.A. 2002. Sedimentology and sequence stratigraphy of a lower Ordovician mixed siliciclastic-carbonate system, Shakopee Formation, Fox River Valley of east-central Wisconsin: Geoscience Wisconsin v. 17, p. 21-33. Rogala, B., Fralick, P.W., Heaman, L.M., Metsaranta, R. 2007. Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada: Canadian Journal of Earth Sciences v. 44, p. 1131-1149. 125 �Evaluating H/V analysis of passive seismic data as a means to map sediment thickness in the Duluth-Superior harbor SAGER, Tyler A. and WATTRUS, Nigel. of Earth and Environmental Sciences, University of Minnesota, Duluth, MN Drilling during the construction of the Interstate High Bridge between Duluth, MN and Superior, WI documented: 1) significant variations in depth to bedrock, showing that the thickness of the sediment cover increases towards Duluth; 2) that there are significant amounts of anthropogenic material (dredge spoil, waste material linked to historical sawmill activity in the harbor) in the areas surrounding the modern day harbor. The seismic velocity structure of near-surface sediments can be determined from H/V analysis of passive seismic data collected with a 3-component geophone (ie the instrument measures ground motion in three orthogonal directions). This technique utilizes the ambient background seismic noise generated by nearby human activities (such as traffic or heavy machinery), winds and other atmospheric phenomena, and ocean waves. The response of the near-surface to excitation by ambient seismic noise varies with its’ seismic velocity structure. Specifically, it exhibits unique seismic resonance frequencies that are linked to the structure of the near-surface. The H/V spectral ratio method was originally introduced by Nogoshi and Igarshi (1971) with further development by Nakamura (1989). At each recording site, amplitude spectra are calculated for each of the three-component records of the ambient seismic noise. The ratio between the amplitude spectra of the horizontal (H) to vertical (V) frequency components identifies the resonance frequencies within the near-surface sediment package (SESAME, 2005). The lowest (fundamental) resonance frequency can be used to determine the thickness of the surface sediment layer. H/V analyses of passive seismic observations made at multiple sites in the Duluth-Superior harbor area, are used to create a map depicting the spatial variation in seismic resonance frequency in the area. A map of sediment thickness in the harbor area is derived from this map by applying an empirical calibration function that links seismic resonance frequency to sediment thickness. This relationship is derived by making resonance measurements at sites with well control. A more detailed interpretation of the near-surface shear-wave velocity structure is undertaken at selected sites. At these sites, the recorded data is used to construct a model of the near-surface velocity structure by iteratively perturbing a trial model, so that its calculated response matches the recorded data. This profile can yield valuable geotechnical information (specifically stiffness) about the near surface sediments (Lai and Rix, 1998; Xia et al., 1999). REFERENCES Lai, C. G. and Rix, G. J.: Simultaneous inversion of Rayleigh phase velocity and attenuation for near-surface site characterization, Georgia Institute of Technology, 1998 Nogoshi M. and Igarashi T. (1971) On the amplitude characteristics of microtremor (part 2) (in Japanese with English abstract). Journal of Seismological Society of Japan, 24, 26-40. Nakamura, Y. (1989) A method for dynamic characteristics estimation of subsurface using microtremor on the ground surface. Quarterly Report of the Railway Technical Research Institue 30 (1), 25-30. SESAME, 2005. Guidelines for the implementation of the H/V spectral ratio technique on ambient vibrations measurements, processing and interpretation. SESAME European Research Project, Deliverable D23.12, 62p. Xia, J., Miller, R. D., and Park, C. B.: Estimation of near-surface shear-wave velocity by inversion of Rayleigh waves, Geophysics, 64, 691–700, 1999 126 �Geochemical and Petrological Comparisons of Peridotite Units in Marquette County, Michigan SASSO, Andrew, and THAKURTA, Joyashish Department of Geosciences Western Michigan University1903 W Michigan Ave Kalamazoo MI 49008-5241 USA This study characterizes the following rock units in Marquette County, Michigan in terms of geochemistry and petrology: (1) Presque Isle Peridotite, (2) Deer Lake Peridotite, and (3) Yellowdog Peridotite. Analyses were conducted to determine if any petrological or geochemical relationships exist between these units, and to assess the potential of these units to host magmatic sulfide deposits. Based on the findings, these units have been separated into the two groups presented below. Deer Lake Peridotite Mineralogical compositions and textural characteristics of Deer Lake Peridotite are similar to those observed in the Presque Isle Peridotite and Yellowdog Peridotite. Prior to alteration, the unit’s mineralogical composition was dominated by olivine and pyroxene. This rock type displays a cumulate texture. In some cases a poikilitic texture in which pyroxene oikocrysts enclose olivine chadacrysts is also present. However, these similarities are not sufficient to conclude that this unit shares a common origin with either of the other peridotite units. Also, it is noteworthy that the degree of serpentinization and hydrothermal alteration observed in the Deer Lake Peridotite is different, being far greater than that observed in either of Marquette County’s other peridotite units. Geochemical comparison of Deer Lake Peridotite with the other three units addressed in this study reveals obvious differences in chemical composition. Geochemical analysis also reveals that the Deer Lake Peridotite crystallized from a magma which formed as a result of shallow melting; whereas, the parent magmas of Presque Isle Peridotite and Yellowdog Peridotite formed as a result of deep melting. Truncation of the Deer Lake Peridotite along its south-western margin by the Great Lakes Tectonic Zone suggests that the unit must have been formed either during, or prior to the formation of the GLTZ (2.7-1.85 Ga). This window of time for this formational event proves that the Deer Lake Peridotite predates the Yellowdog Peridotite’s age (1.1 Ga) by no less than 750 Ma. Trace element, and petrographic and geochemical analysis also suggest that Deer Lake Peridotite may actually represent two separate peridotite units emplaced during two separate events. Type 2 Deer Lake Peridotite displays foliation which is not present in Type 1 Deer Lake Peridotite. Possibly, Type 2 Deer Lake Peridotite was emplaced early in the formation of the GLTZ, and later deformed during the Penokean Orogeny (1.86-1.83 Ga), with Type 1 crystallizing during this later compressive phase. Bornhorst et al. (1993) postulated that the Deer Lake Peridotite may represent the subvolcanic base of the Mona Formation. Deer Lake Peridotite also appears to be associated with the metavolcanics of the Kitchi Formation. It is possible that these successive metavolcanic units may correspond to the two ultramafic units of Deer Lake Peridotite. Based on geochemical and petrographic analysis, in conjunction with the geologic setting of the Deer Lake Peridotite, it is reasonable to conclude that the Deer Lake Peridotite was formed independently, and substantially earlier, than Presque Isle Peridotite and Yellowdog Peridotite. It is likely that this unit crystallized from a parent magma which resulted from shallow melting during the formation of the Great Lakes Tectonic Zone. Presque Isle Peridotite and Yellowdog Peridotite Presque Isle Peridotite and Yellowdog Peridotite, upon first examination, appear to be very different units. Presque Isle Peridotite is more finely grained, and lacks visible plagioclase feldspar crystals, such as those observed in the Yellowdog Peridotite. Presque Isle Peridotite has also undergone a notably higher degree of serpentinization, as is evidenced by a greater density of hydrothermal veins. 127 �Additional examination also reveals that primary mineral assemblages in both units include a large fraction of olivine. Pyroxene, mostly in the form of augite, with a much smaller fraction of enstatite, also constitutes a substantial fraction of both units. They both display a cumulate texture. Additionally, a poikilitic texture in which rounded olivine chadacrysts are partially, or fully enclosed by pyroxene oikocrysts can be observed in both the Presque Isle Peridotite and Yellowdog Peridotite. Petrographic analysis of both units also reveals olivine hosted sulfide inclusions. Such inclusions may indicate the presence of an immiscible sulfide liquid at the time of crystallization. Geochemical comparison of the Presque Isle Peridotite and Yellowdog Peridotite indicates that both formed from a parent melt of mantle origin, and became contaminated by crustal material prior to crystallization. Minor trace element analysis of the two units, reveals that they share a very similar geochemical composition. Limited outcrop exposure of the Yellowdog Peridotite, and the lack of accessibility to host rock contacts at the Presque Isle Peridotite, make it difficult to draw conclusions based strictly on geologic setting relationships at both sites. However, Yellowdog Peridotite has intruded through the Late-Archean granite basement of the “Northern Complex”, and Paleoproterozoic sediments of the Baraga Basin (as confirmed by the exploration teams of both Kennecott and Lundin Mining). Ding et al. (2010), by the use of U-Pb baddeleyite dating, has confirmed that this unit crystallized at 1107.2 ± 5.7 Ma. This date allows for the reasonable conclusion that the Yellowdog Peridotite formed during the Midcontinent Rift event. The only unit which has been observed in direct contact with the Presque Isle Peridotite is the uppermost member of the Keweenawan series, the Jacobsville Sandstone. Here, it is clear that Presque Isle Peridotite is nonconformably overlain. Radiometric dating of zircons from Jacobsville Sandstone confirms that the unit is no younger than 960 Ma (Malone et al., 2015). Observation of the peridotite’s contacts with other units is not possible because they are concealed beneath Lake Superior. However, it can be safely assumed that the Presque Isle Peridotite has also intruded through the Archean granite basement of the “Northern Complex”. No geochronological dates have ever been obtained for the Presque Isle Peridotite. The findings suggest the possibility that the Presque Isle Peridotite may also date to the Mesoproterozoic, at which time, it may have been formed contemporaneously with the Yellowdog Peridotite, during the early stages of the Midcontinent Rift event (1.1 Ga). Geochemical similarities between these units also suggest the likelihood that the plume induced, parent melts of both units were very similar, and may have been directly related. Presque Isle Peridotite is shown to display the following: (1) High nickel content, comparable to that of the Yellowdog Peridotite (as shown by XRF analysis). (2) Sulfide inclusions within olivine. (3) Primary magmatic sulfide assemblages of chalcopyrite and pentlandite. (4) Incompatible element enrichment. Comparison of Ni content with Fo molar percentage, also yields results similar to those observed in both the “Eagle” and “Eagle East” intrusions of Yellowdog Peridotite. All these factors make Presque Isle Peridotite a prime target for future exploration, as they suggest the possibility that the unit has the potential to host magmatic sulfide deposit similar to those hosted by both intrusions of the Yellowdog Peridotite. It can be concluded from the available data, that peridotite units of Keweenawan age, located in the Lake Superior region should be considered high priority targets for magmatic sulfide exploration. References Bornhorst, T. J. and R. C. Johnson, 1993, Geology of Volcanic Rocks in the South Half of the Ishpeming Greenstone Belt, Michigan. United States Geological Survey. USGS Publication Warehouse. Ding, Xin, Chusi Li, Edward M. Ripley, Dean Rossell, and Sandra Kamo, 2010, "The Eagle and East Eagle Sulfide Ore-bearing Mafic-ultramafic Intrusions in the Midcontinent Rift System, Upper Michigan: Geochronology and Petrologic Evolution." Geochemistry Geophysics Geosystems 11.3. Malone, David H., Carol A. Stein, John P. Craddock, Jonas Kley, Seth Stein, and John Malone, 2015, "Maximum Depositional Age of the Neoproterozoic Jacobsville Sandstone, Michigan: Implications for the Evolution of the Midcontinent Rift." Jacobsville/MCR. Illinois State University. 128 �Metal isotopic signatures in the Duluth Complex associated with magmatic Cu-Ni-PGE mineralization SCHARDT, Christian Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 1114 Kirby Dr. Duluth, MN 55812 USA Magmatic Cu-Ni-PGE sulfide mineralization occurs in a series of deposits along the western margin of the Duluth Complex in northeastern Minnesota. While the mineralogy, geochemistry, and formation of these deposits has been investigated in some detail, there is little information available on the metal isotopic signatures (Ni, Cu) of these deposits and if this information allows to I) gain more insight into the genesis and evolution of these deposits and II) holds potential as an exploration tool. Recent studies have demonstrated that Ni isotopic fractionation exists in both high and low-temperature terrestrial material with a range of 2.1 ‰ [1-3]. Fractionation of Cu isotopes in magmatic systems has been established and significant differences between primary and secondary sulfides are noted [4,5]. For the Duluth Complex, recently reported Cu isotopic values range from -0.85 to 0.45 ‰ [6]. Sulfide-bearing drill core samples (massive, disseminated) and sulfide-barren material from all major deposits in the Duluth Complex were collected and analyzed, along with weathered surface material from mineralized outcrops and till samples. δ60/58Ni values in olivine (ave: -0.03 ‰), sulfides (ave: 0.36 ‰), and secondary oxides (ave: -0.50 ‰) are comparable to previous work [1-2] and indicate i) high temperature fractionation of Ni into olivine with bulk silicate earth signature (BSE) during initial crystallization, iii) increasing lighter Ni isotope incorporation as a function of sulfide content, and iii) preferred 58Ni incorporation into secondary oxides and silicates due to weathering processes. 60Ni is presumed to enter solution during weathering, supported by isotopic signatures of garnierite (1.5 ‰) from a magmatic sulfide deposit in Germany. δ65/63Cu values for the Duluth Complex deposits range from -1.28 ‰ to 0.36 ‰ (Figure 1), comparable to published data [6]. Other magmatic Cu-Ni-PGE deposits in the area (Eagle, Tamarack; massive sulfides) show a distinctly heavier isotopic signature (> 0.69 ‰) while disseminated ore material from Eagle is much lighter (-0.16). While both massive and disseminated sulfides from all deposits in the Duluth Complex show similar Cu isotopic values (ave: -0.32 to -0.35 ‰), individual deposits differ in their isotopic signature and are quite distinct from each other. Surface material generally shows variable negative enrichment of 63Cu, attributed to the weathering process and the preferential enrichment of lighter Cu isotopes into the weathering products. Due to the difference in Ni isotopic values between unmineralized material (ave: -0.05) and sulfide-bearing samples (up to -0.97 with increasing sulfide content), Ni isotopes may be useful to distinguish barren magmatic systems (BSE signature) from mineralized systems because of the preferential incorporation of isotopically light Ni into sulfide-bearing rocks. The significance and use of Cu isotopes in the Duluth Complex is less clear since a variety of processes (Cu source, sulfide fractionation) likely influenced the overall Cu isotopic signature. However, if a mantle signature of ~ 0 ‰ is assumed, the relative contribution of sedimentary Cu in the Duluth Complex (ave: 0.97 ‰) and magmatic Cu may be determined and differences in the formation of individual deposits in the Duluth Complex assessed. 129 �Figure 1: Cu isotopic values of primary and secondary sulfide material from Duluth Complex Cu-Ni-PGE ore deposits compared to other magmatic sulfide deposits and selected VMS deposits. Data for orange bars are taken from [6]. References 1. 2. 3. 4. 5. 6. Gueguen B., Rouxel O., Ponzevera E., Bekker A., Fouquet Y. (2013) Ni isotope variations in terrestrial silicate rocks and geological reference materials measured by MC-ICP-MS. Geostandards and Geoanalytical Research, v. 3, p. 297-317 Hiebert RS., Rouxel, O., Houlé, MG., Bekker, A. (2014) Ni isotope fractionation between komatiite and sulfide mineralization at the Neoarchean Hart deposit, Abitibi greenstone belt, Canada. Geological Society of America Abstracts, v. 46: p. 467 Wasylenski, L.E, Howe, Haleigh D., Spivak-Birndorf, L.J., Bish, DL. (2015) Ni isotope fractionation during sorption to ferrihydrite: implications for Ni in banded iron formations. Chemical Geology, v. 400, p. 56-64 Larson, P.B., Maher, K., Ramos, F.C., Chang, Z., Gaspar, M., Meinert, L.D. (2003). Copper isotope ratios in magmatic and hydrothermal ore-forming environments. Chemical Geology 201: 337-350 Markl, G., Lahaye, Y., Schwinn, G. (2006) Copper isotopes as monitors of redox processes in hydrothermal mineralization. Geochemica et Cosmochimica Acta 70: 4215-4228 Ripley, E.M., Dong, S., Li, C., Wasylenski, L.E. (2015) Cu isotope variations between conduit and sheet-style Ni–Cu–PGE sulfide mineralization in the Midcontinent Rift System, North America. Chemical Geology, v. 414, p. 59–68 130 �Acid-Generating and Acid-Neutralizing Potential of Silicate Rocks from the Basal Mineralized Zone of the Duluth Complex, Minnesota SCHULTE, Ruth F.1, PIATAK, Nadine M.1, SEAL, Robert R., II1, and WOODRUFF, Laurel G.2 1 2 U.S. Geological Survey, MS 954, Reston, VA 20192 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 The acid-generating potential (AP) and acid-neutralizing potential (NP) of the basal zone of the Mesoproterozoic Duluth Complex are atypical of most mafic hosted copper-nickel-platinum-group-metal (Cu-Ni-PGM) ore deposits because: 1) the sulfide ores are typically disseminated, and 2) the gangue assemblages generally lack carbonate minerals; thus, any acid-neutralizing potential must be derived from more slowly reacting silicate minerals. Recent technological advances in mineral processing have improved the viability of developing these low grade, large tonnage Cu-Ni-PGM, which represent some of the largest undeveloped deposits of their type on earth (Eckstrand and Hulbert, 2007; Miller and Nicholson, 2013). The potential development of these resources necessitates an understanding of their environmental behavior including assessment of the AP and NP of future mine waste. The 1.1 Ga Duluth Complex is a large (> 5,000 km2) composite body of troctolitic, gabbroic and anorthositic intrusions located in northeastern Minnesota. The complex is part of the Midcontinent Rift system that occurs in the Lake Superior region. The intrusions were emplaced into footwall rocks of Archean granite and greenstone terranes, Proterozoic metasedimentary rocks (Biwabik Iron Formation and Virginia Formation) and penecontemporaneous North Shore Volcanic Group basalts. In order to evaluate the AP and NP of potential mine waste from the Duluth Complex, five drill cores were sampled at the Minnesota Department of Natural Resources Core Library in Hibbing, Minnesota. Drill core in the vicinity of the NorthMet (26075), Local Boy (10089 and 10107), Mesaba (B1-95), and Maturi (DU-14) Cu-Ni-PGM deposits was included in this study. Mineralogical, petrographic, and geochemical analyses were undertaken to evaluate NP contained in carbonate and silicate minerals and AP contained in sulfides. For most of the drill cores, the main analytical focus was on samples in the basal zone hosting disseminated sulfide mineralization, although representative samples from above and below the mineralized zones were also assessed to evaluate variations in AP and NP based on rock type and stratigraphic position. If the Cu-Ni-PGM deposits are developed, mine waste will consist of processed rock from the mineralized zones as well as material adjacent to those zones. The ability of mine waste to generate or neutralize acid is commonly estimated using AP and NP values. In general, if the acid consumption or neutralization value (NP) exceeds the acid producing value (AP), the material will likely not be a source of acidity; whereas, if AP exceeds NP, the material may generate acid. The AP is based on the sulfide content and is typically estimated from total sulfur (S) content and calculated assuming all S occurs as the acidgenerating mineral pyrite (Sobek et al., 1978). The NP of mine waste is typically based on direct or, more commonly, indirect measurement of its carbonate content. Carbonate minerals are considered the most effective neutralizing agents to mitigate the acid generated by sulfide oxidation in solid mine waste. However, this is problematic in deposits without significant amounts of carbonate minerals, such as those in the Duluth Complex. For these cores, we examined the potential for other minerals, particularly silicates, to contribute to the NP; if these deposits are developed, silicate minerals would constitute the bulk majority of the mine and milling waste. The main silicate minerals identified in drill cores from the Duluth Complex include plagioclase feldspar, olivine, pyroxene, mica, amphibole, and secondary minerals, such as serpentine 131 �and chlorite. The NP contribution from the silicate minerals was calculated based on NP values of monomineralic samples from Jambor et al. (2002; 2007) and weighted by their abundances in the drill core samples from powder X-ray diffraction (XRD) analyses. Olivine contributes the most NP of the silicate minerals, at 38 kilograms of calcium carbonate equivalent per ton (kg CaCO3/t) (Jambor et al., 2007). The NP of plagioclase feldspar depends on the anorthite content and ranges from less than 1 up to about 14 kg CaCO3/t (Jambor et al. 2007). The NP of pyroxene is 4.6 kg CaCO3/t (Jambor et al. 2007). As shown in Figure 1, calculated NP values from these cores range from 6.7 to 19 kg CaCO3/t. Because the carbonate content is low in the Duluth Complex core samples, the NP contribution from the silicate minerals is significantly greater than NP from carbonate (Fig. 1). It is important to note that not all of the drill cores are predicted to be acid generating (Fig. 2). Specifically, all drill core samples from 26075 (NorthMet deposit) and most from DU-14 (Maturi deposit) have AP values that are less than their NP values. As a result, the NP in these cores may be sufficient to neutralize acid generated by sulfides. Furthermore, if we assume that 90 percent of sulfides will be recovered during mineral processing, the AP values of mill tailings decrease dramatically; thus, bulk samples previously predicted to generate acid no longer fall into the acid-generating category because most of the sulfides have been removed (Fig. 2). This is particularly important in drill core samples with significant sulfide contents, such as the Local Boy ore zone (drill cores 10089 and 10107). NP and AP also can be discussed in terms of a neutralization potential ratio (NPR), where NPR = NP/AP. For these cores, NPR values vary from 0.1 to 14.2 assuming no sulfide has been recovered during mineral processing. Generally, NPR ratios less than 1 are considered acid producing and ratios greater than 2 are non-acid producing (Price 2009). NPR ratios between 1 and 2 are capable of producing acid rock drainage. Based on these benchmark values, some samples from the Duluth Complex Cu-Ni-PGM deposits could contribute to the production of acid mine drainage (Fig. 2). However, at 10 percent of the AP (assuming 90% sulfide recovery), the neutralization potential ratios are all above 1, with an average NPR value of 49, suggesting the mine waste generated after recovering sulfides from the ores may not be a source of acidity. Although silicates weather more slowly than carbonates, the NP of the silicates can facilitate the neutralization of acids generated in mine waste. Neutralization from silicate minerals is likely to have the greatest impact in tailings storage facilities where groundwater should have longer residence times due to the significantly finer grain size of the material compared to that of typical waste rock. REFERENCES Eckstrand, O.R., and Hulbert, L.J., 2007, Magmatic nickel/copper/platinum-group element deposits, in Goodfellow, W.D., ed., Mineral deposits of CanadaA synthesis of major deposit types, district metallogeny, the evolution of geological provinces, and exploration methods: Geological Association of Canada, Mineral Deposits Division, Special Publication 5, p. 205-222. Jambor, J.L., Dutrizac, J.E., Groat, L.A., and Raudsepp, M., 2002, Static tests of neutralization potentials of silicate and aluminosilicate minerals: Environmental Geology, v. 42, p. 1-17. Jambor, J. L., Dutrizac, J.E., and Raudsepp, M., 2007, Measured and computed neutralization potentials from static tests of diverse rock types: Environmental Geology, v. 52, p. 1019-1031. Miller, J. and Nicholson, S.W., 2013, Geology and mineral deposits of the 1.1 Ga Midcontinent Rift in the Lake Superior region– an overview, in Miller, J., ed., Field guide to the copper-nickel-platinum group element deposits of the Lake Superior Region: Journal of Precambrian Research Center Guidebook, 13-01. Price, W.A., 2009, Prediction manual for drainage chemistry from sulphidic geologic materials: Mine Environment Neutral Drainage (MEND) Report 1.20.1, 579 p. Sobek, A.A., Schuller, W.A., Freeman, J.R. and Smith, R.M., 1978, Field and laboratory methods applicable to overburden and minesoils: Environmental Protection Agency, 600/2-78-054, 203 pp. 132 �The Geochemistry of the Siemens Creek Formation and the Nature of Early Midcontinent Rift Basaltic Magmatism in the Western Lake Superior Region SCHULZ, Klaus J., and NICHOLSON, Suzanne W. U.S. Geological Survey, 12201 Sunrise Valley Drive, MS 954, Reston, VA 20192 The Siemens Creek Formation is the lower member of the Powder Mill Group and represents the earliest (>1108 Ma) exposed lava flows of the Midcontinent Rift in northern Michigan and Wisconsin (Hubbard, 1975). It extends from Sturgeon Falls in Michigan to as far west as at least Atkins Lake in Wisconsin, although it is disrupted by the intrusion of the Mellen Complex near Mellen, Wisconsin (Cannon and others, 2006). It is best exposed in a prominent east-west belt of knobby hills north of U.S. Route 2 between Bessemer, Michigan and Upson, Wisconsin. Near Ironwood, Michigan the Siemens Creek Formation is about 1,340 m thick (Hubbard, 1975). The Siemens Creek Formation lies conformably upon the Bessemer Quartzite; rare quartzite interbeds also occur in the lower 30 m. The basal flows, 35 to 50 m thick, are typically pillowed, strongly chloritized, and at least locally contain hyaloclastite breccia. Basal flows west of a point about midway between Upson and Hurley, Wisconsin typically have olivine and pyroxene phenocrysts and range from picrite to basalt. In contrast, to the east of that point the basal flows mostly lack olivine-pyroxene phenocrysts and range from basalt to andesite. The western basal pillowed flows are locally overlain by a stromatolitic carbonate layer that divides the Siemens Creek into lower and upper members (basalt Types I and II, respectively of Nicholson and others, 1997). The upper lava flows are generally thinner (3 to 14 m) than the lower flows and were subaerially erupted with pahoehoe and sparsely vesicular flow tops. The upper flows range from basalt to andesite with more evolved andesites containing small plagioclase phenocrysts. The western basal flows (Type I olivine-pyroxene-phyric picrites and related basalts) are geochemically distinct from the upper Siemens Creek Formation (Type II basalts). The basal Type I flows are characterized by high MgO (~8 to 17%), FeOt (~13 to 16%), and TiO2 (~2-3%), and low Al2O3 (~8 to 10%). Those samples that are chloritized have high water contents (2 to 3%), and also appear to have lost Na2O (have high K2O/(K2O+Na2O)). Based on the major element characteristics, the Type I flows would be classed as ferropicrites and ferrobasalts (Gibson and others, 2000). Their trace element characteristics also are distinctive with steep, strongly light REE enriched REE patterns, high high-field-strength element and V contents, and primitive mantle normalized patterns that peak at Ta, have a saddle at Zr-Hf, and positive V anomalies (Fig. 1A). Isotopically these rocks have an initial ɛNd of ~0. The trace element characteristics are similar to some ocean island basalts (OIB), particularly alkaline HIMU OIB, and alkaline meymechites (low Al, high Mg rocks) in the Siberian Traps of Russia (Fig. 1A). They are also similar to the basal Keweenawan lava flows at Ely’s Peak near Duluth, in the Grand Portage area, and at the base of the Osler Group in Ontario (Fig. 1B). The eastern basal flows, although also pillowed, are Upper Siemens Creek Type II basalts in composition. Type II basalts are broadly similar in overall composition, although in detail three compositional types can be distinguished that appear to have distinct stratigraphic and lateral distribution (Type IIA, B, C in Fig 1C). The basal flows (Type IIA) extend from about the midpoint between Upson and Hurley eastward to Sturgeon Falls. They have high SiO2 ~50 to 58%, MgO ~5 to 10% (one picrite with 17%), FeOt ~10 to 13%, Al2O3 ~10 to 15%, and TiO2 ~1.5 to 2.5% with low Al2O3/TiO2 ratios (5-8). They show light REE enrichment with positive Th and mostly large negative Nb-Ta anomalies when normalized to primitive mantle (IIA in Fig. 1C). The Type IIB flows are directly above the basal flows and extend from east of Mellen to just east of Lake Gogebic. They show a smaller compositional range with SiO2 ~50 to 55%, MgO ~5 to 8%, FeOt ~10 to 12%, Al2O3 ~14 to 16%, and TiO2 ~ 1.4 to 2% with Al2O3/TiO2 ratios of 8 to 10. They are more enriched in light REE than the Type IIA basalts and most do not show negative Nb-Ta anomalies when normalized to primitive mantle (IIB in Fig. 1C). The last type (Type IIC) occurs as the uppermost Upper Siemens Creek flows in the area between just west of Hurley to Silver Mountain. They overlap in composition with the basalts below but mostly have lower TiO2 contents (~1 to 1.8%) and higher Al2O3/TiO2 ratios (9 to14). Their trace elements are similar to those of 133 �Type IIA basalts but they extend to more enriched Th and also have negative Nb-Ta anomalies when normalized to primitive mantle (IIC in Fig. 1C). The high SiO2 contents, enriched Th, negative Nb-Ta anomalies, and negative initial ɛNd values (-2 to -7) suggest significant, probably lower crustal contamination of the Type II Upper Siemens Creek basalts. The lava flows of the Siemens Creek Formation and correlative units around western Lake Superior appear to record the initial decompressive melting of an enriched mantle plume (Fig. 1D). The first picritic to basaltic flows represent low degree partial melts (~1-3%) derived from considerable depth well in the garnet stability field (>120km) (Fig. 1D). Progressive lithospheric extension, however, appears to have relatively quickly given rise to basaltic melts derived at higher degrees of melting at intermediate to relatively shallow depths (spinel stability field), and that ponded and interacted with the lower crust (Fig. 1D). Figure 1. A. Primitive mantle normalized trace element plot for the western basal Type I Siemens Creek lava flows and field for Siberian meymechites; B. Primitive mantle normalized trace element plot for basal Keweenawan lava flows in western Lake Superior; C. Primitive mantle normalized trace element plot for lower Type I and upper Type II Siemens Creek basalts; D. Ce/Yb vs Sm/Yb plot comparing Siemens Creek basalts with model melts generated by progressive lithospheric extension. Tick marks on model melt curve indicate depth of final melt segregation in km (after Ellam, 1992). Cannon, W.F., Woodruff, L.G., Nicholson, S.W., and Hedgman, C.A., 1996, Bedrock geologic map of the Ashland and the northern part of the Ironwood 30’x60’ quadrangles, Wisconsin and Michigan: U.S. Geological Survey Miscellaneous Investigation Series Map I-2566, scale 1:100,000. Ellam, R.M., 1992, Lithospheric thickness as a control on basalt geochemistry: Geology, v. 20, p. 153–156. Gibson, S.A., Thompson, R.N., and Dickin, A.P., 2000, Ferropicrites–geochemical evidence for Fe-rich streaks in upwelling mantle plumes: Earth and Planetary Letters, v. 174, p. 355–374. Hubbard, H.A., 1975, Lower Keweenawan volcanic rocks of Michigan and Wisconsin: U.S. Geological Survey Journal of Research, v. 3, no. 5, p. 529–541. Nicholson, S.W., Shirey, S.B., Schulz, K.J. and Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504-520. 134 �Potential value of pre-mining baseline oxygen, hydrogen, and sulfur isotopic data from surface waters for proposed large mining projects in northern Minnesota SEAL, Robert R. II1, JONES, Perry M.2, PIATAK, Nadine M.1, & WOODRUFF, Laurel G.2 1 U.S. Geological Survey, Reston, VA 20192, rseal@usgs.gov, npiatak@usgs.gov 2 U.S. Geological Survey, Mounds View, MN 55112, pmjones@usgs.gov, woodruff@usgs.gov Large mining projects, such as those proposed for the disseminated Cu-Ni-platinum group metal ores of the basal zone of the Duluth Complex, face a number of long-term challenges regarding environmental management due to their large size and the duration of mine lives. Leachates from mine waste may reach groundwater and eventually surface water from sources difficult to identify because they may be concealed by extensive piles of waste material. Acid drainage, even from low-sulfide mine waste, is one type of leachate that can be especially problematic to the environment by contributing acid, sulfate, and trace metals to the surrounding watershed. Large mines can have mine lives that span multiple decades. Under such long time spans, climate change, particularly drought, may have unanticipated effects on water quality and quantity in mine settings presently dominated by wetlands, such as those in northern Minnesota. Bacterial sulfate reduction is a common process acting in wetlands that consumes sulfate and organic matter to produce sulfide and alkalinity. The resulting sulfide can combine with iron or other metals to form authigenic sulfide minerals that sequester sulfur under the water-saturated anoxic conditions within wetlands. If these wetlands become unsaturated due to drought conditions, these authigenic sulfide minerals may oxidize and release sulfate to the watershed as surfacewater levels decline, which can confound attempts to identify sources of dissolved sulfate and trace elements in regional watersheds. Pre-mining baseline characterization of the oxygen and hydrogen isotopic composition of water and the sulfur and oxygen isotopic composition of dissolved sulfate provides a potential means of identifying the future influx of both acid drainage leached from mine-waste piles and sulfate derived from oxidation of naturally occurring authigenic sulfide minerals from drying wetlands. Three watersheds, Filson Creek, Keeley Creek, and the Saint Louis River, all of which transect the basal zone of the Duluth Complex, were sampled for baseline water-quality and stable isotope geochemistry over the course of a year focusing on peak-flow (June 2013) and base-flow (September 2012, 2013) conditions. The oxygen and hydrogen isotopic compositions of surface waters reflect the varying amount of evaporation that these waters have experienced. Isotopic analyses from samples from base-flow conditions reflect the influence of significant evaporation (δ18O = -9.4 – -4.1 ‰; δD = -65.5 – -38.8), whereas samples taken during peak flow show minimal deviation from the meteoric water line and imply little or no evaporation. Comparison of the oxygen isotope composition of water and dissolved sulfate reflects the relative importance two pyrite oxidation reactions that primarily differ in terms of their oxidants: FeS2 + 7/2 O2 + H2O  Fe2+ + 2 SO42- + 2 H+ (Reaction 1) FeS2 + 14 Fe3+ + 8 H2O  15 Fe2+ + 2 SO42- + 16 H+ (Reaction 2) Reaction 1 with molecular oxygen as the oxidant is most important at pH values above 4 – conditions of incipient sulfide oxidation. In contrast, Reaction 2 with ferric iron as the oxidant is most important below pH 4 – classic acid-mine drainage conditions. The oxygen isotopic composition of the water and dissolved sulfate can be used to estimate the relative importance of these two reactions in the generation of dissolved sulfate. The oxygen isotopic composition of 135 �dissolved sulfate in surface waters around the basal zone of the Duluth Complex (δ18O = -0.2 – 4.3 ‰) is consistent with greater than 70 percent of its derivation from Reaction 1 (Figure 1). Any future input of acid-mine drainage would be expected to produce sulfate with a distinctly lower oxygen isotopic composition. 6.0 4.0 2.0 δ18O SO4 0.0 ‐2.0 ‐4.0 ‐6.0 Filson Creek ‐8.0 Keeley Creek ‐10.0 ‐12.0 ‐12.0 ‐11.0 ‐10.0 St. Louis River ‐9.0 ‐8.0 ‐7.0 ‐6.0 ‐5.0 ‐4.0 δ18O H2O Figure 1. Oxygen isotope composition of dissolved sulfate and water from surface waters samples in the Filson Creek, Keeley Creek, and Saint Louis River watersheds near the basal zone of the Duluth Complex. The dashed lines correspond to the relative proportions of Reaction 1 and Reaction 2 contributing to the dissolved sulfate through sulfide oxidation. The sulfur isotopic composition of dissolved sulfate has potential for being a sensitive indicator of sulfate derived from future emergence and drying of wetlands due to drought conditions. The sulfate-sulfur isotopic composition of dissolved sulfate (δ34S = 5.6 – 8.6 ‰) overlaps known variations for ore and country rocks sulfides, suggesting that the current dissolved sulfate is derived from near surface oxidation of sulfide minerals from the basal zone of the Duluth Complex and the Virginia Formation. In contrast, the sulfur isotope fractionation caused by bacterial sulfate reduction can produce sulfide with isotopic compositions that can be between 15 and 70 ‰ lower than the corresponding sulfate. The sulfur isotope composition of sulfate derived from emergent wetland sulfide minerals in the future may be markedly different from the current sulfate in the watersheds. Baseline isotopic data for waters and dissolved sulfate can potentially be the basis for sensitive indicators of the future onset of acid-mine drainage, or the release of sulfate from wetlands that emerge due to prolonged droughts. The recognition of either of these processes near their earliest inception will facilitate appropriate responses for environmental protection during mining and after closure. 136 �The Mineralogy, Petrography and Geochemistry of the Anderson Lake Pegmatite Occurrence V. Smith, S. Zurevinski Department of Geology, Lakehead University, Thunder Bay, Ontario Anderson Lake Pegmatite is a S-type granitic pegmatite derived from the Hilma Lake Granite within the Quetico Terrane of the Southern Superior Province. The Anderson Lake pegmatite mineralogy consists of potassium feldspar, muscovite, quartz, beryl, and molybdenite. The pegmatite is crosscut by later quartz veins which occasionally host amethyst. The molybdenite within the pegmatite is syngenetic and occurs directly within quartz rich areas of the pegmatite, as well as within late stage dark gray quartz veins which crosscut the pegmatite. The molybdenite is occurring as coarse-grained euhedral florets, as well as pod-like aggregates. Of the three main trenches within the property, the molybdenite is more abundant in trench A, with minor occurrences in B and C. Within trench C, ferrimolybdenite is present within fractures alongside the molybdenite. Re-Os dating of the molybdenite within the pegmatite produced an age of 2689 +/- 12 Ma, which predates much of the plutonism, metamorphism and subsequent pegmatite injections within the Quetico Terrane (Percival and Sullivan, 1986). The pegmatite is roughly N-S trending along the contact between the host metasedimentary rocks and the Hilma Lake Granite. The associated Hilma Lake granite is classified within this study as a equigranular, coarse-grained monzogranite to syenogranite. The Hilma Lake granite contains plagioclase – potassium feldspar – quartz – biotite – muscovite – titanite – garnet – apatite and the biotite is heavily altered to chlorite with magnetite inclusions between the sheets of the biotite. When plotted on a tectonic discrimination diagram (Rb vs. Y + Nb) (Pearce et al., 1984), the Hilma Lake Granite and the Anderson Lake Pegmatite plot as volcanic arc granites. This would be expected had they formed during the subduction and transpression of the Wabigoon Terrane to the north and the Wawa Terrane to the south, when the Quetico was an accretionary complex composed of sediments derived from the volcanic arcs to the north and south (Percival and Williams, 1989). Figure 2: Tectonic discrimination diagram of the Hilma Lake Granite and the Anderson Lake Pegmatite based on (Pearce et al., 1984). 137 �Figure 3: A) Pegmatitic Beryl crystal found in trench A. B) Amethyst within a quartz vein in trench C. C) Hand sample from trench A showing a large molybdenite crystal. REFERENCES: Pearce, J. A., Harris, N. B., & Tindle, A. G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of petrology, 25(4), 956-983. Percival, J. A., & Williams, H. R., 1989. Late Archean Quetico accretionary complex, Superior province, Canada. Geology, 17(1), 23-25. Percival, J. A., & Sullivan, R. W., 1986. Age constraints on the evolution of the Quetico belt, Superior Province Ontario. Geological Survey of Canada: Radiogenic Age and Isotopic Studies: Report 2, 97-108. 138 �A Preliminary Investigation of Enigmatic Igneous Rocks on Big Powder Island, Northern Lake Superior: A Possible Mesoproterozoic Magmatic Event SMYK, Mark C.1, HOLLINGS, Peter 2 and FRALICK, Philip2 1 Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, Suite B002, 435 James Street South, Thunder Bay, ON P7E 6S7 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada A previously undescribed package of igneous rocks was investigated in 2015 on the southern shore of Big Powder Island (aka Anguros Island) in northern Lake Superior near Rossport. These relatively flat-lying rocks apparently overlie Paleoproterozoic Animikie Group sedimentary rocks of the Rove Formation (ca. 1835 Ma) and appear to be overlain by Mesoproterozoic clastic sedimentary rocks of the Pass Lake Formation, the lowermost formation of the ca. 1.4 Ga Sibley Group. All local rocks are intruded by ca. 1.1 Ga Nipigon diabase sills and dykes of the Midcontinent Rift (MCR). This package of igneous rocks, originally ascribed to the MCR by Giguere (1975) comprise a variety of fragmental units with aphanitic to phaneritic groundmasses. Locally displaying columnar joints, these rocks contain features that suggest that they may have been alternatively emplaced as shallow intrusions, flows and/or pyroclastic units. Three broad varieties of fragmental rocks have been identified:    Polymictic rocks with < 50% fragments; Monomictic rocks with < 30% black, fine-grained fragments; and Fragment-poor, aphanitic rocks with possible vesicles/amygdules. The majority of fragments are sub-angular to sub-rounded and resemble chemical and clastic Animikie Group sedimentary rocks. Baked margins are visible on some fragments, while others show evidence of partial resorption, flow alignment and/or flow-induced folding. The entire recovered zircon population appears to consist of detrital grains that exhibit metamorphic rounding but apparently lack metamorphic overgrowths. U-Pb data are concordant, yield typical Archean greenstone belt (2718-2720 Ma) and granitoid (2678 Ma) ages, and have normal igneous Th/U. These xenocrysts may have come from a xenolith that had been assimilated and heated sufficiently to cause resorption on the zircon surfaces, but not enough to produce melted rims (S. Kamo, personal communication, 2016). Or, alternatively, they represent a detrital zircon population that was abraded during transport and incorporated into a magma prior to overgrowth formation. Unlike nearby Nipigon sills, the primary mineralogy of these rocks has been extensively altered. Plagioclase and microcline have been sausseritized and sericitized, respectively. Primary ferromagnesian minerals have been altered to amphibole and chlorite. Small, irregular patches of quartz are suggestive of partial melting or resorption. Whole rock geochemistry of relatively fragment-free rocks is consistent with andesites to dacites (SiO2 = 57-70 wt%), with generally elevated K2O contents of 1.8 to 6.1 wt%, placing them in the calc-alkaline field on an AFM diagram. Very low Na2O/K2O ratios for most of the samples strongly suggests removal of Na during near-surface weathering, though the significant amount of K in the samples indicates the weathering was not extensive. The Big Powder Island (BPI) rocks are characterized by enriched LREE (La/Smn = 2.3-3.9), weakly fractionated HREE (Gd/Ybn = 1.2-1.7) and strong negative Nb anomalies. 139 �When compared to MCR-related magmatic suites, the BPI rocks have similar La/Smn and Gd/Ybn characteristics to both the Nipigon sills and the Osler volcanic rocks. Two thin dikes that cut a nearby Nipigon sill are geochemically similar to the BPI rocks rather than the sills, suggesting that they could be co-magmatic. However, the BPI fragmental rocks have lower TiO2 contents at a given Mg# than the majority of MCR rocks, suggesting a distinct source region. Interestingly, the only other MCR suite that shows similar TiO2 systematics is from the Moss Lake Intrusion. On a regional scale, the ~1590 Ma Badwater intrusive rocks show similar trends to the BPI with variable Mg#s at broadly constant TiO2 contents (Hinz 2015) whereas the ~1540 Ma English Bay granites (Hollings et al. 2007) do not. Lacking unequivocal primary zircon age data and based on apparent field relationships with other local Proterozoic rocks, it is suggested that the BPI rocks may represent magmatism that occurred between 1.4 and 1.8 Ga. Detrital zircon data from the overlying Sibley Group rocks contain populations that cluster between 1.9 to 1.8 Ga and 1.6 to 1.4 Ga (Rogala et al. 2007). Very few igneous rocks of those ages occur in this part of northwestern Ontario, suggesting that perhaps the BPI rocks could be a possible source of some of these zircons. Further work is required to establish and constrain field relationships and determine the nature and age of these enigmatic igneous rocks. REFERENCES Giguere, J.F. 1975. Geology of St. Ignace Island and adjacent islands, District of Thunder Bay; Ontario Division of Mines, Geological Report 118, 35p. Hinz, S., 2015. Geochemistry of the Badwater Gabbro south of Armstrong, Ontario. Unpublished HBSc thesis, Lakehead University, 98p. Hollings, P., Fralick, P. and Cousens, B., 2007. Geochemistry and sedimentology of the Osler Formation: Evaluating rifting in the Proterozoic. Canadian Journal of Earth Sciences, 44, 389-412. Rogala, B., Fralick, P., Heaman, L.M., and Metsaranta, R., 2007, Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada: Canadian Journal of Earth Sciences, v. 44, pp. 1131-1149. 140 �SEQUENCE STRATIGRAPHY AND BASIN EVOLUTION OF THE MESOPROTEROZOIC NONESUCH FORMATION, ASHLAND SYNCLINE, NORTHERN WISCONSIN KINGSBURY STEWART, Esther1 and MAUK, Jeffrey L.2, 1 Wisconsin Geologic and Natural History Survey, 3817 Mineral Point Road, Madison, Wisconsin 53705- 5100 2 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 The late Mesoproterozoic Oronto Group consists of the Copper Harbor Conglomerate, successively overlain by the Nonesuch Formation and the Freda Sandstone. These sediments were deposited within the Midcontinent Rift above about 20 km of mostly volcanic rocks. The Nonesuch Formation is a significant host for copper and silver, a potential unconventional hydrocarbon resource, and a potential seal for carbon sequestration (Thorleifson, 2011, Bornhorst and Williams, 2013). In addition, the formation is the focus of fundamental research into environmental controls on the expansion of terrestrial life (Cumming et al., 2013). Despite its significance to applied and fundamental geologic questions, the Nonesuch Formation’s detailed sedimentology and stratigraphy remain relatively unstudied, and the formation has never been interpreted within a genetic, sequence stratigraphic framework. We present a sequence stratigraphic framework for the Nonesuch Formation within the Ashland Syncline, northern Wisconsin, based on core description and geochemistry. Because the Midcontinent Rift developed as a series of relatively enclosed basins, we use lacustrine facies associations to interpret a sequence stratigraphic framework (Carroll and Bohacs, 1999). This does not exclude the probability of episodic marine incursions, and lacustrine environments described and discussed herein may be partly to fully marine (Hieshima and Pratt, 1991). This framework ties geologic variability to geologic processes, thus improving predictability of the distribution of physical, geochemical, and biologic characteristics of the Formation. We observe ten lithofacies which we group into fluvial-alluvial, fluctuating-profundal, and fluvial-lacustrine facies associations. Handheld XRF analyses provide geochemical data that help classify these lithofacies. We observe three regional stratigraphic surfaces: two flooding surfaces, and one progradational surface that is a potential sequence boundary. The lithofacies succession we observe records a progression from a fluvial and alluvial depositional environment recorded by the upper Copper Harbor Conglomerate, to a balancefilled and then overfilled lacustrine environment within the Nonesuch Formation, and finally returning to a fluvial and alluvial environment within the uppermost Nonesuch Formation and lower Freda Sandstone. This interpreted evolution of depositional environments suggests that the primary control on sediment deposition and preservation was tectonic rather than climatic. The fluvial-alluvial facies association of the upper Copper Harbor Conglomerate was deposited when tectonic subsidence was relatively low. The basal Nonesuch contact records renewed tectonic subsidence, and the facies associations and stacking within the overlying Nonesuch Formation and lower Freda Sandstone record an evolution of sedimentary environments that developed in response to waning tectonic subsidence. The relative thicknesses and distribution of proximal and distal depositional environments, interpreted from lithofacies, show interpreted isopachs with arcuate shapes that thicken to the east. This indicates greater subsidence to the east. From this, we interpret deposition within an asymmetric half-graben bounded by one or more westdipping normal faults to the east. Growth faults like these resulted in a series of relatively isolated sub-basins within the Midcontinent Rift basin structure. Distinct stratigraphic architecture and basin-bounding faults of these sub-basins may have acted as barriers to 141 �mineralizing fluids and likely explain the difference between the unmineralized Nonesuch Formation in the Ashland Syncline and the ore-grade Cu mineralization in the Western Syncline and White Pine areas. We demonstrate that a modern, sequence stratigraphic approach may be applied to finegrained Precambrian sediments by using traditional rock description techniques and supporting lithogeochemistry. Our identification of a characteristic succession of lithofacies in Mesoproterozoic sediments demonstrates fundamental controls that are commonly interpreted for Phanerozoic lake systems may be extended into the Precambrian; these fundamental controls result in a predictable association of lithofacies, with distinct physical, biological, and geochemical properties. Figure 1: A. Fence diagram showing cores described for this study. Grain size increases to the right in each column. B. Nonesuch Fm. isopach map showing location of inferred basinbounding fault(s). Bornhorst, T.J., Williams, W.C., 2013. The Mesoproterozoic Copperwood sedimentary rock-hosted stratiform copper deposit, Upper Peninsula, Michigan. Economic Geology 108, 1325-1346. Carroll, A.R., Bohacs, K.M., 1999. Stratigraphic classification of ancient lakes: Balancing tectonic and climatic controls. Geology 27, 99-102. Cumming, V.M., Poulton, S.W., Rooney, A.D., Selby, D., 2013. Anoxia in the terrestrial environment during the late Mesoproterozoic. Geology 41, 583-586. Hieshima, G.B., Pratt, L.M., 1991. Sulfur/carbon ratios and extractable organic matter of the Middle Proterozoic Nonesuch Formation, North American Midcontinent Rift. Precambrian Research 54, 65-79. Thorleifson, L.H., 2011. Potential for implementation of mineral carbonation as a carbon sequestration method in Minnesota. Minnesota Geological Survey Open-File Report 11-2. 142 �DISCOVERING HIDDEN FOLDS AND FAULTS IN THE PRECAMBRIAN: NEW INSIGHTS INTO BARABOO-INTERVAL STRATIGRAPHY AND DEFORMATION IN SOUTHERN WISCONSIN KINGSBURY STEWART, Esther1, STEWART, Eric D.1, LAMB, Matthew2 1 Wisconsin Geologic and Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 2 Department of Geology and Geography, University of Wisconsin - Whitewater, 800 W Main St, Whitewater, WI 53190 Laurentia records a complex history of continental collision and construction, accretion, stabilization, and reactivation of continental lithosphere. In the Southern Lake Superior Region, the Baraboo-interval sediments, including the Baraboo Quartzite, Seeley Slate, the iron-rich Freedom Formation (Sauk County) and the Waterloo Quartzite (Dodge and Jefferson Counties), were deposited after ca. 1.71 Ga and then deformed during ca. 1.65- 1.63 Ga Mazatzal accretion (Holm et al., 1998; Medaris et al., 2003). Ca. 1.4 Ga intrusion of the Wolf River Batholith and 1.0 Ga failed Midcontinent rifting subsequently affected the region. Since then, the Southern Lake Superior region has remained tectonically stable for the past billion years and thus uniquely preserves the complex history of Laurentian assembly but also results in subdued topography and a thick cover of Phanerozoic sediments that obscures direct observation of these rocks. We present preliminary results from geologic mapping of the buried Precambrian basement for a ~9000 km2 area in southern Wisconsin that is based on integration of data from new bedrock drill core, sparse outcrops, and existing geophysical and subsurface data sets. Our mapping revises the stratigraphy of the Baraboo interval quartzites and allows us to interpret the deformation history of the area. We describe a two-stage deformation history, beginning with regional folding and ending with thick-skinned thrusting. We integrate regional aeromagnetic data (Snyder and Daniels, 2002) with data from outcrops, existing bedrock geologic maps (Dalziel and Dott, 1970), drill cuttings, and new and existing bedrock drill cores (e.g. Weidman, 1904; Schmidt, 1951) and associated geophysical logs in a manner analogous to the NICE Working Group (2007). Aeromagnetic anomalies reflect the distribution of magnetic minerals in the upper crust and can be used to locate buried faults and characterize the distribution and structure of magnetic rocks. In the Lake Superior Region, Pleistocene glacial deposits and Paleozoic sedimentary rocks are magnetically transparent, so the aeromagnetic map reflects changes in the composition and geometry of the buried Precambrian crust. Importantly, a new drill core collected by the Wisconsin Geological and Natural History Survey in Dodge County recovered Precambrian iron-formation over a strata-bound aeromagnetic high present in the aeromagnetic anomaly map of Snyder and Daniels (2002). We correlate this iron formation to the Freedom Formation (Lamb and Stewart, this volume), allowing the strata-bound aeromagnetic high to be used as a marker unit in the aeromagnetic anomaly map, even in areas of southern Wisconsin with no surface exposure of Precambrian rocks. Patterns observed in the aeromagnetic data suggest the development of broad-scale doubly plunging folds within Baraboo-interval sediments and underlying post-Penokean granites and rhyolites. These folds occur across >9000 km2 of southern Wisconsin. The Baraboo syncline in Sauk County represents the western edge of regional folding. In Dodge County, the Waterloo Quartzite occurs in a broad basin stratigraphically above the aeromagnetic high marker unit interpreted as the Freedom Formation, indicating the Waterloo quartzite is a distinct quartzite unit stratigraphically above the Baraboo quartzite, the Seeley Slate and the Freedom Formation. 143 �Aeromagnetic map patterns further suggest that regional folds were truncated by at least two basement-involved thrust faults. The well-known Baraboo Syncline lies in the hanging wall of one of these thrusts, which placed folded post-Penokean granites and rhyolites, as well as the lower portion of the Baraboo-interval stratigraphy, over stratigraphically higher folded Baraboointerval sediments. Mazatzal-age accretion is interpreted to have caused this two-stage (i.e. folding followed by faulting) deformation history. Figure 1. Preliminary Precambrian basement map. Dalziel, I.W.D. and Dott, R.H., 1970. Geology of the Baraboo District, Wisconsin: A description and field guide incorporating structural analysis of the Precambrian rocks and sedimentologic studies of the Paleozoic strata. Wisconsin Geological and Natural History Survey Information Circular 14. Holm, D., Schneider, D.A., Coath, C., 1998b. Age and deformation of Early Proterozoic quartzites in the southern Lake Superior region: implications for extent of foreland deformation during final assembly of Laurentia. Geology 26, 907–910. Medaris, L.G., Singer, B.S., Dott, R.H., Naymark, A., Johnson, C.M., and Schott, R.C., 2003. Late Paleoproterozoic Climate, Tectonics, and Metamorphism in the Southern Lake Superior Region and Proto-North America: Evidence from Baraboo Interval Quartzites. The Journal of Geology. 111, 243-257. NICE working group, 2007. Reinterpretation of Paleoproterozoic accretionary boundaries of the north-central United States based on a new aeromagnetic-geologic compilation. Precambrian Res. 157, 71–79. Schmidt, 1951. The Subsurface Geology of Freedom Township ni the Baraboo Iron-Bearing District of Wisconsin. UW Msc. Thesis, 40 pp. Snyder, S.L. and Daniels, D.L., 2002. Wisconsin Aeromagnetic and Gravity Maps and Data: A web site for distribution of data. USGS Open File Report 02-493. 144 �A RE-EXAMINATION OF THE KAPUKASING STRUCTURAL ZONE STINSON, Victoria R., PAN, Yuanming, GAMELIN, Gleceria, and NADEAU, Matthew 114 Science Place, Saskatoon, Saskatchewan, S7N 5E2; vis211@mail.usask.ca The Kapukasking Structural Zone in the Wawa-Abitibi terrane crosscuts the Superior province and exposes amphibolites, granulites, and migmatites of the middle and lower crust. The Wawa Gneiss Domain west of the Kapukasing Structural Zone is composed of various types of felsic intrusive bodies that crosscut diatexite migmatites. The diatexite migmatites transition to metatexites into the Kapukasing Structural Zone and are migmatized amphibolites, granulites, and metaconglomerate. The transition from the Wawa Gneiss Domain into the western Kapukasing Structural Zone is gradational and is defined by a systematic pattern of regional to microscopic pinch and swell and boudinage structures. The boudin necks are defined by penetrative, tightly-spaced schistosity and gneissosity, dominantly east-west to northeast-southwest striking, steeply to moderately dipping south to southeast or north to northwest, steeply to moderately plunging dominantly east or west mylonites in dextral ductile to brittle-ductile shear zones. The rheologically competent boudins, acting as regional to microscopic lithons, display varying types of boudinage structures, and have wider-spaced gneissosity, strike east-west to northeastsouthwest, are horizontal to gently dipping north or south, and horizontal to gently plunging east or west. The formation of the Kapukasing Structural Zone can be explained by continental collision to dextral transpression between the Minnesota River Valley and the Wawa-Abitibi terranes in the Neoarchean to Paleoproterozoic. The oblique continental collision produced lithospheric-scale boudinage, creating and exposing the Kapukasing Structural Zone and metamorphic core complex of the Wawa Gneiss Domain, which were subsequently brittlely deformed throughout the Proterozoic. 145 �Source of Native Iron in Canadian Arctic Artifacts SVENSSON, Matthew, and KISSIN, Stephen Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada A collection of fourteen samples of metal made from iron artifacts recovered from the Canadian Arctic Archipelago was received from the Canadian Conservation Institute, Department of Canadian Heritage in order to examine the origin of the materials. In Greenland, three sources of iron in artifacts have been determined: terrestrial native iron, manmade iron of European origin and meteoritic iron. The present study was undertaken in order to determine the origin of the metal in the Canadian artifacts. Observation of polished surfaces made from the samples by reflected light microscopy revealed the presence of the Widmanstätten structure deformed in all cases by cold working, clearly indicating the meteoritic origin of all of the artifacts. The Widmanstätten structure is an oriented intergrowth of kamacite (α-Fe) lamellae in taenite (γ‐Fe).  This structure forms as the result of the exsolution of kamacite during the diffusion of the Ni into the taenite phase in octahedrites. The Ni distribution and concentration of the Ni in the kamacite phase can be used as an indicator of the equilibrium state of the specimen and in the classification of the meteorite. The results of a step scan across the entire bandwidth of the kamacite lamellae can yield an Mshaped profile indicating the diffusion of Ni. Although kamacite and taenite can be found in terrestrial iron sources, the Widmanstätten texture is unique to meteoritic iron and is commonly found in octahedrites such as the Cape York meteorite. The Cape York meteorite landed close to Thule and Dorset territory in southern Greenland. It is reasonable to hypothesize that the Cape York meteorite is the source of iron for these artifacts because of its close proximity to Thule and Dorset territory. A comparison was done of the Ni-content and the bandwidth of the kamacite lamellae in the artifacts with that of the Cape York meteorite in order to precisely test this hypothesis. Here, we report the results of the mineral chemistry and compare these results to those reported for the Cape York meteorite in order to draw conclusions. Only 0.5 to 2.0mm sized samples were taken for use in SEM-EDX analysis in order to preserve as much of the artifacts as possible. A step-scan and line scan were performed roughly perpendicular to the kamacite taenite interface to acquire a representative sample of the artifact’s whole composition. The standard used for this analysis was the Filomena meteorite. Filomena is comprised almost entirely of kamacite, making it an excellent standard for the analysis of the kamacite lamellae. Both the deformation of the kamacite lamellae caused by the cold forging of the artifacts, and the small sample size prevent direct observation of the full bandwidth of the kamacite lamellae. The bandwidth of the kamacite lamellae may be calculated through the use of Goldstein’s (1965) plot of average half-width of kamacite vs the average Ni-content in the kamacite, which also related mean kamacite bandwidth to Ni-content in the meteorite The M-profiles in this study are nearly flat, reflecting the approach to equilibrium in diffusion of Ni from kamacite lamellae. The kamacite bandwidth as measured in SEM images and by determination from Goldstein’s (1965) figure 9 yield a range of widths of 0.042 to 0.0196mm . These bandwidths are significantly smaller than Buchwald (1975) reported for Cape York at 1.20±0.2mm. However, Buchwald’s measurements were made on the Savik I mass of Cape York, which has the lowest reported Ni-content of the Cape York masses as reported by Esbensen et al. (1982). The Thule mass (table 1) has a much higher Ni-content and therefore a 146 �narrower kamacite bandwidth. However, the bandwidths for other than Savik I have not been reported, and kamacite bandwidths can vary widely in a given meteorite. We conclude that this study reveals that the Canadian Arctic metal artifacts are made from meteoritic iron and further, that the likely source is the Cape York meteorite. Table 1: Ni-content of Cape York Meteorite Masses Fragment Ni (wt %) Savik I 7.46 Savik II 7.54 Ahnighito East 7.46 Ahnighito West 7.63 Woman 7.65 Dog 7.89 Agpalilik 8.25 Thule 8.52 Data from Esbensen et al. (1982) REFERENCES Buchwald V.F., 1975. Handbook of Iron Meteorites. University of California Press Buchwald V.F., Mosdal G., 1985. Meteoritic iron, telluric iron and wrought iron in Greenland. Monographs on Greenland, Man and Society 242: 3-49 Esbensen K.H., Buchwald V.F., Malvin D.J., Wasson J.T., 1982. Systematic compositional variations in the Cape York iron meteorite. Geochimica et Cosmochimica 46: 1913-1920 Goldstein J.I., 1965. The formation of the kamacite phase in metallic meteorites. Journal of Geophysical Research 70: 6223- 6232 147 �The Badwater gabbro as an analogue for the weathering of Martian basalts SVENSSON, Matthew1, FRALICK, Philip1 1 Department of Geology, Lakehead University, 955 Oliver Rd. Thunder Bay, ON P7B 5E1 Canada The best way to study the surficial processes of other planets apart from indirect observation through rovers and orbiters is through the use of analogues. Studying the surface of Mars has become a popular topic as more sophisticated direct measurements are undertaken using instruments such as those on board the Curiosity rover. It is known from the study of Martian meteorites that the crust is entirely basaltic with no evidence of felsic igneous rocks. Through weathering, the basaltic crust was broken down and now comprises the material analyzed by Curiosity. Therefore, investigating the possible effects of weathering of basalts in a low oxygen atmosphere is important in order to better understand data retrieved from rovers such as Curiosity. Here, we analyze the petrography, mineral chemistry and whole-rock geochemistry of the Badwater gabbro with a focus on the effects of weathering. The results were interpreted in order to assess the viability of the uppermost weathered zone of the Badwater gabbro as an analogue for the planet Mars. The Badwater gabbro is a 1598 ± 1.1Ma, coarse-grained, intrusive unit that is disconformably overlain by the pillowed Pillar Lake volcanics. The gabbro has a striking paleoweathering profile developed in its upper five meters. The age of this weathering profile is poorly constrained, but it is Mesoproterozoic, constrained between the age of the Badwater gabbro and an 1100 Ma Midcontinent Rift related sill, which cuts the Pillar Lake volcanics. Sandstones present interbedded with the volcanics are similar to those of the ~1400 Ma Sibley group, which outcrops to the south. Data was collected from drill core samples obtained by East West Resources Corporation in 2004 during their search for PGE mineralization, near Armstrong, Ontario. Trends in the whole rock geochemistry are some of the most telling features of the Badwater gabbro. Potassium was found to increase from 0.482%, 1612cm below the contact with the Pillar Lake volcanics to 2.508%-3.573% towards the Pillar Lake volcanics contact (Figure 1a). Similarly, the magnesium was found to increase from 7.319% to 14.602% towards the contact (Figure 1b). These are thought to be the influence of weathering due to the presence of a saline lake such as those for which the neighboring Sibley group is known. Aluminum and sodium were found to show little variance near the contact. An average aluminum composition of 13.341±1.889% (Figure 2a), and an average sodium concentration of 4.325±1.502% (Figure 2b) were found. The aluminum data reflects the lack of overall loss or gain in the major elements. Therefore it can be concluded that the trends in magnesium and potassium do not reflect mass loss in the system. Similarly, sodium follows the same trend as Al, but is a much more mobile element and is therefore usually lost during weathering in subaerial or freshwater environments. This suggests that weathering took place in a saline lake environment thus restricting the mobility of Na. The upward increase in K and Mg is unusual as these elements, in particular K, are commonly depleted during weathering. The upward enrichment without similar enrichment in Al means they were being added from above during weathering. This implies the presence of a saline water mass or groundwater system near a saline waterbody. The overlying pillowed basalts reinforce the probability of an overlying saline lacustrine system, which had abundant K and Mg that was incorporated into the loose sediment of the weathering profile. The saline 148 �lacustrine sediments of the Sibley group attest to the development of the appropriate climatic conditions in the Mesoproterozoic for the formation of saline systems. b) MgO vs Depth 4 20 3 15 MgO K2O a) K2O vs Depth 2 10 1 5 0 0 0 500 1000 1500 2000 0 500 Depth (cm) 1000 1500 2000 Depth (cm) Figure 1: 1a) K2O vs depth. 1b) MgO vs depth. Both these oxides increase significantly in concentration through the weathered zone towards the contact. b) Na2O vs Depth 20 8 15 6 Na2O Al2O3 a) Al2O3 vs Depth 10 5 4 2 0 0 0 500 1000 1500 0 2000 Depth (cm) 500 1000 1500 2000 Depth (cm) Figure 2: 2a) Al2O3 vs depth. 2b) Na2O vs depth. Neither of these oxides show any significant increase or decrease in concentration towards the contact. Saline lakes would have low amounts of sulfur thus invoking less of a restriction on dolomite production. The magnesium derived from the dolomite is the likely source of the anomalous magnesium values. These conclusions are consistent with the widely accepted magnesium concentrations and dolomitic evaporites in the Sibley lakes. Given that weathering is dominated by the presence of saline water in a low oxygen environment it has these in common with weathering on the red planet as its water mass decreased. However, during this period on Mars the water became acidic resulting in iron being transported and iron sulfates commonly forming, whereas the non-acidic waters studied were capable of forming potassic clays and magnesium evaporites. It is reasonable to suggest that the Badwater gabbro may be as close to an analogue for Mars that can be studied on Earth, but would not make a totally appropriate analogue. REFERENCES Hinz, 2015. Geochemistry of the Badwater Gabbro south of Armstrong, Ontario. Lakehead University Mezger, K., Debaille, V., Kleine, T. 2012. Core formation and mantle differentiation on Mars. Space Science Reviews, 174: 27-48 Nesbitt, 2003. Petrogenesis of siliciclastic sediments and sedimentary rocks. Geochemistry of Sediments and Sedimentary Rocks: Evolutionary Considerations to Mineral Deposit-Forming Environments, 4: 39-51 149 �Geologic mapping of Neoarchean and Proterozoic rocks near Kekekabic Lake, northeastern Minnesota, by students of the Precambrian Research Center’s 2015 field camp Abstract refers to poster entitled: “Bedrock geologic map of the Knife Lake Group and related intrusions near Kekekabic Lake, Lake County, Minnesota.” First author: “Christenson” UPTON, Margaret1, PUZEL, Ryan1, CHRISTENSON, Jaron1, KENT, Morgan1, SPREITZER, Steven1, and JIRSA, Mark2 1 2015 Field Camp Students, Precambrian Research Center, University of Minnesota-Duluth, 5013 Miller Trunk Highway, Duluth, Minnesota 55811 2 Minnesota Geological Survey, University of Minnesota, 2609 W. Territorial Rd., St. Paul, Minnesota 55114 (jirsa001@umn.edu) The University of Minnesota-Duluth’s Precambrian Research Center conducted its ninth annual field camp in 2015, and this presentation shows results of one of several “capstone” mapping projects. The projects test student skills by creating new geologic maps in areas of poorly known geology. This benefits both students and mentor organizations, and contributes to our collective understanding of Minnesota geology. The capstone project described here involved mapping an area of ~12 mi2 in the Boundary Waters Canoe Area Wilderness (BWCAW), centered on the western part of Kekekabic Lake, but included all or parts of Spoon, Dix, Skoota, Missionary, Pickle, Kek, Strup, and Wisini Lakes (Fig. 1). The resulting map provides details about the complex depositional, magmatic, and tectonic history of a Neoarchean metavolcanic and metasedimentary terrane that is part of the Wawa subprovince of Superior Province, the basal Mesoproterozoic Duluth Complex and diabase dikes. Compared with the other 8 capstones mentored by Jirsa (2007-2014), this one presented the greatest lithologic diversity and logistical challenges. Figure 1. Generalized bedrock geologic map of NE Minnesota showing the Kekekabic Lake capstone area (solid black polygon). The Neoarchean unit labeled “Supracrustal Rocks” encloses both older volcanic sequences and younger, largely sedimentary ones. Outline of Boundary Waters Canoe Area Wilderness is dashed. The Neoarchean rocks in the central BWCAW comprise a Timiskaming-type extensional basin deposit that consists of a broad array of sedimentary (terrestrial and shallow marine), volcanic, and intrusive rocks in close proximity. The Kekekabic Lake map area (Fig. 2) provides a window into this complex terrane. Highlights of the mapping include the following:  Tightly folded graywacke and mudstone (Fig. 2, unit Aks), locally containing thin lenses and layers of iron-formation. The latter implies deposition in marine or restricted basin settings.  A unique sequence of unsorted oligomictic conglomerate composed of hornblende-, pyroxene-, and plagioclase-phyric trachyandesite clasts that grades up stratigraphic section to sandstone and gritstone containing abundant mafic minerals and chlorite presumably derived from them (unit Akg). Ripple marks and trough cross-bedding in sandy portions indicate fluvial transport (Fig.3).  Both units are cut by a large, lithologically diverse intrusion (unit Akp) composed of porphyritic rocks that vary from monzonite to hornblende-, pyroxene- and biotite-bearing lamprophyre. 150 � Heterogeneous gabbro and troctolite of the basal Duluth Complex intruded the Knife Lake rocks, creating a Hornfels contact aureole that displays metamorphic recrystallization, thermally augmented deformation features, and magmatic brecciation of the host Neoarchean rocks. Figure 2. Gray-scale version of the 14-map unit Kekekabic Lake geologic map highlighting pertinent units discussed above. NAD83, Zone 15 UTM coordinates. Figure 3. Field photographs of unit Akg. Left: unsorted volcanic conglomerate; Right: ripple-marked volcanic sandstone. See discussion above. North shore of Kekekabic Lake. The results of this and other capstone mapping projects can be viewed at www.d.umn.edu/prc. 151 �Influence of mineral liberation on metal leaching and dissolution rates in ore material and associated host rock VANDERWAAL, Gerrit and SCHARDT, Christian Department of Earth and Environmental Sciences, University of Minnesota Duluth, 1049 University Drive, Duluth, MN 55812, United States Current methods of sulfide ore comminution produce harmful waste in the form of leached metals and acidic runoff. Conventional ore processing techniques leave some ore minerals with the tailings, the efficiency of which depends on many factors. In the case of sulfide mining, these ore minerals oxidize to form sulfate and free H+ ions, lowering the pH of the system and accelerating sulfide oxidation and the release of metals, a phenomenon known as “acid rock drainage” (Nicholson et al., 1990; Lapakko et. al, 2013). A technology that has recently become commercially available, electric pulse disaggregation (EPD), breaks apart rock at mineral grain boundaries, void spaces, and other cavities within the rock. A given sample is submerged in water and variable amounts of electricity (200 kV in this experiment) are pulsed through the water, generating plasma streamers which break apart the rock by shockwaves formed by rapid heating and expansion (Cabri et al., 2008). By breaking the rock apart into individual mineral grains, the surface area of both sulfides and silicates is drastically increased, improving the recovery of ore minerals and increasing the buffering capacity of silicates, thereby limiting the potential production of acidic runoff. EPD should increase ore recovery efficiency while simultaneously decreasing the amount of waste released into the environment by separating ore minerals from the gangue at grain boundaries. An ongoing experiment utilizing Cu-Ni-PGE sulfide ore material from Minnesota’s Duluth Complex is currently testing this hypothesis. Six different experiments, three with material processed via EPD (material stored in water at room temperature, at 50°C, in a solution with a starting pH of 4) and three with conventionally crushed material (same conditions as above) were set up and run over the course of 8+ weeks. Temperature, pH, conductivity, and mass measurements are being taken on a weekly basis. Samples of the aqueous solution will be taken twice over the course of the experiment and will be subjected to geochemical analysis. The reaction product will eventually be analyzed via x-ray diffraction to determine the final composition of the rock. Preliminary data indicate a slight increase in free ions in the material processed via EPD compared to the mechanically treated material but other parameters are similar. All experiments exhibit a decrease in mass to the current date and the pH of 4 experiment quickly reached neutral conditions (within five weeks) as expected, while the other experiments showed no signs of significant change. REFERENCES Cabri, L. J., Rudashevsky, N. S., Rudashevsky, V. N. and Oberthür, T. 2008. Electric-pulse disaggregation (EPD), hydroseparation (HS) and their use in combination for mineral processing and advanced characterization of ores. In Proceedings of the 40th Annual Canadian Mineral Processors Conference. p. 213. Ontario, Canada: Canadian Mineral Processors. Lapakko, K. A., Olson, M. C., and Antonson, D. A. 2013. Duluth Complex tailings dissolution: Ten-year laboratory experiment. Minnesota Department of Natural Resources. p. v. Nicholson, R., Gillham, R., and Reardon, E. 1990. Pyrite oxidation in carbonate-buffered solution: 2. Rate control by oxide coatings. Geochimica Et Cosmochimica Acta, 54(2), p. 395-402. 152 �Small-Scale Petrographic Variations in a Nipigon Diabase Sill WALLRICH, Blake M. and ZIEG, Michael J. Department of Geography, Geology, and the Environment, Slippery Rock University, Slippery Rock, PA, 16057 In this study, we report ongoing results describing the petrographic characteristics of a Nipigon diabase sill, part of the 1.1 Ga Midcontinent rift magmatic suite of ultramafic to mafic sills located in the Lake Superior region (Hollings et al., 2007). A ~250 m continuous core (BSE-07-01) drilled by RPT Uranium Inc. (2007) in the vicinity of Black Sturgeon Lake, southwest of Lake Nipigon has been archived in the geology department at Slippery Rock University (SRU). Previous research by SRU students and faculty has identified several reinjection horizons where new magma has intruded older, partially solidified magma within a crystal mush zone (e.g., Zieg, 2014). This investigation focused on a detailed petrographic and textural analysis of a 20-meter section (65-85 m above lower contact) to provide a better understanding of petrographic signatures of the reinjection process within this interval. The primary data for this study were variations in modal mineralogy (obtained by point counting) and plagioclase mean length (determined by manually tracing crystals in photomicrographs). Using this data, we have identified a zone where plagioclase sizes deviate from a normal coarsening-inward trend. This deviation is accompanied by variations in modal olivine: anomalous decreases in plagioclase mean length are associated with sudden increases in olivine abundance. The decreases in mean plagioclase size are thought to represent soft internal chills. The olivine accumulations are interpreted as suspended crystals that settled out of the reinjected magma soon after emplacement, collecting on top of the more viscous crystal mush at the base of the freshly injected magma (Hayes et al. 2015; Zieg, 2014). In addition to grain size we also examined plagioclase orientation. In the section with anomalous textures, we also found plagioclase grains with a preferred orientation, possibly induced by shear stresses related to the inflowing magma, parallel to reinjection margins (perpendicular to the core axis). The combination of plagioclase alignment, finer-grained textures, and the sudden increase in olivine abundance supports the hypothesis that magma was emplaced into the existing mush at this location as an olivine-rich slurry (Fig. 1, Fig. 2; Zieg, 2014). The results of this project will contribute to a larger study (Zieg, this volume) that evaluates the processes of emplacement and identifies diagnostic criteria for recognizing reinjection horizons within sills. This data will then be used to refine model parameters for magma chamber evolution within the Earth’s crust. Future work will focus on further defining the petrologic characteristics of reinjection horizons using textures and modal mineralogy, and eventually geochemical analysis, to further constrain the physical and chemical consequences of reinjection in sills. References Hayes, B., Bédard, J. H., and Lissenberg, C. J., 2015, Olivine slurry replenishment and the development of igneous layering in a Franklin sill, Victoria Island, Arctic Canada: Journal of Petrology, v. 56, p. 83-112. Hollings, P., Hart, T., Richardson, A., and MacDonald, C. A., 2007, Geochemistry of the Mesoproterozoic intrusive rocks of the Nipigon Embayment, Northwestern Ontario: Evaluating the earliest phases of rift development: Canadian Journal of Earth Sciences, v. 44, p. 1087-1110. Zieg, M. J., 2014, Petrologic evolution of a Nipigon diabase sill, Ontario, Canada: Insights from compositional and textural profiles: Economic Geology, v. 109, p. 1383-1401. 153 �Figure 1(left to right): Variations in modal olivine, mean plagioclase length, and plagioclase alignment factor ( ) 85 m 75 m 72 m 71 m Figure 2: Variations in olivine abundance, mean plagioclase length, and alignment factor with height. FOV in micrographs 4.1 x 3.3 mm 154 65 m �Assessment of Undiscovered Nickel-Copper-Platinum Group Element (Ni-CuPGE) Resources Related to Conduit-Type Mineralization in the Midcontinent Rift System, Michigan, Minnesota, Ontario, and Wisconsin ZIENTEK, Michael L.1, SCHULZ, Klaus J. 2, WOODRUFF, Laurel G.3, CANNON, William, F. 2, NICHOLSON, Suzanne W. 2, ZÜRCHER, Lukas4, PARKS, Heather L. 1, and DICKEN, Connie2 1 U.S. Geological Survey, 904 West Riverside Avenue, Spokane, WA 99201 2 U.S. Geological Survey, 12201 Sunrise Valley Drive, MS 954, Reston, VA 20192 3 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 4 U.S. Geological Survey, 520 North Park Avenue, Tucson, AZ 85719 The U.S. Geological Survey (USGS) uses a geology-based three-part mineral resource assessment approach. The three parts for a specific deposit type are: 1) delineation of areas permissive for undiscovered mineral resources permitted by the geology; 2) estimation of the number of undiscovered deposits within each delineated area; and 3) estimation of the amount of resources contained in the undiscovered deposits, using appropriate ore characteristics and metal contents defined by worldwide grade and tonnage models. The USGS is completing such an assessment for conduit-type Ni-Cu- PGE sulfide deposits (defined as magmatic sulfide mineralization restricted to small- to medium-sized mafic and ultramafic dikes and sills derived from picritic and tholeiitic basaltic magma; Schulz and others, 2014) in rocks related to the Midcontinent Rift System (MRS) in Michigan, Minnesota, Ontario, and Wisconsin. The name of this deposit type, conduit-type, emphasizes the relation of these Ni-Cu-PGE sulfide-rich deposits to small- to medium-sized mafic and ultramafic dikes and sills that served as pathways for the flow-through of picritic and tholeiitic basaltic magmas. These intrusions are orders of magnitude smaller than many layered intrusions that host contact-type Ni-Cu±PGE sulfide (e.g., Duluth Complex) or reef-type PGE deposits (e.g., Stillwater Complex, Montana). Critical processes of conduit-type deposit ore formation are used to identify essential criteria appropriate for a regional-scale assessment and to define proxies for these system components. For this assessment, we emphasized a metal source (MRS-related mafic and ultramafic magmas), magma pathways (mafic and ultramafic dike swarms and small intrusions emplaced during development of the MRS), and a source of sulfur in country rocks that, if assimilated, could result in sulfur saturation of MRS magmas (in this setting, sulfide-bearing Paleoproterozoic metasedimentary and metavolcanic rocks). It is difficult to predict, however, if a particular mafic or ultramafic intrusion emplaced into favorable country rock served as a high-flux magma pathway and thus, could contain a magmatic conduit-type sulfide deposit. Also, at the regional scale of an assessment, there are no proxies for predicting if and where sulfide minerals might accumulate in sills and dikes. As a result, all areas where MRS-related mafic and ultramafic dikes or sills are inferred to be present in sulfur-rich Paleoproterozoic rocks are considered to be permissive for the occurrence of conduit-type magmatic Ni-Cu-PGE sulfide mineralization. Thus, for this assessment, five permissive tracts are identified: 155 �1) Animikie tract – strata of the Proterozoic Animikie Group in northern Minnesota (one known deposit (Tamarack) but no other known occurrences) 2) Little Falls tract – metamorphosed and deformed metasedimentary and volcanic rocks that are part of the Penokean-age fold and thrust belt in Minnesota (no known deposits or occurrences) 3) Michigamme tract – the known and inferred distribution of the Michigamme Formation, Michigan and Wisconsin (one known deposit (Eagle mine) and multiple occurrences (e.g., BIC, Roland Lake, Eagle East)) 4) Pigeon Point tract – strata of the Animikie Group in northeast Minnesota and southern Ontario (one known deposit (Great Lakes Nickel deposit) and several occurrences) 5) Pembine-Wausau tract – volcanic and sedimentary rocks of the Pembine-Wausau magmatic terrane, Wisconsin (no known deposits or occurrences) After a discussion of conduit-type deposit requirements and the favorable (or unfavorable) geology of the each tract, assessment team members made separate estimates of the numbers of undiscovered deposits. Estimators were asked for the least number of deposits of a given type that they believed could be present at three specified levels of certainty (90 percent, 50 percent, and 10 percent). Each person made initial estimates without sharing their results until everyone was finished; and then the results were compiled and discussed. Following the discussion, individual scores were adjusted and a single mean estimate of the number undiscovered deposits in each tract down to 2 km depth was determined. This final estimate reflects both uncertainties in what could exist and in the favorability of a tract. As might be expected from the number of known deposits and occurrences, the Michigamme tract was assessed to have the highest number of possible undiscovered deposits, with a mean estimate of five undiscovered deposits. The Animikie, Pigeon Point, and PembineWausau tracts each had mean estimates of two undiscovered deposits, whereas the Little Falls tract had a mean estimate of one undiscovered deposit. New grade and tonnage models for conduit-type sulfide deposits were used in Monte Carlo simulations to obtain estimated probability distributions of undiscovered metals in each tract. The current known and, here estimated, undiscovered resources represented by MRS-related conduit-type deposits are significantly less than the known Ni-Cu±PGE resources in the undeveloped contact-type deposits along the western margin of the Duluth Complex; conduittype deposits, however, are typically much higher grade, and therefore remain very attractive exploration targets. Reference Schulz, K.J., Woodruff, L.G., Nicholson, S.W., Seal, R.R., II, Piatak, N.M., Chandler, V.W., and Mars, J.L., 2014, Occurrence model for magmatic sulfide-rich nickel-copper-(platinum-group element) deposits related to mafic and ultramafic dike-sill complexes: U.S. Geological Survey Scientific Investigations Report 2010–5070–I, 80 p. http://dx.doi.org/10.3133/sir20105070I 156 �Geochemistry and Petrography of the Volcanic Strata Hosting the Flambeau Cu-Zn-Au Deposit in Rusk County, WI: A Re-examination of Wisconsin’s Only Past-Producing Volcanogenic Massive Sulfide Mine. ZENS, Zacharie A., and LODGE, Robert W.D. Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI Volcanogenic massive sulfide (VMS) deposits are polymetallic mineral deposits and are the source of many important base (e.g., Cu, Zn, Pb) and precious metals (e.g., Au, Ag, Ga) (LaBerge, 1996; DeMatties, 1994). The Flambeau Cu-Zn-Au Mine and other VMS deposits are hosted within the Paleoproterozoic juvenile and continental volcanic arc sequences of the Wisconsin Magmatic Terrane of the Penokean Orogeny (Schulz and Cannon, 2007). Despite the Flambeau Mine being the only partially extracted VMS deposit in the Penokean orogeny (only the supergene-enriched zone was mined), the volcanic and tectonic setting of the rocks hosting the deposit are poorly constrained. Research on the deposit essentially ceased after the mine closure in 1997. The mine site has since been successfully reclaimed and all bedrock exposures were covered. This study revisits the volcanic strata hosting the Flambeau VMS deposit through examination of historic drill cores to describe the geologic, geochemical and alteration characteristics of the deposit in light of almost 20 years of advances in the fields of geochemistry, economic geology, and the tectonic evolution of the region. Due to the absence of outcrop in the area surrounding the past-producing mine, Flambeau Mine drill cores were obtained and re-logged at the Wisconsin Geological and Natural History Survey core repository, in Mount Horeb, Wisconsin. Core samples were analyzed using X-ray Fluorescence and Inductively Coupled Plasma Mass Spectrometry at the Materials Science Center at UW-Eau Claire. This new geochemical data was compiled with historic mine maps and cross sections to develop a coherent scientific model describing the nature and evolution of the volcanic and hydrothermal system that hosts and formed the deposit. The metamorphosed and recrystallized altered strata hinder the interpretation of the intensity of hydrothermal alteration. Presently, the altered rocks of the Flambeau consist of biotiteandalusite±sericite schists and quartz-sericite-pyrite stringer zones. There is little visual correlation between porphyroblast abundance/size and the degree of alteration. However, major element mobility and geochemical alteration indices emphasize the variable intensities that are present throughout the stratigraphy and potentially highlight several fluid pathways and new ore horizons (Figure 1). Trace elements reveal that the protoliths of the hanging wall strata consist of primarily dacites with interlayered mafic units. The protoliths of the foot wall consist of thick dacite and rhyolite units with local thin mafic units increasing up section. Based on these preliminary geochemical characteristics of the stratigraphy, the Flambeau deposit is likely a series of stacked ore lenses in a rifting arc geodynamic setting where a submarine volcanic arc was undergoing extension and likely developing a back-arc rift (Zens et al., 2015). This is evidenced by the dacite-dominated volcanic pile with increasing abundance of mafic protoliths toward the ore horizon and stratigraphic hanging wall. This data provides invaluable constraints on the petrogenesis of the volcanic assemblages in this part of the Penokean Orogeny. 157 �Figure 1: Downhole profiles of several hydrothermal alteration indices for the volcanic strata hosting the Flambeau Mine compiled using representative holes of the hanging wall (hole 22-2) and foot wall (hole 22-60). AI – Ishikawa Alteration Index (Ishikawa et al., 1976); CCPI – Chlorite-carbonate-pyrite index (Large et al., 2001). REFERENCES DeMatties, T.A., 1994, Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Economic Geology, 89: 1122-1151. Ishikawa, Y., Sawaguchi, T., Iwaya, S. and Horiochi, M. 1976. Delineation of prospecting targets for Kuroko deposits based on models of volcanism of underlying dacite and alteration halos; Mining Geology, v. 26, p.105-117. Large, R.R., Gemmell, J.B., Paulick, H. and Huston, D.L. 2001. The alteration box plot—A simple approach to understanding the relationship between alteration mineralogy and lithogeochemistry associated with volcanic-hosted massive sulfide deposits; Economic Geology, v. 96, p.957–971. LeBerge, G.L. (ed), 1996, Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume. Institute on Lake Superior Geology, Proceedings, 42nd Annual Meeting, Cable, WI, v. 42, part 2, 179 p. Schulz, K.J. and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region. Precambrian Research, 157: 4-25. Zens, Z.A., Helmuth, S.L., and Lodge, R.W.D., 2015, Geochemistry and petrography of the strata hosting the Flambeau Cu-Zn-Au Deposit: Revisiting Wisconsin’s only past-producing volcanogenic massive sulfide mine. Geological Society of America, Abstracts with Programs, 47: 5. 158 �Evidence for Episodic Emplacement History of a Nipigon Diabase Sill ZIEG, Michael J. and WALLRICH, Blake M. Department of Geography, Geology, and the Environment, Slippery Rock University, Slippery Rock, PA 16057 The emplacement of magma into a magma chamber is one of the least-understood parts of the igneous cycle. Melting, crystallization, and differentiation can all be simulated experimentally as well as modeled theoretically. Emplacement, however, is an inherently large-scale, dynamic, and complex process that doesn't easily lend itself to experimental simulation or straightforward modeling. By the time an intrusive body is exposed at the surface, emplacement effects are commonly masked by subsequent processes including differentiation and textural reequilibration, making it more difficult yet to recognize and reconstruct emplacement conditions. Despite this difficulty, research on the timing and consequences of magma reinjections into an existing chamber is a dynamic topic of ongoing research. It is now well-accepted that magma chambers grow through incremental inflation (e.g., Menand et al., 2015, and references therein). By supplying new material, including more primitive liquids with or without suspended crystals, they can reset differentiation trends, and potentially deposit cumulate layers (Brandriss et al., 2014). Through a combination of thermal and compositional effects, reinjections into existing magma chambers have also been shown to be crucial in establishing conditions for the formation of magmatic ore bodies (e.g., Charlier et al., 2010; Maier et al., 2013). We are attempting to develop a better understanding of the nature and consequences of these reinjection events using a continuous drill core profile through a ~250 m thick diabase sill from Nipigon, Ontario. A moderately thick sill was chosen for this study because it is large enough to have experienced a complex emplacement history, while also cooling quickly enough to avoid the pervasive textural re-equilibration effects associated with larger intrusions. Using petrographic (plagioclase grain sizes, alignment factors, modal mineralogy, Fig. 1) and compositional (modal mineralogy and bulk-rock geochemistry, Fig. 2) trends in this sill, we have identified multiple horizons that can be tentatively identified as reinjection sites. The intervals interpreted as reinjection horizons are marked by an association between a sudden increase in olivine content and a decrease in plagioclase grain size. Additional characteristics include reversals in fractionation trends and changes in the composition of olivine populations. These horizons are believed to represent the influx of a new pulse of magma that arrives bearing a suspended load of olivine phenocrysts. These phenocrysts settle toward the base of the newly-injected liquid layer, producing an olivine accumulation. Meanwhile, the older, partiallycrystalline mush undergoes textural coarsening due to the heat from the fresh magma, enhancing the textural contrast between rocks crystallized from the old and new magma pulses. Future work will focus on three key questions: 1) what exactly are the diagnostic petrographic and geochemical characteristics of a reinjection horizon? 2) what processes control the development of these signatures? and 3) what can these reinjection horizons tell us about the thermal evolution of mafic sills in general, and this sill in particular? 159 �Figure 1. Petrographic data. Dashed lines represent the locations of inferred reinjection horizons. Figure 2. Compositional data. Dashed lines are the same proposed reinjection horizons, which were identified based on petrographic (primarily textural) variations. REFERENCES Brandriss, M.E., Mason, S., and Winsor, K., 2014, Rhythmic layering formed by deposition of plagioclase phenocrysts from influxes of porphyritic magma in the Cuillin Centre, Isle of Skye: Journal of Petrology, v. 55, p. 1479–1510. Charlier, B., Namur, O., Malpas, S., de Marneffe, C., Duchesne, J.C., Vander Auwera, J. and Bolle, O., 2010, Origin of the giant Allard Lake ilmenite ore deposit (Canada) by fractional crystallization, multiple magma pulses and mixing: Lithos, v. 117, p. 119–134. Maier, W.D., Barnes, S.J. and Groves, D.I., 2013, The Bushveld Complex, South Africa: Formation of plRatinum– palladium, chrome- and vanadium-rich layers via hydrodynamic sorting of a mobilized cumulate slurry in a large, relatively slowly cooling, subsiding magma chamber: Mineralium Deposita, v. 48: p. 1–56. Menand, T., Annen, C., and de-Saint Blanquat, M., 2015, Rates of magma transfer in the crust: Insights into magma reservoir recharge and pluton growth: Geology, v. 43, p. 199–202. 160 � https://digitalcollections.lakeheadu.ca/files/original/363792e794c7e2a9a4b835f47d25c32d.pdf 4e56613284f2d80676e46ab8fcfdb9aa PDF Text Text INSTITUTE ON LAKE SUPERIOR GEOLOGY DULUTH, MINNESOTA MAY 4-8, 2016 HOSTED BY: JAMES D. MILLER JR., UNIV. OF MINN. DULUTH MEETING CHAIRPERSON CHRISTIAN SCHARDT, UNIV. OF MINN. DULUTH DEAN M. PETERSON, PETERSON GEOSCIENCE LLC PROCEEDINGS VOLUME 62 Part 2 Field Trip Guidebook Compiled by Dean M. Peterson (Peterson Geoscience LLC) i �TABLE OF CONTENTS Proceedings Volume 62 Part 2 – Field Trips PRE-MEETING FULL DAY FIELD TRIPS, WEDNESDAY, MAY 4 1) Glacial Geology of the Laurentian Uplands 1 Trip Leaders: Phil Larsen (Vesterheim Geoscience PLC) and Howard Mooers (UMD) 2) Neoarchean Geology of the Western Vermilion District 14 Trip Leaders: Mark Jirsa, Terry Boerboom, and Amy Radakovich (MGS) 3) Cu-Ni-PGE Deposits of the Duluth Complex 27 Trip Leaders: Mark Severson (Teck American), Andrew Ware (PolyMet Mining), and Kevin Boerst (Twin Metals Minnesota), Steve Geerts (UMD-NRRI) PRE-MEETING AFTERNOON FIELD TRIPS, WEDNESDAY, MAY 4 4) Duluth Stream Geomorphology and the June 2012 Flood – CANCELLED 79 Trip Leader: Karen Gran (UMD) 5) Geology of the Endion Sill, Duluth 80 Trip Leader: Jim Miller (UMD) POST-MEETING EVENING FIELD TRIP, FRIDAY, MAY 6 6) Geology and Trout Fishing along Amity Creek, Duluth 102 Trip Leaders: Dean Peterson (Peterson Geoscience) and George Hudak (UMD-NRRI) POST-MEETING FULL-DAY TRIPS, SATURDAY, MAY 7 (and SUNDAY, MAY 8 for TRIP 7) 7) Archean and Proterozoic Geology of the Western Gunflint Trail (Two-Day Trip) 110 Trip Leader: Mark Jirsa (MGS) 8) Keweenawan Geology of the Hovland Area 137 Trip Leaders: Terry Boerboom (MGS) and John Green (UMD) 9) Duluth Harbor Geologic History Boat Cruise: Quaternary to Anthropocene Trip Leaders: Irv Mossberger, Mehgan Blair, Eric Dott (Barr Engineering), Andy Breckenridge (UW-Superior), Todd Kremmin (UMD) ii 160 �FIELD TRIP 1 Wednesday, May 4, 2016 GLACIAL GEOLOGY OF THE LAURENTIAN UPLANDS Phil Larson, Vesterheim Geoscience PLC Howard Mooers, University of Minnesota Duluth with contributions by Margretta Meyer, University of Minnesota Duluth INTRODUCTION The glacial geology of NE Minnesota has been the subject of study for over 100 years and our knowledge of glacial history and chronology have steadily evolved. Within the area that is the focus of this field trip there has been relatively little study. Wright (1956) was first to describe in detail the glacial sediments and landforms. The only detailed investigations since then are the MS theses of Friedman (1981), Fenalon (1986), and Meyer (2009); the map compilation Hobbs et al. (1988); a study of the Rögen moraine by Kryzer (2013). However, the recent availability of high-resolution digital elevation LiDAR has changed the way we map and interpret glacial geology. This field trip will highlight heretofore unrecognized landforms that significantly change our understanding of glacial landform genesis, the history glacial recession, and the nature of subglacial erosional processes. Widespread occurrence of ice-walled lake plains and other subtle ice-stagnation landforms reveal a complex history of ice recession punctuated by numerous ice-marginal stabilizations or minor readvances. These features suggest that the retreating ice must have been relatively debris poor and thin (2-3 meters) sheets of stagnant ice existed over large areas. Recognition of large tracts of Rögen moraine within the Toimi drumlin field the suggest an evolving subglacial erosional landscape that is interpreted as an indication that the subglacial system switched from depositional to erosional at or near the Last Glacial Maximum. The orientation of individual ridges within tracts of Rögen moraine and their association with eskers suggest that these features formed late during deglaciation in area of the glacier bed that were well drained. Lastly, we will highlight aspects of the scoured bedrock surface that allow interpretation of the nature of the depth of glacial scour of bedrock. Saprolite of varying thickness is exposed sporadically throughout the region. These occurrences, often in fractures or other protected settings, indicate that in large part scour of the Precambrian shield was limited by the depth of preglacial weathering. GLACIAL HISTORY Northeastern Minnesota was continuously covered by ice from the earliest Late Wisconsin ice advance approximately 27-29ka (Clayton and Moran, 1982; Mooers and Lehr, 1997) until about 11ka by the Rainy and Superior lobes of the LIS (Fig. 1.). The earliest formal studies of glacial deposits in northeastern Minnesota were conducted by Upham (1894) who identified a series of moraines across Minnesota. He identified the Vermillion moraine, as the 12th moraine, although he did not define its entire length. Winchell (1900), as the first Minnesota state geologist, organized systematic mapping of the glacial geology of Minnesota. Along with Upham and others, Winchell (1899) was one of the first to map large portions of northeastern Minnesota and describe the surficial deposits. Todd (1898) postulated two lobes 1 �of ice, the Lake Superior lobe, which flowed along the axis of Lake Superior, and the Red River lobe, which advanced from the west. Elftman (1898) suggested two lobes for the northeastern portion of Minnesota because of observed till differences and provenances; he named these the Superior and Rainy lobes; the Rainy lobe referring to the ice flowing from the Rainy River area. Figure 1. a-c, General sequence of glaciation modified from Mooers and Lehr (1996). d, landforms of NE Minnesota modified from Lehr and Hobbs (1992). Leverett (1932), based mostly on the work of his predecessors, proposed that northeastern Minnesota was glaciated by three separate lobes of ice. He recognized that the earliest drift in the area was the result of ice flowing from the Patrician [Labradoran] ice center located in the Hudson Bay Lowlands between the Keewatin and Labradorean ice accumulation centers. In terms of the overall glacial history of northeastern Minnesota, the modern understanding began with Wright (1956) who was the first to conduct systematic fieldwork in the area between the border lakes and Lake Superior. Wright (1964) recounted the glacial history of Minnesota as phases of the different ice lobes and was the first to identify and interpret tunnel valleys, drumlins, and eskers (Wright, 1972). In addition to laying out the general glacial geologic framework, Wright (1964, 1972) established the first regional chronology. The late Wisconsin maximum limit of the Rainy lobes was placed at the St. Croix moraine in central Minnesota (Wright 1964) ca. 20,500 BP. Wright (1972) then suggested that the Rainy lobe retreated and readvanced to the Vermilion moraine by about 18,000 BP, however, no evidence of ice recession between these two phases was presented. Little additional work was done in this area until the MS theses of Friedman (1981) and Fenelon (1986) followed by the compilation of the surficial geologic map of the Isabella area by Hobbs et al. (1988). Lehr and Hobbs (1992) outline the glacial history and landforms of the area and 2 �describe the stratigraphy of the Independence Till from three rotasonic cores in the Toimi drumlin field. Since the work of Lehr and Hobbs (1992) the only significant investigations of the glacial geology are those of Meyer (2009) and Kryzer et al. (2013), which focused on the distribution and genesis of Rögen moraine. CHRONOLOGY The chronology of Wright (1972) mentioned above has long been in question. The date of 20,500 for the St. Croix phase is based on the basal radiocarbon date at Wolf Creek, an interdrumlin swale behind the St. Croix moraine (Wright, 1972; Birks, 1976). The Vermilion moraine was correlated with the Mille Lacs/Highland moraine system which was dated ca. 18,000 BP; a date inferred as intermediate between the 20,500 date at Wolf Creek (Wright, 1972; Birks, 1976) and a date of 16,150 BP at Kotiranta Lake associated with the Split Rock phase of the Superior lobe (Wright and Watts, 1969). There are, however, basal radiocarbon dates from lakes in the Toimi drumlin field. Florin and Wright (1969) and Banerjee et al. (1979) report basal dates on aquatic mosses of 14,690 BP at Weber Lake and 16,500 Kylen Lake, respectively. Lowell et al. (2009) got a similar date of 14,050 BP on aquatic moss from the base of a core of nearby Salo Lake. Despite the general agreement of radiocarbon ages from the Toimi drumlin field, the possibility of significant carbonate error exists as the Independence Till is calcareous at depth. DESCRIPTION OF FIELD TRIP STOPS Figure 2. Location of field trip stops. 3 �STOP 1 – Ice-walled Lake Plain in Highland Moraine 564355E/5201830N (UTM Zone 15, NAD83 datum) Fredenberg 7.5’ USGS Quadrangle This site is located at a prominent ice-walled lake plain situated on the crest of the Highland Moraine (Fig. 3). Although composed predominantly of sand and gravel, this sediment is technically glaciolacustrine, deposited in an ice-dammed basin located in an icecored end moraine. The Highland Moraine is composed of hundreds of similar ice-walled lake plains, coalesced to form a belt over 100 km long and as much as 109 km wide; the massive volume of the moraine is a factor of the considerable length of time the Superior Lobe margin stood at this margin, the high sediment flux of the Superior Lobe, or both. Figure 3. Ice-walled lake plain in Highland moraine north of Duluth. 4 �STOP 2 - Superior Lobe Outwash Mantled Over Fluted Rainy Lobe Drift 567180E/5210690N (UTM Zone 15, NAD83 datum) Thompson Lake 7.5’ USGS Quadrangle This site is located on an outwash plain extending westward from the junction of the Rainy and Superior Lobes (Fig. 4). The intersection of two ice lobes forms a trough on the ice surface that focuses both surface and subglacial meltwater and sediment discharge. Sedimentladen meltwater discharge deposited a relatively flat, westward sloping outwash surface. In the vicinity of Stop 2, the outwash plain is partially collapsed, revealing fluted subglacial topography associated with northeast- to southwest-flowing ice of an older Rainy Lobe phase. These relationships indicate that retreat of the active Rainy Lobe ice margin in this area was accompanied by stagnation of a large area of the marginal zone, rather than gradual retreat of the ice margin. Sediment-poor clear stagnant ice rapidly melted until incipient melting of the sediment-rich basal debris layer formed an insulating blanket of supraglacial sediment. The outwash plain was deposited over this relatively thin layer of stagnant ice, later collapsing as the last of the buried ice melted. Figure 4. Collapsed outwash overlying Rainy Lobe stagnant ice topography. Ice margins indicated by dashed line, meltwater flow direction by blue arrows, and Rainy Lobe ice flow direction by black arrow. 5 �The relative elevation difference between the intact surface of the outwash plain and the lowest parts of the collapsed area places an important constraint on the relative thickness of the basal debris layer of the Rainy Lobe ice. The apparent thickness – about 2 m – is consistent with basal debris layer thickness observed in the modern Greenland and Antarctic ice sheets. STOP 3 - Rögen Moraine Superimposed on Drumlins 568050E/522170N (UTM Zone 15, NAD83 datum) Boulder Lake Reservoir NE 7.5’ USGS Quadrangle Stop 3 is at a road cut through one of a series of ice flow-perpendicular sediment ridges known as Rögen moraine (Fig. 5). Rögen are a common occurrence in the Toimi Drumlin field and adjacent up-ice Rainy Lobe terrane. They appear to be the product of remobilization of older subglacial sediment by sliding glacial ice, in this case the underlying drumlins. The ridges themselves are characterized by about 5 m of relief, and are spaced at a characteristic 100 m along flow lines. Here and elsewhere in the Toimi Drumlin field, Rögen moraine shows a close spatial association with subglacial meltwater discharge (the esker). Ice flow-parallel elongate corridors of Rögen moraine commonly flank tunnel valleys and eskers. This suggests that Rögen formed under warmed-bedded conditions near the ice margin, a conclusion distinctly at odds with other models for their formation. It further suggests that periodic fluctuations in basal shear stress associated with annual variations in subglacial meltwater discharge may play a role in Rögen formation, in particular cyclic stick-slip coupling of basal ice to subglacial sediment. Figure 5. Ice-marginal Rögen moraine superimposed on drumlins. Esker is sinuous feature in southeast quadrant; ice flow direction indicated by arrow. 6 �STOP 4 - Toimi Drumlin Field 595550E/5250390N (UTM Zone 15, NAD83 datum) Mount Weber 7.5’ USGS Quadrangle Stop 4 is located at a road cut through one of the hundreds of drumlins that collectively form the Toimi Drumlin field (Fig. 6). This particular drumlin is located near the center of the field. Toward the east and northeast, in the up-ice flow direction, the drumlins show increasing frequency of bedrock cores and are perhaps better described as cragand-tail features. Toward the southwest, in the down-ice flow direction, the drumlins display the characteristic streamlined form. In the near vicinity, borehole records indicate drift thicknesses of 30 to 65 m, predominantly composed of till (Lehr and Hobbs, 1992). Relief on the drumlins is about 30 m, suggesting the drumlin field is composed of a somewhat discontinuous layer of drift. Significantly, deeper tills encounrtered in borehole are weakly to moderately calcareous. Figure 6. Typical Toimi drumlins, crag and tail features in the southeast quadrant. 7 �STOP 5 - Thermokarst in Sediment-poor Stagnation Moraine 602510E/5250800N (UTM Zone 15, NAD83 datum) Mount Weber 7.5’ USGS Quadrangle This stop is located in an ice flow-perpendicular belt characterized by thin ice-walled lake plains (pannukaku) (Fig. 7). Pannukaku typically show as little as 1 m of relief with their surroundings, and are commonly draped over subglacial topography. Belts of pannukaku in the Toimi Drumlin field define minor surgestagnation moraines deposited by the retreating Rainy Lobe. The relatively small volume of sediment contained in these moraines is a reflection of relatively low sediment flux on the part of the warmedbedded Rainy Lobe. Figure 7. Rainy Lobe surge-stagnation moraine, defined by belt of pannukaku. Ice margin defined by dashed line, pannukaku by thin dashed outlines. 8 �STOP 6 - Superior Lobe Outwash 603980E/5280650N (UTM Zone 15, NAD83 datum) Slate Lake East 7.5’ USGS Quadrangle Stop 6 is located at a road cut through Superior Lobe sand and gravel, deposited by meltwater flowing from southeast to northwest (Fig. 8). Following retreat of the Rainy Lobe ice margin north of the Laurentian Upland, meltwater from the Superior Lobe was able to flow northwest into Glacial Lake Dunka and ultimately Glacial Lake Norwood. This meltwater deposited a distinct tongue of Superior Lobe-provenance sand and gravel cross-cutting continuous Rainy Lobe drift. Figure 8. Superior Lobe meltwater channel cross-cutting older Rainy Lobe drift. 9 �STOP 7 - Relict Pre-glacial Saprolite 601510E/5289820N (UTM Zone 15, NAD83 datum) Bogberry Lake 7.5’ USGS Quadrangle Much of what is commonly thought of as typical ‘glacially sculpted’ rugged shield topography is better explained as the morphology of the base of a pre-glacial saprolite (Feininger, 1971). This stop highlights relict pre-glacial saprolite exposed in a road cut during recent (2013) highway reconstruction. STOP 8 - Ice-contact Outwash Fan 615910E/5284460N (UTM Zone 15, NAD83 datum) Mitawan Lake 7.5’ USGS Quadrangle Stop 8 is located on an ice-contact outwash fan formed on a stagnant Rainy Lobe ice margin (Fig. 9). The fan itself is composed largely of sand and gravel, with some gravels at the proximal head of the fan approaching >1 m in mean grain size. Esker-like segments in the fan suggest deposition on stagnant ice. The volume and coarse-grained nature of the sediment suggest a highly energetic, high discharge meltwater system. A number of such fans are evident in the area, suggesting frequent reorganization of a broad zone of stagnant ice at the Rainy Lobe margin. Figure 9. Ice-contact Rainy Lobe outwash fan. 10 �STOP 9 - Rögen Moraine 627670E/5284350N (UTM Zone 15, NAD83 datum) Sawbill Landing 7.5’ USGS Quadrangle Stop 9 is located in a road cut through a single Rögen moraine, described by Meyer (2009) (Fig. 10). The Rögen is composed of exceedingly dense till, and is part of a system of Rögen moraine characterized by 5-10 m of relief spaced 300-400 m apart along flowlines. In this respect, these Rögen, situated at the very up-ice limit of the Toimi Drumlin field, are significantly different from the lower amplitude, shorter wavelength Rögen characteristic of the southwestern portion. Ground-probing radar profiles through this and other Rögen ridges in the vicinity display structures evocative of northeast-dipping foresets, onlapping to the south (Fig. 11). These features suggest that Rögen may form by erosion of lee-side sediment and re-deposition on the stoss side of down-ice ridges, akin to migrating dune forms in fluvial systems. In this respect, Rögen may ‘migrate’ up-ice under flowing ice as the overall subglacial surface is lowered by net erosion. Figure 10. Rögen moraine near Sawbill Landing. 11 �Figure 11. Ground penetrating radar (GPR) profile of a Rögen ridge near Sawbill Landing. Vertical scale is approximate based on radar two-way travel time. Ice flow direction is from left to right. REFERENCES Banerjee, S.K., and Lund, S.P., 1979, Geomagnetic record in Minnesota lake sediments - Absence of the Gothenburg and Erieau excursions: Geology, v. 7, p. 588–591. Birks, H.J.B., 1976, Late Wisconsinan vegetational history at Wolf Creek, central Minnesota: Ecological Monographs v. 46, p. 395-429. Birks, H.J.B., 1981, Late Wisconsin vegetational and climatic history at Kylen Lake, northeastern Minnesota: Quaternary Research, v. 16, p. 222–355. Björck, S., 1990, Late Wisconsin History North of the Giants Range, Northern Minnesota, Inferred from Complex Stratigraphy: Quaternary Research, v. 33, p. 18–36. Brown, T.R., 1988, Eskers and heavy mineral prospecting, northeastern Minnesota: M.S. Thesis, p. 103 p. Buchheit, R.L., Malmquist, K.L., and Niebuhr, J.R., 1989, Glacial Drift Geochemistry for Strategic Minerals; Duluth Complex, Lake County, Minnesota: Minnesota Department of Natural Resource Division of Lands and Minerals Project, v. 262, no. part I, p. 95 p. Clayton, L, and Moran, S.R., 1982, Chronology of late Wisconsin glaciation in middle North America. Quaternary Science Reviews, v.1, pp55-82. Elftman, A.H., 1898, The geology of the Keweenawan Area in northeastern Minnesota, Part I.: The American Geologist, v. 21, p. 90–109. Eyles, N., Putkinen, N., Sookhan, S., and Arbelaez-Moreno, L., 2016, Erosional origin of drumlins and megaridges: Sedimentary Geology, doi: 10.1016/j.sedgeo.2016.01.006. Fenelon, J.M., 1986, Glacial geology of the Cramer quadrangle, northeastern Minnesota. [M.S. Thesis]: Milwaukee, University of Wisconsin, 76pp. Florin, M.B. and Wright, H.E., 1969. Diatom evidence for the persistence of stagnant glacial ice in Minnesota, Geological Society of America Bulletin, 80(4), 695-704. Friedman, A.L, 1981, Surficial geology of the Isabella quadrangle, northeastern Minnesota [M.S. thesis]: Minneapolis, University of Minnesota, 66pp. 12 �Friedrich, H.G., 2011, Assessment of Sand and Gravel and Clay Deposits in Parts of Northern St. Louis and Lake Counties: Minnesota Department of Natural Resource Division of Lands and Minerals Project, v. 380, p. 1– 47.Fries, M., 1962, Pollen Profiles of Late Pleistocene and Recent Sediments from Weber Lake, Northeastern Minnesota: Ecology, v. 43, no. 2, p. 295–308. Hobbs, H.C., Friedman, A.L., Fenelon, J.M., and Stark, J.R., 1988, Surficial Geologic Map of the Greenwood Lake, Isabella, and Cramer Quadrangles, Minnesota: Minnesota Geological Survey Open-File Report, v. 88-02. Johnson, M.D., Adams, R.S., Gowan, A.S., Harris, K.L., Hobbs, H.C., Jennings, C.E., Knaeble, A.R., Lusardi, B.A., and Meyer, G.N., 2016, Quaternary lithostratigraphic units of Minnesota: Minnesota Geological Survey Report of Investigations RI-68. Kryzer, R., Mooers, H.D., and Larson, 2013, Rögen moraine as a transitional bedform in an erosional subglacial system: Geological Society of American Abstracts with Programs, v. 45, no. 7, p. 119. Lehr, J.D., and Hobbs, H., 1992, Field Trip Guidebook for the Glacial Geology of the Laurentian Divide Area, St. Louis and Lake Counties, Minnesota: Minnesota Geological Survey Guidebook Series 18. Leverett, F., 1932, Quaternary Geology of Minnesota and Parts of Adjacent States: USGS Professional Paper, v. 161, 149 pp. Lowell, T. V., Fisher, T.G., Hajdas, I., Glover, K., Loope, H.M., and Henry, T., 2009, Radiocarbon deglaciation chronology of the Thunder Bay, Ontario area and implications for ice sheet retreat patterns: Quaternary Science Reviews, v. 28, no. 17-18, p. 1597–1607, doi: 10.1016/j.quascirev.2009.02.025. Lund, S.P., and Banerjee, S.K., 1985, Late Quaternary paleomagnetic field secular variation from two Minnesota Lakes: Journal of Geophysical Research: Solid Earth, v. 90, no. B1, p. 803–825, doi: 10.1029/JB090iB01p00803. Martin, D., and Eng, M., 1985, Esker Prospecting Over the Duluth Complex in Northeastern Minnesota: Minnesota Department of Natural Resource Division of Lands and Minerals Project, v. 246, p. 27. Meyer, M.S., 2009, Paleoglaciological context of Rögen moraine, northeastern Minnesota: University of Minnesota Duluth, 71 p. Mooers, H.D., and Lehr, J.D., 1997, Terrestrial record of Laurentide ice sheet reorganization during Heinrich events: Geology v. 25, p. 987-990. Todd, J.E., 1898, A revision of the moraines of Minnesota: American Journal of Science (ser. 4), v. 6, p. 469-477. Upham, W., 1894, Preliminary report of the field work during 1893 in northeastern Minnesota, chiefly relating to the glacial drift, in Winchell, N. H., Geological and Natural History Survey of Minnesota, 22nd Annual Report, for the year 1893, pp.18-86. Winchell, N.H., 1899, The geology of the north part of St. Louis County, in Winchell, N.H., The geology of Minnesota, Volume IV of the final report: Geological and Natural History Survey of Minnesota, p. 222265. Winchell, N.H., 1900, Glacial Lakes of Minnesota. Geological Society of America Bulletin v. 12, p. 109-128. Winter, T.C., 1971, Sequence of Glaciation in the Mesabi-Vermilion Iron Range Area, Northeastern Minnesota: USGS Professional Paper, v. 750-C, p. C82–C88. Wright, H.E., and Watts, W.A., 1969, Glacial and Vegetational History of Northeastern Minnesota: Minnesota Geological Survey Special Publication, v. 11, 59 pp. Wright, H. E. Jr., 1964, The classification of the Wisconsin Glacial Stage. Journal of Geology 72, 628-637. Wright, H.E., 1972, Quaternary history of Minnesota, in Sims, P.K. and Morey, G.B. eds., Geology of Minnesota: A Centennial Volume, St. Paul, Minnesota, p. 515–547. Wright, H.E., 1956, Sequence of glaciation in eastern Minnesota: Geological Society of America Guidebook for Field Trips, v. 3, p. 1–24. 13 �FIELD TRIP 2 Wednesday, May 4, 2016 NEOARCHEAN GEOLOGY OF THE WESTERN VERMILION DISTRICT Mark Jirsa, Terry Boerboom, and Amy Radakovich Minnesota Geological Survey Figure 1. Regional bedrock geologic map showing locations of field trip stops (squares numbered 1-10), published geochronologic sample sites (stars), and pertinent geologic features. Modified from Jirsa and others, 2012. INTRODUCTION This field trip takes a preliminary look at stratigraphic relationships between older largely volcanic rocks inferred to be equivalent to the Ely Greenstone (~2720 Ma) and the apparently 30 Ma younger, volcanic, volcaniclastic, and epiclastic sedimentary strata of the Lake Vermilion Formation. New mapping supports recently acquired geochronologic data that indicates sediments of the Lake Vermilion Formation were deposited unconformably on the variably weathered, “deep-water” volcanic rocks. The significant hiatus, contrasting depositional environments, evidence for magmatism synchronous with sedimentation, and local unconformable contacts between the two units implies that the Lake Vermilion Formation formed in a late-tectonic extensional basin. 14 �The field guide is brief, primarily for expediency, but also reflecting the tentative nature of newly acquired outcrop information on which it is based. The field work was the first phase of a multi-year effort by the Minnesota Geological Survey to create geologic atlases of St. Louis and Lake Counties— two of the largest counties in Minnesota. It was supported by a grant from the U.S. Geological Survey STATEMAP element of the National Geologic Mapping program, and by the Minnesota Environmental and Natural Resources Trust Fund. Although geochronologic analyses were conducted as part of this mapping, the new data are not yet ready for publication. GEOLOGIC SETTING The traditional definition of the Neoarchean Lake Vermilion Formation describes the unit as complexly interbedded strata that vary from felsic volcaniclastic rocks, to rocks having evidence of reworking, to mixed-source graywacke-siltstone. The informally named Gafvert Lake sequence (Fig. 1) consists of quartz- and plagioclase-phyric, dacitic to rhyodacitic breccia and tuff that yielded a 207Pb/206Pb age of 2689.7±0.8 Ma from magmatic zircons (Lodge and others, 2013). The sequence has been inferred to lie disconformably atop the Soudan Iron Formation member of the Ely Greenstone. Previous mapping (Jirsa and others, 2001) demonstrated that the iron-formation is transitional with metabasaltic rocks of the Lower Ely Greenstone, and a felsic unit within the greenstone yielded an age of 2722±0.9 Ma (Peterson and others, 2001). Thus, the Gafvert Lake sequence is approximately 30 Ma younger than the subjacent metabasalt- and iron-formation-bearing rocks. Quartzofeldspathic sediments apparently derived from the Gafvert Lake sequence make up a variable, but locally large proportion of the detritus in the Lake Vermilion Formation. Regionally, a series of outcrops from Gafvert Lake westward shows an irregular transition from proximal, perhaps subaerial deposition on the east, to distal submarine turbiditic fan deposition to the west. On this basis, the Gafvert is now considered part of the Lake Vermilion Formation, which by extension is also 30 Ma younger than greenstone. Recent field work was conducted in part to explore lithologic attributes of the Lake Vermilion Formation and ascertain the nature of contacts between these strata and the older greenstone. This field trip examines outcrop evidence that we believe documents the unconformable nature of these strata outboard of the Ely Greenstone. The evidence acquired to date is consistent with the inference that the Lake Vermilion Formation represents deposition in a Timiskaming-type successor basin, much like the equivalent Knife Lake Group to the northeast (Driese and others, 2011; Jirsa and others, in prep.) and the Midway sequence to the southwest (Jirsa, 2000). The Ely Greenstone and Lake Vermilion Formation—as defined here—are part of the Wawa subprovince of the Superior Province. Temporal distinctions between various geologic components of this terrane are evolving with new geochronologic analyses. Nevertheless, they remain largely based on fabrics and structures that resulted from three major phases of deformation, denoted D1, D2, and D3. The D1 event involved generally pre-lithification deformation of graywacke sequences (in some localities forming large nappe structures), and tilting, broad folding, and thrust imbrication of the thick, more rigid volcanic strata. D2 deformation accompanied regional metamorphism to greenschist to amphibolite facies, and produced pervasive metamorphic foliation and lineation, folding, and strike-slip faulting. U-Pb dates of intrusions bracket the D2 event between about 2,674 and 2,685 Ma (Boerboom and Zartman, 1993). The TowerSoudan anticline (Fig. 1) is considered a D1 structure because both limbs of the complex fold are transected by D2 cleavage that trends more northeasterly than bedding. Deposition of the Lake Vermilion Formation is bracketed to an approximately 10 million year period between volcanism of the Gafvert Lake sequence at ca. 2690 Ma, and its deformation during D2 at ca. 2680 Ma. D3 is assigned to partitioned deformation that produced crenulation and faults within rocks affected by D2. All three deformation events can be attributed to variably north-northwest—south-southeast-directed transpression. 15 �FIELD TRIP STOP DESCRIPTIONS NOTE: all UTM coordinates are given in NAD 83, Zone 15 STOP 1 – Felsic pyroclastic breccia and tuff—Gafvert Lake sequence; Lake Vermilion Formation Location: UTM: 0553467E/5294482N, Highways 1 and 169, west edge of village of Tower. Description: The Gafvert Lake sequence consists of dacitic to rhyodacitic lava flows and pyroclastic rocks to the northeast that are more or less transitional with volcaniclastic strata to the west. This outcrop lies in the transition zone. It is poorly sorted and contains angular to subrounded clasts of dacitic composition (plagioclase and quartz-phyric) that range in size from several millimeters to 20 cm. Rare angular to ornate clasts of pyrrhotite and pyrite imply a pyroclastic origin. However, the presence of siltstone clasts locally, the subrounded nature of some dacitic fragments, and the rare appearance of bedding imply reworking has occurred. Figure 2. Poorly sorted fragmental rock containing clasts of quartz- and plagioclase-phyric dacite, rusty sulfides, and rare black mudstone. [Field station LS046] DIRECTIONS: Continue west ~4 mi. along highways 1 and 169 to highway 77, turn north and proceed ~ 0.5 miles and cross Pike River to pull-off on west (left) side of highway. STOP 2 – Feldspathic greywacke-mudstone of Lake Vermilion Formation—the “classic” outcrop Location: UTM: 0547272E/5293368N; west side of Hwy. 77, north side of Pike River near dam. Description: This glacially scoured outcrop exposes a nearly perfect cross-section of tabular-bedded, variably graded, feldspathic graywacke and dark gray slate. The feldspar-rich, dacitic composition of the sandy textured beds is presumed to represent derivation from the Gafvert Lake sequence exposed to the east (Stop 1). The beds contain numerous “soft-sediment” deformation features including load structures, flames, intrafolial slump folds, and growth-faults. Bedding is nearly vertical, and graded beds indicate younging to the south. This topping direction is consistent with a position on the south limb of a large, south-overturned regional D1 fold structure—inferred to be the western extension of the Tower-Soudan Anticline (Jirsa and Boerboom, 2003). Northeast-trending kink bands, fault zones, and quartz veins traversing the outcrop are Figure 3. Road cut exposing white tonalite assigned to the latest, D3 deformation event. dike with large inclusions of adjacent black siltstone. Note 40 cm-long hammer in leftcenter of photo for scale. [Field station LS051] DIRECTIONS: Walk/drive north ~700 feet to road cut on east (right) side of Highway 77 16 �STOP 3 – Tonalitic dikes cutting sedimentary strata of the Lake Vermilion Formation Location: UTM: 0547193E/5293565N; road cut on east side of County Rd 77 Description: On first glance, this outcrop appears to represent chaotic dacitic (arkosic) sedimentary strata interbedded with gray sandstone and black siltstone, compositionally similar to that at Stops1 and 2. Graded sandstone-siltstone couplets indicate stratigraphic facing to the south, as at stop 2. Closer inspection reveals the “dacitic” rocks are tonalitic dikes emplaced more or less along bedding planes of the enclosing sandstone-siltstone. Discordance is only evident locally. Tonalitic dikes are fine- to mediumgrained, equigranular, and contain abundant quartz. The intrusion suspended numerous angular to subangular xenolithic inclusions of mudstone and siltstone as large as 1m in diameter. Many of the inclusions have dark, presumably contact metamorphic rinds. A smaller tonalitic intrusion exposed in the northern part of the outcrop is cut by a 2-meter wide lamprophyric dike. Figure 4. Graded beds of white arkosic sandstone and feldspathic graywacke, black siltstone and mudstone. Stratigraphic facing is up in the photo (southward). DIRECTIONS: Return to vehicle near Pike River; drive north on Hwy 77 approximately 1.3 mi. to County Road 104 (Bois Forte or Vermilion Reservation Rd.); drive east ~0.7 mi. to Waters if Vermilion Rd. on the south; drive south ~0.2 mi. to STOP 4. STOP 4 – Arkosic sandstone and siltstone of Lake Vermilion Formation Location: UTM: 0548130E/5294880N; PRIVATE PROPERTY! Description: This apparently blasted and cleaned outcrop consists of white, coarse to fine grained arkose, with rare gray-black siltstone layers. The arkose is inferred to have been sourced from the Gafvert Lake sequence. The exposure provides a 3-dimensional view of some inferred depositional, dewatering, and compaction features. Where it can be determined, stratigraphic facing is southward as at previous stops 2 and 3. The remarkable similarity between the arkose here and the tonalitic dikes at stop 3 invite correlation and a preliminary interpretation that magmatism was synchronous with sedimentation. A. B. Figure 5. A. White arkose with rare lenses, layers, and fragments of black siltstone (Amy Radakovich for scale). B. Ball and pillow structures inferred to have formed during settling and dewatering of inferred mass flow deposit. [Field station LS056, “Majestic Rocks” development] 17 �DIRECTIONS: Return to Highway 77, turn north (right) and proceed 0.6 miles to road cut on east side of highway. STOP 5 – Mixed-source graywacke of Lake Vermilion Formation cut by quartz-plagioclase (tonalitic) dikes Location: UTM: 0547145E/5296230N; Long road cut on east side of Highway 77. Description: Complex exposure of mixed-source graywacke cut by several fine- to mediumgrained quartzofeldspathic (tonalitic) dikes. The southern third of the roadcut consists entirely of tonalite. Normal faulting is apparent locally. Stratigraphic facing is northward (Fig. 6A)—the inverse of that at prior stops 2-4—which reflects a geographic position north of an inferred D1 antiformal nappe structure that may be the western extension of the Tower-Soudan anticline (see Jirsa and Boerboom, 2003). A B. Figure 6. A. Graded bedding in mixed-source feldspathic graywacke-siltstone; stratigraphic facing is to the left in photo (north in outcrop). B. Mixed source graywacke-siltstone cut by one of several tonalitic dike dipping to right (south in outcrop). [Field station LS058] DIRECTIONS: Continue north on Highway 77 for ~0.5 miles to road cut on east side of highway. STOP 6 – Large D2 folds and shear zones in mixed-source graywacke of Lake Vermilion Formation Location: UTM: 0546970E/5297030N; Road cut on east side of Highway 77. Description: Long road cut that exposes complex shearing and folding in mixed-source graywackesiltstone beds. Shearing is manifest as ankeritesericite-chlorite phyllite near the south end of the cut. Folds exposed farther north along the road cut are tight to isoclinal synforms and antiforms having D2 fold axes that are steeply dipping. Considerable rodding lineation of more competent sandy beds indicates shallow plunge (~47º) to the east (away from the viewer). Figure 7. D2 folds (dashed white lines) of thinly bedded graywacke-siltstone [Field station LS059] 18 �DIRECTIONS: To STOP 7A—U-turn to head south on Highway 77 for ~1.4 mi. to Lost Lake Road; turn west (right) and proceed 0.8 miles to gated trail entry. To STOP 7B—drive west on Lost Lake Road ~0.2 mi. and turn south (left) on Holter Road and proceed 0.2 mi. to shallow gravel pit on east side of road. NOTE: BOTH STOPS ON PRIVATE PROPERTY. STOPS 7A, 7B –Conglomerate with abundant volcanic and sulfidic clasts similar to Stop 1 Locations: A. UTM: 0545650E/5294580N; B. 0545365E/5294515N; both exposures on floors of shallow gravel pits—PRIVATE PROPERTY! Description: These two stops (7A and B) expose slightly different versions of conglomerate in an area essentially surrounded by arkosic and mixed-source sedimentary strata similar to stops 2-6. Both stops are glacially polished outcrops of clast-supported conglomerate. The conglomerate contains abundant felsic to intermediate volcanic fragments, together with clasts of layered siliceous rock, dacitic porphyry, and sulfide-rich rock. Clasts vary from rounded to subangular. The overall composition of fragments and the presence of sulfide clasts indicates a potential correlation with volcanic strata of the Gafvert Lake sequence as seen at stop 1, which we interpret to lie stratigraphically beneath the exposures at stops 2-6. On this basis, we infer that the conglomerate represents a localized uplift of the basin floor on which other sediments of the Lake Vermilion Formation were deposited. This is consistent with a structural position near the axis of an antiformal D1 nappe structure inferred to be the western extension of the Tower-Soudan anticline (as shown on Jirsa and Boerboom, 2003). Figure 8. Lithologically diverse conglomerate (including abundant sulfide clasts on right side) at STOP 7A. [Field station LS141] DIRECTIONS: Continue south on Holter road ~1 mi. to Highway 1; turn east on Hwy 1 and proceed 1.5 miles to junction with Highway 169; turn southwest (right) on 169 and travel 0.8 miles to Peyla Road; turn east and travel to STOP 8 described below. Access will depend on road conditions, and several scattered exposures off Peyla Road will be examined. STOP 8 – Peyla sequence – basalt, conglomerate, and sandstone Location: UTM: 0549822E/5292233N and environs, Tower quadrangle, Peyla Road east of Highway 169. Description: Near the end of Peyla Road is a series of outcrops of pillow basalt overlain by mafic conglomerate interbedded with sandstone (Figure 9). Published geologic maps (Ojakangas and others, 1978, Sims and Southwick, 1985, Southwick, 1993) show the basalts but do not distinguish the conglomerate and sandstone from the typical graywacke of the Lake Vermilion Formation. The Peyla conglomerate and interbedded sandstone overlie variably variolitic pillow basalt (Figure 9). Clasts in the conglomerate range from less than 1 cm to as much as 40 cm in size and are very angular; in fact the term ‘sedimentary breccia’ might be more appropriate in most cases. Topping indicators in both the basalts and sediments are difficult to find. 19 �The conglomerate (Figure 10A) is dominated by fine-grained basalt clasts that include weakly porphyritic, variolitic, and amygdaloidal phases. Medium-grained clasts of metagabbro/lamprophyre are common and in a few places predominant; other less common clast types include fine-grained possibly tuffaceous felsic rocks, and rare clasts of sulfides (pyrite) and hornblende-phyric andesitic hypabyssal intrusive rocks (Figure 10B). The matrix is similar to the adjacent sandstone. The varied types of basalt clasts (amygdaloidal, massive, variolitic, and porphyritic) imply reworking of the basalt substrate, and the polymictic nature of the conglomerate leads to the inference that this is a “Timiskaming-type” sedimentary package. Figure 9. Preliminary geologic map of the Peyla sequence on a 1 m lidar base. Most of fringing area has not been re-examined as part of recent field work, and the full extent of Timiskaming-type sandstone and conglomerate has not been established. The units shown here as green sandstone and conglomerate had not been distinguished on prior published maps, thus the fringing area is subject to revision by future mapping endeavors. 20 �B A Figure 10. Photograph (A; scale in cm) and photomicrograph (B) of typical Peyla conglomerate. Photomicrograph shows the edge of a basalt clast (B), a clast of hornblende-plagioclase porphyry (HAP), and hornblende (H) and plagioclase crystals (P) in the sandy matrix, which also contains minor quartz (Q). Lithic clasts outlined by white dashed line. Scale bar= 1 mm. The photomicrograph is from a different sample than that shown in the left photo. The ‘green sandstone’ (field term) interbedded with the conglomerate contains detrital plagioclase which commonly exhibits blocky and broken shapes, minor quartz derived from a volcanic source, blocky to euhedral detrital hornblende, small mafic to felsic volcanic rocks fragments, and rare detrital sphene and apatite along with metamorphic epidote, hornblende, biotite. The sandstone is typically quite massive and poorly bedded. The Peyla basalts are commonly variolitic, with irregular pillow shapes that commonly don’t yield reliable topping indicators. Locally the interiors of the pillows exhibit an incipient state of brecciation (Figure 11), outlining fragments similar in size and shape to those in the conglomerate. The reason for this brecciation is not known, but speculatively may be due to weathering and paleosol development prior to or during deposition of the overlying conglomerate and sandstone. Further development of this fracturing/brecciation process may have produced disaggregated fragments of angular basalt that were then shed into the adjacent sediments. Figure 11. Incipient brecciation in basalt. The sizes and shapes of these fragments are similar to the angular basalt clasts in the overlying mafic conglomerate. The white veining between the fragments is composed of clinozoisite, quartz, and carbonate. 21 �Mafic lamprophyric dikes, typically less than 2 meters wide and with sharp straight edges, intrude the sedimentary rocks, but none have been noted in the basalts so far. Multiple generations of lamprophyric dikes are visible on some outcrops, as well as rare instances of apparent lamprophyric peperite (Figure 12). The presence of mafic peperite and the clasts of hornblende-rich mafic intrusive rocks in the conglomerate imply that the lamprophyres may be related to subalkalic igneous activity contemporaneous with deposition of the sedimentary rocks. In contrast, very few if any mafic/lamprophyric dikes were noted in the Peyla basalts, based on the mapping completed thus far. The reason for this is not clear, but could be that the basalts were relatively impermeable to the dikes compared to the overlying sedimentary rocks. Figure 12. Thin dark green lamprophyric intrusion; on one side is in sharp, straight contact with the conglomerate and on the other side is diffuse and peperitic. DIRECTIONS: Return to Highway 169 and continue south. Drive ~2 miles from Peyla road to Flaim road, which is at the south end of a long road cut. Turn right (west) on Flaim road and park (STOP 9). STOP 9 – Sandstone, pillowed basalt, and lamprophyric dikes. Location: UTM: 0546116E/5287910N (Intersection of Flaim Road and Hwy. 169; Park on Flaim Road). Note: the road shown on the published Biwabik NW quadrangle is the old road which has been straightened and now cuts further west. Description: This is a new road cut that runs north-south parallel to Highway 169, and along the north side of Flaim Road west of Highway 169. It exposes metasedimentary strata on the south, and metavolcanic strata on the north, and mafic dikes emplaced into both rock types. Start on the westernmost outcrops along the north side of Flaim road and work east toward the highway. This glacially polished flat outcrop consists primarily of light tan-colored, fine-grained sandstone, having weakly graded beds that indicate southward stratigraphic facing. The sedimentary strata were intruded by 2 lamprophyric/dioritic dikes. Both the dikes and the bedding in the sedimentary strata dip steeply. One of the dikes contains xenoliths of varied rock types that range from 5 to 60 cm in size, and include layered 22 �tonalitic gneiss, porphyritic dacitic volcanic rocks, granodioritic to tonalitic intrusive rocks, and mafic schist. Adjacent to (south of) the inclusion-rich lamprophyre is a >30 cm thick dike of hornblende- and plagioclase-phyric diorite. This dike contains abundant phenocrysts of equant hornblende and blocky plagioclase as large as 3mm (Figure 13A), and small apatite crystals visible only in thin section. A B Figure 13. Plane light photographs of thin sections of samples from this stop (sections are approximately 2.5 cm wide). A – Hornblende- and plagioclase-phyric lamprophyric dike that intruded sedimentary strata. B – Lamprophyric dike that intruded pillow basalt showing trachytoid alignment of euhedral hornblende crystals. Continue walking east to Highway 169 along the sandstone outcrop, then go north along the freshly blasted roadcut. It is difficult to distinguish between the sandstone and subjacent basalt on the fresh rock faces; however, the transition is marked by a rusty interval that is visible from a distance (look to the outcrops across the highway). Two cherty horizons occur within the basalt– a layer that is orange in color, and further north a black chert horizon that contains thin layers of pyrite. Numerous dark green 0.5 – 10 m thick lamprophyric/mafic dikes are present within the basalt. Continue to north end of roadcut, climb up to top, and walk back south. Here on the glacially polished exposure one can see pillowed metabasalt and the mafic dikes emplaced into them. The dikes are weakly hornblende-phyric (Figure 13B), have chilled margins, and appear to be less deformed than the basaltic host. One highly unusual feature in the pillow basalt is the presence of two 7 cm rounded xenoliths of pink, medium-grained granite (Figure 14). The nearly vertical contact between the pillowed metabasalt to the north and adjacent metasedimentary strata to the south is fairly straight, abrupt, and lacks evidence for shearing that might indicate a fault origin. Although the contact zone is quite rusty (presumably pyritic), the contact is unremarkable. Nevertheless, the contrast in apparent depositional environments, and the abrupt termination of pillowed basalt (i.e., no gradation of pillows to flow-top breccia) implies that the contact is an unconformity. 23 �Figure 14. Rounded xenolith of pink, medium-grained granite in pillow basalt, on upper flat outcrop surface. UTM coordinates notes as 546129, 5288055 (Nad83 zone 15). STOP 10 – Conglomerate and metagraywacke, mafic dikes. Location: UTM: 0546327E/5286624N (Hwy. 26 0.6 mi east of Hwy. 169). Description: Highly flattened/lineated multi-lithic conglomerate interbedded with tightly folded graywacke, which is cut by multiple north-dipping, weakly foliated mafic/lamprophyric dikes. The conglomerate contains 1-10cm (longest dimension in end view) clasts that are also lineated (foliation N85°E, 70°N; lineation plunge 45° to N70°E). The clasts show weak size grading, and vary from felsic (light grayish-tan to pink and fine-grained) to mafic (dark green plagioclase-phyric, hornblende-rich; Figure 15). The matrix contains little if any quartz, and is generally dark green and amphibolitic; however it is commonly difficult to distinguish between matrix and pseudomatrix (i.e., flattened, less competent clasts). The surface of the conglomerate outcrops has a patchy pink staining, which is likely due to abundant microcline. In thin section the fine-grained light-colored clasts are composed predominantly of finely granoblastic microcline, with some larger (but less than 1mm) grains of irregularly-shaped plagioclase that may have been phenocrysts, and little if any quartz. Other clasts cannot be distinguished from matrix/pseudomatrix. Some of the clasts are more coarse-grained and microcline-rich, and may in part be granoblasticrecrystallized syenite. The matrix is composed of granoblastic plagioclase and metamorphic amphibole, lesser K-feldspar, and minor carbonate, and has small blocky, saussuritized plagioclase crystals. The metagraywacke in the western portion of this outcrop exhibits some weakly graded beds which mostly indicate south topping, thus are slightly overturned. Fine-grained amphibolitic lenses are present in the graywacke, and overall it has a slight orange stain due to finely disseminated pyrite. The mafic dikes here are likely of the same timing as those which cut the pillow basalts at stop 9. 24 �Figure 15. Strongly flattened and lineated dark green mafic-intermediate plagioclase-phyric clasts and smaller fine-grained, light-colored clasts in conglomerate. REFERENCES Boerboom, T.J., and Zartman, R.E., 1993, Geology, geochemistry, and geochronology of the central Giants Range batholith, northeastern Minnesota: Canadian Journal of Earth Science, v.30, p. 2510-2522. Driese, S.G., Jirsa, M.A., Ren, M., Sheldon, N.D., Brantley, S.L., Parker, D., and Schmitz, M., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, v. 189, p. 1-1. Jirsa, M.A., 2000, The Midway sequence: a Timiskaming-type, pull-apart basin deposit in the western Wawa subprovince, Minnesota: Canadian Journal of Earth Sciences, 37:1-15. Jirsa, M.A., and Boerboom, T.J., 2003, Bedrock geology of the Vermilion Lake 30’ X 60’ quadrangle, northeast Minnesota: Minnesota Geological Survey Miscellaneous Map M-141; scale 1:100,000. Jirsa, M.A., Boerboom, T.J., and Chandler, V.W., 2012, Geologic map of Minnesota: Precambrian bedrock: Minnesota Geological Survey State Map Series S-22, scale 1:500,000. Jirsa, M.A., and Boerboom, T.J., and Peterson, D.M., 2001, Bedrock geologic map of the Eagles Nest quadrangle, St Louis County, Minnesota: Minnesota Geological Survey Miscellaneous Map M-114; scale 1:24,000. Jirsa, M.A., Starns, E., and Schmitz, M.D., in prep., Bedrock geologic map of the Cavity Lake fire area, northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map M-193, scale 1:24,000 [in preparation—in the interim, refer to MGS Open-File Report OF-08-05, or contact lead author for details]. 25 �Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., and Jirsa, M.A., and Hamilton, M.A., 2013, New U-Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research v. 235, p. 264-277. Ojakangas, R.W., Sims, P.K., and Hooper, P.R., 1978, Geologic map of the Tower quadrangle, St. Louis County, Minnesota: U.S. Geological Survey geologic quadrangle map GQ-1457, scale 1:24,000. Peterson, D.M., Gallup, C., Jirsa, M.A., and Davis, D.W., 2001, Correlation of Archean assemblages across the U.S.-Canadian border: Phase geochronology: Institute on Lake Superior Geology, 47th Annual Meeting, Madison, Wisc., Proceedings v. 47, Part 1., p.77-78. Sims, P.K, and Southwick, D.L, 1985, Geologic maps of Archean rocks, western Vermilion District, northern Minnesota: U.S. Geological Survey Miscellaneous Investigation Series Map I-1527, scale 1:48,000. Southwick, D.L., 1993, Geologic map of Archean bedrock, Soudan-Bigfork area, northern Minnesota: Minnesota Geological Survey Miscellaneous Map Series map M-79, scale 1:100,000. Southwick, D.L., 1994, Assorted geochronologic studies of Precambrian terranes in Minnesota: A potpourri of timely information: in Southwick, D.L., (Ed.), Short contributions to the geology of Minnesota: Minnesota Geological Survey Report of Investigations 43, p. 1-19. [age reference in Figure 1] 26 �CU-NI-PGE DEPOSITS OF THE DULUTH COMPLEXGEOLOGY AND DEVELOPMENT Figure from Dean Peterson. (Note that resources are exclusive of reserves, and that various mineral resources were added to various mineral reserves just to allow comparisons in orders of magnitude. Note also that some of the categories only consider mineral in the ground and do not take into consideration the costs, recoveries, and other relevant factors associated with extraction and recovery of the metal or mineral) Mark Severson (Teck American Incorporated) Andrew Ware (PolyMet Mining) Kevin Boerst (Twin Metals) Stephen Monson Geerts (Natural Resources Research Institute) Portions previously written by: Richard L. Patelke (formerly with PolyMet Mining) Tim Jefferson (formerly with Teck-Cominco Inc.) Dean M. Peterson (formerly with Duluth Metals Limited) 27 �EXPLORATION AND DEVELOPMENT BACKGROUND By the late Richard Patelke (PolyMet Mining) with modifications by Mark Severson Large resources of low-grade copper-nickel sulfide ore that locally contain PGE concentrations are well documented by drilling in the basal zones of the Partridge River, Bathtub, and South Kawishiwi intrusions. At least eleven occurrences of significant mineralization have been delineated in the basal 300 to 1000 feet of these intrusions. Of these eleven occurrences, two projects have undergone fairly recent definition drilling, including the Mesaba deposit (Teck American) and Maturi deposit (Twin Metals). Definition drilling at the Birch lake deposit (Twin Metals) took place up to four years ago. A fourth project, the NorthMet deposit (PolyMet Mining) is currently undergoing environmental review and mine permitting. Recent exploration drilling has taken place at the Serpentine deposit (Encampment Resources). Overall, the copper-nickel mineralization consists predominantly of disseminated sulfides that historically are estimated to contain over 4.4 billion tons of material averaging 0.66% Cu and 0.20% Ni at a 0.5% Cu cut-off, according to an earlier study (Listerud and Meineke, 1977): note that this estimate is historic and does not follow reporting guidelines as established by CIM Definition Standards). As outlined in Miller et al. (2002a), serious exploration for Cu-Ni deposits at the base of the Duluth Complex began in 1948, about 8 miles to the southeast of Ely, MN, when strongly mineralized rocks were uncovered in an excavation used to source road material for Spruce Road. Local prospector Fred S. Childers of Ely noted copper stains in the material and he, along with Roger V. Whiteside of Duluth, began searching along the basal contact in the vicinity of the Kawishiwi River. In 1951, they diamond drilled a 188 foot deep hole and intersected mineralized gabbro that averaged 0.36% Cu and 0.13% Ni. In 1952, both Bear Creek Mining Company (BMC) and the International Nickel Company (INCO) began intensive exploration efforts along a 38 mile-long zone that coincided with the basal contact. INCO eventually picked up the Childers-Whiteside properties (Spruce Road and Maturi deposits); whereas, BMC concentrated most of their effort near the town of Babbitt which resulted in the discoveries of the Babbitt (formerly called Minnamax and now known as Mesaba) and Serpentine deposits. By 1960, these exploration efforts indicated that very large tonnages of disseminated Cu-Ni mineralization were present; however, the low-grade nature of the deposits and the unavailability of state-owned mineral lands at the time led to suspension of activities. In 1966, state mineral leases were offered by the Minnesota Department of Natural Resources (DNR) and were awarded to successful bidders, resulting in renewed exploration activity (including the return of BMC and INCO). Since 1966, over 20 companies have been actively involved in exploration for Cu-Ni and Fe-Ti-V deposits along the basal contact of the Complex. Over 2,600 holes, totaling over 4.5 million feet of core, have been drilled. Exploration efforts during this period also defined several more deposits including: Dunka Road (now NorthMet) and Wyman Creek (United States Steel Corp.), Birch Lake (Duval Corporation and Newmont Mining), South Filson Creek (Hanna Mining), Dunka Pit (Erie Mining, BMC, and Exxon), and Wetlegs (BMC and Exxon). AMAX Exploration Inc. leased the Babbitt deposit from BMC and renamed it the Minnamax deposit in 1973. During mid to late 1970s, the Spruce Road and Minnamax deposits came closest to development. Mining plans were submitted, test shafts were sunk (one each at the Maturi and Minnamax/Mesaba deposits), surface bulk samples were collected from three sites, and various land-use and water-use permits were requested from State and Federal agencies. In 1974, the Minnesota Environmental Quality Board required that a regional Environmental Impact Statement (EIS) be conducted prior to acceptance of any site-specific EIS mining-related proposals. The DNR discontinued lease sales of State lands (1974-1982) until completion of the regional EIS. However, by the time the regional EIS was submitted in 1979, development of the Cu-Ni deposits was put on hold by the most of the mining companies involved due to weakened copper and nickel markets and the inability to make marketable (i.e., "smeltable") separate copper and nickel concentrates. Amax abandoned their plans to develop an underground high grade ore zone within the Minnamax/Mesaba deposit (known as the Local Boy ore body) in late 1982. 28 �Then starts the “PGE” era. During the early period of drilling (prior to 1980), all of the exploration companies recognized that the Cu-Ni deposits had some potential for hosting PGEs. Based on very limited sampling, the companies assumed that the typical Cu-Ni ore contained no more than a few hundred parts per billion (ppb) combined platinum and palladium. In 1985, the DNR and Minerals Resource Research Center (MRRC of the U of M) conducted a geochemical evaluation of portions of a Duval drill hole (DU-15), from the Birch Lake area, and found significant values of up to 9 parts per million (ppm) combined Pt and Pd (Sabelin and Iwasaki, 1985, 1986). This was at a time when demand for these elements was increasing due to their use in automotive catalysts. A short time later, Morton and Hauck (1987) compiled all of the known PGE data for the Complex and reported the presence of anomalous PGE values, often associated with high Cu values, at several other Cu-Ni deposits. These discoveries sparked renewed interest in the Cu-Ni deposits as potential polymetallic deposits (Miller et al., 2002; and references therein). E.K. Lehman and Associates of Minnesota obtained mineral leases from the state of Minnesota and began drilling wedges off the discovery hole (DU-15W) in the Birch Lake area. These Lehmann leases were later incorporated into Franconia Minerals holdings. Additional drill holes were sampled and analyzed for PGEs by several other companies throughout the Duluth Complex, and as a result, significant PGEs were found at many deposits. The occurrences of PGE mineralization for each deposit will be more thoroughly discussed later in this guidebook. Enter the “Hydromet” era. Early development of the deposits was hampered both by state leasing issues, complex metallurgy that resulted in an inability at the time to make marketable separate Cu and Ni concentrates, and by general environmental concerns regarding sulfide mining and conventional pyrometallurgical processes. In the mid to late 1990s, the potential of developing the Cu-Ni deposits using hydrometallurgical techniques once again sparked renewed activity in the Duluth Complex. PolyMet plans to use the PlatSol technique, developed and patented by SGS Lakefield on NorthMet ores, to recover Cu, Ni, Co, and PGE at the NorthMet deposit. REGIONAL GEOLOGIC SETTING, DULUTH COMPLEX The Duluth Complex and associated intrusions of Keweenawan age (~1.1 billion years) in northeastern Minnesota constitute one of the largest mafic intrusive complexes in the world, second only to the Bushveld Complex of South Africa (Miller et al., 2002). These rocks cover a 2,200 square mile (5,700 square km) arcuate area associated with the two strongest gravity anomalies (+50 and +70 milligals) in North America, that imply intrusive roots more than 8 miles (13 km) deep (Allen and others, 1997). The co-magmatic flood basalts and intrusive rocks underlying much of northeastern Minnesota were emplaced during development of the Mesoproterozoic Midcontinent rift, which can be traced geophysically from exposures in the Lake Superior region along a 1250 mile (2,000 km) long, segmented, arcuate path to Kansas and Lower Michigan. The Duluth Complex is defined as the more or less continuous mass of mafic to felsic plutonic rocks that extends for >170 miles (275 km) in an arcuate fashion from Duluth nearly to Grand Portage (Fig. 3-1). It is bounded by a footwall of Paleoproterozoic sedimentary rocks and Archean granite-greenstone terranes (Peterson and Severson, 2002), and a hanging wall largely of comagmatic, rift related flood basalts and hypabyssal intrusions of the Beaver Bay Complex (Fig. 3-1). In genetic terms, the Duluth Complex is composed of multiple discrete intrusions of mafic to felsic tholeiitic magmas that were episodically emplaced into the base of a volcanic edifice between 1108 and 1098 Ma. The geology of the Duluth Complex and adjacent areas has recently been described in two major publications by the Minnesota Geological Survey (MGS). These include a 1:200,000 scale regional bedrock geological map of northeastern Minnesota (Miller et al., 2001), and a comprehensive written description of the geology depicted on this map (Miller et al., 2002), commonly referred to as the “bible” by geologists working on Duluth Complex geology. Readers’ interested in more detailed descriptions of 29 �the geologic setting of the Duluth Complex should begin their quest for knowledge by downloading these publications from the MGS website (ftp://mgssun6.mngs.umn.edu/pub2/). Within the nearly continuous mass of intrusive igneous rock forming the Duluth Complex, four general rock series are distinguished on the basis of age, dominant lithology, internal structure, and structural position within the complex. Felsic series—Massive granophyric granite and smaller amounts of intermediate rock that occur as a semi-continuous mass of intrusions strung along the eastern and central roof zone of the complex, that were emplaced during an early stage magmatism (~1108 Ma). Early gabbro series—Layered sequences of dominantly gabbroic cumulates that occur along the northeastern contact of the Duluth Complex, emplaced during early stage magmatism (~1108 Ma). Anorthositic series—Structurally complex suite of foliated, but rarely layered, plagioclase-rich gabbroic cumulates emplaced throughout the complex during main stage magmatism (~1099 Ma). Layered series—Suite of stratiform troctolitic intrusions that comprises at least 11 variably differentiated mafic layered intrusions that occur mostly along the base of the Duluth Complex. These intrusions were emplaced shortly after the Anorthositic series (~1099 Ma). Figure 3-1. Generalized geologic map of northeastern Minnesota (modified from Miller et al., 2002). Rock Type and Unit Classification Igneous rock types in the Duluth Complex are classified at each of the deposits by visually estimating the modal percentages of plagioclase, olivine, and pyroxene, and using a rock classification scheme (Figure 3-2) modified from Phinney (1972). Using this classification, the majority of rocks at the various deposits consist of troctolite, augite troctolite, anorthositic troctolite, and norite (near the basal contact) with local 30 �ultramafic layers consisting of melatroctolite to dunite. Due to subtle changes in the percentages of the estimated minerals, there can be subsequent variations in the defined rock types within a specific igneous stratigraphic rock unit on a hole by hole basis, on an interval by interval basis, and even on a geologist by geologist basis. Figure 3-2. Modified Phinney (1972) diagram for rock type classification. Overall, stratigraphic unit definitions are based on: dominant rock type; textural relationships; mineralogy; sulfide content; and context with respect to bounding surfaces (i.e., ultramafic horizons, oxide-rich horizons). Unit definitions are not always immediately clear in logging, but are usually clarified when drill holes are plotted on cross-sections. In other words, to correctly identify a particular stratigraphic unit, the context of the units directly above and below should also be considered. LOCAL GEOLOGIC SETTING-PARTRIDGE RIVER, SOUTH KAWISHIWI, AND BATHTUB INTRUSIONS By: Mark Severson The three deposits under review for this trip are located in three of the oldest intrusions in the Duluth Complex. The NorthMet deposit and parts of the Mesaba deposit are in the Partridge River intrusion, the majority of the Mesaba deposit in the newly defined Bathtub intrusion, and Maturi deposit in the South Kawishiwi intrusion (Fig. 3-3). Partridge River intrusion The Partridge River intrusion (PRI) consists mainly of troctolitic cumulates, dips gently to the southeast, and is exposed in an arc-shaped area that extends from the Water Hen deposit, on the southwest, to the southern edge of the Mesaba/Babbitt deposit, on the northeast (Fig. 3-3). Footwall rocks include the Paleoproterozoic Virginia Formation and locally the Biwabik Iron Formation. The basal 3000 ft. (900 meters) are known in great detail from studies of abundant drill core (Severson and Hauck, 1990) and are subdivided into seven or more units that can be traced over a strike-length of 15 miles (24 kilometers). 31 �Figure 3-3. Location of Cu-Ni±PGE sulfide deposits, Fe-Ti±V oxide deposits (Oxide-bearing Ultramafic Intrusion - OUI), and other exploration areas along the western edge/base of the Duluth Complex. Note that the NorthMet deposit was referred to as the Dunka Road deposit and the Mesaba deposit was referred to as the Babbitt deposit; the most recent names for these two deposits are used in this guidebook. The Birch Lake deposit will not be discussed in this guidebook. The units of the Partridge River intrusion (PRI) are recently described in Miller and Severson (2002) and are depicted in Figure 3-4. At the base of the PRI is Unit I which consists of a suite of heterogeneous- 32 �textured troctolitic rocks that contain the vast majority of disseminated sulfide-mineralized zones. The top of Unit I is marked by a fairly persistent ultramafic horizon, which in actuality is at the base of Unit II. Within Unit I are several laterally-discontinuous ultramafic horizons and abundant footwall sedimentary inclusions of the Virginia Formation. Noritic rocks are common at the basal contact and adjacent to the inclusions due to silica contamination from assimilated footwall rocks. Unit II consists of more homogenous-textured troctolitic rocks with minor sulfide-bearing zones. However, at the Wetlegs deposit, both Units I and II contain abundant laterally-discontinuous ultramafic horizons, interbedded with troctolitic rocks that are collectively referred to as the Wetlegs Layered Interval (Fig. 3-4). Figure 3-4. Generalized stratigraphy of the basal zone of the Partridge River intrusion (modified from Severson, 1994). Roman numerals (I through VIII) denote igneous units in the Partridge River intrusion; BT1 and BT4 denote igneous units in the Bathtub intrusion; and OUI denotes Oxide-bearing Ultramafic Intrusions. Overlying Unit III in the PRI are units IV through VIII. Unit IV varies from a troctolite to augite troctolite, often contains an ultramafic base, and commonly grades upward into Unit V which is coarsergrained and varies from a troctolite to troctolitic anorthosite. Units VI and VII, and additional units above VII, are generally homogenous-textured troctolitic to anorthositic troctolitic rocks; each with a persistent ultramafic base that record magma injection events. Unit III is a major marker bed throughout much of the PRI (Wetlegs to Mesaba deposits - Figs. 3-3 and 34) in that it is characterized by a poikilitic leucotroctolite with olivine oikocyrsts that are randomly dispersed throughout the rock giving it a mottled appearance. This mottled-appearance, and the relatively fine-grained nature of Unit III, give it a distinct appearance in drill core and it is easily identified. Unit III pinches out to the west of the Wetlegs deposit and is present on only the southern fringe of the Mesaba deposit. The rapid pinch-out of Unit III to the north within the Mesaba deposit appears to be related to emplacement of a distinctly different sub-intrusion herein referred to as the Bathtub intrusion (see discussion below). Bathtub intrusion The Bathtub intrusion (BTI) is wholly contained in the central portion of the Mesaba (Babbitt) deposit. It has recently been singled out as a separate intrusion to explain the abrupt change from typical Partridge 33 �River intrusion stratigraphy in the extreme southern part of the deposit to a completely different stratigraphy, to the north in the remainder of the deposit (Severson and Hauck, 2008). There are three structural features that are pertinent to understanding the intrusive history of the BTI that include (Fig. 35): 1. an east-west trending paired syncline and anticline in the footwall rocks referred to as the Bathtub Syncline and Local Boy Anticline; 2. a zone that is closely associated with the Local Boy Anticline, referred to as the “Hidden Rise,” that separates the PRI and BTI; and 3. a north-trending fault zone, referred to as the Grano Fault, that is situated on the extreme eastern portion of the Mesaba deposit – the fault has been postulated to have been the feeder zone for the BTI and footwall-injected massive sulfides of the Local Boy ore zone. The “Hidden Rise” is a loosely-defined zone wherein scattered hornfels inclusions, and associated noritic rocks, are fairly common. When viewed collectively, the inclusions in “The Hidden Rise” define an eastwest trending “ridge” that is coincident with the Local Boy Anticline and roughly positioned at the contact between the PRI and BTI. Thus, “the “Hidden Rise” is used to both define this hornfels-bearing “ridge” and to artistically, and conveniently, divide the BTI from the PRI (Fig. 3-5). The morphology of this feature suggests that it may have originally served as the southern edge of an earlier intruded BTI and later served as a wall along the floor and north edge of the PRI as its upper units were emplaced. The BTI has been subdivided into two main units, BT1 and BT4, each of which contain several internal subunits (Fig. 3-5). In the vicinity of the Bathtub Syncline, ultramafic layers and modally-bedded rocks are extremely common within the BT4 Unit and have been collectively referred to as the Bathtub Layered Interval (BTLI). Figure 3-5. Schematic “type-section” looking east through the Mesaba deposit that crudely displays the spatial distribution of most of the igneous units in the Bathtub intrusion and pertinent structural features. Note that not all of the PRI units are shown on the right side of the figure. Cu-rich massive sulfides are locally present at the Mesaba deposit in a small zone referred to as the Local Boy ore zone. Local Boy is positioned along the crest of the Local Boy Anticline, in close to proximity to the “Hidden Rise,” and just west of the Grano Fault. Most of the massive sulfides are associated with either hornfelsed sedimentary inclusions above the basal contact or with footwall rocks below the contact while the interfingering intrusive rocks (mostly norite) are relatively barren of massive sulfides (Severson and Barnes, 1991). This suggests that the massive sulfide ores were not formed by the gravitational settling of sulfides, but rather, the ores formed by injection of an immiscible sulfide melt into structurally prepared areas within the footwall rocks along the Local Boy anticline in a vein-like setting. A possible 34 �feeder vent for the sulfide injection event may have been the Grano Fault, which was repeatedly reactivated during emplacement of the Complex. West-directed increases in Cu and PGE, associated with the massive sulfides at Local Boy, suggest that the immiscible sulfide melt fractionally crystallized and became progressively enriched in Cu and PGE as it was deposited in an east-to-west direction. Partridge River and Bathtub intrusion footwall rocks Because the footwall at NorthMet and Mesaba is so similar, the following is a generic description appropriate to both deposits. The drilled footwall rock types at Mesaba and NorthMet consist mainly of the Virginia Formation and Biwabik Iron Formation. Both are Paleoproterozoic in age (approximately 1.9-1.8 Ga) and are the two upper units of the Animikie Group. Any discussion on these two formations must include a description of their type-section on the Mesabi Range, as well as, a description of them as related to the metamorphism and partial melting that was produced during emplacement of the Complex. Lying beneath the Biwabik Iron Formation, but encountered in only a few drill holes are the Paleoproterozoic Pokegama quartzite (also of the Animikie Group), along with granitic rocks of the Archean Giant’s Range Batholith. Biwabik Iron Formation The Biwabik Iron Formation (BIF) exposed on the nearby Mesabi Range has typically been subdivided into four informal lithostratigraphic members (Wolff, 1917) that are, from the bottom up: Lower Cherty, Lower Slaty, Upper Cherty, and Upper Slaty. Diamond drill holes at Mesaba and NorthMet generally pierce the top submembers of the Upper Slaty, and end in submember C or D. Submember A is comprised of chert and marble, submember B is characterized by alternating bands of green diopside and chert with very coarse-grained hedenbergite, and submember C is a thin-bedded, green rock consisting of chert-fayalite-ferrohypersthene with black magnetite-rich bands. Virginia Formation below the PRI and BTI The Virginia Formation is a thick sequence of argillite, siltstone, and graywacke at the top of the Animikie Group. In close proximity to the Complex the effects of partial melting are profound and portions of the hornfelsed Virginia Formation no longer even remotely resemble a sedimentary rock. Severson et al. (1994a) subdivided the hornfelsed Virginia Formation, in both the footwall and in inclusions within the Duluth Complex, into at least five informal units based largely on metamorphic attributes, which are each related to varying degrees of partial melting. These members, and a pre-Duluth Complex sill, are described below and are schematically portrayed in Figure 3-6 - although in real occurrence, this idealized metamorphic progression is more erratic, often with rapid lateral and vertical changes between the four metamorphic units discussed below. 35 �Figure 3-6. Schematic cross-section showing the general relationships of the metamorphosed footwall rocks beneath the Duluth Complex at the Mesaba, NorthMet, Wetlegs, and Serpentine deposits. Cordieritic hornfels Directly beneath the basal contact of the Duluth Complex, the adjacent Virginia Formation typically consists of massive/non-foliated, cordierite-rich hornfels that display a bluish-gray color in drill core. The rock is generally fine-grained, granoblastic, and biotite-poor (due to loss of water into the Complex) and locally may contain porphyroblastic and/or poikiloblastic cordierite. Original bedding planes are preserved in some localities, but mostly the bedding planes have been obliterated by contact metamorphism. Recrystallized unit (RXTAL) Beneath the cordieritic “capping” the next metamorphic variant of the Virginia Formation nearest to the Duluth Complex is a rock that is referred to as the RXTAL unit. The RXTAL unit is properly classed as a diatexite and is characterized by fine- to medium-grained cordierite, plagioclase, biotite, quartz, and Kspar with lesser amounts of Opx and opaques. Bedding planes of the original argillaceous rocks are obliterated and what remains is a massive recrystallized rock with decussate biotite that contains enclaves (blocks and folded boudins) of more structurally competent calc-silicate hornfels and thin-bedded siltstone Disrupted unit (DISRUPT) With increased distance from the Complex, the RXTAL unit progressively grades into the DISRUPT unit which is a thin-bedded rock that is visibly deformed and underwent less degrees of partial melting. Textures that characterize the DISRUPT unit are bedding planes that are extremely chaotic and random in orientation due to pervasive small-scale folding, faulting, and brecciation. Superimposed on this chaotic pattern are abundant zones of leucocratic partial melts that are also chaotic and folded. The rock consists of varying amounts of quartz, cordierite, K-spar, biotite, plagioclase, and muscovite with leucosome veins and patches containing quartz, K-spar (microperthite), plagioclase, and muscovite (Duchesne, 2004). The DISRUPT unit is properly classed as a metatexite. 36 �Graphitic argillite and Bedded Pyrrhotite (BDD PO) units Carbonaceous argillite of the lower portion of the Virginia Formation is commonly preserved as either the BDD PO unit, or graphitic argillite, in close proximity to the Duluth Complex. This rock commonly contains over 5% disseminated pyrrhotite and/or extremely thin-bedded pyrrhotite laminae (hairlinethick), and variable amounts of graphite, staurolite(?) and sillimanite. Wherever the unit contains conspicuous and regularly-spaced laminae of pyrrhotite (0.5-3.0 mm thick at 1-20 mm spacings) it is informally referred to as the bedded pyrrhotite unit (BDD PO unit). In some areas, the BDD PO served as a local sulfur source to both disseminated and massive sulfide occurrences at the base of the Duluth Complex. VirgSill The VirgSill is generally present in the bottom 0.5-130 feet of the Virginia Formation, and as local apophyses into the top of the Biwabik Iron Formation. The VirgSill was intruded along the contact between the Virginia Formation and Biwabik Iron Formation and exhibits a granoblastic texture indicating that it was metamorphosed by the Duluth Complex (and thus the VirgSill is pre-Duluth Complex in age). On this basis, the VirgSill is inferred to be equivalent to the Logan sills (circa 1,109 Ma); as is another sill, the BIFSill, in the C submember of the Biwabik Iron Formation (Hauck et al., 1997). However, the VirgSill and BIFSill are different chemical entities (the VirgSill is much more Crenriched), and thus, these two sills may be related to at least two different intrusive events. Identification of the VirgSill in drill core is hampered by the fine-grained granoblastic texture that makes it difficult to distinguish from the enclosing hornfelsed Virginia Formation rocks. The VirgSill is subdivided into two textural varieties (Severson et al., 1994a; Park et al., 1999) referred to as: 1. the Massive Gray unit (MG unit); and 2. a coarser-grained interior with obvious hornblende and/or olivine. South Kawishiwi intrusion The South Kawishiwi intrusion (SKI) consists mainly of troctolitic cumulates and dips gently to the southeast. The SKI is exposed in an arc-shaped area that extends from the Serpentine deposit, on the southwest, to the Spruce Road deposit, on the northeast (Fig. 3-3). Footwall rocks include the Paleoproterozoic Virginia Formation, Biwabik Iron Formation and Archean Giants Range Batholith, the latter is the dominant footwall rock type. The presence of Biwabik Iron Formation as inclusions, from the Birch Lake deposit to as far north as the Spruce Road deposit, indicates that the majority of Paleoproterozoic units were assimilated and removed from the footwall during emplacement of the South Kawishiwi intrusion (Severson et al., 2002). The basal stratigraphic section of the SKI is known in great detail from studies of abundant drill core and is subdivided into 17 different units (Fig. 3-7) that are present over a strike-length of 19 miles (31 kilometers). The lowermost units are unevenly distributed along the strike length of the intrusion in a “compartmentalized” fashion, suggesting a complicated intrusive history (Miller and Severson, 2002). A few salient features to keep in mind regarding the igneous stratigraphy of the SKI include: • The vast majority of sulfide mineralization is confined to the BH (Basal Heterogeneous Unit), BAN (Basal Augite Troctolite and Norite Unit), UW (Updip Wedge Unit), and U3 (Ultramafic 3 Unit) – the latter three of these units are combined and referred to as the BMZ (Basal Mineralized Zone) by Twin Metals at their Maturi deposit; • Major marker beds, at specific areas in the SKI, include three horizons that contain abundant cyclic ultramafic layers (U1, U2, and U3 Units) and a pegmatite-bearing unit (PEG Unit originally recognized by Foose, 1984). The U1, U2 and U3 Units represent periods of rapid and 37 �continuous magma replenishment that crystallized more primitive ultramafic layers before mixing with the resident magma (Severson et al., 2002); • The U3 Unit is unique in that it contains several massive oxide pods (titanomagnetite-rich), as well as, recognizable inclusions of bedded Biwabik Iron Formation; especially at the Birch Lake deposit. The spatial correspondence between the U3 Unit and footwall iron-formation suggests that most of the massive oxide pods are iron-rich “restite” produced by assimilation and partial melting of the iron-formation (Muhich, 1993; Severson, 1994; Severson et al., 2002); • The U3 Unit contains the vast majority of high PGE values, especially within the Birch Lake area and possibly at the Maturi deposit. However, high PGE values are also present in the PEG Unit (Birch Lake area and Maturi deposit), the top of the BH Unit (Maturi deposit), and very locally in troctolitic rocks situated well above the basal contact (South Filson Creek deposit); and • A large inclusion/pillar of anorthosite is present at the Maturi deposit. This pillar, and possible proximity to a vent area and magma flow paths (see discussion for Maturi deposit) are some of the inferred reasons for high PGE values at the Maturi deposit. Figure 3-7. Generalized stratigraphy of the basal zone of the South Kawishiwi intrusion (modified from Severson, 1994; and included in Miller and Severson, 2002). The lowermost igneous units are: BAN = Basal Augite Troctolite and Norite; BH = Basal Heterogeneous; U3 = Ultramafic 3; PEG = Pegmatitic unit of Foose (1984); U2 = Ultramafic 2; U1 = Ultramafic 2; AT-T = Anorthositic Troctolite to Troctolite; UW = Updip Wedge; Main AGT = Main Augite Troctolite. 38 �PART 3A: POLYMET NORTHMET DEPOSIT By: the late Richard Patelke with modifications by Andrew Ware NORTHMET PROJECT SUMMARY NorthMet, located in the Partridge River intrusion of the Duluth Complex, is a large, disseminated sulfide deposit in heterogeneous troctolitic rocks associated with the 1,100 million year old Mid-Continent rift. Metals of interest are copper, nickel, cobalt, platinum, palladium, and gold. The majority of the metals are concentrated in four sulfide minerals: chalcopyrite, cubanite, pentlandite, and pyrrhotite, with platinum, palladium and gold also found in bismuthides, tellurides, and alloys. NorthMet is one of eleven coppernickel-PGE deposits along the northern margin of the Complex (PGE: platinum, palladium, gold). All of these share grossly similar geologic settings–disseminated sulfides with minor local massive sulfides in heterogeneous rocks forming the basal unit of the Duluth Complex along the contact with older rocks. The deposit is on the southern flank of the Mesabi Iron Range, which is host to six large operating taconite mines, the closest of which is less than two miles (3.2 km) north of the planned NorthMet pits (Figure 5A-1). Ore from NorthMet will be processed at a rate of 32,000 short tons per day through the former LTV Steel Mining Company iron ore concentration plant (“Erie Plant) with new facilities for processing of the NorthMet copper-nickel-PGE concentrates through a hydrometallurgical method to produce copper metal and various hydroxide and concentrate products of nickel-cobalt-PGE (Figure 5A2). PERMITTING and THE FINAL ENVIRONMENTAL IMPACT STATEMENT After 10 years of environmental review, the Final Environmental Impact Statement (FEIS) for the NorthMet Project was released in November 2015. The Minnesota DNR deemed the FEIS adequate in March 2016, and this DNR decision initiated the permitting process. The two other regulatory decision documents on the FEIS (Records of Decision from the US Forest Service and the US Army Corp of Engineers) are expected in 2016. The timelines for submitting and obtaining permits will be different for each permit. EXPLORATION and DEVELOPMENT There have been four major drilling programs since 1969, re-sampling for PGE began in 1989, three PolyMet joint ventures were pursued and dissolved in the 1990's, processing technology was developed in the late 1990's, the former LTV Steel Mining Company concentrator and other property was optioned in 2003, and the metallurgical process was refined in 2005-2008. Drilling programs have been conducted by United States Steel (USS, 1969-1974) and PolyMet Mining Inc. (Reverse Circulation or “RC” drilling and core drilling in 1998-2000 & two phases of core drilling in 2005 and 2007), plus two (actually two pairs of twins) holes by NERCO Minerals Company in 1991. This drilling encompasses 285,756 feet over 371 holes as of May 2008. Over 35,973 acceptable assays have been taken from this drilling (216,344 feet assayed). Table 3A-1 gives a breakdown of years, footages, and number of assays for all project drilling. United States Steel (USS) began core drilling at NorthMet (as the Dunka Road project) in 1969. Drilling targeted a conductor that turned out to be in the footwall metasedimentary rocks, but the first drill hole hit massive sulfide in the Duluth Complex. Drilling continued over five years for 112 holes with 133,716 feet of intercept. The working assumption was to mine the deposit from underground, sampling was limited to the most continuous zones with strong visible copper-nickel mineralization, and only about 2,200 samples representing about 22,000 feet were taken. USS assayed only for copper, nickel, sulfur, and iron. PGE 39 �presence was known from sampling on concentrates, but the economics of PGE recovery were apparently not pursued. Project work stopped while apparently incomplete and was not restarted. USS did not do much follow-up, but kept their land ownership, core, pulps, coarse rejects, and records for the project. In the mid 1980's the Minnesota Department of Natural Resources (MDNR) began sampling various historic drill core intervals in the Duluth Complex for PGE and got some good, but localized, results. In 1989 Fleck Resources (Fleck) leased the Dunka Road property from USS and began a program of re-assaying USS pulps and coarse rejects with a much more extensive multi-element suite, as well as adding in some new samples from existing core through cooperative work with the Natural Resources Research Institute (NRRI). The results were very positive in showing elevated PGE values in the deposit and confirming the previous copper-nickel assays. Fleck partnered with NERCO in 1991 for some bulk sample work, mine plans, environmental reviews etc., done through Fluor Daniel Wright engineers, but the partnership was eventually dissolved. In 1995 Fleck joined with Argosy Mining Corp. to do more work on the project, again with no major progress towards production. In June 1998, Fleck became PolyMet and focused their resources on Dunka Road, which was renamed NorthMet. Without partners, except for a brief venture with North Mining (North), PolyMet drilled and sampled 87 holes in 1998-2001, and sent two large bulk metallurgical samples to Lakefield Laboratories (now SGS) in Lakefield, Ontario for development and refinement of the PlatSol hydrometallurgical process and began some environmental background work. In the summer of 2000, North was taken over by Rio Tinto. The joint venture agreement was terminated upon consideration that NorthMet appeared to be a low priority to Rio Tinto. However, much of the North funding was already in place and was used to partially finance the 2001 pre-feasibility study. After release of the pre-feasibility study (2001), a brief hiatus, and a major re-evaluation of how the project should proceed, PolyMet became active again in 2003 with new management and a new development plan. This plan involves integrating the former LTV Steel Mining Company iron ore concentration plant (“Erie Plant) with new facilities for processing of the NorthMet copper-nickel-PGE concentrates through a hydrometallurgical method at rate of 32,000 short tons of ore per day to produce copper metal and various hydroxide and concentrate products of nickel-cobalt-PGE. Geologic work towards this end began in 2004 and first focused on a careful and total re-compilation of the historic NorthMet project drill hole related data. This effort organized and verified all drilling metadata, location, downhole survey, lithology, and assay data, and cataloged all paper (and digital) records for the project. Of note is that this resulted in an increase in the number of acceptable assays from 12,000 to around 17,200 and an improved geologic picture from careful consolidation of existing records. This work was used as background for a revised resource estimate in January 2005 and planning of a drill program for 2005. The 2005 program entailed drilling and sampling 109 holes (77,000 feet), collection of a forty ton metallurgical bulk sample for pilot scale testwork, geotechnical (oriented core) drilling, in-fill sampling of previously drilled core, and extensive collection of waste characterization data. The 2005 drilling program added 13,450 multi-element assay records to the existing database. A PolyMet report covers the details of historic drilling and assaying (Patelke & Geerts, 2006). 40 �Figure 3A-1. PolyMet NorthMet project site. 41 �Figure 3A-2. Detail of Erie Plant site showing existing facility and new construction. 42 �Drilling in 2007 for 24.530 feet with 3,546 assays concentrated on defining mineralization in the upper units in the west part of the deposit (the “Magenta Zone”). This drilling and the subsequent re-modeling of the deposit turned about 50 million tons of material previously classed as waste to ore. There is also over 34,000 feet of hydrogeology drilling and “stratigraphic holes” (drilling by other companies not done as part of the NorthMet project). No assays are in use from these 44 holes which are used for geologic control. Approximately 89.5% of Unit 1 and about 57% of the upper units have been sampled across the deposit. The sampled percentages are higher in the anticipated area of mining. Table 3A-1. Total drilling and assaying for NorthMet project. No. of drill holes Total footage for group No. of assay intervals used in “accepted values” tables Assayed footage used in final database Assay Laboratories 1969-1974 1969-1974, 19891991, 19992001, 20052006, 2008 112 133,716 11,259 73,303 USS, ACME, ALS-Chemex 1991 1991 2 (4) 842 165 822 ACME 1998-2000 1998-2000 52 24,650 4,765 23,767 ACME PolyMet core drilling 1999-2000 2000-2001, few in 2005 32 22,156 4,058 20,727 ALS-Chemex PolyMet RC drilling deepened with AQ core tail 2000 2000 3 2,696 524 2,610 ALS-Chemex PolyMet core drilling 2005 2005-2006 109 77,166 11,656 71,896 ALS-Chemex PolyMet core drilling 2007 2007 61 24,530 3,456 23,310 ALS-Chemex 371 285,756 35,973 216,344 Company US Steel NERCO PolyMet reverse circulation drilling Drilling years Assaying years Totals for Exploration Drilling:    US Steel stratigraphic holes 1970's? none 6 9,647 none none 1956 none 3 2,015 none none Humble Oil / Exxon 1968-1969 none 3 9,912 none none Bear Creek / AMAX 1967-1977 none 11 8,893 none none PolyMet / Barr Engineering (hydrologic testing) 2005-2007 none 21+ 3,459+ none none INCO Sampling in Unit 1 (the main mineralized zone) is now mostly continuous through the zone for all generations of drilling. The PolyMet RC and core holes have continuous sample through the upper waste zones (which do have some intercepts of economic mineralization). Work in 2005 through 2008 43 �essentially completed the sampling of historic USS core within the area likely to be mined. This broad sampling limits the possibility of location bias in the sample set. The entire USS core footage has not been sampled, however there is no known un-sampled mineralized intervals. Table 3A-2. Large metallurgical samples collected at NorthMet. Bulk Sample Year Tons Location of sample USS Bulk sample pit No. 1 1971 Unknown, but small Pit in center of property USS Bulk sample pit No. 2 1971 300 Pit at east end of property USS Bulk sample pit No. 3 1971 20 Pit at east end of property NERCO PQ drill core 1991 Argosy Mining 1995 Unknown, but small Composited from USS coarse rejects PolyMet RC drill cuttings 1998 26 One composite, mostly from what is now considered east part of 10 year pits PolyMet RC drill cuttings 2000 33 One composite, mostly from what is now considered east part of 10 year pits PolyMet 4 inch and PQ core and coarse reject 2005 10.5, 21.5, and 10.7 Three composites from within ten year pits across property PolyMet coarse reject 2006 4.2 and 4.94 One composite from 10 year east pit, one from 20 year pit across property PolyMet ¼ core from 2005 and 2007 Drilling 2007 500  kg One composite,  from east and west pit areas PolyMet ¼ core from 2005 and 2007 Drilling 2008 4.44 One composite,  from east and west pit areas PolyMet ¼ core from 2005 and 2007 Drilling 2008 4.48 One composite,  from east and west pit areas Estimated at 4.5 One PQ drill hole from each tons or less by drill end of property core size There have been numerous bulk samples taken at NorthMet. Samples have been representative by Unit and rock type. Agreement between calculated grades (based on core sampling) and analyzed grades of final sample has been excellent. Earlier bulk samples represented the first ten years of production, more 44 �recent samples used material from across the deposit. Each bulk sample has built upon the previous, and work has progressed to the point where PolyMet has confirmed the ability to make separate, saleable, copper and nickel concentrates. This will allow the company to develop cash flow from sales much earlier in production while completing construction of the hydrometallurgical facility. The planned hydrometallurgical process (PlatSol) was developed on NorthMet ores. The process uses pressure oxidation (225°C, over 30 atmospheres) in the presence of chloride to capture all base and precious metals in the concentrate. Hydromet process recoveries are all over 98%. Other geologic data collected includes: recovery and RQD measurements on all core, over 7,000 specific gravity measurements, over 900 whole rock analyses, over 300 Rare Earth Element packages and a large amount of microprobe data collected for waste characterization purposes. GEOLOGY OF THE NORTHMET DEPOSIT NorthMet consists of seven igneous units that dip southeast, with most economic sulfide mineralization in the top parts of the lowermost unit (Unit 1). The following is a summarized description of the geology of the deposit, based on observations from drill core and limited outcrop mapping. Quaternary Geology In general the Quaternary geology of the region is a thin (0-30 feet or 0-10 meters, but locally thicker) blanket of glacial deposits including till, lacustrine materials, and outwash. Low spots are usually peat bog or open wetland. Topography is subdued and drainage is poor. Site specific geologic studies of the drift have not been done, though a series of geophysical soundings were carried out in 2006 to better define drift thickness outside the area to be mined (Ikola, 2006). Structural Geology The general structure of the NorthMet deposit, as defined by igneous contact dips, foliation in serpentinized zones, bedding trends in the Biwabik Iron Formation (BIF) and in the Virginia Formation, is dominated by an overall dip ranging from 15-25° to the southeast, striking about N56°E. Dips in the seven igneous units are grossly similar, but dips of the mineralized zone are up to 60° in the east pit area. Dips in both the Animikie and the Duluth Complex rocks can be attributed to crustal loading, associated with the input of large volumes of magma originating from the Mid-continent Rift System (Sims and Morey, 1972). Numerous faults have been proposed across the NorthMet Deposit, based largely on reconciling dips in the footwall rocks. There is insufficient evidence, based on drilling to indicate with certainty the exact location of offsets or faulting within the igneous rock units or the footwall rocks on a hole-to-hole basis. Clearly however, pre-intrusion offset or faulting probably exists within the footwall rocks, due to substantial offsets in the BIF (assuming an average 20° dip) as evidenced between drill holes portrayed in cross-sections. Many of these same offsets can be correlated in adjacent cross-sections. Fault zones are apparent in drill core and show up as brecciated intervals (up to several feet thick), including gouge mineralization (clay, calcite, quartz, etc.), slickensides on serpentinized fracture faces, and/or severely broken (rubble) core. Extensive angle drilling in 2005 and 2007 (142 of 170 holes) brought no great clarity to this issue (virtually all previous drilling was vertical). The current geological model and working cross-sections are therefore constructed with minimal faulting influence, especially within the igneous rock units of the Partridge River intrusion. 45 �Logging and Mapping Units A summary of the general stratigraphy of the NorthMet Deposit is outlined below. Rock units and formations are listed in descending order, as would be observed from top to bottom in drill hole. NorthMet units are labeled as Units 1 through 7 (Units I through VII in Severson’s terminology), bottom to top. Unit 3 is probably the oldest, the intrusion sequence of the other units is not clear. The broad picture is of a regular stratigraphy of troctolitic to anorthositic rock units, dipping southeast at 20° to 25°, with basal ultramafic units defining the boundaries of some of these units. The basal ultramafic zones tend to have diffuse tops, sharp bases, and are commonly serpentinized and foliated. Geologists have generally picked the unit boundaries at the base of these ultramafics though there are local exceptions. Economic sulfide mineralization is ubiquitous in the basal igneous unit (Unit 1) and is locally present, but restricted, in the upper units. Unit Definitions and Descriptions Descriptions of the general igneous Stratigraphy for the NorthMet deposit is described below and presented in a stratigraphic column in Figure 3A-3. Unit 7 Unit 7 is the uppermost unit intersected in drill holes at the NorthMet Deposit. It consists predominantly of homogeneous, coarse-grained anorthositic troctolite and troctolitic anorthosite, characterized by a continuous basal ultramafic subunit that averages 20 ft. thick. The ultramafic consists of fine- to mediumgrained melatroctolite to peridotite and minor dunite. The average thickness of Unit 7 is unknown due to erosion removing the upper parts. Unit 7 is generally not mineralized. Unit 6 Very similar to Unit 7, Unit 6 is composed of homogeneous, fine- to coarse-grained, troctolitic anorthosite to troctolite. It averages 400 ft. thick and has a continuous basal ultramafic subunit that averages 15 ft. thick. Overall, sulfide mineralization is minimal, although a number of drill holes in the southwestern portion of the NorthMet Deposit contain significant sulfides and associated elevated PGEs (Geerts 1991, 1994). Sulfides within Unit 6 generally occur as disseminated chalcopyrite/cubanite with minimal pyrrhotite. This mineralized occurrence, the “Magenta Zone”, transitions into Units 3, 4, and 5, and is discussed in greater detail below. Unit 5 Unit 5 exhibits an average thickness of 250 ft. and is composed primarily of homogeneous, equigranulartextured, coarse-grained anorthositic troctolite. Anorthositic troctolite is the predominant rock type, but can locally grade into troctolite and augite troctolite towards the base of the unit. The lower contact of Unit 5 is gradational and lacks any ultramafic subunit; therefore the transition into Unit 4 is a somewhat arbitrary pick. Due to the ambiguity of this contact, thicknesses of both units vary dramatically. However, when Units 5 and 4 are combined, the thickness is fairly consistent deposit-wide. Aside from Magenta Zone mineralization in the west, Unit 5 is not mineralized. Unit 4 Being somewhat more mafic than Unit 5, Unit 4 is characterized by homogeneous, coarse-grained, ophitic augite troctolite with some anorthosite troctolite. Unit 4 averages about 250 ft. thick. At its base, Unit 4 may contain a local thin (usually no more than 6 inch) ultramafic layer or oxide-rich zone. The lower contact with Unit 3 is generally sharp. Unit 4 is rarely mineralized outside the Magenta Zone. 46 �Figure 3A-3. Generalized stratigraphic column for NorthMet units (modified after Geerts, 1994) Unit 3 Unit 3 is used as the major “marker bed” in determining stratigraphic position in the PRI. It is composed of fine- to medium-grained, poikilitic and/or ophitic, troctolitic anorthosite to anorthositic troctolite. Characteristic poikilitic olivine gives the rock an overall mottled appearance. On average Unit 3 is 300 ft. thick. As with Units 4 and 5, the thickness of Units 2 and 3 tend to be highly variable, whereas if combined into one unit, it is more consistent deposit-wide (though not as consistent as Units 4 & 5). Unit 2 Unit 2 is characterized by homogeneous, medium- to coarse-grained troctolite and augite troctolite with a consistent basal ultramafic subunit. The continuity of the basal ultramafic subunit, in addition to the relatively uniform grain size and homogeneity of the troctolite, makes this unit distinguishable from Units 1 and 3. Unit 2 has an average thickness of 100 ft. The ultramafic subunit at the base of Unit 2 is the 47 �lowermost continuous basal ultramafic horizon at the NorthMet Deposit, averages 25 ft. thick, and is composed of melatroctolite to peridotite and minor dunite. Unit 1 Of the seven igneous rock units represented within the NorthMet Deposit, Unit 1 is the only unit that contains significant deposit-wide sulfide mineralization. Sulfides occur primarily as disseminated interstitial grains between a dominant silicate framework and are chalcopyrite > pyrrhotite > cubanite >pentlandite. Unit 1 is also the most complex unit, with internal ultramafic subunits, increasing and decreasing quantities of mineralization, complex textural relations and varying grain sizes, and abundant sedimentary inclusions. It averages 450 ft. thick, but is locally 1,000 feet thick and is characterized lithologically by fine- to coarse-grained heterogeneous rock ranging from anorthositic troctolite (more abundant in the upper half of Unit 1) to augite troctolite with lesser amounts of gabbro-norite and norite (becoming increasingly more abundant towards the basal contact) and numerous sedimentary inclusions. By far the dominant rock type in Unit 1 is medium-grained ophitic augite troctolite, but the textures can vary wildly. Two internal ultramafic subunits occur in drill holes in the southwest, and have an average thickness of 10 ft. Footwall rocks are covered in the Partridge River intrusion description. Inclusions Two broad populations of inclusions occur at NorthMet: hanging wall metabasalts (Keweenawan) and footwall metasedimentary rocks. Basalts are fine-grained, generally gabbroic, with no apparent relation to any mineralization. Footwall inclusions may carry substantial sulfide (pyrrhotite) and often appear to contribute to the local sulfur content. Footwall inclusions are all Virginia Formation, no iron-formation, Pokegama Quartzite, or older granitic rock has been recognized as an inclusion at NorthMet. Sedimentary inclusions make up about 4% of the logged rock types, and basalt inclusions sum to less than 1% of the drilling footage. Other Igneous Units Quadrangle scale outcrop mapping indicates that other igneous stratigraphic units are present above Unit 7. These units are similar to Units 6 and 7 in that they consist of homogeneous-textured troctolitic rocks with basal ultramafic members. There are minor, mineralized, pre-Complex sills in both the Virginia Formation and Biwabik Iron Formation at NorthMet (VirgSill and BIF Sill in footwall descriptions above). In neither case is there any apparent relation to Duluth Complex mineralization. Alteration The vast majority of rock within the NorthMet Deposit would be considered fresh and is unaltered or only weakly altered. Types of alteration most commonly observed in NorthMet rocks are serpentinization / chloritization of olivine, sericitization and saussuritization of plagioclase, and uralitization of pyroxenes. Most alteration is related to close proximity of fractures and/or joints that cross-cut the troctolitic rocks. Likewise, on a microscopic level the center of alteration is focused around microfractures. This pattern suggests that both fracturing and accompanying alteration of the rock occur as a result of the migration of late-stage deuteric fluids during the cooling phase. The vast majority of sulfide mineralization is independent of alteration. Nickel in Silicates (Lab Assay Nickel vs. Recoverable Nickel) It has been characteristic of NorthMet and other Duluth Complex deposits to show lower nickel recoveries in process test work than would be expected from laboratory assays on drill core. Generally there is a loss of about 25-35% of the nickel compared to drill core assays when concentrating sulfides. From previous work, it is known that small amounts of unrecoverable nickel occur as a magnesium-iron- 48 �nickel silicate [(Mg,Fe,Ni) SiO ] that is tied up in the mineral olivine, which is one of three significant gangue minerals that occur across the NorthMet deposit.. 2 4 Figure 3A-4. Geologic map of NorthMet Deposit, all units dip southeast, Magenta Zone is projected upward, does not actually subcrop. 49 �Figure 3A-5. Cross section 35700 at west end of property and 45600 at east end. Purple shading indicated ore zones, bar graphs along holes indicate grades expressed as dollar values, where red = $7.42 cut-off to average grade (~$14.39), and purple shows above average grade, blue are zones of potential lean ore should metals prices rise. 50 �ECONOMIC MINERALIZATION The majority of economic mineralization (copper, nickel, cobalt, platinum, palladium, and gold) at NorthMet occurs in the upper parts of basal Unit 1, with copper and nickel in chalcopyrite, cubanite, and pentlandite, all in the presence of pyrrhotite. Cobalt is contained in sulfides. Platinum, palladium, and gold, while showing good correlation with sulfur and the other metals, are also in a variety of tellurides, bismuthides, and alloys, as well as associated with the major and minor sulfides. Table 3A-3 shows correlation of metals values in drill core data. Table 3A-3. Simple correlation ® table for economic metals and sulfur Cu % Ni % S % Pt ppb Pd ppb Au ppb Pt+Pd+Au Co ppm Zn ppm Cu % 1.000 Ni % 0.860 1.000 S% 0.541 0.572 1.000 Pt ppb 0.568 0.508 0.195 1.000 Pd ppb 0.750 0.635 0.292 0.673 1.000 Au ppb 0.591 0.472 0.250 0.482 0.699 1.000 Pt+Pd+Au 0.760 0.645 0.292 0.778 0.983 0.755 Co ppm 0.544 0.704 0.621 0.217 0.281 0.241 0.288 1.000 Zn ppm -0.021 -0.004 0.286 -0.041 -0.037 -0.017 -0.039 0.093 1.000 1 The simple correlation table above (number of samples=19,516) shows the strong relation of copper, nickel, and palladium, and a somewhat surprising relation of cobalt to sulfur. Zinc’s low factor is probably related to its multiple origins as either magmatic or derived from assimilation of footwall rock, hence representing two populations of data. The sulfur vs. metal correlation is probably greatly affected by iron, the presence of which is not shown here, but is in excess in all rocks. Grades are highest at the top of Unit 1 and fade going down hole. Grades appear to be higher down-dip though this may be an artifact of less dense sampling. There is a smaller zone of economic mineralization (about 50 million tons) at the western end of the property in the upper units, known as the “Magenta Zone.” This zone is generally copper and PGE-rich (sulfur-poor relative to metals) and of “average” reserve grade. The minerals of interest from a waste characterization perspective are the same as above, but pyrrhotite is expected to be the main mineral affecting water quality in regards to waste rock, though the traces of chalcopyrite, cubanite and pentlandite are studied for waste rock storage. Trace pyrite and pyrrhotite are the main sulfide minerals found in the tailings. Most sulfide mineralization at NorthMet is thought to be of a distant source (magmatic?), some is locally modified by sulfur derived from footwall metasedimentary rocks (Virginia Formation). Minor veins and other cross-cutting relations indicate some movement of sulfides within the deposit, but there is no evidence recognized for large scale relocation of sulfides, nor any macroscopic evidence for any hydrothermal event that may have remobilized PGE’s or sulfides. Element distributions, on a single section through the west pit in figure 5A-6 are, located on a single page at the end of the PolyMet Section. The Magenta Zone mineralization cutting across upper intrusive units, is illustrated in this section. 51 �RESOURCE The PolyMet resource and reserve (Table 3A-4) models have been done in cooperation with several consultants, most recently PEG Mining of Toronto. PolyMet supplies the geologic solids model, database, and block model geometry. Geostatistics and population of the block model, and hence the resource estimate, are done in consultation, with finalized resource block models then sent forward to engineers for reserve calculation and mine planning. Table 3A-4. The NorthMet resource and reserve values work was done by Wardrop Engineering 2007. Cut-off based on “Net Metals Value” per ton, accounting for grade, average flotation and Hydromet recovery, realization costs, metal prices, and other factors. See Desaultels and Patelke, 2008 for resource calculation details. Enough reserve has been shown for 20 years of permit constrained production. RESERVES-2007 Proven Probable Proven and Probable Cut-off value $7.42 $7.42 $7.42 Million Tons 118.1 156.5 274.6 Cu % 0.30 0.27 0.28 Cut-off value $7.42 $7.42 $7.42 $7.42 Million Tons 202.5 491.7 694.2 229.7 Cu % 0.285 0.256 0.265 0.273 Ni % 0.09 0.08 0.08 Co Ppm 75 72 73 Pt ppb 75 75 75 Pd ppb 275 248 260 Au ppb 38 37 37 RESOURCES-2007 Measured Indicated Measured & Indicated Inferred Ni % 0.083 0.075 0.077 0.079 Co ppm 74 70 71 56 Pt ppb 71 66 68 73 Pd ppb 258 231 239 263 Au ppb 36 34 35 37 ASSUMPTIONS Metal and Units Assumed Metal Price Average % recovery, as used in DFS Cu % $1.25 lb. 92.33 Ni % $5.60 lb. 70.34 Co Ppm $15.25 lb. 40.75 Pt ppb $800 oz. 75.74 Pd ppb $210 oz. 72.69 Au ppb $400 oz. 67.04 In the center of the deposit the highest, near surface, Unit 1 grades transition into the middle of the unit, while in the east, mineralization is strong and vertically persistent throughout the unit. The top of the merged Unit 1 and Unit 2 mineralized domain (domain 1) forms a hard boundary that, combined with the bedrock ledge (depth to bedrock) surface, forms the bottom and top estimation boundaries for the upper units (exclusive of the “Magenta Zone”, which is internal to this domain). There is no conclusive relation between specific Unit 1 specific rock type and presence or grade of mineralization except that noritic rocks are generally of lower grade. Units 2 and 3: These units are treated as one unit in the geologic model, with PolyMet geologists considering them as a single package grading from an ultramafic base to an anorthositic top for modelling purposes. The thickness of the package stays relatively constant, though the thickness of the two individual units varies, primarily due to Unit 2 locally thinning. While generally barren, Unit 2 has mineralization at its base in the western half of the deposit. These zones may not be strictly equivalent to Unit 1 type mineralization. Copper and nickel values are lower, as is pyrrhotite, but behavior of other metals is inconsistent, with PGE (Pt + Pd +Au) content varying locally relative to nearby grades at the top of Unit 1. Above the basal zone of Unit 2 it is usually barren, medium-grained, and homogenous in texture. Average PGE in Unit 2 is slightly above that of Unit 1. 52 �Unit 3 shows mineralization in the west, in the middle of the unit and near the top. This occurrence is merged into the Magenta Zone. Units 4 and 5 are also modeled as a geologic package. There is no compelling geologic reason to fully separate these units, the boundary between them being an arbitrary pick based on overall changes in texture from homogenous to heterogeneous, grain size, and plagioclase content, but without a welldefined bounding horizon. The top boundary of Unit 5 is the basal ultramafic of Unit 6, which is an unused hard boundary in grade modelling. The bottom boundary of Unit 4 is a discontinuous ultramafic horizon. There are also discontinuous oxide-rich zones along the contact between Units 3 and 4. Metals and sulfur grades in Unit 4 are proportional to Unit 1, but consistently lower. Unit 4 has few high copper or sulfur assay intervals. There is some near surface mineralization, modelled as a part of the Magenta Zone, described below. Otherwise there is only low grade, discontinuous material at the base. Unit 6 and Unit 7: These units are very similar in nature. Both are homogenous anorthositic troctolite with well-defined ultramafic bases. No top for Unit 7 has been seen in drill hole. Units 3, 4, 5 and 6 host a zone of mineralization, modeled as the Magenta Zone. Unit 6 mineralization was described by Geerts (1994) as the “Magenta Horizon” when originally found in six drill holes. Further drilling has extended these copper rich, sulfur poor zones (of moderate overall grade) into more than fifty drill holes in Units 3, 4, 5, and 6. The zone transitions across the ultramafic base of Unit 6 and into Units 3, 4 and 5, (i.e., does cross the igneous stratigraphy) which is problematic if the emplacement model of these units representing individual pulses of magma is correct. There is no gross evidence for this mineralization being hydrothermal, which could cross boundaries, but would presumably alter large masses of rock. Unit 7 has a few good assay intercepts, but no apparent continuity for sulfides. Copper, nickel, and sulfur values in Table 3A-5 are calculated after removing samples with less than 0.05% copper. 53 �Table 3A-5. Average values for assays by unit after removal of the less than 0.05% copper intervals (drill core samples). Unsampled zones not accounted for here. Data complete through 2006. Cu Ni % % S Co Cu+Ni Total % of unit Pt+Pd+Au ppb Cu/Ni Cu/S % ppm % sampled Average sample length-feet Unit 1 0.3 0.09 0.83 349 76 0.39 3.35 0.43 90 5.3 Unit 2 0.2 0.07 0.39 365 73 0.27 2.74 0.61 80 5.6 Unit 3 0.19 0.05 0.5 286 62 0.25 3.19 0.53 71 7.2 Unit 4 0.21 0.06 0.58 269 66 0.28 3.40 0.44 51 7.6 Unit 5 0.27 0.07 0.54 398 65 0.35 3.64 0.54 41 7.8 Unit 6 0.33 0.08 0.48 532 69 0.41 3.74 0.69 27 7.2 Unit 7 0.2 0.06 0.32 330 83 0.26 3.60 0.72 11 8.4 • Gatehouse (North Mining) did report some geochemical cyclicity in unit 1, but this has not been revisited with the larger data set; • Poor assay grades in the noritic rocks are related to footwall assimilation and contamination, otherwise there is little connection between grades and specific rock type. About 83% of the igneous rocks at NorthMet are troctolites, 6% anorthositic rocks, 4% ultramafic rocks, and 4% footwall inclusions. The remainder are norite, gabbro and others; • Within Unit 1 copper:sulfur ratio tends to be highest at top, then diminishes with depth, following the pattern of PGE’s; • The upper units have higher copper:sulfur ratios than Unit 1 (i.e., more chalcopyrite rich), but lower overall copper values; • Ratio of PGE to copper is lowest in Unit 1, but Unit 1 has greatest quantities of both; • Chalcopyrite is the dominant sulfide in the upper units regardless of total sulfur content; Sulfide (Ore) Mineral Proportions Various metallurgical test programs have been conducted on NorthMet ores since the 1970's. Reported sulfide mineral proportions have not been entirely consistent between these tests. Sulfide mineralogy within the NorthMet Deposit has been described in detail through petrographic observations and microprobe analysis. Approximately 95-98% of all sulfide mineralization consists of 4 predominant species, in decreasing order of abundance: chalcopyrite (cp) > pyrrhotite (po) > cubanite (cb) > pentlandite (pn). In general, Po:Cp+Cb ratios increase towards the basal contact or in proximity to sedimentary inclusions. Likewise, Cp:Cb ratios increase with increased distance away from the footwall rocks. In core logging and other work, chalcopyrite is often not distinguished from cubanite. Mining Mining at NorthMet will begin with contractor clearing and overburden stripping of the pit and stockpile areas. Engineered stockpile bases and liner systems must be in place before mining begins, as does the overall water collection system for treatment and pumping to the tailings basin. Ore and waste production will start in the east pit, with production from the west pit ramping up soon afterwards. Up through about year 11 or 12 production from both pits will be equal until the east pit is mined out. At that point, backfilling of the east pit will begin, with the ultimate goal of constructing a wetland in that pit. The central pit area will be mined last. Ore will be moved at a rate of 32,000 tons per day. Waste to ore strip ratio will be about 1.46:1. Ore will be moved by truck to the “superpocket” and loaded to 100 ton capacity side dump rail cars by pan feeder. There will be twenty trains per day of 16 cars each. 54 �Ore and waste categorization (“ore control”) will be by assay of core and / or blast holes and careful pit mapping. Waste material will be sorted to stockpiles, and stockpile liners will be built, according to the sulfur and metals content of the waste rock. Fig. 3A-6. Economic element distribution, sulphur and NSR in the west pit. Section is orthoganal to the NE-SW strike of the Virginia Fm - Duluth Complex contact. 55 �PART 3B: TECK AMERICAN MESABA DEPOSIT By: Mark Severson (portions originally by Tim Jefferson) BACKGROUND Previous exploration and development work The Mesaba deposit was first discovered along the base of the Duluth Complex in 1958 by Bear Creek Mining Company (BMC). Between 1958 and 1960, BMC completed 55 shallow drill holes for 43,000 feet (13,952 meters). BMC renewed drilling activities in 1967-1971 completing 149 additional holes. Drill hole B1-105 intersected substantial amounts of semi massive to massive sulfide mineralization between 1,400 and 1,800 feet (425 and 550 meters) below surface in footwall rock. Subsequent drilling defined a high grade zone appropriately named the Local Boy ore zone after BCM geologist Stuart Behling, the “local boy,” who encouraged BMC to continue drilling this site. In late 1973, AMAX Exploration, Inc. agreed to take over BMC’s state and private leases. During the next four years (1974-1978) AMAX continued drilling the deposit (completing 228 drill holes). In particular their focus was drawn to the Local Boy ore zone (Watowich and others, 1981), and following successful permitting, they sank a shaft in 1976-1977. Four drifts totaling 3,800 feet (1,160 meters) were developed and 218 underground holes were completed. Based on this work, AMAX reported an overall underground estimate of 364 million tons (330.2 million tonnes) averaging 0.84% Cu and 0.19% Ni, with a Local Boy-only resource of 5 million tons (4.54 million tonnes) grading 1.89% Cu, 0.36% Ni (Watowich, 1978). Both underground resources were estimated based on a 0.60% Cu cut-off (Watowich, 1978). Due to weakening copper and nickel markets and the inability to produce separate high grade Cu and Ni concentrates, AMAX abandoned their plans to develop the deposit in late 1981. Rhude and Fryberger obtained leases and, along with the NRRI, evaluated the PGE potential of the Local Boy ore zone circa 1990. Arimetco Inc., picked up the Babbitt deposit leases, renamed it the Mesaba deposit, and evaluated the property circa 1994-1996. They did not complete any drilling but collected two bulk samples for metallurgical test work. Arimetco reported a resource estimate to 3,300 million tons (2,993.7 million tonnes) grading 0.46% Cu, 0.12% Ni, cut-off 0.38% Cu (Miller et al, 2002). Arimetco Inc. declared bankruptcy in late 1996. Present exploration and development work Teck American Incorporated acquired a package of state and private leases covering the Mesaba deposit in 1998. Teck drilling began in 2007-2008 for a total of 67,430 feet (20,560 meters) in 64 drill holes (Fig. 3B-1). This drilling was concentrated on the western portion of the deposit to complete a 400 foot (120 meter) grid infill program. In 2012-2013, Teck conducted three additional drilling campaigns (Fig. 3B-1). 56 �Figure 3B-1. Drill hole location map for Mesaba Deposit. Grid north is about 33° west of north. Re-logging of historic holes at the Mesaba deposit over an 18 year period, in addition to information gained from logging of holes completed in 2007-2008 and 2012-2013, indicates that the deposit is primarily hosted by a previously unrecognized intrusion within the Duluth Complex (Severson and Hauck, 2008). It is believed that this intrusion, informally named the Bathtub intrusion (BTI) lies between the Partridge River intrusion (PRI) to the south, and the South Kawishiwi intrusion (SKI) to the north and east. This intrusion is believed to have been fed by a vent in the Grano Fault area on the east side of the Mesaba deposit. The BTI is believed to pre-date the SKI and to be coeval to slightly older than the PRI. It is further believed, based on drill hole evidence, that the upper igneous units of the PRI overlap specific BTI units. Supporting evidence for this new interpretation is based on igneous units that are unique to either the PRI or BTI, and different styles of sulfide mineralization between the two. The following geologic discussion is largely based on the work of Severson and Hauck (2008) but is condensed and summarized. 57 �Figure 3B-2. Preliminary geologic map of the Mesaba deposit showing major geologic units of the Bathtub, Partridge River, and South Kawishiwi intrusions. Major structural features associated with the deposit are also shown. 58 �GEOLOGIC SETTING Footwall Rocks As there is great commonality between the footwall rocks at the NorthMet Deposit and the Mesaba Deposit, they are discussed in the regional geology section. Structure As discussed earlier in this guidebook, there are three structural features that are pertinent to understanding the intrusive history of the BTI that include (Fig. 3B-2): 1. an east-west trending paired syncline and anticline in the footwall rocks referred to as the Bathtub Syncline and Local Boy Anticline; 2. a zone referred to as the “Hidden Rise” that separates the PRI and BTI; and 3. a north-trending fault zone, referred to as the Grano Fault, that has been postulated to have been the feeder zone for the BTI and footwall-injected massive sulfides of the Local Boy ore zone. The “Hidden Rise,” as discussed earlier, is a loosely-defined zone wherein scattered hornfels inclusions of footwall Virginia Formation are fairly common. When viewed collectively, the “Hidden Rise” defines an east-west trending “ridge” that is roughly positioned at the contact between the PRI and BTI. Along the far eastern edge of the Mesaba deposit is the north-trending Grano Fault, so named for the abundant and sometimes voluminous amounts of late granitoid and oxide rich pyroxenitic lenses (OUIs) associated with the fault zone (Severson, 1994). The late intrusive lenses are interpreted to have vertical configurations and were injected along subsidiary fault zones parallel to, and immediately west of, the Grano Fault. The late OUI and granitoid bodies cut the troctolitic rocks and thus demonstrate that the fault was active during and after emplacement of the BTI and PRI. BATHTUB INTRUSION The newly named Bathtub intrusion (BTI) is wholly contained in the central portion of the Mesaba (Babbitt) deposit. The BTI has recently been singled out as a separate intrusion to explain the abrupt change from typical Partridge River intrusion (PRI) stratigraphy, in the southern part of the deposit, to a completely different stratigraphy to the north in the remainder of the deposit. The BTI has been divided into two major units, BT1 and BT4, each of which contain several subunits. These units, in addition to footwall rocks, and structural features, are portrayed in Fig. 3-5. BT1 Unit The lowermost unit of the BTI is referred to as the BT1 Unit. It is very similar to Unit I of the nearby PRI in that it is heterogeneous-textured at all scales, contains abundant hornfels inclusions near the basal contact, and is the main sulfide-bearing unit at Mesaba. However, there are some important differences between Units I and BT1 that include: • Massive sulfide occurrences are more common near the basal contact in the BT1 than in Unit I (excluding the unique Local Boy ore zone) indicating that sulfide settling may have been a more important mineralization mechanism in the BTI; • Coarse- to very coarse-grained disseminated sulfides (up to several centimeters across) are exceedingly common in the lowermost portions of BT1; whereas, this same relationship is not so obvious in Unit I – this again implies the importance of a sulfide settling origin, and; • Ultramafic horizons and patches are very common in portions of the BT1; whereas, similar ultramafic horizons are not as common in Unit I of the PRI. 59 �The BT1 Unit has been further subdivided into several internal subunits that are discussed below. BT1-a This subunit of the BT1 is a heterogeneous-textured augite troctolite grading to olivine gabbro. The BT1a subunit is more common in the bottom half of the BT1 Unit and increases up dip (to the north) at the expense of most other subunits of the BT1. BT1-c At the base of the BT1 there is significant silica contamination of the magma, due to assimilation of the footwall rocks, and noritic rocks (norite to gabbro norite), with common hornfels inclusions, are the dominant rock types. Graphite occurrences are also commonly found in various rock types of BT1-c. The BT1-c subunit spatially occurs as a rind or coating along the basal contact. BT1-uz Wherever olivine-rich ultramafic rocks are common over appreciable intervals in the BT1 Unit, this subunit is used to designate ultramafic zones. The ultramafic rocks in these zones range from welldefined layers to zones where irregular ultramafic patches are presumably peppered throughout a troctolitic host rock. BT1-at This subunit of the BT1 is used to denote areas where anorthositic troctolite is the dominant rock type. BT-sli A few holes in the western end of the BTI exhibit well-defined modally-bedded rocks consisting of alternating troctolitic and ultramafic rocks. These intervals are designated as BT-sli for the Bathtub Side Layered Interval. The BT-sli subunit occurs about in the center of BT1 unit in close proximity to the “Hidden Rise.” BT4 Unit The uppermost unit of the BTI is referred to as the BT4 Unit. It was originally correlated with Unit IV of the PRI. However, the BT4 Unit is distinctly different from Unit IV in that the BT4 Unit at Mesaba is: • Often more heterogeneous-textured at all scales and composed of many alternating rock types; • commonly contains local sulfide-bearing zones; whereas, Unit IV is mostly sulfide-barren – the sulfides in BT4 are generally chalcopyrite-rich in comparison to chalcopyrite/cubanite ores in the underlying BT1 Unit; • floored by a semi-persistent ultramafic layer termed the "± Picrite" (see discussion below) in the central portion of the Bathtub ore zone; and • ultramafic layers and modally-bedded zones, termed the Bathtub Layered Interval (BTLI), are common in the central portion of the Bathtub ore zone. The BT4 Unit has been further subdivided into several internal subunits based on the presence of a dominant rock type. The various subdivisions of the BT4 Unit are briefly discussed below. BT4-a This subunit of the BT4 on the cross-sections denotes areas where heterogeneous-textured augite troctolite is the dominant rock type. BT4-at 60 �This subunit of the BT4 is used to denote areas where anorthositic troctolite is the dominant rock type. "± Picrite" At the base of BT4 is a semi-persistent olivine-enriched ultramafic horizon referred to as the "± Picrite." It is present in about 70% of the drill holes in the BTI-portion of the Mesaba deposit. The "± Picrite" is generally absent in the up dip direction (to the north) and is variably present to the south in the contact zone between the PRI and BTI. Where present, the "± Picrite" is about 1-15 feet thick, but exceptions are locally present. In some areas, the "± Picrite" consists of several stacked ultramafic horizons, or modal beds, that are interlayered with troctolitic rocks, and thus, the zone represents a collection of several cyclic layers. In other areas of the Mesaba deposit, the "± Picrite" is not always easily singled out as it occurs in close proximity to a downward thickening BTLI with similar ultramafic layers and modal beds. Therefore, in some instances it is difficult to pick the "± Picrite" out of a myriad of ultramafic horizons associated with either the BTLI or BT-sli. Bathtub Layered Interval (BTLI) In the vicinity of the Bathtub Syncline and the “Hidden Rise,” ultramafic layers and modal-bedded zones are extremely common within the BT4 Unit. In the eastern half of the Mesaba deposit the BTLI appears to be present in a subhorizontal saucer-shaped morphology. Conversely, in the western half of the deposit, the BTLI is confined to one or two cylinder-shaped zones, albeit with irregular edges, that is positioned in close proximity to the “Hidden Rise.” Overall, the ultramafic rock types of the BTLI are characterized by alternating assemblages of either/or: melatroctolite (picrite), feldspathic peridotite, peridotite, dunite (minor), olivine-rich troctolite, and troctolite with modal beds of olivine-rich layers. One or more of these rock types may be stacked above the other in no particular order, and the thickness of this assortment may be highly variable between drill holes. The number of individual ultramafic layers present within the BTLI for any particular drill hole varies drastically. The range in thickness for each of the individual ultramafic beds also shows considerable variation, ranging from a few inches to over tens of feet thick. Although the BTLI can be correlated as a package of alternating troctolitic and ultramafic layers, each of the individual ultramafic layers cannot be correlated on a hole by hole basis. This situation indicates that the ultramafic layers either: 1) bifurcate/divide into many thin ultramafic layers; 2) pinch out or have very limited spatial extent; 3) some may actually represent dike-like features (filter pressed crescumulates?); or 4) combinations of the above. Further complicating the picture, the inclination of contacts and modal bedding associated with the ultramafic layers are highly variable, ranging from 5°-80° (with localized overturned beds). This variation in inclinations can even be present in a single drill hole. For the most part, the bedding and contact inclinations in the BTLI are steeper higher up in the drill hole and gradually shallow with depth. The shallow to steep angles exhibited by the BTLI may reflect that the ultramafic layers originated via a variety of mechanisms that include: 1) crystal settling to form subhorizontal layers (dominant in the eastern half of the deposit); 2) filter-pressing to form localized dike-like morphologies; 3) slumpage and folding of the beds took place before they were fully crystallized to form highly irregular and overturned beds; 4) compaction differences took place during lithostatic loading of the crystal pile to form steep and irregular beds; 5) cooling and crystallization, or size/density sorting, took place along, and parallel to, the southern wall of the BTI (up against “The Hidden Rise”); or 6) combinations of all of these mechanisms. Whatever their origin, the steep beds displayed by the BTLI in the western half of the Mesaba deposit are inordinately associated with the “Hidden Rise” and the southernmost edge of the BTI. PARTRIDGE RIVER INTRUSIVE (PRI) AT MESABA Many of the igneous rock units that are present at the nearby NorthMet deposit are also present along the southern edge of the Mesaba deposit and are believed to represent units of the Partridge River Intrusion (see previous geologic setting discussion). Additionally, Units IV through VI of the PRI appear to extend 61 �northward and overlie the heterogeneous-textured BT4 Unit. This relationship, also depicted in Figure 3B-2, suggests that the BTI was eventually over-ridden/overlain by the upper units of the PRI. The overall timing of emplacement for the lower PRI units versus the BTI is unknown but correlations in cross-sections crudely suggest the following: • Units I through III were intruded first along the southern edge of the Mesaba deposit with a vent area located somewhere to the southwest. The “Hidden Rise” generally marks the northern extent of this intrusive activity and originally formed as part of the floor to these units. Unit III may have been intruded as thin lenses across and north of the “Hidden Rise” – this may explain the local presence of Unit III-like inclusions in the BTI. • Concurrent with or after the above activity, the BT1 Unit was intruded from a vent area located somewhere to the east, possibly from the Grano Fault area. The “Hidden Rise” formed the southern wall of this particular magma chamber. • The BT4 Unit was intruded into the same magma chamber but was emplaced above the BT1 Unit. • Concurrent with or after the above activity, Units IV through VII+ of the PRI were intruded from a vent area located somewhere to the southeast. These upper units were emplaced over the BT4 Unit. MINERALIZATION AT MESABA The Mesaba deposit is characterized by disseminated sulfide mineralization, which occurs most commonly as accumulations of chalcopyrite, cubanite, pyrrhotite, and pentlandite. Additionally, common occurrences of talnakhite have been noted in close proximity to the “Hidden Rise.” Short intercepts of semi-massive to massive sulfide mineralization are locally encountered in the BT1-c. Sulfur isotope analyses have indicated that the source of the sulfur used in the formation of the sulfides at Mesaba is the Virginia Formation (Ripley, 1986). The model of sulfide deposition entails turbulent injection of units of the BTI wherein immiscible sulfide droplets coalesce within the silicate melts and attract the chalcophile elements (chiefly copper and nickel) through magma mixing. Thus, the most contaminated magma (from assimilation of footwall Virginia formation) hosts basal sulfides that contain excess sulfur and iron relative to intrusive units higher above the footwall. The sulfide content of the rock increases, often dramatically, as the footwall is approached. This sulfide content increase with depth is accompanied by the increasing presence of pyrrhotite and a subsequent change in the copper bearing sulfides (cubanite is dominant over chalcopyrite with depth). The disseminated mineralization is generally composed of 1-4% sulfides, but can reach upwards of 8-12 % sulfides as the footwall is approached. The most important mineralized zone at Mesaba is the basal zone, starting at the footwall Virginia Formation contact, and ranges between 200 and 400 feet thick (60 and 125 meters thick). Higher up in the intrusive package, often overlapping the BT1-BT4 unit boundary, is a second zone of disseminated sulfide mineralization that is more erratic and discontinuous in nature but contains markedly lower amounts of cubanite and pyrrhotite. The mineralized footprint of the Mesaba deposit is oblong to arcuate in shape, 3,000 by 13,000 feet (925 by 4,000 meters) in approximate dimensions, crops out at the surface on the northern/up-dip side and extends to approximately 1,650 feet (500 meters) below surface in the southern/downdip direction. The strongest basal mineralization is often localized within the Bathtub Syncline. Here, concentration of sulfides by gravitational settling into the footwall depression has likely occurred. 62 �PGE Mineralization at Mesaba Platinum group element (PGE) mineralization in the BTI at Mesaba occurs to a lesser degree than the other deposits and intrusions. Generally, the highest PGE values at Mesaba are associated with Cu-rich massive to semi-massive sulfides in the Local Boy ore zone. Analyses from sampled intervals (5-15 feet thick) record values as high as 11.1 ppm Pd, 8.3 ppm Pt, 13.1 ppm Au, and 62 ppm Ag in the sulfide-rich ores (Severson and Barnes, 1991; Hauck and Severson, 2000). The majority of the anomalous PGE values are spatially distributed along the axis of the Local Boy anticline with the highest Cu and PGE values occurring in the west half of Local Boy. The Grano Fault may have served as a feeder zone to the massive sulfides that were injected into the footwall rocks along the Local Boy Anticline as an immiscible sulfide melt. This melt fractionally crystallized in an east-to-west direction and progressively became enriched in PGE towards the west (see discussion below). MASSIVE SULFIDES AT THE LOCAL BOY ORE ZONE OF THE MESABA DEPOSIT Cu-rich massive sulfides near the basal contact of the Complex are locally present at the Mesaba deposit in a small zone referred to as the Local Boy ore zone. In 1976, AMAX Inc. completed a 1,700-foot-deep exploratory shaft (Minnamax shaft), and in 1977, completed four drifts (A, B, C, and D; Figures 3B-3 and 3B-4). Underground Fan drilling (217 holes) was completed in 1978 to further define the massive sulfide distribution. Sulfide minerals include pyrrhotite, pentlandite, chalcopyrite, talnakhite, cubanite, maucherite (nickel arsenide), sphalerite, bornite and late mackinawite, chalcocite, covellite, godlevskite, and native silver (Severson and Barnes, 1991). Footwall Structures in the Local Boy Ore Zone Several investigators have recognized that pre-existing structural conditions in the footwall rocks strongly influenced the basal contact of the Duluth Complex (Mancuso and Dolence, 1970; Watowich, 1978; Holst et al., 1986; Martineau, 1989; Severson and Barnes, 1991). Major irregularities in the basal contact are generally related to folds in the underlying country rock indicating that intrusion proceeded more or less along bedding planes in the footwall rocks (Holst et al., 1986). This is readily expressed by a major eastwest -trending trough and ridge in the basal contact at Mesaba that coincides exactly with a synclineanticline that is defined by the top of the Biwabik Iron Formation (BIF). The thickness of preserved Virginia Formation between the Complex and the BIF is variable due to the amount of material assimilated by the Complex. The Local Boy ore zone is also situated over this anticlinal ridge. The majority of massive sulfide ore zones, hosted mainly by the Virginia Formation (Severson and Barnes, 1991), are broadly coincident with the axis of the anticline. The contoured top of the BIF in the Local Boy area is shown in Figure 3B-3 (left). Similar anticline geometries are also present for the basal contact as shown in Figure 3B-3 (right). All the data indicate that an EW-trending anticline is the major structural feature present within the footwall rocks of the Local Boy area. 63 �Figure 3B-3. Contoured top of the Biwabik Iron Formation at Local Boy (left), and the contoured top of the basal contact between the footwall Virginia Formation and the Partridge River intrusion at Local Boy (right). Diagrams from Severson and Hauck, 2008. Mineralization Trends in the Massive Sulfide at the Local Boy Ore Zone The vast majority of massive sulfides at Local Boy are contained within the Paleoproterozoic Virginia Formation. Even though the massive sulfides straddle the basal contact, most of the massive sulfides are associated with either hornfelsed sedimentary inclusions above the contact or with footwall rocks below the contact while the interfingering intrusive rocks are relatively barren of massive sulfides (Severson and Barnes, 1991). This suggests that the massive sulfide ores were not formed in this area by the gravitational settling of sulfides, but rather, the ores formed by injection of an immiscible sulfide melt into structurally prepared areas within the footwall rocks along the Local Boy anticline in a vein-like setting. A similar mechanism is proposed for the Norilsk-Talnakh deposits in Russia. Even though the basal contact of the Complex with the Virginia Formation is highly undulatory, the massive sulfides exhibit a definite top and bottom. The ore is distributed such that most of it is contained within a zone between 20 feet and 300 feet above the top of the Biwabik Iron Formation. The geologic constraint for the bottom of the ore zone generally corresponds to the top of the VirgSill (a structurally competent unit). The constraints for the upper portion of the ore zone are unknown and may have been obliterated during emplacement of the Complex. Figure 5B-7 is an attempt to show, in a plan view, where massive sulfide zones are present. Also shown in the figure are the different massive sulfide types (ranging from pyrrhotite-dominant to Cu-rich) relative to structural features. The relationships shown in Figure 3B-4 indicate that the massive sulfides show a progressive change in an east-to-west direction from Cu-poor massive sulfides to Cu-rich massive sulfides in the vicinity of the Local Boy anticline. These relationships suggest that the injected immiscible sulfide melt underwent fractional crystallization 64 �and progressively became more Cu and PGE enriched as it moved through the footwall rocks in an eastto-west direction. Figure 3B-4. Potential distribution of semi-massive to massive sulfide types (Cu-poor versus Cu-rich) at the Local Boy ore zone (left); and an isopach map of the cumulative thickness of the massive sulfide zones at the Local Boy ore zone (right). Note that the massive sulfides are not present as a continuous blanket, but rather, as one or more stacked disjointed/separated multiple horizons near the basal contact. Diagrams from Severson and Hauck, 2008. A possible feeder vent for the sulfide injection event may have been the Grano Fault, which was repeatedly reactivated during emplacement of the Complex. Other data that indicates that the Grano Fault was a potential feeder vent include: 1) the massive sulfides are more common, and thicker (Figure 3B-4), close to the Grano Fault (feeder) and along the axis of the Local Boy anticline (structurally-prepared site); 2) the VirgSill rarely contains significant amounts of disseminated sulfides – except near the Grano Fault; and 3) the Biwabik Iron Formation rarely contains sulfides – except near the Grano Fault. In summary, the massive sulfides at the Local Boy ore zone are interpreted to be structurally controlled in that they are situated along the axis of the Local Boy anticline. The massive sulfides are locally Cu-rich (5-25% Cu; Severson and Hauck (2008) – based on historical assay data on file at MDNR) and are almost exclusively hosted by the Virginia Formation. Sulfide textures suggest that the massive sulfides were injected as an immiscible sulfide melt into the footwall rocks. The overall pattern of sulfide types and PGE contents suggest that the sulfides formed via a process of fractional crystallization of an immiscible sulfide melt as it migrated into the footwall rocks. The Grano Fault is inferred to represent the potential feeder zone in this scenario. 65 �PART 3C: TWIN METALS MINNESOTA’S MATURI DEPOSIT Kevin D. Boerst – Twin Metals Minnesota (Portions previously written by Dean M. Peterson – formerly with Duluth Metals Limited) INTRODUCTION Twin Metals Minnesota’s Maturi deposit is the largest and highest-grade classified Cu-Ni-PGE deposit in the 1.1 Ga. Duluth Complex of northeastern Minnesota. The deposit is located near the north end of the South Kawishiwi intrusion (SKI) west-southwest of the junction of the Nickel Lake Macrodike (NLM) and the SKI (Peterson et al., 2006; Peterson and Albers, 2007; Tharlason et al., 2007; Peterson, 2008; Gal, 2008; and White, 2010). The deposit was discovered utilizing a genetic ore deposit model that identified channelized magma flow within the SKI under a large xenolith/pillar of anorthosite. The model led to exploratory drilling in 2006; deposit discovery and resource estimations in 2007 & 2008; a joint venture with Antofagasta plc in 2010; and significant resource expansion and classification upgrades in 2012 and 2014 (Fig. 3C-1). The company is currently optimizing its business case with the aim of completing a Mine Plan of Operation (MPO) in the future. RESOURCES The Mineral Resource estimate for the Maturi deposit (completed by AMEC) incorporate assay data from 564 drill holes totaling 1,466,641 feet (excluding wedges) drilled on the Maturi deposit that includes 75 legacy holes also in the geologic data base. The April 2014 Resource Estimates for the Maturi (as well as the Birch Lake and Spruce Road deposits) is based on a 0.3% copper cut-off grade to define the resource model. Based on AMEC`s review of metal prices, process recoveries, refining costs and underground mine operating costs likely to apply at the Twin Metals site, the 0.3% copper cut-off grade (highlighted) is considered the base case for the statement of Indicated and Inferred Mineral Resources at this time. The estimates at the cut-off grades higher and lower than the base case are provided to show sensitivity of the cut-off grade (Table 3C-1). Inferred Resources Indicated Resources Measured Resources Table 3C-1. April 2014 resource estimate for TMM’s Maturi Deposit. Cu% Million Cu Ni Pt Cut-off Tons % % ppm 0.2 327 0.61 0.20 0.141 0.3 308 0.63 0.20 0.146 0.4 275 0.66 0.21 0.5 237 0.70 0.22 0.6 183 0.74 0.2 881 0.3 0.4 Au Ag Co ppm ppm ppm ppm 0.328 0.080 2.2 105 0.339 0.083 2.3 107 0.155 0.359 0.088 2.4 110 0.165 0.383 0.093 2.5 113 0.24 0.177 0.411 0.100 2.7 116 0.56 0.18 0.148 0.336 0.080 2.0 102 822 0.58 0.19 0.155 0.350 0.083 2.1 104 716 0.61 0.20 0.166 0.375 0.089 2.2 106 0.5 547 0.66 0.21 0.186 0.420 0.099 2.4 109 0.6 379 0.71 0.22 0.205 0.461 0.108 2.6 111 0.2 768 0.42 0.13 0.116 0.262 0.059 1.6 81 0.3 531 0.49 0.16 0.138 0.314 0.070 1.8 98 0.4 358 0.57 0.19 0.167 0.376 0.083 2.0 110 0.5 235 0.63 0.20 0.202 0.449 0.099 2.3 112 0.6 127 0.69 0.21 0.246 0.545 0.118 2.5 111 66 Pd �Figure 3C-1. Map of Twin Metals Minnesota’s deposits and resource classification. 67 �GEOLOGY OF THE MATURI DEPOSIT The Maturi Deposit is located within the South Kawishiwi Intrusion (SKI), a shallow dipping (~24º eastsoutheast) sill-like troctolitic intrusion exposed in an 8- x 32-kilometer arcuate band along the northwestern margin of the Duluth Complex. Lithologic units within the Maturi deposit include Mesoproterozoic rocks of the SKI and Anorthositic Series of the Duluth Complex as well as basalt xenoliths of the North Shore Volcanic Group. At Maturi, SKI magmas intruded between hanging wall anorthositic rocks and footwall granitic rocks of the Neoarchean Giants Range batholith (Fig. 5C1). Brief descriptions of the lithostratigraphic units within the Maturi Deposit are given in Table 3C-2. Table 3C-2. Lithostratigraphic units within the Maturi deposit. Duluth Complex and related rocks (1.1 Ga.) Anorthositic troctolite to troctolite (ATA Series) - Medium to coarse-grained, homogeneous, wellfoliated and locally layered anorthositic troctolite, troctolite, and ophitic troctolitic rocks. In the field, this unit is commonly referred to as the “sea of troctolite”. SKI Augite-bearing troctolite (Main AGT) - Homogenous, coarse-grained, subophitic to ophitic, poorly foliated augite troctolite characterized by scattered augite-rich pegmatitic clots and patches. Commonly capped by hanging wall inclusions (HB & Ai) and interpreted to be the solidified basaltic liquid that carried the BMZ crystals and sulfides. Sulfide-bearing troctolite (BMZ) - Heterogeneous, sulfide-bearing, varitextured troctolite, augite troctolite, anorthositic troctolite, and olivine gabbro with 0.5 - 5% disseminated chalcopyrite, cubanite, talnakhite, pentlandite and pyrrhotite. Anorthosite (An-Series & Ai) - Undifferentiated Anorthositic Series inclusions. Includes wellfoliated very coarse-grained anorthosite, troctolitic-anorthosite, poikilitic troctolitic anorthosite, gabbroic anorthosite, gabbro, and locally troctolite. Inclusions range from a few cm’s to elongate bodies measured in km’s. Xenoliths in the SKI Anorthositic gabbro to gabbro (Upper Gabbro) - Mixed group of Anorthositic Series rocks that occur in the central portion of the map area. Includes well-foliated anorthositic gabbro, gabbro, oxide gabbro, anorthosite, and augite troctolite. Basaltic hornfels (Upper Basalt, HB) - Fine-grained, granoblastic to poikiloblastic basaltic hornfels; consists of variable amounts of plagioclase, augite, olivine, hypersthene, and inverted pigeonite. Commonly associated with Anorthosite xenoliths (unit Ai). Giants Range Batholith (2.68 Ga.) Footwall Porphyritic quartz monzonite (GRB) - Pink, coarse-grained, hornblende-phyric, porphyritic quartz monzonite with large (1-2 cm) orthoclase phenocrysts. Also contains irregular zones of aplite, lamprophyre, and supracrustal xenoliths. Strongly recrystallized and partially melted locally adjacent to the contact with the SKI. Early modeling of the Maturi Deposit was limited due to wider-spaced drill holes across the deposit area, though Severson (1994) did a remarkable job in defining the igneous stratigraphy of the SKI along a 19 mile strike length. Severson recognized the fact that the rocks below the lower pegmatite (PEG) unit typically contained sulfides and that rocks above the PEG unit were a monotonous sequence of sulfidebarren anorthositic troctolite to troctolite (AT/T), and augite-troctolite (Main AGT). Two detailed geologic cross sections through the Maturi Deposit are presented in Figure 3C-2. These sections display the continuity of the basal mineralization as well as the differences in the hanging wall 68 �stratigraphy from west to east through the deposit. In the east, the deposit is located under an extremely thick (>1,000m) megaxenolith of Anorthosite Series rocks, and in essence the basal SKI can be viewed as a thin sill-like body. To the west, the anorthosite xenolith ends and the immediate hanging wall rocks to the deposit are sulfide-barren troctolites of the Main AGT unit. We interpret that the Main AGT as the solidified troctolite melt that carried the crystals and sulfide droplets of the magmatic slurry. Figure 3C-2. Geologic cross sections through the Maturi Deposit. 69 �In 2008, geologists from Duluth Metals came to the realization that the initial basaltic composition SKI magmas that ultimately solidified to create the Maturi deposit intruded as sulfide-bearing, crystal-laden (olivine and plagioclase crystals), magmatic slurries. Based on this new interpretation, the company reinterpreted Severson’s (1994) regional basal stratigraphy (units U3, BH, BAN) of the SKI (Fig. 3C-3) at Maturi into the Basal Mineralized Zone, or BMZ. The company believes that the geometry of the system (sill-like sub-horizontal intrusion) and the inherent crystallinity of the basaltic melts (phenocrysts of plagioclase and olivine) led to crystal sorting and melting of the footwall granitic rocks to create the heterogeneous lithologies and textures of the BMZ. Figure 5C-3. Simplified crystal-liquid slurry model for the SKI in the Maturi area. MATURI ORE DEPOSIT MODEL In 2012, the geology of the mineralized portions (the BMZ) of the Maturi deposit were reevaluated by the geologic staffs of Duluth Metals, Twin Metals, and geologists from AMEC utilizing a significant volume of new, high-quality geochemical and geological data during the completion of an updated mineral resource classification by the consulting firm AMEC. Mineralization in both the BMZ and footwall at Maturi were reclassified based on patterns in the physical distribution of mineralization as projected on down-hole plots. Sulfide mineralization is characterized by several distinct patterns, including (1) very low grade mineralized intervals showing low variability (Stage 1), (2) moderate grade mineralized intervals showing low variability (Stage 2), and 3) higher grade mineralized intervals showing higher variability and commonly bounded by low grade selvages (Stage 3) (Fig. 3C-5). Significantly, the contacts between different mineralized intervals are typically quite abrupt. A single hole might contain one or several distinct mineralized intervals within the BMZ, including higher grade intervals with the highest grade occurring at the top, middle, or bottom of the section. Based on these criteria, four intrusive subunits, characterized by common grade profiles, were defined in the BMZ. In addition, two distinct suites of mineralization were identified in the footwall 70 �rocks, including Ni-Co enriched semi-massive to massive sulfide zones and disseminated Cu-PGE enriched zones deep in the footwall granitoids. The classifications derived from this exercise were validated by multivariate statistical analysis of multielement geochemical data, including principal component analysis (Fig. 3C-6) and factor analysis. This investigation revealed a significant correlation of multi-element geochemistry to mineralization within the BMZ as well as several possible subdivisions of the BMZ based on both the physical distribution patterns of mineralization and the geochemistry of the host rocks. The Maturi subunits so defined and validated were determined to occur in a consistent stratigraphic order, and are correlative across the deposit. Figure 3C-5. Revised igneous stratigraphy of the BMZ and adjacent rocks within the Maturi deposit. Figure 3C-6. Multi-element principal component analysis plot of MEX-Series drill hole geochemical data. 71 �Typical geochemical plots of Maturi drill holes are presented in Figure 3C-7 and display several of the patterns that were originally identified in the development of the revised geological model of the Maturi deposit. As well, an idealized intrusive sequence model for the SKI in the Maturi deposit area is given in Figure 3C-8 and a NW to SE cross section of the modeled units of the BMZ is presented in Figure 3C-9. Figure 3C-7. Downhole geochemical and principal component plots of typical drill holes within the Maturi deposit. 72 �Figure 3C-8. Idealized intrusive sequence model of the SKI in the Maturi deposit area. Figure 3C-9. Northwest (left) to Southeast (right) cross section through the Maturi deposit depicting the recently modeled geological units of the BMZ. Detailed descriptions of the seven units modeled for the Maturi deposit is well beyond the scope of this field trip guidebook. However, brief descriptions are provided in Table 3C-3 below and geochemical plots of copper, nickel, and precious metals are given in Figure 3C-10. One of the most important outcomes of the reinterpretation of the geology and mineralization within the Maturi deposit has been the identification of the higher-grade Stage 3 (S3) intrusive unit of the SKI (Table 3C-4). S3 has the highest grade and is the most widely distributed of the four BMZ units (Fig. 3C11). Cu, Ni, and PGEs and are all significantly elevated in S3 relative to the other BMZ units and the 73 �mineralized GRB. Stage 2 (S2) mineralization is overall much lower grade than S3, but locally S2 is well mineralized, and will likely contribute significantly to the deposit economics. Mineralization in the GRB is overall low grade and discontinuous. However, local zones of the unit G-N (where Cu and especially Ni locally occur as massive and semi-massive sulfides) are very high grade and may contribute to the resource. Table 3C-3. Interpreted lithostratigraphic-chemostratigraphic units within and adjacent to the BMZ within the Maturi deposit. Unit UH Description Stage 3 Continuous, higher-grade, PGE-enriched, heterolithic troctolite and melatroctolite Stage 2 Continuous, moderate-grade, heterolithic, oxide-bearing, augite-troctolite to troctolite Stage 1 Discontinuous, barren to very low-grade, homogeneous troctolite, gabbro, anorthosite and/or norite G-N Irregular, locally high-grade, Ni- and Co-enriched semi-massive to massive sulfide pods and veins at or immediately below the basal SKI contact. G-M Discontinuous, low to moderate-grade, disseminated Cu and PGE enriched mineralization. G-B Continuous, barren granitoid footwall rocks Discontinuous, barren to low-grade, highly variable troctolite Figure 3C-10. Geochemical boxplots of composited drill hole geochemistry for units within and adjacent to the BMZ within the Maturi deposit. 74 �The current lithostratigraphic model of the Maturi deposit effectively discriminates between higher- and lower-grade mineralization and provides a realistic geological model. The new data allowed correlation of units from hole-to-hole and section-to-section resulting in a very robust geologic model upon which to build mine plans and further our understanding of the magmatic processes that occurred to generate TMM’s Maturi Deposit. Figure 3C-11. Plan views of Stage 3 Cu, Ni, and TPM grades from the TMM Maturi deposit block model. 75 �REFERENCES Desaultels, P., and Patelke, R., 2008, Updated Technical Report on the NorthMet Deposit, Minnesota, USA, by Wardrop Engineering, (Toronto) for PolyMet, 109 p. Duchesne, L., 2004, Partial melting in microstructures associated with the contact aureole of the Duluth Igneous Complex: Unpublished M.S. thesis [in French], University of Quebec at Chicoutimi, 217 p. Foose, M.P., 1984, Logs and correlation of drill holes within the South Kawishiwi intrusion, Duluth Complex, northeastern Minnesota: United States Geological Survey, Open-file Report 84-14. Gal, B., 2008, The South Filson Creek deposit, a Masters of geology thesis: Eotvos Lorand University, Budapest, Hungary. Geerts, S. D., 1994, Petrography and geochemistry of a platinum group element-bearing horizon in the Dunka Road prospect, (Keweenawan) Duluth Complex, northeastern Minnesota: Unpublished M.S. thesis, University of Minnesota, Duluth, 100 p. Geerts, S.D., 1991, Geology, stratigraphy, and mineralization of the Dunka Road Cu-Ni prospect, Northeastern Minnesota: Natural Resources Research Institute, Univ. Minn., Duluth, Tech. Rpt., NRRI/TR-91/14, 63 p. Hauck, S.A., Severson, M.J., Zanko, L.M., Barnes, S.-J., Morton, P., Aliminas, H.V., Foord, E.E., and Dahlberg, E.H., 1997, An overview of the geology and oxide, sulfide, and platinum-group element mineralization along the western and northern contacts of the Duluth Complex, in Ojakangas, R.W., Dickas, A.B., and Green, J.C., eds., Middle Proterozoic to Cambrian rifting, central North America: Geological Society of America Special Paper 312, p. 137-185. Hauck, S.A., and Severson, M.J., 2000, Platinum-group element-gold-silver-copper-nickel-cobalt assays primarily from the Local Boy ore zone, Babbitt deposit, Duluth Complex, northeastern Minnesota: An update: Natural Resources Research Institute, University of Minnesota, Duluth, Report of Investigations NRRI/RI-2000/01, 48 p. Holst, T.B., Mullenmeister, E.E., Chandler, V.W., Green, J.C., and Weiblen, P.W., 1986, Relationship of structural geology to the Duluth Complex to economic mineralization: Minnesota Department of Natural Resources, Division of Minerals Report 241-2. Ikola, R.J., 2006, Vertical Electrical Soundings - NorthMet Project - St. Louis County, Minnesota, Prepared for PolyMet Mining Corp.: R.J. Ikola and Associates, Inc., Hibbing, Minnesota, 115 p. Jennings, C.E., and Reynolds, W.K., 2005, Surficial Geology of the Mesabi Iron Range, Minnesota: Minnesota Geological Survey, Miscellaneous Map M-164, scale 1:100,000. Lehr, J.D., and Hobbs, H.C., 1992, Glacial geology of the Laurentian divide area, St. Louis and Lake Counties, Minnesota: Minnesota Geological Survey Field Trip guidebook series, 18, Prepared for the 39th Midwest Friends of the Pleistocene Field Trip, Biwabik, Minnesota, 1992, 73 p., map, 1:250,000. Listerud, W.H., and Meineke, D.G.., 1977, Mineral resources of a portion of the Duluth Complex and adjacent rocks in St. Louis and Lake Counties, northeastern Minnesota: Minnesota Department of Natural Resources, Division of Minerals, Report 93, 74p. Mancuso, J.D., and Dolence, J.D., 1970, Structure of the Duluth Gabbro Complex in the Babbitt area, Minnesota (abs.): 16th Annual Institute on Lake Superior Geology, Thunder Bay, Ontario, p. 27. Martineau, M. P., 1989, Empirically derived controls on Cu-Ni mineralization: A comparison between fertile and barren gabbros in the Duluth Complex, Minnesota, U.S.A.: in Prendergast, M. D., and Jones, M. J., eds., Magmatic Sulphides - The Zimbabwe Volume: Inst. Min. Metall., London, p. 117-137. Miller, J.D., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.M., 2001, Geological map of the Duluth Complex and related rocks, Northeastern Minnesota; Minnesota Geological Survey, Miscellaneous Map M119, scale 1:200,000. Miller, J.D., Jr. and Severson, M.J., 2002, Geology of the Duluth Complex in Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., 2002a, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations RI-58, p. 106-143. 76 �Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., 2002a, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations RI-58, 207 p. Morton, P.M., and Hauck, S.A., 1987, PGE, Au and Ag contents of Cu-Ni Sulfides Found at the Base of the Duluth Complex, Northeastern Minnesota: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report NRRI/GMIN-TR-87-4, 94 p. Muhich, T.G., 1993, Movement of titanium across the Duluth Complex–Biwabik Iron Formation contact at Dunka Pit, Mesabi Iron Range, northeastern Minnesota: Duluth, Minn., University of Minnesota Duluth, M.S. thesis, 154 p. Park, Y.R., Ripley, E.M., Severson, M.J., and Hauck, S.A., 1999, Stable isotopic studies of mafic sills and Proterozoic metasedimentary rocks located beneath the Duluth Complex, Minnesota: Geochimica et Cosmochimica Acta, v. 63, no. 5, p. 657-674. Patelke, R.L., 2003, Exploration drill hole lithology, geologic unit, copper-nickel assay, and location database for the Keweenawan Duluth Complex, northeastern Minnesota: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report, NRRI/TR-2003/21, 96 p. Patelke, R.L., and Geerts, S.D.M., 2006, PolyMet NorthMet Drill Hole / Geological Database Recompilation: Location, Downhole Survey, Assay, Lithology, Geotechnical Data, and Related Information, 2004 to 2006, PolyMet Mining Inc. Internal report, 86 p. Peterson, D.M. and Albers, P.B., 2007, South Kawishiwi Intrusion Cu-Ni-PGE mineralization in association with the Nickel Lake Macrodike, Institute on Lake Superior Geology, 53rd Annual Meeting, Field Trip Guidebook, Lutsen, Minnesota, Volume 53. Peterson, D.M., 2001a, Development of a conceptual model of Cu-Ni-PGE mineralization in a portion of the South Kawishiwi Intrusion, Duluth Complex, Minnesota: Laurentian University – Society of Economic Geologists, Second Annual PGE Workshop, Sudbury, Ontario. Peterson, D.M. and Severson, M.J., 2002, Chapter 4, Archean and Paleoproterozoic rocks forming the footwall of the Duluth Complex, in Geology and mineral potential of the Duluth Complex and related intrusions of northeastern Minnesota, Minnesota Geological Survey, Report of Investigations 58, pp. 76-93. Peterson, D.M., 2008, Bedrock geologic map of the Duluth Complex in the northern South Kawishiwi intrusion and surrounding area, Lake and St. Louis Counties, Minnesota: Natural Resources Research Institute, Map Series NRRI/MAP-2008-01, scale 1:20,000. Phinney, W.C., 1972, Duluth Complex, history and nomenclature, in Sims, P. K., and Morey, G. B., eds., Geology of Minnesota: A Centennial Volume: Minn. Geol. Survey, pp. 333-334. Ripley, E.M., 1986, Origin and concentration mechanisms of copper and nickel in Duluth Complex sulfide zones – a dilemma: Economic Geology, v. 81, p. 974-978. Sabelin, T., and Iwasaki, I., 1985, Metallurgical evaluation of chromium-bearing drill core samples from the Duluth Complex (Contract report of Minnesota Department of Natural Resources, Division of Minerals): Minerals Resources Research Center, University of Minnesota, Minneapolis, MN, 58 p. Sabelin, T., and Iwasaki, I., 1986, Evaluation of platinum group metal occurrence in Duval 15 drill core from the Duluth Complex: Internal report, Minerals Resources Research Center, University of Minnesota, Minneapolis, MN, 23 p. Severson, M.J., 1994, Igneous stratigraphy of the South Kawishiwi intrusion, Duluth Complex, northeastern Minnesota: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report NRRI/TR 93/34, 210 p. (with plates) Severson, M.J., and Barnes, R.J., 1991, Geology, mineralization and geostatistics of the Minnamax/Babbitt Cu-Ni deposit (Local Boy area), Minnesota, Part II: Mineralization and geostatistics: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report, NRRI/TR-91/13b, 216 p. Severson, M.J., and Hauck, S.A., 1990, Geology, geochemistry, and stratigraphy of a portion of the Partridge River intrusion: Natural Resources Research Institute, University of Minnesota-Duluth, Technical Report, NRRI/GMINTR-89-11, 236p. (with plates). 77 �Severson, M.J., and Hauck, S.A., 2003, Platinum group elements (PGEs) and platinum group minerals (PGMs) in the Duluth Complex: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-2003/37, 296 p. Severson, M.J., and Hauck, S.A., 2008, Finish logging of Duluth Complex drill core (And a reinterpretation of the geology at the Mesaba (Babbitt) deposit): University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR-2008/17, 68 p. Severson, M.J., Miller, J.D., Jr., Peterson, D.M., Green, J.C., and Hauck, S.A., 2002, Mineral potential of the Duluth Complex and related intrusions in Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., 2002a, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations RI-58, p. 164-200. Severson, M.J., Patelke, R.L., Hauck, S.A., and Zanko, L.M., 1994a, The Babbitt copper-nickel deposit, Part B: Structural datums: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report NRRI/TR-94/21b, 48 p. Severson, M.J., Patelke, R.L., Hauck, S.A., and Zanko, L.M., 1994b, The Babbitt copper-nickel deposit, Part C: Igneous geology, footwall lithologies, and cross-sections: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report NRRI/TR-94/21c, 79 p. Sims, P.K., and Morey, G.B., eds. 1972, Geology of Minnesota: A Centennial Volume, Minnesota Geological Survey. Tharalson, E., Sweet, G., Boisjoli, T., Lentz, B., Fellows, T., and Peterson, D., 2007, Geological Map of the Nickel Lake Macrodike and Northern Bald Eagle Intrusion, Lake County, Northeastern Minnesota: Precambrian Research Center Map Series PRC/MAP-2007-01, scale 1:10,000. Watowich, S.N., Malcomb, J.B., and Parker, P.D., 1981, A review of the Duluth Gabbro Complex as a domestic source of critical and strategic metals: Society of Mining Engineers of AIME, preprint 81-351, 9 p. White, C., 2010, The Nokomis Deposit, a Masters of Geology thesis: University of Minnesota, Duluth. Wolff, J.F., 1917, Recent geologic developments on the Mesabi range, Minnesota: American Institute of Mining and Metallurgical Engineers Transactions, v. 56, p. 142-169. Zanko, L.M., Severson, M.J., and Ripley, E.M., 1994, Geology and mineralization of the Serpentine copper-nickel deposit, Duluth Complex, Minnesota: Natural Resources Research Institute, University of Minnesota, Duluth, Technical Report, NRRI/GMIN-TR-93-52, 90p. 78 �FIELD TRIP 4 May 4, 2016 DULUTH STREAM GEOMORPHOLOGY AND THE JUNE 2012 FLOOD Karen Gran Department of Earth and Environmental Sciences University of Minnesota Duluth FIELD TRIP CANCELLED 79 �FIELD TRIP 5 May 4, 2016 GEOLOGY OF THE ENDION SILL EXPOSED ALONG DULUTH’S LAKEWALK James D. Miller Department of Earth and Environmental Sciences and the Precambrian Research Center University of Minnesota Duluth Introduction The Endion Sill is a mafic to felsic, hypabyssal intrusion that was emplaced into the North Shore Volcanic Group lavas during the 1.1 Ga Midcontinent Rift (Miller and Green, 2002). The sill is the lowest hypabyssal intrusion in the volcanic rocks overlying the Duluth Complex at Duluth. This gently east-dipping, approximately 425m-thick intrusion is semi-continuously exposed along a half-mile stretch of Lake Superior shoreline in wave-washed outcrops adjacent to Duluth’s Lakewalk path between 16th and 28th avenues east. This field trip will highlight the mineralogic and textural attributes of the diverse lithologies exposed in these shoreline outcrops and consider the various ideas for its emplacement and crystallization history. All who have studied the Endion Sill (Schwartz and Sandberg, 1940; Ernst, 1955, 1960; Oestrike, 1983; Gardner, 1987; and Jerde, 1991) have noted its being generally composed of a lower gabbroic zone and an upper intermediate to felsic zone. The sill was emplaced between basaltic to basaltic andesite flows comprising the upper part of the Leif Erickson Park Lavas (unit Mnb, Fig. 1) and the 300m-thick, Congdon Park rhyolite (unit Mnc, Fig. 1) which forms the lowermost flow unit of the Lakeside Lavas of the NSVG (Green and Miller, 2008). The emplacement of mafic intrusions beneath felsic lavas and intrusions is a commonly observed phenomenon in igneous rocks of the Midcontinent Rift and likely is triggered by the felsic rocks serving as a density barrier to mafic magma (Miller, 2015). The gradational contacts typically observed between of felsic rock types and underlying mafic intrusions further suggests partial melting and assimilation of the felsic hanging wall is involved in the evolution of the mafic intrusion. This contamination process appears to have occurred in the Endion Sill as well. Previous Studies Schwartz and Sandberg (1940) were the first to report on the petrology of mafic and felsic components of the Endion Sill and of related sills in the Duluth area, the Northland and Lester River sills. From their field observations and petrographic studies and major and minor element wet chemical analysis of seven samples, they considered five different hypotheses to explain the bimodal distribution of rock types, and particularly the occurrence of felsic rock types (commonly referred to as granophyre or “red rock”): 1) as a composite felsic intrusion, 2) by metamorphism of the rhyolite, 3) by hydrothermal alteration, 4) by a mix of assimilation and differentiation, and 5) by differentiation alone. They concluded “that many factors had some effect, but that differentiation within each sill accounts for the major part of the segregation of rock types into two distinct facies.” Nevertheless, they acknowledge that simple fractional crystallization of the diabase could not produce the large amount of felsic material in the sills. 80 �Figure 1. Geologic setting of the Endion Sill (brown unit Mne) in the Duluth 7.5’ quadrangle (from Green and Miller, 2008). Footwall rocks (to the west) comprise the basaltic Leif Erickson Park Lavas and hanging wall rocks are composed of the Congdon Park Rhyolite (Unit Mnc) and overlying mixed lavas of the Lakeside lavas. The Mep unit is a granophyre that occurs at the diabase-rhyolite contact. Geochemical analyses of Esamples of Jerde (1991) are given in Table 1. For his MS thesis at the University of Minnesota under the advisement of SS Goldich, Ernst (1955, 1960) conducted a detailed petrographic study of 27 samples collected from lakeshore exposures and streambed exposures up Tischer Creek (Fig. 2). Ernst (1955) defined several stratiform map units within the sill. The lower half of the sill is composed of medium-grained, subophitic to ophitic diabase. The diabase is overlain over a narrowly gradational contact by a medium-grained, intermediate “mottled” granodiorite which gradually transitions upward into granophyre/”red rock” which he calls a “sodalite-adamellite”. This “medial granophyre” then abruptly grades upward into a fine-grained diabase. The upper 10 meters of the sill is capped by a fine-grained “upper granophyre” which is in sharp irregular contact with the diabase. The upper granophyre-Congdon Park rhyolite contact is observed to be abrupt in sea cliff exposures and in streambed outcrops along Tischer Creek (where it crosses East 2nd Street, Fig. 2). The transition into rhyolite is marked by the abrupt decrease in grain size. Ernst (1955, 1960) concluded that much of the mafic and intermediate rock types in the lower part of the sill (about 40% of the total thickness) were likely related by fractional crystallization. However, like Schwartz and Sandberg (1940), he recognized that there was too much felsic material to have been produced by in situ differentiation. He offered two possible explanations for the large volume of granophryre in the upper part of the sill - 1) additional granophyre generated by differentiation from a downdip, thicker section of the sill had injected itself into an up-dip section of the sill, where it is currently exposed or 2) the granophyre represents a separate intrusion of felsic magma unrelated to the mafic component. 81 �Figure 2. Geology of the Endion Sill from Ernst (1955) showing his sample locations. For his University of Illinois MS thesis, Oestrike (1983) conducted a more thorough petrographic and mineral chemical study of the Endion Sill along the shoreline and Tischer Creek sections (Figs. 3 & 4). He subdivided the sill into 3 main units that he termed the Gabbroic Zone, the Acidic Zone and the Intermediate Zone. Grading abruptly up from a basal chill zone, the Gabbroic Zone is largely composed medium-grained, subophitic to ophitic, oxide gabbro to olivine gabbro with variable concentrations of “red spots” (interstitial concentrations of granophyre) and felsic dikelets. About 100 meters above the basal contact, Oestrike noted that the gabbro is in sharp contact with a medium-grained, subprimatic ferromonzodiorite – a rock type commonly found in the Acidic Zone. The next exposure upsection of the ferromonzodiorite is more subophitic gabbro, but the contact is not exposed. Oestrike calls this interval the Gabbroic Zone Acidic Layer (Figs. 3&4). The contact between the Gabbroic and Acidic Zone is about 220 meters above the basal contact and is gradational over several 10’s of meters. It is marked by the transition from a subophitic to subprismatic pyroxene texture and an increase in interstitial granophyre to above 10%. Within the Acidic Zone this trend continue to where the rock becomes a subprismatic to prismatic quartz ferromonzonite and develops a very deep salmon red color. Oestrike notes that these rocks have tridymite paramorphs of quartz. The upper 50 meters of the sill, an interval Oestrike calls the Intermediate Zone, is largely composed of a gray, fine-grained ferrodiorite that locally contains granophyric clot and dikes. Its lower contact with the Acidic Zone s gradational over several meters and its upper contact is characterize by an irregular mixture of ferrodiorite and massive red granophyre over a 3 meter interval. This Intermediate Zone is not observed in the Tischer Creek section. The granophyre-rhyolite contact is inaccessible along the shore due to sea cliffs. 82 �| Int Zn | Acidic Zone | Gabbroic Zone Figure 3. Geology and lithostratigraphy of the Endion Sill from Oestrike (1983) showing his sample locations. Figure 4. Cryptic variation of An content in plagioclase and En% (relative to Fs and Wo) in pyroxene through shoreline exposure of the Endion Sill. Mineral compositions were determined by electron microprobe (from Oestrike, 1983). Citing constant and distinct mineral compositions within the Gabbroic and Acidic Zones (Fig. 4), Oestrike (1983) disagreed with the conclusions of Schwartz and Sandberg (1940) and Ernst (1955, 1960) that the compositional variations in the Endion Sill represent mostly the effects of magmatic differentiation. Instead, he concluded that the Endion Sill was formed by the composite emplacement of mafic and intermediate-felsic magmas in rapid succession. He speculated that the two magmas did not readily mix due to the effects of double-diffusive convection. He did not have a satisfactory explanation for the origin of the Intermediate Zone ferrodiorite. 83 �A lithogeochemical study of the Endion Sill, including Sm-Nd isotopes, was conducted by Gardner (1987) for his MS thesis at Washington University. (I have not been able to track down a copy of his 350 page thesis, so I will paraphrase the conclusions of his study from a Lunar and Planetary Science Conference Abstract (Gardner et al., 1987) and from a summary by Jerde (1991)). Although Jerde reports that Gardner (1987) conducted 9 analyses of Sm-Nd isotopes for his thesis study, Gardner et al. (1987) reported Sm-Nd analyses of only one Gabbroic Zone sample and one Acidic Zone sample. Gardner’s (1987) INAA analyses of 26 samples through the sill show incompatible trace element concentrations increasing uniformily through the sill. He notes that REE patterns and concentrations of Gabbroic Zone samples resemble andesitic NSVG compositions. He further notes that the 30m-thick granophyre interval, which he calls the altered rhyolite zone and which occurs between the top of the sill and the base of the unaltered rhyolite, is depleted in REE and other incompatible trace elements (Ta, Rb, Th) and enriched in more compatible elements such as Sr, Ba, Co, and Sc relative to unaltered rhyolite. He suggests that devolution of fluids from the rhyolite may have mobilized trace element into the underlying sill. Gardner et al. (1987) reports ƐNd (1100Ma) values for the Gabbroic Zone sample of -0.6 and -2.0 for the Acidic Zone sample (Fig. 5). Gardner concludes that if the Congdon Park Rhyolite has an ƐNd value of <8, its partial melting and assimilation along with fractional crystallization of the Endion sill may explain its Nd isotopic compositions. Gardner did not analyze the Sm-Nd isotopic compositions of the Congdon Park Rhyolite, but Vervoort and Green (1997) report a ƐNd (1100Ma) value of -4.1 (Fig. 5). Figure 5. ƐNd (1100Ma) and 1/Nd values for samples from the Gabbroic and Acidic zones of the Endion Sill compared to primitive NSVG basalt, the Congdon Park Rhyolite, and 5% and 10% model partial melts from Superior Province crust. The values show for the Congdon Park Rhyolite are from Vervoort and Green (1997). All other data are from Gardner et al. (1987). Gardner (1987) conclude from his geochemical data that the Gabbroic and Acidic zones represent two separate intrusions into the sill and are not related by in-situ fractional crystallization. Both show contamination which may have been contributed by the Congdon Park rhyolite or by partial melts of Archean crust during the intrusion of mantle-derived magmas (Fig. 5). He interpreted the more contaminated Acidic Zone to have been intruded later and above the semi-molten gabbroic zone. Gardner (1987) further interpreted the upper Intermediate Zone to be a more contaminated upper contact of the Gabbroic Zone. 84 �Jerde (1991) studied the lithogeochemistry of the Endion Sill as a part of a larger study of hypabyssal mafic intrusions into the NSVG for his PhD dissertation at UCLA. Locations of the 17 samples he analyzed by INAA and microprobe analyses of fused glass are shown in Figure 1 and the data are presented in Table 1. By his own admission, he conducted only reconnaissance field investigations and instead relied on the previous studies of Oestrike (1983) and Gardner (1987). Table 1. Geochemical analyses of Endion Sill samples from Jerde (1991). Locations shown in Figure 1. Table 1 (cont.) 85 �The chemostratigraphic variations of major and trace elements of Jerde’s (1991) data are shown in Figures 6 and 7, respectively, relative to the lithostratigraphy of Oestrike (1983). The major element data (SiO2, TiO2, mg# and Na2O+K2O) clearly show abrupt compositional changes between the gabbroic and acidic zones across 50 meter-thick interval that Jerde (1991) termed the transitional zone (AZ/GZ, Fig., 6). Oestrike (1983) recognized this hybrid transitional zone petrographically and interpreted this as a mixing zone between Acidic Zone magma that was compositely emplaced above the partially solidified Gabbroic Zone. Interestingly, the Intermediate Zone (IZ, Fig. 6) has major element compositions that are comparable with the transitional zone compositions suggesting that it is the upper hybrid chill zone of the Acidic Zone rather than a remnant chill of the Gabbroic Zone as interpreted by Gardner (1987). The transitional major element compositions of the recrystallized rhyolite (xRhy, Fig. 6) between the unaffected rhyolite and the Intermediate Zone rocks suggests that the recrystallized rhyolite has been contaminated by and provided contaminants to the Endion Sill as suggested by Gardner (1987). 86 �Figure 6. Chemostratigraphic variation of major elements through the Endion Sill. Lithogeochemical data from Jerde (1991), see Table 1. Lithostratigraphic units modified from Oestrike (1983); see Fig. 3 – GZGabbroic Zone; GZa- acidic layer in the Gabbroic Zone; AZ/GZ – transitional zone between the Acidic and Gabbroic Zones; AZ – Acidic Zone; IZ – Intermediate Zone; xRhy – recrystallized rhyolite; Rhy – unaffected rhyolite. Darker shades indicate felsic compositions. Jerde’s (1991) trace element data (Zr, Ba, Ni, and Cr, shown in Figure 7) are more consistent with the Acidic and Gabbroic Zones being formed by composite intrusions, as opposed to in situ fractional crystallization. Moreover, the similarity of trace element abundances between the transitional zone (AZ/GZ, Fig. 7) with the Intermediate Zone is consist with the IZ being an upper hybrid chill of the Acidic Zone magma. However, the very uniform incompatible trace element ratios of Ce/Yb and Zr/Hf indicate that both zones were formed from a common parent magma which presumably differentiated at depth. The subtle increase in the Th/Hf ratio in the Acidic zone is consistent with Gardener’s (1987) Sm-Nd isotopic data that implies contamination of the acidic zone magma by the overlying Congdon Park rhyolite. Most Midcontinent Rift rhyolites are interpreted from negative ƐNd isotopic compostions to have formed by partial melting of Archean to Paleoproterzoic crust, which should also be enriched in Th (Vervoort and Green, 1997). 87 �Figure 7. Chemostratigraphic variation of trace elements through the Endion Sill. Lithogeochemical data from Jerde (1991), see Table 1. Lithostratigraphic units modified from Oestrike (1983) – GZ-Gabbroic Zone; GZa- acidic layer in the Gabbroic Zone; AZ/GZ – transitional zone between the Acidic and Gabbroic Zones; AZ – Acidic Zone; IZ – Intermediate Zone; xRhy – recrystallized rhyolite; Rhy – unaffected rhyolite. Darker shades indicate more felsic compositions. The REE data reported by Jerde (1991), and plotted in normalized spidergrams in Figure 8, also provide some insight into the petrogenetic relationship between the Gabbroic and Acidic Zones. That the general slopes of the REE curves from the Endion Sill samples are general coparallel are again consistent with the Gabbroic and Acidic Zones being evolved from a common magma. However, the GZ and AZ samples clearly cluster in two distinct REE abundance groups. The basal chill sample (black dot in Fig. 8) shows enriched REE abundances relative Gabbroic Zone samples, but show a similar trend and lack of Eu anomaly. This suggests that the gabbroic zone samples may have cumulate tendencies (i.e. concentrations of primocrysts over parental magma) as suggested by Jerde (1991) and supported by his noting higher than normal concentrations of olivine in his sample E6 from the Gabbroic Zone (Table 1). The Acidic Zone samples have distinctively higher REE concentration with a similar slope to the GZ samples, but a moderate negative Eu anomaly. Interestingly, two sample from the transitional (AZ/GZ) zone have distinctive compositions with one lining up with the Gabbroic Zone samples and the other similar to the Acidic Zone samples, though with a less pronounced negative Eu anomaly. Interestingly, the REE-enriched transitional zone sample is most similar to the REE compositions of the upper 88 �Rock/Primitive Mantle Intermediate Zone. This lends additional evidence to the suggestions of Jerde (1991), Gardner (1987) and Oestrike (1983) that the acidic zone is a later composite intrusion above the semi-crystallized Gabbroic Zone and that the Intermediate Zone and transitional zone are upper and lower margins of that later impulse of intermediate magma. The REE patterns of the overlying Congdon Park Rhyolite and its recrystallized lower interval in contact with the Intermediate Zone are somewhat similar, but have steeper LREE slopes and a much more pronounced negative Eu anomaly. That Eu values for Acidic Zone straddle the range between Intermediate Zone and transition zone samples and those of the rhyolite and recrystallized rhyolite are also consistent with Gardner (1987) Nd isotope data that suggest that the Acidic Zone assimilated a modest amount of anatectic melt from the rhyolite. Figure 8. REE normalization plot of samples from the Endion Sill analyzed by Jerde (1991); see Table 1. Samples are color coded to the lithostratigraphic column modified from Oestrike (1983). REE abundance normalized to primitive mantle composition of Sun and McDonough (1989). 89 �Recent Studies The shoreline exposures of the Endion Sill, which will be the focus of this field excursion, were mapped in detail in September 2007 by the author as the final element of bedrock mapping integrated into the geologic map of the Duluth 7.5’ quadrangle (Green and Miller, 2008); See Figure 1. A total of 27 sample were collected during the mapping and thin sections were made for 25 of these. These sections were only recently investigated and photographed earlier this year (nothing like leading a field trip to finally getting around to it). Field stop descriptions given in the next section will incorporate the field and petrographic observations and will be illustrated with field photos and photomicrographs from these recent studies. Based on these field and petrographic studies, a revised lithostratigraphy for the Endion Sill is proposed and summarized in Figure 9. This figure highlights the textures, mineralogy and reddish hemitization of the various lithologies with thumbnail scans of thin sections and summary petrographic descriptions. The main differences with the stratigraphic column proposed by Oestrike (1983) is to expand the base of the AZ/GZ transition zone from 200 meters down to 100 meters. This corresponds to the sharp intrusive contact between gabbroic rocks and ferromonzodiorite of his “acidic layer in the Gabbroic Zone” (Fig. 3). Oestrike (1983) resumes the Gabbroic Zone over the acidic layer based on interpreting his sample E-24 as being gabbro. However, sample MD769 is observed to be varitextured oxide gabbro with about 10% granophyre and no evidence of olivine. Indeed olivine is only observed in the lower 100 meters of the sill. Field description of this outcrop notes that this sample was taken from the least granophyric sample. As such, it bears greater resemblance to the variably granophyric olivine-barren gabbros ranging in pyroxene texture from subophitic to intergranular to poikiloprismatic that are common in the transition zone. Unfortunately, Jerde (1991) did not sample this exposure to see if it maintains a Gabbroic Zone chemistry or starts to show evidence of mixing with the Acidic Zone magma. (Don’t know if Gardner (1987) sampled here). Other elements of the lithostratigraphic column are as portrayed by Oestrike (1983). Questions to Consider As we progress up section through the excellent shoreline exposures of the Endion Sill, some questions to consider include: • What is the petrogenetic relationship between the Gabbroic Zone and the Acidic Zone and particularly theAZ/GZ transition zone between them? • Do the field relations (and geochemical data) give any indication that fractional crystallization is involved? • What is the evidence for composite emplacement of the GZ and AZ, and what is the relative timing of their emplacement? • What does the ferrodioritic Intermediate Zone represent? - an upper chill of the Acidic Zone? a remnant of the upper chill of the Gabbroic Zone? or something unrelated to either? • What is the source of granophyric material in the sill? – partial melting of the overlying rhyolite or fractional crystallization in situ or in a remote staging chamber? 90 �Figure 9. Revised lithostratigraphy of the Endion Sill based on recent field and petrographic studies. 91 �Field Stop Descriptions The field trip will start at a large glacially polished and striated outcrop (location 2) just several meters east of the Lakewalk where 17th Ave. East projects to the shore and an elevated walkway crosses the interstate highway. Proceed to the SW edge of the outcrop to observe the basal contact (Location 1). Area A – Gabbroic Zone of the Endion Sill Location 1: Basal Chill of the Endion Sill UTM (NAD83): 570490_5183340 Description: Exposed here is a sharp contact between an intermediate volcanic (icelandite?) and an aphanitic gabbroic rock with sparse small phenocrysts of plagioclase and an altered mafic (olivine?). The volcanic occurs in an 5 4 outcrop just SW of the fine gabbro and as a 0.5 m thick 3 slab/septum of similar intermediate volcanic about 1 meter above the (unexposed) basal contact (Fig. 11). Fragments 2 of the volcanic also occur as blocks in the fine gabbro. The 1 contact between the volcanic slab and the very chilled hackly fractured surface of the gabbro has a strike and dip of N30°W/18°NE. A small (unexposed) fault is evidenced by an apparent left lateral offset of the volcanic slab across a N15E trending gap in exposure. Given the NE dip of the contact, this offset may also indicate east-side up displacement, which would imply it formed during late rift compression. IV F BC Figure 11. Panoramic photo of the basal chill (BC) of the Endion Sill with a slab/septum of intermediate volcanic (IV) offset by a reversed? fault (F). Location 2: Glacially polished whale-back outcrop granophryic gabbro. UTM (NAD83): 570620_5183305 Description: Progressing northeasterly across this glacially smoothed, polished, and striated outcrop, it is easy to observe the progressive coarseing in texture and the increase in granophyre clots, which stand out in relief, from about 5% to over 30% (Fig. 12A) Samples MD765D and MD765E (Fig. 10) show a transition from a medium fine-grained, felty subophitic slightly granophryic oxide gabbro to a mediumgrained, ophitic apatitic, granophryric olivine gabbro with 1-2 cm clots of micrographic Ksp+Qtz and free quartz (Fig. 13). All igneous phase show moderate degrees of alteration to bowlingite (Ol), uralite (Cpx), and sericite (Pl). Small inclusion of fine-grained, locally amygaloidal basalt occur in the gabbro and tend to concentrate granophyre at their margins (Fig. 12B). Also, the outcrop is cut by alteration veins trending N15E (parallel to the fault) along which liesegang redox bands are developed. 92 �A B Figure 12. A) Glacially polished and striated subophitic gabbro with irregular granophyre clots standing out in relief; B) Basaltic inclusion with concentrations of granophyre in surrounding gabbro. B A Cp Ol gp B’ A’ Figure 13. Photomicrographs of felty subophitic gabbro texture of sample MD765D (A & A’) and subophitic granophyric olivine gabbro texture of MD765E (B & B’). Poikilitic augite (Cp), bowlingite-altered olivine (Ol) and granophyre-rich (gp) area noted in B. All photos at 1.25x, scale bar = 3mm Location 3: Ophitic olivine gabbro UTM (NAD83): 570665_5183350 Description: Medium fine-grained, poorly foliated, ophitic olivine gabbro. Augite oikocrysts are about 0.5cm diameter (Fig. 14A). Granular olivine tends to cluster in the inter-ophite areas (Fig. 14B) and preferentially weather out on the bedrock surface to form pits. 93 �B A Figure 14. Photomicrographs from Sample 766. A) Ophitic augite with plagioclase chadacryts. B) Altered (bowlingite) olivine clusters in interophite areas. Both photos at 4x, scale bar = 0.5mm Location 4: Columnar jointed, ophitic olivine diabase UTM (NAD83): 570725_5183402 Description: The best developed columnar jointing in the Endion Sill are exposed here (Fig. 15A) as is obvious ophitic texture (Fig. 15B). Cpx oikocryts range from 0.5 to 2 cm diameter over the outcrop. Thin sections show this gabbro to contain less than 5% granophyric mesostasis. B A Figure 15. Field photos of Location 4 outcrop of ophitic olivine diabase displaying columnar jointing (A) and <1cm oikocrysts of pyroxene on a weathered surface (B). Location 5: Contact between Gabbroic Zone and AZ/GZ Transitional Zone UTM (NAD83): 570851_5183420 Description: Exposed in the outcrop at this point is a sharp contact between medium-grained, intergranular oxide gabbro and intergranular to subprismatic ferromonzodiorite (Fig. 16B). This contact has been recognize by all workers and marks the lower contact of Oestrike’s acidic layer in the Gabbroic Zone (Fig. 3) . The gabbro near the contact is intergranular, moderately granophyric (5-10%) and devoid of olivine (Fig. 16C), but the westernmost exposures grade into a subophitic to ophitic, mildly granophyric (<5%) olivine gabbro typical of the Gabbroic Zone (Fig. 16D) . Given the textural zoning of the gabbro against the rather homogeneous ferromonzodiorite, one could argue that the gabbro is intrusive 94 �into the ferromonzodiorite. This of course, is counter to interpretations of Oestrike (1983), Gardner (1987) and Jerde (1991) that the intermediate rocks intruded the gabbro. A B D C Figure 16. Progression of rock types at Location 5. A) subprismatic ferromonzodiorite, B) sharp contact between ferrodiorite and intergranular granophyric gabbro, C) intergranular oxide gabbro 1 meter form contact, subophitic olivine gabbro 5 meters from the contact. Area B – Transitional AZ/GZ Zone of the Endion Sill Location 6: Heterogeneous granophyric gabbro UTM(NAD83): 570995_5183600 Description: Although Oestrike (1983) shows this exposure as belonging to the Gabbroic Zone, this medium-grained, variably granophyric (15-25%) and varitextured (poikiloprismatic to subophitic to intergranular) oxide gabbro seems better grouped in the transitional AZ/GZ zone (Fig. 17). B’ B gp A Cp Cp Figure 17. Field and petrographic photos of sample MD769 from Location 6. A) Field photo of the variably granophyric composition of this exposure. B) Plane and cross-polar images of intergranular to subophitic augite (Cp) and interstitial areas rich in K-feldspar and quartz (granophyre). Photomicrographs at 1.25x, scale bar = 3mm. 95 �Location 7: Gradation from Granophyric Gabbro to Ophitic Gabbro UTM(NAD83): 571125_5183608 Description: Across this point, a very gradational transition in rock type is observed. At the SW end of the exposure, the rock is a subprismatic to subophitic granophyric (5-10%) oxide gabbro (Fig. 18A) . Crossing the point to the southeast, the granophyre abundance drops to below 5% and an ophitic texture is clearly developed with Cpx oikocrysts up to 3 cm across evident (Fig. 18B). At one location, a softballsized anorthosite inclusion is present. Plagioclase phenocrysts are present throughout the ophtic gabbro. A B gp gp A’ B’ Figure 18. Photomicrographs from samples at Location 7. A) Subophitic granophyric oxide gabbro with granophyric mesostasis (gp) (Sample MD770A). B) Ophitic oxide gabbro (Sample MD770B). All photos at 1.25X; scale bar = 3mm. Location 8: Foliated, Subprismatic, Apatitic Granophyric, Oxide Gabbro. UTM (NAD83): 571185_5183635 Description: Across a 5 meter gap in exposure from the ophitic gabbro at the east end of Location 7, the rock type abruptly changes to a medium fine-grained, subprismatic to intergranular, apatitic oxide gabbro with up to 20% granophyre. Locally, the rock displays a crude foliation of plagioclase and pyroxene (Fig. 19). This is the first significant occurrence of more than 1% apatite as slender needles, which persists upsection throughout the Acidic Zone. 96 �B B’ A Figure 19. Textures of foliated intergranular granophric gabbro in outcrop (A) and thin section (B) at Location 8. Location 9: Subprismatic Apatitic Quartz Ferromonzodiorite UTM(NAD83): 571245_5183675 Description: The stretch of Endion ledges marks the consistent onset of subprismatic to prismatic texture of mafic phases, the increase in granophyric mesostasis to greater than 15%, the presence of quartz paramorphs after tridymite, and the occurrence of 2-3% apatite (Fig. 20). This interval marks the upper part of the AZ/GZ transitional zone. The rock has a pinkish hue, but its does not develop the deep red color characteristic of the Acidic Zone until the next outcrop at Location 10. A B C Figure 20. Photomicrographs of subprimatic, apatitic quartz ferromonzodiorite sample MD770D(A) and MD 770E(B & C). Quartz paramorphs of tridymite are commonly associated with granophyre patches. Scale bar = 3 mm. 97 �Area C – Acidic Zone of the Endion Sill Locations 10 - 13: Hematized Prismatic Quartz Ferromonzonite UTM (NAD83): 571320_5183720 to 571607_5183960 Description: Homogeneous exposures along this stretch of shore display the strongly hematized rock that characterizes the Acidic Zone (or Red Rock of Schwartz and Sandberg, 1940). Granophyre content through this section is consistently between 30 and 50%, 20-35% plagioclase laths are commonly bleached white, and altered pyroxene becomes is consistently subprismatic to prismatic (Fig. 21). In thin section, tridymite paramorphs are ubiquitous. Apatite ranges from 0.5 to 2 %. B A C Figure 21. Prismatic to subprismatic textures of ferromonzonite exposures at Locations 10, 11 and 12. A B Figure 22. Strongly hematized and granophyre+tridymite mineralogy of samples MD771B (A) and MD772A (B) from locations 10 and 12, respectively. Scale bars = 3mm. 98 �Area D – Intermediate Zone and Upper Contact of the Endion Sill Location 14: Heterogeneous mix of ferromonzodiorite and granophyre UTM (NAD83): 571615_5183990 to 571660_5184125 Description: Along this continuous stretch of shoreline outcrop, the prismatic quartz ferromonzonite grades into a less granophyric composition of ferrodiorite to ferromonzodiorite that becomes mixed with irregularshaped masses of granophyre (Fig. 23). The host ferromonzodiorite still locally displays a very prismatic texture (Fig. 23A). Progressing north across this mixed zone, the granophyric and dioritic components become more strongly contrasted (Fig. 23B  Fig. 23C) with the intermediate component becoming dominant over granophyre and more ferrodioritic in composition (less granophyric mesostasis). A B C Figure 23. Exposures of complexly mixed granophyre and ferromonzodiorite observed at Location 14 progressing north. A) irregular masses of granophyre mixed with prismatic ferromonzodiorite. B) Broader scale view of complexly mixed granophyre with ferromonzodiorite. C) Granophyre occurring as irregular dikes and dikelets in more homogeneous ferrodiorite. Location 15: Ferrodiorite of the Intermediate Zone UTM(NAD83): ~ 571685_5184160 Description: At this point along the shore, the dominant rock is a dark ferrodiorite with only local occurrences of irregular granophyre masses. This is the main rock type of what other workers (Oestrike, 1983; Gardner, 1987; Jerde, 1991) have termed the intermediate zone. Petrographically, it is a mediumgrained, subprismatic to poikiloprismatic quartz ferrodiorite to ferromonzodiorite with significant 99 �amounts of apatite (4%), tridymite paramorphs (15%), and primary brown amphibole as rims on pyroxene prisms (Fig. 24). Strongly zoned pyroxene and plagioclase display preferentially altered cores and fresh rims (Fig. 24B). A B A’ B’ Figure 24. Photomicrographs of Samples MD773E (A) and MD773F (B) displaying the subprismatic to poikiloprismatic texture of the quartz ferromonzodiorite/ferrodiorite that forms the Intermediate Zone of the Endion Sill. Scale Bar = 3mm. The ferrodiorite is cut here by a 0.5-1.5 meter wide diabase dike trending to the NE and dipping steeply to the NW. The diabase shows well-developed columnar joints that are curved indicating fault motion before the dike completely cooled (Fig. 25). Figure 25. Columnar-jointed diabase dike cutting ferromonzodiorte at location 15. Location 16: Contact of Intermediate Zone and Recrystallized Rhyolite UTM(NAD83): ~ 571710_5184240 100 �Figure 26. Panoramic view of the contact between ferrrodiorite and granophyre (recrystallized rhyolite) at Location 16. Description: Exposed in the sloping ledge at this location is the subhorizontal (slightly NE-dipping) contact between the medium fine-grained ferrodiorite and medium fine-grained granophyre (Fig. 26). In detail, the contact is very irregular and interfingering with granophyre occurring as irregular masses to dike-like bodies in the ferrodiorite. In the overlying granophyre, some lobate masses of fine-grained ferrodiorite suggestive of two magma mixing textures. The granophyre transitions into flow banded rhyolite in the seacliffs that line this bay (inaccessible except by boat). This transition is observed in Tischer Creek, where it is observed to occur over several meters (Ernst, 1955; Oestrike, 1983). References Ernst, W.G., Jr., 1955, Petrology of the Endion Sill. M.S. thesis, University of Minnesota-Minneapolis, 31p. Ernst, W.G., Jr., 1960, Diabase-granophyre relations in the Endion Sill, Duluth, Minnesota. Journal of Petrology, v. 1, p. 286-303. Gardner, J.E., 1987, Origin of the Endion Sill. M.A. thesis, Washington University, St. Louis, 359p. Gardner, J.E., Haskin, L.A., Brannon, J.C., 1987, Possible assimilation by a mafic magma: the Endion Sill, Duluth, Minnesota. Lunar Planetary Science Conference Abstracts, v. 18, p. 312-313. Green, J.C. and Miller, J.D., 2008, Bedrock geology of the Duluth quadrangle, St. Louis County, Minnesota. Minnesota Geological Survey Miscellaneous Map M-182, scale 1:24,000 Jerde, E. A., 1991, Geochemistry and petrology of hypabyssal rocks associated with the Midcontinent Rift, northeastern Minnesota; Appendix D – The Endion Sill. Ph.D. dissertation, University of California - Los Angeles, p. 278-288. Oestrike, R.W., 1983, The Endion Sill, Duluth, Minnesota: mineralogy and petrology of a composite intrusion. M.S. thesis, University of Illinois, Urbana-Champaign, 141 p. Schwartz, G.M. and Sandberg, A.E., 1940, Rock series in diabase sills at Duluth, Minnesota. Geological Society of America Bulletin, v. 51, p. 1135-1172. Sun, S, -S. and McDonough, W.F., 1989, Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: A.D. Saunders and M.J. Norry (eds.), Magmatism in Ocean Basins. Geol. Soc. London, Special Publ., v. 42, p. 313-345. 101 �FIELD TRIP 6 May 6, 2016 GEOLOGY AND TROUT FISHING ALONG AMITY CREEK, DULUTH Dean Peterson Peterson Geoscience LLC George Hudak UMD-NRRI Figure 1. Bedrock geology and field trip stop map of Amity Creek, Duluth, Minnesota. 102 �Introduction: Circa 1985…. was the time, as seemingly all of the University of Minnesota Duluth (UMD) geology undergraduate and graduate students were aware, of monthly handouts of the infamous 5-pound bricks of government “eutectic” cheese as well as bags of rice and other edible things that we cannot remember. As a poor geology undergraduate student living on these government handouts and potatoes at the UMD, the first author spent many many days on Amity Creek fishing for meat (brook trout) for dinner after class. Most trips were successful back then and still are today. This short field trip will include the geology of the Seven Bridges Road portion of Amity Creek as well as trout fishing tips and perhaps some stories and seemingly tall tales from two proud products of the geology department of UMD. We will end our trip with a BBQ where additional tales may be told. History of Seven Bridges Road: This field trip largely follows the path of Duluth pioneer and former mayor Samuel Snively, who was instrumental in the development (late 1890s) and construction of our roadway, Seven Bridges Road. Seven Bridges Road is one of many parkways in Duluth, a city renowned for its parks and outdoor amenities. We will travel up Seven Bridges Road to Hawk Ridge and onto Duluth's famous hilltop boulevard, Skyline Parkway. Winding our way up Amity Creek through a mixed forest of pine, fir, maple, and birch you will see how the road intertwines and crosses with the creek over seven stone-arch bridges, and the reason for the road’s current name. Each bridge is faced with local 1.1 billion year old basalt and diabase collected from the creek bed or blasted from nearby outcrops and the pink cap rocks consist of granite quarried in St. Cloud, Minnesota. The following historical description of the development of Seven Bridges Road is modified from Ryan (1999). Work on the original section of the road was begun in 1899, and opened for use in the next year, though it took over three decades before the upper connection across Hawk Ridge was finally completed. The drive was built by Samuel Snively, a Duluth pioneer who owned a large 400 acre farm in the back country above Duluth suburbs of Lester Park and Lakeside. His hilltop farm was well-known for its thoroughbred stock and the beauty of its layout, which included a glimpse of Lake Superior through the nearby Amity Creek valley. Snively apparently often hiked the valley and explored the woods on both sides of the ridge line overlooking the east end of Duluth. During these strolls, Snively began to envision a park drive that would rival any other in Duluth. Donating sixty acres of his own property, Snively set to work contacting all the other landowners in the area, successfully garnering donations of the necessary rights-of-way for his road, as well as some of the necessary monies to build it. A crew of workers from the surrounding countryside was hired and construction on Snively's road began in the late fall of 1899 and continued into the following summer. The road crew started its work at the junction of Oriental and Occidental boulevards near Lake Superior, two carriage paths that ran through and alongside the boundaries of Lester Park, a city park in east Duluth. Despite its popularity as a scenic parkway, the city of Duluth neglected to maintain Snively's road, and within a decade all the wooden bridges had fallen to ruin, making the road impassable to vehicle traffic. In 1910, the road's destiny changed for the better when it was handed over to the Duluth's park commission, and a new plan for its rejuvenation was developed. The park board hired an architectural landscaping firm to design a new series of bridges for the road. In the fall of that year, the firm of Morell & Nichols of Minneapolis presented the park commissioners with sketches and blueprints for nearly a dozen new stone-arch bridges to replace the wooden ones Snively had built. During 1911, the roadway was regraded and graveled, and several first class stonemasons from the Duluth area were hired to build the bridges simultaneously. When completed, the newly refurbished road would become an official part of the Duluth's boulevard system. News of the park board's intentions delighted Snively, for the plans were exactly what he had in mind when he first built the road. 103 �When Snively's road reopened on July 6th, 1912, it was renamed Amity Parkway and added nearly 6 miles to the city's growing boulevard system. The new Amity Parkway became a popular destination for tourists and locals alike. Winding its way up into the eastern hills, the route presented many scenic sights of the landscape and rushing creek. Flowers lined many of the drive's turns and curves, and in the autumn, the poplar and birch forests presented spectacular colors for the tourist. "When the park board decided to take over and improve this roadway, it greatly pleased me, for it assured the consummation of the very purpose I had in view, the appropriation by the city for park and boulevard purposes of some of the scenic and natural park property in and about the city...Our possible park system rightly developed will be the city's greatest asset and advertisement." Samuel Snively Satisfied for the time being, Snively moved on again to other things in his life, but would return some two and a half decades later, (this time as the mayor of Duluth) to build the final leg to his road, the segment of the road leading to Hawk Ridge. Two of the bridges (#8 & #9) fell into disuse (by automobiles) when the eastern extension Winter Bridge to and along Hawk Ridge was completed in the thirties. Remnants of these two closely-built bridges still stand along a pedestrian pathway that shoots off from the main road near the last bridge crossed before the road ascends toward Hawk Ridge Nature Reserve. The remaining seven bridges are still used by vehicle traffic, but years of weather, combined with vandalism, and vehicle accidents had taken their toll. In the mid-1990s the city of Duluth, realizing the historic significance of the bridges, initiated a program to repair and restore the structures. Bridge #2, just south of the Lakeview hockey rinks, being the most damaged of the lot, was the first to be restored. The original blueprints were consulted with the work beginning in late 1996, and completed the following summer. The bridge was restored to its original condition, and the project was hailed a success. In the spring of 1998, the Duluth Preservation Alliance awarded the restoration with a plaque at its annual awards ceremony. Bridge #6 restoration work was begun the following year, a century after Samuel Snively began construction on the original road. Repairs to the remaining five bridges are slated to take place over the next few years. Although in fairly rough condition and its paved section in need of resurfacing, Seven Bridges Road remains one of Duluth's more idyllic drives. Traveling the road, you'll often meet hikers and bicyclists, equestrians, and automobiles. Fisherman can be spotted angling for trout along the creek bank, and during the warmer days of summer, swimmers are often seen cooling themselves in pools such as the one situated just beneath the falls near Historic Bridge #6. During the winter months, snowmobilers share the route with hikers, cross-country skiers, and snowshoe enthusiasts. 104 �One day, back in early November of 1934, Sam Snively stood along Hawk Ridge overlooking eastern Duluth and the blue expanse of Lake Superior. Much of his celebrated farm, less than a mile away, had been destroyed sixteen years before in a devastating forest fire that had swept through the area. He sold the property soon after. Now, as he stood there, fast approaching his 75th birthday, and well into in his last term as mayor, he contemplated his long life in Duluth. "Sometimes, when I become discouraged, I say to myself, I should have gone to another city to seek my fortune. But when I look over these hills and see the great natural beauties of our community, I console myself and wonder--where in all this wide world could I find such a view as this?" Samuel Snively Geologic Setting: As Amity Creek decends the steep hillside of Duluth to Lake Superior, it flows over mafic and felsic lava flow sequences and hypabyssal diabase dikes and sills of the 1.1 Ga. North Shore Volcanic Group as well as locally over red glacial rift. The variability in resistance to weathering and erosion of these rocks has lead to the varied character of Amity Creek and its most famous inhabitants, the brook trout. We will spend a late afternoon exploring the rocks, the character of the creek, and the possibility of hooking some trout. The authors wish to let all of the field trip participants know that neither of us have ever formally mapped the geology we will be looking at during this short Duluth-centric field trip. The experts on the geology of this area are University of Minnesota Duluth geology professors Dr. John Green and Dr. James Miller, and thus we will be critiqueing their geology (Fig. 1) as shown on the Minnesota Geological Survey Miscellaneous Map M-182 (Green and Miller, 2008) during this field trip. However, as two experienced NE Minnesota geologists we can perhaps answer most questions and will explore the geologic details during the field trip simply along with all of the participants. Angling Setting: Unlike the trout fishery of southeastern Minnesota’s driftless area, the streams of Minnesota’s north shore of Lake Superior generally are only fair trout streams. These waterways depend on unstable runoff for their flow and surge after large rain events and during spring snowmelt which can dwindle to a trickle during drought and the winter season. In the summer some stretches (especially the upper portions of the stream systems) can get warmer than is best for trout. The 1.1 Ga. volcanic bedrock over which the North Shore streams flow contain few of the water-soluble minerals that help keep the water alkaline and the aquatic invertebrate population large. Consequently, these streams tend to be soft, slightly acidic to neutral, and not particularly productive. Because they generally lack spring water, the streams get very cold in winter and can form "anchor ice" on the bedrock of the streambed, destroying aquatic life and habitat. However, North Shore streams have two things in their favor for the development of a trout fishery. The first is the cool Lake Superior-moderated climate and the second is the deep shade provided by the ubiquitous forest bank cover, which generally keeps these streams just cool enough to support trout. Trout are not native to the upper reaches of the North Shore streams. Coaster brook trout occupied Lake Superior and ascended the rivers as far as the first barrier falls-usually less than a mile from the lake. Only during the last century have brook trout, rainbow trout, and locally brown trout been stocked above the barrier falls of North Shore streams. 105 �FIELD TRIP STOP DESCRIPTIONS NOTE: all UTM coordinates are given in NAD 83, Zone 15 and keyed to Figure 1. STOP 1—Lakeside lavas and the confluence of Amity Creek and the Lester River at Lester Park Location: UTM NAD83 coordinates 575795E, 5187900N General description: The confluence of Amity Creek and the Lester River occurs at Duluth’s Lester Park, where these two stream systems flow on undifferentiated mafic flows of the Lakeside Lavas (Green and Miller, 2008). The lavas consist of dark gray to brown, aphyric to sparsely porphyritic basalt and basaltic andesite lava flows. Individual flows are generally 5 to 30 meters thick, with an amygdaloidal upper zone and smooth (rarely rubbly) upper surface. Phenocrysts, where present, are predominantly plagioclase, with minor altered olivine, magnetite, and augite. Angling tip: Lester Park provides numerous angling opportunities (young steelhead and occasional brown or brook trout) with the best chances of success in pools immediately downstream of the massive interiors of lava flows. The park’s streams also seasonally (spring and fall) receive runs of large adult salmon, steelhead, and kamloops rainbow from Lake Superior. Anglers should be aware that only lures with single hooks are allowed at Lester Park. Figure 2. The lower Lester River at Lester Park. STOP 2—Lakeside lavas waterfalls of Amity Creek at “The Deeps” Location: UTM NAD83 coordinates 575470E, 5188350N General description: A very short walk downstream from our parking spot leads us to numerous waterfalls and deep pools associated with the erosion of the Amity Creek diabasic basalt. This huge (100 meters thick) lava flow is a dark gray to brown, fine- to medium-grained, locally seriate to porphyritic, heterogeneously textured basalt flow. The flow contains up to 10 percent phenocrysts of thin plagioclase tablets and minor phenocrysts of ilmenite, magnetite, and oxidized olivine. Groundmass is intergranular/intersertal to ophitic to felty and locally diktytaxitic, and contains plagioclase, augite, oxidized olivine, tabular ilmenite, and magnetite, and a mesostasis including K-feldspar, quartz, apatite, and chlorite. Diktytaxitic cavities contain chlorite, quartz, calcite, and laumontite. The massive nature of the Amity Creek diabasic basalt has lead to the formation of numerous waterfalls and deep pools, which are favorite swimming/jumping/diving spots for local Lakeside area kids. 106 �Angling tip: Within these pools prowl some of the largest stream-bound rainbow trout of the whole North Shore of Lake Superior in Minnesota. The summer angler must get up early in the morning (and bring a net) to fish these pools as swimmers always abound later in the day. The bridge we will cross and immediately park after is the Posted Boundary of Amity Creek. Below this bridge anglers are only allowed to fish lures with single hooks, while above the bridge treble hooks are allowed. Figure 3. Waterfall of Amity Creek into “The Deeps”. STOP 3—Resistant diabase dike in icelandite lavas forming the “Rainbow Trout Pool” Location: UTM NAD83 coordinates 575410E, 5188000N General description: A steep hike down from Seven Bridges Road is needed to observe first hand the diabasic Lakeside intrusion. The intrusion here is a black, fine- to medium-grained, plagioclase-phyric, intergranular olivine diabase with abundant apatite, interstitial quartz, and K-feldspar. This exposure shows it to be a vertical dike cutting the Lester Park icelandite lava flow. On lakeshore exposures, the western contact is crosscutting the Lester Park icelandite. These relationships are interpreted to indicate that the diabase came in along a high-angle reverse fault. The resistant diabase dike forms a beautiful waterfall and large pool in Amity Creek immediately downstream from the dike. Angling tip: Small to moderate size (8-13 inches) rainbow trout are abundant in the pool and several casts of a 1/32 oz panther martin spinner into this pool should always be tried by anglers. Anglers should always remember to lower your rod tip when the hooked rainbow trout jumps out of the water. There have been days when the first author hooked rainbow trout on his first 12 casts. Figure 4. Subvertical diabase dike and the “Rainbow Pool”. 107 �STOP 4—Flaggy weathering of icelandite lava and the meandering of Amity Creek Location: UTM NAD83 coordinates 575300E, 5189465N General description: A stop to observe the large Lester Park icelandite, a pink, red, or tan, porphyritic icelandite lava flow, about 180 meters (590 feet) thick. Mostly massive, but the upper part is strongly flow-laminated and folded, with local large, round vugs containing calcite, quartz, barite, and fluorite. Phenocrysts (5 to 10 percent) are of plagioclase, oxidized Fe-silicate, magnetite, zircon, and apatite. The groundmass contains poikilitic quartz (paramorphs after tridymite), K-feldspar, oxidized Fe-silicate, and magnetite, calcite, and chlorite. Angling tip: Perhaps we as a group we can locate the best brook trout pool since the first author has never had much luck in the past angling in this portion of Amity Creek. STOP 5—Waterfalls and pools off of the Northland Sill and the excellent brook trout fishing trail along the upper portions of Amity Creek Location: UTM NAD83 coordinates 574970E, 5190135N General description: Brook Trout Central!! This portion of Amity Creek is, in the opinion of the first author, the best place to first start angling on Amity Creek. The geology is dominated by the contact of the highly resistant Northland Sheet intrusion and the Amity Creek Icelandite lava flow, which culminates for the angler in a series of waterfalls and excellent brook trout pools. The Northland sheet is a brown, fine- to mediumgrained, diktytaxitic intergranular/intersertal diabase grading to augite quartz ferromonzonite. The diabasic intrusion contains minor traces of low-Ca pyroxene, olivine, and primary hornblende and abundant magnetite, ilmenite, and apatite. The sheet-like intrusion is variable in thickness and cuts across more than 500 meters (1,640 feet) of volcanic flows of the Lakeside lavas. Angling tip: At this stop, we’ll investigate waterfalls and excellent brook trout fishing pools situated at the uppermost portions of the Amity Creek icelandite lava flow as well as those within the basal portions of the diabasic Northland Sheet. An excellent horse/hiking trail heads upstream from our parking spot and anglers should be aware that numerous large hook-jawed brook trout swim in these waters. If one is a believer in “take a kid fishing”, then perhaps this is one place to begin (Fig. 5). Figure 5. His smile will last forever, Nathan’s first trout, 2015 (nephew of Dean M. Peterson). 108 �REFERENCES Green, J.C. and Miller, J.D., 2008, Bedrock geology of the Duluth quadrangle, St. Louis County, Minnesota: Minnesota Geological Survey, Miscellaneous Map M-182, scale 1:24,000. Ryan, Mark, 1999, The history of Duluth, Minnesota’s Seven Bridges Road, Samuel Snively and the building of a Northern Minnesota parkway: http://www.amitycreek.com/sevenbridges/index.html. 109 �FIELD TRIP 7 May 7-8, 2016 ARCHEAN AND PROTEROZOIC GEOLOGY OF THE WESTERN GUNFLINT TRAIL Mark Jirsa Minnesota Geological Survey Figure 1. Regional geologic map of northeastern Minnesota showing the location of western Gunflint Trail area (modified from Jirsa and Miller, 2005). Inset box outlines Figure 2. INTRODUCTION This field trip along the western end of the Gunflint Trail explores Neoarchean, Paleoproterozoic, and Mesoproterozoic rocks, and a diversity of well displayed unconformable and intrusive contact relationships. The trip is modified from the ill-fated one attempted for the 2007 ILSG meeting in Lutsen (Jirsa and Weiblen, 2007) that was canceled due to outbreak of forest fire. For expediency, some of that field guide is repeated here. In addition, this guide borrows heavily from two other field trips: a 110 �workshop on iron-formation hosted by the Precambrian Research Center (Jirsa and Fralick, 2010), and a Geological Society of America field guide that focused on the Sudbury Impact Layer (Jirsa and others, 2011). Readers should consult those publications for additional information and references. The order and number of stops that will be visited during this trip will be determined by time, weather, and access issues. The bedrock geology of the field stops is portrayed on a recent map of the Western Gunflint Trail area published as Minnesota Geological Survey Miscellaneous Map M-191 (Jirsa, 2011; Fig. 2). In this area, the Neoarchean greenstone-granite terrane of the Wawa subprovince of Superior Province is represented by a succession of metavolcanic rocks (~2720 Ma) known informally as the Paulson Lake volcanic sequence, intruded by the Saganaga Tonalite (~2690 Ma) and Paleoproterozoic diabasic dikes. The Neoarchean and diabasic rocks are unconformably overlain by Paleoproterozoic sedimentary strata of the Animikie Group (~1870-1830 Ma), which includes the Gunflint Iron Formation. The stratigraphic top of the iron-formation is marked by seismically deformed and brecciated strata and ejecta—known collectively as the Sudbury Impact Layer—that resulted from a meteorite impact near Sudbury Ontario (~1850 Ma). A disconformity separates the impact layer from overlying siltstone and graywacke of the Paleoproterozoic Rove Formation (~1835 Ma). Mesoproterozoic rifting is manifest in hypabyssal dikes and sills of the Logan intrusions (~1115 Ma), and several phases of the Duluth Complex (~1100 Ma), emplaced into both the Archean and Paleoproterozoic rocks. Figure 2. Geologic map (Jirsa, 2011) of the western Gunflint Trail (Highway 12 dashed), showing pertinent features of geology and field trip stops. Note that stops 2 and 3 lie just off the northwest corner of the map. Sudbury Impact Layer is reddish; Proterozoic dikes are shown as thin red lines; Logan intrusions are shades of purple. Image reduced from 1:24,000 map scale. 111 �GEOLOGIC SETTING Neoarchean The oldest rocks exposed in the region are Neoarchean metavolcanic and metasedimentary strata that are part of the Wawa subprovince of the Superior Province. They are probably equivalent to, but not demonstrably continuous with, the Ely Greenstone and Newton Lake Formation. Although the supracrustal successions are dissected by faults and intrusions, some correlation can also be made with adjacent terranes in Ontario. Most recent regional mapping in this terrane was conducted by Jirsa and Miller (2004) in the 1:100,000-scale, Ely-Basswood Lake map sheet that lies just west of the Gunflint Trail. A more detailed map of the Cavity Lake forest fire area that lies just west of the Gunflint is still in review and production (Jirsa and others, in prep.). Rocks of the terrane are divided into a number of faultbounded segments, each having distinct geologic characteristics that cannot be easily correlated from place to place. The Gunflint Trail area exposes the eastern edge of what Gruner (1941) referred to as the Gabimichigami segment. Supracrustal rocks in this segment include an older suite of variably pillowed, mafic to ultramafic flows and hypabyssal intrusions of the Paulson Lake volcanic sequence (stop 1), and a younger suite of hornblende-bearing andesitic to dacitic pyroclastic and volcaniclastic rocks that comprise the Knife Lake Group that are exposed just east of Fig. 2. Based on stratigraphic facing directions established from pillowed metabasalt flows, the Paulson Lake sequence forms an east-trending and steeply south dipping and younging homocline. The Saganaga Tonalite (stops 2-6) was emplaced into metabasaltic rocks and defines the northern edge of the supracrustal succession. The metamorphic grade of greenstone along this contact is locally increased from greenschist facies that is typical of much of the belt, to amphibolite grade, and foliation in both tonalitic and volcanic rocks is well-developed near the contact. The western edge of the Gabimichigami segment is terminated by a north-northeast-trending fault that juxtaposed greenstone against sedimentary and volcanic rocks of the Knife Lake Group. The Knife Lake Group includes the informally named “Ogishkemunce conglomerate” that contains detrital clasts of the Saganaga Tonalite, along with clasts of iron-formation, metabasalt, metagabbro, and gneiss. This distinctive sequence of conglomerate, sandstone, and alkalic rocks is interpreted to have been deposited in a complex array of successor basins developed along early-formed faults at some time after emplacement of the Saganaga Tonalite at ca. 2690 Ma (Driese and others, 2011; Jirsa, 2016). Lacking detailed geochronologic data for this immediate area, much of the temporal distinction between various geological elements of the Neoarchean bedrock is based on the correlation of U-Pb zircon dates acquired elsewhere with regionally developed fabrics and structures that resulted from three major phases of deformation, denoted D1, D2, and D3. All three deformation events are the result of N-S- to NW-SEdirected compression. The timing of D1 deformation is bracketed between deposition of the metabasaltic and associated rocks of the Wawa subprovince at ca. 2722 Ma (Peterson and others, 2001), and emplacement of the Saganaga Tonalite at ca. 2690 Ma. Folds attributed to D1 deformation in the Ely Greenstone and related rocks are truncated by faults associated with Knife Lake strata, indicating that the latter is synchronous with or post-dates deposition and early deformation of the Ely. As such, the Knife Lake Group is inferred to be a Timiskaming-type extensional basin sequence temporally equivalent to the Shebandowan assemblage exposed in adjacent parts of Ontario (Lodge and others, 2013; Jirsa and others, 2016). D2 deformation and metamorphism affected all of the Archean supracrustal units and can be crudely bracketed by U-Pb dates of intrusions in the Giants Range batholith to the southwest that place the regional deformation and metamorphic event between about 2674 Ma and 2685 Ma (Boerboom and Zartman, 1993). Thus, the age of sedimentary and volcaniclastic rocks is confined to a 10 million year duration between the emplacement of tonalite at ca. 2690 Ma and D2 metamorphism and deformation at ca. 2680 Ma. In this region, D3 deformation is manifest as crenulations and faults in rocks affected by D2. 112 �Paleoproterozoic Mafic dikes Mafic dikes emplaced into the Saganaga Tonalite are prominent on aeromagnetic maps as positive linear anomalies trending northward and eastward. Exposures mapped along the magnetic trajectories (thin red lines on Fig. 2) indicate that the dikes vary from diabasic to lamprophyric, and from a meter or less in width, to more than 30 meters. The roadcut at stop 4 exposes diabase inferred to lie along one of the anomalies. The precise age of the dikes is unclear, and there may be suites representing more than one age. At least one of the northwest-trending dikes is unconformably overlain by, and shed fragments into conglomeratic strata of the basal Paleoproterozoic Animikie Group, indicating a syn- to prePaleoproterozoic age. Animikie Group Sedimentary rocks of the Paleoproterozoic Animikie Group are exposed in an east-trending belt that extends from Thunder Bay on Lake Superior to a point 12 miles (19km) west of the Gunflint Trail, where the belt is truncated by the Mesoproterozoic Duluth Complex (Fig. 1). The Animikie Group in this area consists of locally developed basal conglomerate and sandstone, and iron-bearing strata of the Gunflint Iron Formation. The stratigraphic top of the iron-formation is marked by a major unconformity and ejecta that resulted from a meteorite impact near Sudbury Ontario ca. 1850 Ma (Krogh, 1984; Davis, 2008). Of the 188 known and scientifically verified terrestrial impacts, the Sudbury event is the third largest (based on crater size) and forth oldest (TUwww.unb.ca/passc/ImpactDatabaseUT). The resulting ejecta blanket has been identified in Ontario (Addison and others, 2005), Michigan (Pufahl and others, 2007; Cannon and others, 2010), and here in Minnesota near Gunflint Lake (Jirsa and others, 2011), and in drill core along the Mesabi Iron Range (Addison and others, 2005). The overlying Rove Formation—a mudstone and turbiditic sandstone unit—was deposited directly on the ejecta. The sequence is broadly correlative with Animikie strata exposed to the southwest along the Mesabi Iron Range. The rocks form a homocline that dips gently southward, except where deformed by folding, faulting, and emplacement of Mesoproterozoic intrusions. For example, local folding, inferred to be associated with emplacement of the Logan intrusions, may be the product of magmatic delamination and shouldering of the sedimentary country rocks. Sequence "inflation" by the Logan sills may explain the observation that the dip of Paleoproterozoic rocks increases from 10º away from the contact with Duluth Complex, to 60º in some places near it. Faults are present locally, but few have displacements greater than 50 feet. A notable exception is the Lookout fault (Fig. 2) that crosses the Gunflint trail. As much as 200 feet of uplift on the west and south is speculated (Morey and others, 1981; Jirsa, 2011). In addition, the dip of Animikie strata on the west side of the fault is much steeper than that on the east, which explains in part the difference in the widths of map units apparent from the geologic map. Much of the complex-looking fold pattern east of the fault is an artifact of moderately high topographic relief and shallowly dipping units. The Lookout fault displaces Animikie strata; however the relationship between faulting and the emplacement of Logan and Duluth Complex intrusions is unclear. Gunflint Iron Formation (field trip stops 7-10) As on the Mesabi Iron Range, the Gunflint Iron Formation has historically been subdivided into so-called cherty (granular) and slaty (argillaceous) informal subunits or members. The members are denoted lower cherty, lower slaty, upper cherty, and upper slaty (Wolff, 1917). Although this terminology has some descriptive utility in the field, the subdivision employed here is based instead on a sedimentalogical model (after Pufahl and Fralick, 2004). In this model, depicted in Figure 3, the lower cherty and parts of the lower slaty members represent deposition during a single marine transgression. This was followed by a regression that deposited the lower part of the upper cherty member—the resulting sedimentary strata are collectively and informally termed lower sequence here. The upper part of the upper cherty represents the onset of a second transgression that continued through deposition of the thick upper slaty member, 113 �and is collectively termed the upper sequence here. The contact between the two sequences is a diastem inferred to represent a period of maximum regression. The initial stages of the second transgression is marked by intraformational conglomerate containing oncoliths, fragments of what appear to have been semi-lithified grainstone derived from the lower sequence, and both in-place and dislodged stromatolites. The uppermost strata of iron-formation are variably brecciated and/or chaotically folded, carbonatebearing, and capped by granular ejecta from the ca. 1850 Ma Sudbury meteorite impact event, collectively termed Sudbury Impact Layer here. More detailed descriptions are given below: Lower sequence—Irregularly graded sequence recording marine transgression, followed by regression. It grades from conglomerate and sandstone at the base, unconformably overlying Neoarchean bedrock; to locally stromatolitic, siliceous grainstone; to interlayered, laminated to massive chert, to iron-rich mudstone, and finally to siliceous grainstone. Total thickness is approximately 50 m. The basal part of the sequence is marked by discontinuous conglomerate and minor fine- to medium-grained quartzofeldspathic sandstone that is typically thinner than 1 m. Conglomerate contains pebbles to small cobbles of quartz, Saganaga Tonalite, metabasalt, and diabase. Thicker sections of this facies exposed in Canada are known as the Kakabeka Conglomerate. The uppermost siliceous grainstone forms prominent ridges. It appears to have been partially lithified prior to deposition of, and contributed grainstone fragments to, the basal part of upper sequence. Upper sequence—Siliceous grainstone and laminated chert; locally contains stromatolitic and intraclastic conglomerate at base of the sequence; which grades irregularly up-section to increasingly mudstone-rich; and typically parallel-laminated to wavy-bedded. Total thickness is approximately 45-55 m. Reworked volcaniclastic zircons from the upper sequence exposed in Ontario yielded a U-Pb age of 1878±1 (Fralick and others, 2002). Sudbury impact layer (SEE Discussion below)—Brecciated and complexly deformed iron-formation as much as 10 m thick, overlain locally by less than 1 m of mesobreccia and granular ejecta. Both deformed (seismically shattered and chaotically folded) iron-formation and ejecta are inferred to be related to the Sudbury meteorite impact event (Jirsa, and others, 2011). The macroscopically most apparent feature of ejecta is the presence of 0.1-1.0 cm, concentrically zoned spheres inferred to be accretionary lapilli. Microscopic evidence that this material has an impact origin includes rare occurrence of quartz fragments marked by planar deformation features. Metamorphism here in the contact aureole of the superjacent Duluth Complex presumably has obscured or obliterated other diagnostic attributes (e.g., French and Koeberl, 2010). 114 �Figure 3. Schematic stratigraphic section of Gunflint Iron Formation and adjacent rocks comparing older stratigraphic nomenclature in left column with that used informally here on right. Siliceous grainstone (cherty) units in the iron-formation are represented by pale color; iron-rich mudstone is darker. Approximate stratigraphic positions of field trip stops are shown in boxes. DISCUSSION OF SUDBURY IMPACT LAYER (Field Trip stops 11-15) The Gunflint Lake exposures lie some 480 miles (770 km) west of Sudbury, making this one of the most distant sites known to contain what is considered “proximal ejecta” from the ca. 1850 Ma Sudbury meteorite impact event. Similar deposits have been discovered in Thunder Bay, Ontario, and Michigan that are well documented by Addison and others (2005), Cannon and others (2010), and Pufahl and others, (2007). Deposits near Gunflint Lake appear to be consistently thicker than in other areas, even though these sites are more distal than those in Michigan and Ontario. The impact deposits at sites further away from the crater than Gunflint Lake are much thinner and lapilli are only rarely present. This has led Addison and others (2010) to hypothesize that the Gunflint Lake deposits may represent thick ramparts, 115 �as described for end-of-flow Martian base-surge deposits (Kenkmann and Schonian, 2006; Osinski, 2006; Mouginis-Mark and Garbeil, 2007; Fralick and others, 2012). It should be noted at the onset, that only a small portion of the material described here can be considered true ejecta; i.e., air-borne detritus derived from the impact site. The great majority of the 7 meter-thick deposit is breccia that consists of thoroughly disheveled fragments that appear to have been derived from subjacent iron-formation. Like the deposits near Thunder Bay, the breccia is sandwiched between Gunflint Iron Formation and sedimentary strata of the Rove Formation. Unlike deposits near Thunder Bay, the breccia lies within the metamorphic aureole of the Mesoproterozoic Tuscarora and Poplar Lake Intrusions of the Duluth Complex (ca. 1100 Ma), and is intruded by diabasic sills of the Logan Intrusions (ca. 1115 Ma; Heaman and Easton, 2005). Pervasive carbonate mineralization and metamorphism has overprinted and obscured much of the original, delicate mineralogic features, but macroscopic textures and geochemical content that convey information about protolith and depositional mechanisms are preserved. In the following discussion, the term Sudbury Impact Layer (SIL) is applied to all facies of sedimentary rocks inferred to have been formed or deformed in response to the ca. 1850 Ma Sudbury meteorite impact. By that definition, it includes autochthonous material interpreted to be seismically folded and shattered iron-formation (ejecta-absent), and overlying strata composed largely of allochthonous material derived from the impact site (ejecta-bearing). In no single outcrop are all facies present; however, an approximation of temporal relationships can be inferred from the juxtaposition of two or more facies in individual outcrops (Fig.4). Figure 4. Stratigraphic framework derived from 8 exposures along a 2 mile strike-length near Gunflint Lake; hung from the contact (bold dashed line) between ejecta-bearing (upper) and ejecta-absent (lower) facies of the Sudbury Impact Layer. Facies below are described in apparent stratigraphic order from oldest to youngest: Ejecta-Absent UContorted iron-formation facies:U The uppermost layers of iron-formation are chaotically folded and exhibit both ductile and brittle behavior in close proximity at the scale of individual outcrops. The rheologic response depends on the apparent rigidity of material at the time of deformation. Silica-rich layers display brittle, shattered to semi-ductile, boudinage-like textures. By contrast, much of the iron-silicate mudstone layers behaved in a ductile fashion, locally showing evidence of 116 �fluidization and injection into superjacent strata. Folds are non-systematic in trend and style, and multiple hinge detachments occur. These attributes counter-indicate a regional tectonic origin, and instead are best viewed in the context of impact-generated seismicity imposed on semi-lithified substrate. UParautochthonous breccia facies:U At several locations, straight-bedded iron-formation passes laterally along strike into irregular zones in which the silica-rich layers have been broken and disheveled, while still retaining some semblance of jigsaw-puzzle fit. UMegabreccia faciesU: This term is used for breccia composed of unsorted slabs (as large as 5 m), blocks, and smaller fragments of iron-formation. The fragments are angular and in most places have random orientations. Fragments of green, iron-silicate mudstone typically show some evidence of semi-ductile behavior, and locally this material was fluidized to form irregular matrix and clastic (muddy) dikes. Ejecta-Bearing Mesobreccia facies: This is fragmental rock containing angular, subrounded, and amoeboid clasts (up to 5 cm long) of dark green material and scattered accretionary lapilli. Petrography shows that much of the original structure of the clasts has been metamorphically recrystallized and annealed; however, relicts of amygdaloidal and fluid-looking textures remain—implying a glassy protolith. Lapillistone-gritstone facies: Accretionary lapilli as large as 1.5 cm occur as irregular masses and layers interbedded with sandy to silty gritstone. In areas least affected by metamorphism, several grains of shocked quartz with planar deformation features have been identified in thin section. Spherule, pellet, small lapilli facies: The upper parts of ejecta horizons locally contain layers, lenses, and interbeds of accretionary or relict glass grains that are smaller than typical lapilli and generally lack concentric zonation. These apparently accreted particles may represent waning ejecta plume deposition. Ejecta-bearing conglomerate facies: In a few localities, the uppermost part of impactite deposits consist of conglomerate containing subrounded fragments of iron-formation (in contrast to anglular fragments typical in breccias described above), and matrices containing variably abraded lapilli. U U U U U U U U The apparent contrast in rheologic response to seismicity between siliceous and interbedded “muddy” strata indicates that the relative competancy of the two sediment types was significantly different at the time of impact. One explanation is that the silicification process occurred very early, perhaps just beneath the sediment/water interface, and produced the more cohesive (but not yet fully lithified) siliceous layers. Seismic deformation brecciated those layers selectively, while folding and liquefacting the interbedded muds. This interpretation is relevant to the understanding of depositional environment. It implies that upper layers of the iron-formation were either in a shallow submarine setting or only recently emergent at the time of impact. The arrangement of facies described above and depicted on Figure 4 can be interpreted in the context of experimental evidence and observations from lunar and smaller terrestrial impacts. Using calculations from Collins and others (2005), based primarily on estimated crater dimensions, one can predict arrival times for various effects of the impact here, some 480 miles west of the impact site as follows: EVENT 1) Fireball 2) Earthquake 3) Ejecta Ground Surge 4) Air blast 5) Tsunami * APPROXIMATE ARRIVAL TIME 13 seconds (the modern equivalent of 3rd degree burns) 2-3 minutes (>10.9 at epicenter) 5-10 minutes (predicts ejecta 1-3 meters thick, grain sizes ~1cm) 40 minutes (sonic boom) 1-3 hours P 117 P �[*The latter is speculation, as the arrival time and effects of tsunami are dependent on pre-impact position relative to strand line, and basin bathymetry which is nearly impossible to establish.] Intuitively, only three of these events are likely to have produced a record in the rocks: earthquake, ejecta surge, and tsunami. Nearly all contacts between individual facies are gradational, with one very important exception—in all exposures, the boundary between ejecta-absent and ejecta-bearing facies is extremely sharp. This is inferred to reflect a fundamental shift in geologic processes from intense seismic perturbation of uppermost iron-formation represented by the ejecta-absent facies, to deposition by the passing ejecta plume. The uppermost conglomerate facies is inferred to represent mixing of local and exotic detritus, presumably by tsunamis or other post-impact fluvial or marine processes. Rove Formation The Rove Formation consists of carbonaceous, thinly bedded argillite to slate, and fine- to mediumgrained graywacke (stops 15 and 16). Primary sedimentary structures indicate turbidity current flow was dominantly to the south. The basal several meters of the formation are irregularly bedded, carbonate-rich, and locally conglomeratic (stop 15). Detrital zircons taken from lower parts of the formation in Ontario yielded ages as young as1827±8 (Addison and others, 2005), indicating some considerable hiatus separated deposition of the Rove from that of the underlying 1850 Ma Sudbury Impact Layer. Contact metamorphism The iron-formation, SIL, and overlying argillaceous strata of the Rove Formation were variably replaced by carbonate and metamorphosed by the Duluth Complex to amphibole and pyroxene hornfels. Floran and Papike (1978) delineated irregularly northwest-striking metamorphic zones recognized on the basis of the dominant iron-silicate mineral present in iron-formation. From least metamorphosed on the northeast, to most metamorphosed on the southwest, these indicator minerals are greenalite+minnesotaite, grunerite, hedenbergite, fayalite, and ferrohypersthene. Despite this metamorphism, macroscopic sedimentary textures are well preserved in most outcrops. Metamorphic effects adjacent to the Logan Intrusions are minor. Mesoproterozoic Mesoproterozoic mafic intrusions comprise the remaining exposures in the Gunflint Trail area. The rocks represent early magmatic stages of the Midcontinent Rift. The apparently earliest of these are diabasic sills and dikes emplaced into the Animikie strata and collectively referred to as the Logan intrusions. A baddeleyite age of 1115 ± 1 Ma is reported from a sample of a Logan sill near Thunder Bay (Heaman and Easton, 2005). The sills are intruded, with slight angular discordance, by medium to coarse-grained gabbro and troctolite of the Poplar Lake and Tuscarora intrusions of the Duluth Complex. The Poplar Lake is part of the early gabbroic series of the complex. A basal gabbroic unit of the intrusion yielded a date of 1106.9 ±.8 (Miller and Severson, 2002). Field relationships indicate that the Poplar Lake is intruded by the Tuscarora intrusion, which is considered to be part of the layered series of Duluth Complex. This is consistent with a U-Pb age of 1098.81±0.32 from a sample just west of the Gunflint Trail area (Hoaglund and others, 2010). Logan Intrusions The Logan intrusions (stops 11, 15, 16) are exposed along a series of prominent, east-trending ridges formed by the differential erosion of diabase sills and sedimentary rocks, particularly the Rove Formation. Individual sills are as much as 1100 ft (33m) thick, and can be traced along strike for several kilometers. Branching and merging of individual sills is common, and many sills thicken and thin down-dip. Some 118 �sills terminate against joints and inferred faults. Locally, fractures in the Rove Formation are occupied by thin dikes, which give a box-work configuration to the hypabyssal intrusions. Rock types include aphyric basalt, fine- to medium-grained diabase with ophitic clinopyroxene enclosing plagioclase, plagioclase cumulates, and granophyre (Jones, 1984). Outcrops in the field trip area commonly have plagioclasephyric phases (stop 11). Chilled margins form sharp contacts with, and locally contain inclusions of, the country rocks. Diabase coarsens to medium-grained near the center of individual sills, and clinopyroxene is ophitic throughout. Minor differentiation is manifest in accumulations of plagioclase or granophyric intergrowths (quartz, sodic plagioclase, and orthoclase) in upper parts of sills. Duluth Complex The Duluth Complex (Fig. 5) is a sequence of generally discordant plutonic rocks consisting of many separate intrusions. Two of these intrusions, the Poplar Lake and Tuscarora, are exposed along the Gunflint Trail. The Poplar Lake intrusion, formerly referred to as Nathan's layered series, consists of interlayered gabbroic cumulates, with minor amounts of troctolitic and anorthositic cumulates. Rocks of the Poplar Lake have reversed magnetic polarity, and thus are broadly correlative with lower lavas of the North Shore Volcanic Group and with the Logan intrusions. The Poplar Lake intrusion is composed of at least 27 sheet-like units of mafic cumulates and intermediate to felsic rocks (the so-called Nathan’s layered series). The Tuscarora intrusion irregularly cuts across the layered gabbro of the Poplar Lake intrusion (Morey and Nathan, 1978). The basal part of the Tuscarora intrusion consists of a fine-grained, augite-poikilitic, olivine gabbro (stops 17 and 18). Within 0.3 mi (0.5km) of the basal contact, finegrained troctolite coarsens to medium-grained. The troctolite units consist of 65-70 percent plagioclase and 10-15 percent olivine. Relative amounts of poikilitic augite and iron-titanium oxides vary locally. Orthopyroxene mantles olivine and occurs in symplectic intergrowth with plagioclase. Biotite is locally present in association with iron-titanium oxides. Modal layering is well developed and generally concordant with unit boundaries that dip gently to the south—typically more steeply dipping than the subjacent Animikie Group strata. The basal part commonly contains chalcopyrite, pyrrhotite, and minor pentlandite interstitial to plagioclase and olivine. The sulfide concentrations are subeconomic, but locally form mappable zones (stop 18). 119 �FIELD TRIP STOP DESCRIPTIONS NOTES: 1) Many of these outcrops are scientifically important—the subjects of on-going research, and all are on private or National Forest lands. For these reasons, we ask that you refrain from hammering and sampling without first checking with the leader. 2) This group is unlikely to visit all of the stops described below during this trip. A number of stop descriptions are included in this guide to provide context and for future visits to the region. 3) The stops are presented in general geochronologic order from oldest to youngest. 4) At the time of this writing, many of these stops had not been visited for several years. Regrowth after forest fires and other factors may preclude visiting some stops, and alternative locations may be substituted. 5) Some descriptions below show locations on images extracted from published 7.5-minute U.S.G.S. quadrangles. 6) All locations are given in NAD 83, Zone 15N UTM coordinates. Figure 5. Geologic map and schematic cross-section (A-A') showing the approximate geographic and stratigraphic positions of field trip stops (geology modified from Jirsa, 2011). Note that the horizontal distances on the cross section are much greater than those on the map, and the section is vertically exaggerated by 3.5 X, resulting in apparent dips of contacts steeper and units thicker than true. 120 �STOP 1—Neoarchean pillowed basalt and basal Gunflint Iron Formation Location: UTM locations of several substops along the Kekekabic Hiking Trail west of Hwy 12— Gunflint Trail (Fig. 5) are given below: 1a = Outcrops along the irregular unconformable contact between metabasalt and iron-formation in the vicinity of UTM 661970E/5328460N 1b = Outcrops near the junction of Kekekabic and Lookout trails at UTM 661230E/5328410N 1c= "Paulson Mine" at UTM 660980E/5328320N; about 900 feet west of junction of Kekekabic and Lookout trails General description: Follow the Kekekabic Hiking Trail from the parking lot, westward to several stops listed above, and perhaps others enroute. The Kekekabic trail parallels the base of the Gunflint Iron Formation where it rests unconformably on Neoarchean metabasalt. Several small outliers of the lower sequence (lower cherty member) of the iron-formation containing iron-silicates and magnetite can be found along the route, implying that this gently south-dipping surface is very near the unconformable contact between Neoarchean and Paleoproterozoic rocks. Exposures of metabasalt vary from massive to pillowed, autobrecciated, and locally variolitic. Analyses of a fine-grained hypabyssal intrusion associated with the metabasalt indicate that it has a komatiitic composition (Jirsa and Weiblen, 2007). Intrusions of this composition are also found in the Newton Lake Formation, some 30 miles to the southwest, and high-Mg tholeiitic basalt flows and pyroxenitic to peridotitic sills were described by Vervoort (1987) in the JAP Lake area two miles along strike to the west. Pillow shapes indicate moderate flattening by regional D2 deformation. Bedding trends to the east-northeast, and is steeply southward dipping and facing. The steep north-facing slope immediately south of the trail contains exposures of the lower sequence (lower slaty member). The iron-formation has been strongly metamorphosed in this area and now consists of various assemblages of quartz-grunerite-fayalite-magnetite and quartz-cummingtonitegrunerite-pyroxene-magnetite. Several test pits and shafts can be seen along more than a mile of the trail, including one that is fenced and labeled "Paulson Mine 1893" (stop 1c). In reality, this and other scattered shafts collectively made up the Paulson “mines.” They are developed in the lower sequence of iron-formation, and waste piles contain abundant pyrrhotite, other sulfide minerals, and magnetite. Despite construction of a rail line to Port Arthur (now Thunder Bay), only one train car of “ore” was ever shipped. Presumably the low iron and large sulfide content precluded further work, though the 1893 “financial panic” may also have played a role. STOP 2—Felsic phase of the Neoarchean Saganaga Tonalite cut by a small mafic dike. Location: UTM: 656333E/5335730N; End of the Trail Campground; Campsite #18 Description: Most of the outcrops in this area consist of gray, massive to trachytoid-foliated, medium to coarse-grained tonalite, having plagioclase in much greater abundance than microcline. Large quartz phenocrysts, or eyes, as much as 1 cm in diameter are characteristic of this phase, which is typical of 90% of the batholith (Fig. 6). Quartz eyes are polycrystalline aggregates, in which each crystal has a different optical orientation. Quartz also occurs as an interstitial mineral to subhedral plagioclase (An20-28). Small amounts of microcline occur as antiperthitic exsolution in plagioclase, as rims on plagioclase, and as small interstitial grains. Hornblende is the dominant ferromagnesian mineral, together with minor amounts of augite, biotite, epidote, and chlorite. 121 �Figure 6. Texture typical of the Saganaga Tonalite, including lineated quartz "eyes" (medium gray, trachytoid-aligned). The small dike of aphanitic mafic rock in this exposure has not been analyzed, but is inferred to be related to larger north-trending diabasic and lamprophyric dikes that form prominent north-trending anomalies on aeromagnetic maps. The pronounced foliation in the Saganaga Tonalite was inferred by Grout (1933) as a primary flow fabric (trachytoid). The tonalite is inferred to have been emplaced into Archean metavolcanic rocks shortly after early (D1) deformation, based a U-Pb date of 2689±1 Ma in Canadian exposures (Corfu and Stott, 1998), and 2690.83±0.26 Ma (Driese and others, 2011) just west of the Gunflint Trail. As such, it experienced major regional metamorphism and transpression associated with D2 deformation at ca. 2680 Ma. As with many large plutons in such terranes, the debate remains unresolved about whether the fabrics are wholly magmatic, wholly tectonic, or some hybrid of the two. STOP 3—Neoarchean Saganaga Tonalite with rounded dioritic to granodioritic inclusions Location: UTM: 656219E/5336079N; Campsite #13. Description: Equigranular tonalite with characteristic quartz eyes, containing a wide variety of inclusions. The term "inclusion" has a tortured usage—we prefer to use the term to apply to material that has a contrasting composition or appearance from its host, regardless of origin, size, shape, degree of assimilation, or extent of equilibration with the enclosing host magma. Inclusions may represent blocks of county rock (xenoliths) incorporated into the Saganaga Tonalite, or cognate phases of the intrusion (autoliths). Note that each inclusion contains the same mineralogic components (hornblende, plagioclase, quartz), but the components occur in varied proportions. Although these inclusions have not been studied in detail, field work in the region has identified multiple phases of the intrusion that are compositionally identical with the inclusions, and they are therefore considered autholiths. 122 �STOP 4—Granodioritic phase of Neoarchean Saganaga Tonalite with inclusions; cut by diabase dike Location: UTM 658575E/5335837N; Roadcuts on both sides of Gunflint Trail (Fig. 2) Description: Exposures on the west side of the road consist of pinkish to gray hornblende granodiorite to granite inferred to be a border phase of the Saganaga Tonalite and containing inclusions that are both more felsic and more mafic than the enclosing rock (Fig. 7). For example, the large angular block shown in the photo consists of quartz-eyebearing tonalite, much like rock that is typical of the main phase of the intrusion. Foliation is poorly developed and likely magmatic in origin. This exposure demonstrates that the Saganaga is a composite intrusion that, despite its apparent homogeneity, consists of quite varied magmatic phases, particularly near its border. Tonalite on the east side of the road is cut by a fine-grained diabasic dike several meters in width. Although this location lies some distance east of the prominent north-trending aeromagnetic trends associated with dikes that are described in the text above, the dikes presumably are related. Figure 7. Granodioritic phase of Saganaga Tonalite, containing many and varied inclusions. Large, lighter colored block in center of exposure is inferred to be an autolith of quartz-eye-bearing tonalite similar to the major phase of the batholith. 123 �STOP 5—Border phase Neoarchean Saganaga Tonalite with flattened inclusions and well-developed foliation in contact zone with Neoarchean metabasalt Location: UTM 661834E/5329257N; Roadcut on east side of Gunflint Trail (Fig. 5) Description: This location exposes the border phase of the Saganaga batholith, characterized by a granodioritic composition, general lack of quartz eyes, and an abundance of dioritic inclusions consisting of varied proportions of hornblende, pyroxene, biotite, plagioclase, and minor quartz (Fig. 8). The irregular ovoid and discoid shape of inclusions is oriented subparallel to well developed, steeply dipping and east-trending foliation. Hornblende crystals and aggregates define a prominent lineation plunging shallowly to the east. Petrology indicates that much of this fabric appears magmatic, yet foliation may be a hybrid of approximately coaxial magmatic flow and regional tectonic deformation (D2). This is typical of the border zone of the intrusion against Archean metabasaltic country rocks, which presumably lie in the low ground just to the south. A preliminary comparative geochemical study, summarized in Jirsa and Weiblen (2007), indicates that the mafic inclusions are not partially assimilated, recrystallized, and tectonically deformed country rock volcanic xenoliths as implied by some earlier workers. They have the common chemical characteristics of the sanukitoid suite of Archean granitoid rocks. Thus, we infer that the inclusions studied are autoliths derived from a separate, primitive sanukitoid-magma. Figure 8. Granodioritic border phase of Saganaga Tonalite containing mafic inclusions. 124 �STOP 6—Contact zone of Neoarchean Saganaga Tonalite and metabasalt; unconformably overlain by Paleoproterozoic Kakabeka conglomerate and lower sequence of Gunflint Iron Formation Location: 3 exposures along bush path off Gunflint Trail [individual UTM coordinates given below] (Fig. 5) Description: Stop 6a [UTM 661925E/5329065N] Archean metavolcanic rocks containing abundant granitic sheets and dikes, presumably related to border phases of Saganaga Tonalite (Fig. 9). The boundary between tonalite and metabasalt has been mapped in many places as a fault (Weiblen and others, 1971). These exposures do not preclude that possibility, but they imply that passive emplacement of the intrusion has also occurred, at least locally. Figure 9. Outcrop of border zone of Saganaga Tonalite. Dark gray material probably represents various phases of the intrusion; lighter gray, wedge-shaped area in the foreground is inferred to be metavolcanic country rock. Stop 6b [UTM 661965E/5329062N] Conglomerate developed at the gently southward dipping unconformity between Neoarchean intrusive and metavolcanic rocks and the overlying basal part of the Paleoproterozoic Animikie Group. The unit, regionally known as the Kakabeka Conglomerate, is present only locally on the western end of the Gunflint. In most places, the Lower cherty member of ironformation lies directly on eroded Archean surfaces. This small outcrop is one of the few places where the conglomerate is exposed and accessible along the contact. The conglomerate is greenish gray, poorly bedded, and contains subangular to subrounded fragments of Saganaga Tonalite and related granitoid rocks, metabasalt, and quartz, in a granular siliceous matrix. Stop 6c [UTM 661965E/5329037N] Walking southward from the basal conglomerate is a low "step-up" onto southward dipping strata of the Lower cherty member of Gunflint Iron Formation. 125 �Figure 10. Outcrop photograph of Kakabeka conglomerate lying unconformably on vertically foliated and eroded Archean Saganaga Tonalite (uniform light gray area on right side of photo). Note diversity of fragment types, including Archean metabasalt (dark) and tonalite (light), quartz pebbles (light, subrounded), and Paleoproterozoic diabase (dark gray). STOP 7—Lower Sequence of Gunflint Iron Formation. Location: UTM 661844E/5328896N; Road cut on Gunflint Trail (Highway 12) just north of parking lot for west end of Magnetic Rock Hiking Trail. Description: Gently southward-dipping, thinly interbedded granular and argillaceous iron-formation typical of the lower sequence. In earlier parlance, this stratigraphic position is the lower part of the Upper Cherty member of the Gunflint Iron Formation (Fig. 3). STOP 8—Stromatolitic grainstone at the diastem separating lower and upper sequences of Gunflint Iron Formation. Location: 3 exposures—specific coordinates given below; all adjacent to Magnetic Rock Hiking Trail (Fig. 5). CAUTION and ADVICE: This is a fairly long hike, approximately 1 mile round-trip; please be prepared with water and other field needs. Although this is not in the BWCAW, it does lie within Superior National Forest and is frequented by hikers. For this reason, and to preserve scientific value, please be respectful in matters of hammering and sampling. Description: Stop 8a [UTM 662034E/5328885N] Thin-bedded, fine-grained, chert-amphibole-magnetite-bearing strata assigned to the upper part of the lower sequence of Gunflint Iron Formation (“Lower slaty member”). Beds strike ENE and dip generally less than 8 degrees southward. Stop 8b [UTM 662401E/5329093N] Stromatolites lie within and just above a major regressiontransgression boundary (diastem) that is marked by intraformational conglomerate containing fragments of the underlying granular chert that appear to have been cohesive (though likely not lithified) at the time of incorporation, and in-situ and dislodged stromatolites. Irregular domal and laminar stromatolite forms are present. Note the presence of granules, intraclasts, and oncoliths—the latter consist of intraclasts coated with what likely was biogenic material, now composed largely of silica. Stop 8c [UTM 662583E/5329257N] Crest of ridge exposes the same boundary described above, here with abundant 3-dimensional views of stromatolites, intraformational conglomerate, and stromatolite "hash," all in a peloidal to ooidal, siliceous grainstone matrix. Given the apparent mineralogic replacement and moderate metamorphic grade, little of the original carbon-based material is likely present. Despite this, examples of nearly all morphological forms of stromatolites can be found, including columnar-digitate, domal, and laminar (Fig. 11). 126 �Figure 11. Cherty, stromatolitic Gunflint Iron Formation. A=Oblique surface of small columnar-digital stromatolites; B=Horizontal surface of irregular domal or laminar stromatolites; C=Vertical section of laminar stromatolites; D="Stromatolite hash." STOP 9—The “Magnetic Rock” Location: UTM 0663740E/5329670N; approximately 1 mile walk east of stop 8 on the Magnetic Rock Hiking Trail (Fig. 5) Description: Although most of the trail has magnetic iron-formation underfoot, the trail’s actual namesake lies about a mile walk to the east. The “Magnetic Rock” is a slab of iron-formation in which bedding is essentially vertical and standing nearly 30 feet above the surrounding land surface (Fig. 12). The appearance of this tombstone-shaped block raises the question of how glaciers could have up-ended it, but left the delicately balanced slab intact during ablation. My answer invokes glacial rotation of a “cube” of rock, followed by spalling along bedding planes during repeated cycles of freeze/thaw (frost-heaving). Figure 12. Slab of iron-formation. (Blue-handled hammer against lower 1/3rd of the rock is 40 cm long). 127 �STOP 10—Upper-most, largely argillaceous, Gunflint Iron Formation Location: UTM 663754E/5328212N; Gravel pit north of Gunflint Trail on U.S. Forest Service road #1347 (Fig. 5). Description: This dip-slope exposure consists of interbedded granular (cherty) and laminated (slaty) strata of the uppermost Gunflint Iron Formation. The slope defines the southern limb of a large, shallowly east-plunging anticline. The gentle dip of this limb illustrates the observation that open folding and moderate-relief topography are responsible for the complex map pattern. Note that large ridge visible to the south represents the basal Mesoproterozoic Duluth Complex. The bedding surface is marked by what have been referred to in earlier literature as “syneresis cracks”. The cracks, now filled with quartz, occur both concentrically and radially around a central, apparently raised core within a single granular layer of siliceous iron-formation (Fig. 13). Syneresis cracks are defined generally as shrinkage cracks formed by dewatering in a gel or colloidal suspension. They differ from septarian cracks that may develop in a similar way, in that the latter typically occur in concretions. Surprisingly diverse interpretations can be found in the literature about syneresis cracks (summarized in Pratt, 2001). There is, however, general agreement that they represent localized tensional failure during sediment dewatering. The explanation for localized semi-brittle response to what likely were formationwide stresses—caused by compaction or vibration due to syn-sedimentary earthquakes—is more contentious. It has been ascribed variously to the localization of cements, locally increased pore pressure, or zones of granular sediment made coherent by "microbial glue." It is interesting to note that syneresis structures are more prevalent in Precambrian and Cambrian rocks than younger ones. This may be due in part to more uniform organic bonding of clays in younger strata, which reduced the occurrence of stresslocalization. Recent mapping by the author (Jirsa, 2011) indicates that these quartz-filled cracks occur only in some granular siliceous layers that lie near the stratigraphic top of the iron-formation. This stratigraphic position, and their enigmatic structural attributes, may indicate an origin by impact-induced seismic wave passage through cohesive, semi-rigid chert during the Sudbury meteorite impact event. Figure 13. Polygonal quartz veining on eroded bedding surface of thinly bedded Gunflint Iron Formation. STOP 11—Paleoproterozoic ejecta and breccia from the Sudbury meteorite impact, intruded by sill and dikes of the Mesoproterozoic Logan Intrusions. Location: UTM 664785E/5329200N Description: This traverse provides a cross-section through diabase of the Logan Intrusions and underlying deposits of iron-formation, breccia, and ejecta. The diabase is medium- to coarse-grained in its core to the south, and grades to finer grained and more porphyritic near its base to the north. The northernmost outcrops lie along a steep cliff that exposes the upper Gunflint Iron Formation overlain by a 128 �thick sequence of iron-formation breccia that represents the ejecta-absent facies (Fig. 14.A) of the Sudbury Impact Layer. This is overlain by irregular lenses of bedded lapillistone, mesobreccia (Fig. 14.B), and reworked breccia containing rounded fragments of iron-formation in a matrix composed largely of accretionary lapilli that collectively represent the ejecta-bearing facies. The precise stratigraphic position of the latter two rock types is not entirely clear, though the strata containing accretionary lapilli (true ejecta) appear to lie near the top of the deposit. B. A. Figure 14. A. Large-fragment breccia (ejecta-absent); B. Bedded lapillistone and mesobreccia (ejecta-bearing). STOP 12—Sudbury Impact Layer—folded iron-formation overlain by ejecta. Location: UTM 663700E/5328967N Off Magnetic Rock Hiking Trail Description: Folded siliceous and argillaceous iron-formation overlain by a thin, discontinuous layer of mesobreccia containing scant accretionary lapilli. Note the structural detachment at the base of the outcrop that separates gently dipping, planar-bedded iron-formation layers from the overlying meter or so of folded strata. The chaotic fold style (Fig. 15A) indicates soft-sediment deformation prior to deposition of ejecta, which lends credence to the inference that iron-formation was not yet fully lithified at the time of impact. The walk from here to Stop 11 crosses several exposures of variably deformed iron-formation, all considered part of the ejecta-absent facies of SIL. These outcrops demonstrate the rheologic contrasts of substrate during deformation, and highlight the interpretation that at least some components of ironformation were unlithified at the time of impact deformation (Fig. 15B and 15C). 129 �Figure 15. Outcrop photographs of soft-sediment deformation in the ejecta-absent facies of SIL, locally overlain by ejecta, and demonstrating that deformation occurred during and after silicification of mudstones, but prior to complete lithification. A. Folded siliceous (light-colored) and iron-silicate (darker) iron-formation overlain by thin skin of ejecta containing accretionary lapilli (Bill Addison and Bevan French for scale). B. Irregularly layered siliceous and iron-silicate mudstone cut by a mudstone “clastic” dike (darkest narrow feature running up-down in center of photo). C. Folded and brecciated iron-formation, in which the siliceous layer (light gray) is attenuated and shattered in contrast with the enclosing iron-silicate mudstone that is ductily folded. 130 �STOP 13—Sudbury Impact Layer—deformed substrate, mesobreccia, gritstone and lapillistone. Location: UTM 663535E/5329100N Off Magnetic Rock Hiking Trail Description: This small outcrop provides a complete cross section of the SIL, and some unique sedimentalogical features not seen elsewhere. The stratigraphic sequence is shown in Fig. 16A. Of particular importance are the scoured (channelized) appearance at the base of lapillistone, and the presence of larger fragments of gritstone in lapillistone (Fig. 16B). Both indicate moderately high energy delivery of detritus—presumably by the passing ejecta plume or ground surge. Figure 16A. Photograph and graphic sedimentological analysis of stop 13. Black angular polygons represent fragments of iron-formation; black circles represent lapilli. White box shows approximate location of photo Fig. 16B. In detail, the basal part of this deposit consists of disorganized-bedded boulder “megabreccia”, with clasts composed of rock types characteristic of the underlying Gunflint Iron Formation. The megabreccia is overlain by a decimeters-thick, matrix-supported, pebble “mesobreccia” and massive, pebbly sandstone— here termed gritstone due to its content of moderately sorted, but primarily angular grains. Scattered accretionary lapilli occur in this unit locally, implying that it may be a mixture of ejecta and locally derived detritus. The mesobreccia and gritstone are overlain by lapillistone, composed of tightly packed accretionary lapilli. These fill shallow scours in the top of the mesobreccia and gritstone, or deeper scours that remove strata all the way down to megabreccia locally. The bases of the scours are commonly overlain by a one-centimeter-thick wisp of coarse-grained gritstone, followed vertically by the accretionary lapilli. The scours give a paleocurrent direction of 260 degrees—the bearing from Sudbury 131 �to Gunflint Lake is 280 degrees. At other locations, where individual smaller scours at the base of the lapillistone are not present, the basal, clast-supported lapillistone bed drapes shallow erosive scours. The lowermost accretionary lapillistone is massive-textured, as are overlying accretionary lapilli-rich beds, except where rare, small-scale, low-angle cross-stratification dipping towards 060 degrees is visible. The diameter of accretionary lapilli in the bed at the base of the lapilli-rich interval average 0.7 to 0.8 cm, and those higher in the section and interbedded with sandstone range from 0.2 to 0.4 cm. Gritstone beds become more dominant in the upper few decimeters. Here they are medium- to finegrained with stringers and patches of small accretionary lapilli. Some beds are massive with abundant, isolated lapilli. Parallel lamination to undulating parallel lamination is common in the non-massive beds. Approximately 10 cm of thinly laminated siltstone caps the impact deposit. Figure 16B. Close-up view of lapillistone containing entrained fragment of layered gritstone. STOP 14—Sudbury impact layer—deformed Gunflint Iron Formation overlain by thin ejecta layer that includes small spherules. Location: UTM 663628E/5329186N Off Magnetic Rock Hiking Trail Description: This cliff and ridge-top exposure includes a 7m-thick breccia, abruptly overlain by mesobreccia (Fig. 17A), and capped by strata composed of small (2-5mm) accretionary pellets and slightly larger, concentrically zoned lapilli (Fig. 17B). Some of these small particles may be relict glass spherules; however, metamorphism precludes definitive identification. A. B. Figure 17. A. Megabreccia sharply overlain by mesobreccia and other ejecta. B. Layers composed of accretionary pellets, small lapilli, and inferred relict spherules. 132 �STOP 15—Paleoproterozoic Sudbury impact layer, basal Rove Formation, and Mesoproterozoic Logan Intrusion. Location: UTM 665200E/5329300N Description: This outcrop affords a great number and variety of views of the ejecta and breccia (Fig. 18) because the exposed surface is nearly parallel with strike and dip of formations. The stratigraphic sequence is similar to that at Stop 11; however, this site lies along the top of the deposit, showing the relationship between ejecta and breccia more clearly. Just to the south is the eastern extension of the Logan sill traversed at Stop 11. In the intervening 0.25 mi., the basal contact of the sill cross-cut stratigraphic units to here overlie about 10 feet of slate and graywacke inferred to be the basal section of the Rove Formation. Figure 18. Breccia irregularly overlain by a “skin” of lapillistone. STOP 16—Mesoproterozoic Logan sill and Paleoproterozoic slate of the Rove Formation at Cross River Location: UTM 0665120E/5328890N Description: Outcrops on the north shore of Cross River and in it lie at the top of the same Logan sill that capped basal Rove Formation at stop 15. The cliff on the south shore exposes thinly bedded graywacke and mudstone of the Rove. Crossing the river may not be possible at this time. 133 �STOP 17—Mesoproterozoic Tuscarora intrusion of Duluth Complex—atypical border phase Location: UTM 662074E/5327457N; Roadcut on Cross Lake road (CR#47) south of Hwy 12 Description: A confusing exposure of the lower units of the Tuscarora intrusion. The outcrop consists of intergranular to ophitic gabbro and augite troctolite, with pods and veinlets of coarse mafic pegmatite and shear bands containing sulfide mineralization. Spheroidal weathering produced "core stones" locally. STOP 18—Mesoproterozoic Tuscarora intrusion of Duluth Complex Location: UTM 666638E/5327433N; roadcut on Gunflint Trail east of CR#50. (Fig. 5) Description: Just south of the parking pull-off is the rather poorly exposed intrusive contact between Paleoproterozoic Rove Formation and the Mesoproterozoic Tuscorara Intrusion (Fig. 5 explanation; Morey and others, 1981). The basal unit is a typical example of Cu-sulfide mineralized augite troctolite that is found at the base of the Duluth Complex here and in the Hoyt Lakes-Kawishiwi area to the southwest. It contains disseminated pyrite, pyrrhotite, and chalcopyrite. In the 1970's, International Nickel Company (INCO) drilled 7 holes in the basal Duluth Complex (Tuscarora and western Poplar Lake intrusions) to evaluate potential for Cu-Ni mining. All of these holes lie along the basal part of the intrusion within a few miles east and west of this stop. The archived drill cores were studied by Mogessie (1976) and Mogessie and others (1976). REFERENCES Addison, W.D., Brumpton, G.R., Davis, D.W., Fralick, Philip W., and Kissin, S.A., 2010, Debrisites from the Sudbury impact event in Ontario, north of Lake Superior, and a new age constraint: Are they base-surge deposits or tsunami deposits? Geol. Soc. Am., Special Paper 465. Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N. J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005, Discovery of distal ejecta from the 1850 Ma Sudbury impact event: Geology: 33:193196. Boerboom, T.J., and Zartman, R.E., 1993, Geology, geochemistry, and geochronology of the central Giants Range batholith, northeastern Minnesota: Canadian Journal of Earth Sciences, 30:2510-2522 Cannon, W.F., Schulz, K.J., Horton, J.W. Jr., Kring, D.A., 2010,The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA: Geological Society of America Bulletin v. 122, p. 50-75. Corfu, F., and Stott, G.M., 1998, Shebandowan greenstone belt, western Superior Province: U-Pb ages, tectonic implications, and correlations: Geological Society of America Bulletin 110:1467-1484 Davis, D.W., 2008, Sub-million-year age resolution of Precambrian igneous events by thermal extraction-thermal ionization mass spectrometer Pb dating of zircon: Application to crystallization of the Sudbury impact melt sheet: Geology, v. 36, p. 383-396. Driese, S.G., Jirsa, M.A., Ren, M., Sheldon, N.D., Brantley, S.L., Parker, D., and Schmitz, M., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, v. 189, p. 1-17. Floran, R.J., and Papike, J.J., 1978, Mineralogy and petrology of the Gunflint Iron Formation, Minnesota-Ontario: Journ. Petrology, 19:215-288. Fralick, P.W., Davis, D.W., and Kissin, S.A., 2002, The age of the Gunflint Formation, Ontario, Canada: single zircon U-Pb age determinations from reworked volcanic ask: Can. J. Earth Sci., 39:1085-1091. Fralick, P., Grotzinger, J., and Edgar, L., 2012, Potential recognition of accretionary lapilli in distal impact deposits on Mars: A facies analog provided by the 1.85 Ga Sudbury impact deposit: Society for Sedimentary Geology Special Publication 102, p.211-227. French, B.M., and Koeberl, Christian, 2010, The convincing identification of terrestrial meteorite impact structures: what works, what doesn’t, and why: Earth Science Reviews, v. 98, p. 123-170. Grout, F.F., 1933, Structural features of the Saganaga granite of Minnesota-Ontario: Report of XVI International Geological Congress, Washington, p. 255-270. Gruner, J.W., 1941, Structural geology of the Knife Lake area of northeastern Minnesota: Geological Society of America Bulletin 52:1577-1642. 134 �Heaman, L.M., and Easton, R.M., 2005, Proterozoic history of the Lake Nipigon area, Ontario: Constraints from UPb zircon and baddeleyite dating: in Easton, M., and Hollings, P., eds., Institute on Lake Superior Geology Proceedings, 51st Annual Meeting, Nipigon, Ontario, Proceedings and Abstracts, v. 51, part 1, p. 24-25. Hoaglund, S.A., Miller, J.D., Crowley, J.L., and Schmitz, M.D., 2010, U-Pb zircon geochronology of the Duluth Complex and related hypabyssal intrusions: Investigating the emplacement history of a large multiphase intrusive complex related to the 1.1 Ga Midcontinent rift: Institute on Lake Superior Geology Proceedings, v. 56, part 1, p.25-26. Jirsa, M.A., 2011, Bedrock geology of the western Gunflint Trail area, northern Minnesota: Minnesota Geological Survey Miscellaneous Map M-191, scale 1:24,000. Jirsa, M.A., 2016, Nine years of capstones: A summary of Precambrian Researcch Center field camp capstone mapping projects in the Newoarchean Knife Lake Group and associated rocks, central Boundary Waters Canoe Area Wilderness, Minnesota: (abs.) in Institute on Lake Superior Geology, Part 1—Abstracts and Program; abs., this volume. Jirsa, M.A., Boerboom, T.J., and Radakovich, A.L., 2016, Neoarchean geology of the western Vermilion District: Field Trip 2 in Institute on Lake Superior Geology, Part 2—Field trip guidebook; this volume. Jirsa, M.A., and Fralick, P.W., 2010, Field Trip 4: Geology of the Gunflint Iron Formation and Sudbury Impact Layer, in Field Guide to the Geology of Precambrian Iron Formations in the western Lake Superior region, Minnesota and Michigan: Precambrian Research Center Guidebook 10-01, p.77-92. Jirsa, M.A., Fralick, P.W., Weiblen, P.W., and Anderson, J.L.B., 2011, Sudbury impact layer in the western Lake Superior region, in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 147–169. Jirsa, M.A., and Miller, J.D., Jr., 2004, Bedrock geology of the Ely and Basswood Lake (U.S. portion) 30’ x 60’ quadrangles, northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map Series M-148, scale 1:100,000. Jirsa, M.A., Starns, E., and Schmitz, M.D., in prep., Bedrock geology of the 2006 Cavity Lake fire area: Minnesota Geological Survey Miscellaneous Map M-193, scale 1:24,000. [in preparation; in the interim, refer to OpenFile Report OF-08-05] Jirsa, M.A., and Weiblen, P.W., 2007, Geology along the Gunflint Trail: Field Trip 6 in Miller, J.D., and Peterson, D.M., compilers; Institute on Lake Superior Geology, Part 2-Field Trip Guidebook; p. 143-168. Jones, N.W., 1984, Petrology of some Logan sills, Cook County, Minnesota: Minnesota Geological Survey Report of Investigations 29, 40p. Kenkmann, T., and Schonian, F., 2006, Ries and Chicxulub: Impact craters on Earth provide insights for Martian ejecta blankets: Meteoritics and Planetary Science, v. 41, p. 1587-1603. Krogh, T.E., Davis, T.W., and Corfu, F., 1984, Precise U-Pb zircon and baddeleyite ages for the Sudbury area: in The geology and ore deposits of the Sudbury Structure, E.G. Pye, A.J. Naldrett, and P.E. Giblin, eds, Ontario Geological Survey Special Volume 1, p. 431-446. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., and Jirsa, M.A., and Hamilton, M.A., 2013, New U-Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province: Precambrian Research v. 235, p. 264-277. MillerJ.D., Jr., and Severson, M.J., 2002, Geology of the Duluth Complex: in Miller, J.D., Jr., Green, J.C., Severson, J.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations 58, p. 106-143. Mogessie, A., 1976 Petrologic study of copper-nickel mineralization in the Tuscarora intrusion Duluth Complex, northeastern Minnesota: M.S. thesis, University of Minnesota, Twin Cities, 137p. Mogessie, A., Stumpfl, E.F., and Weiblen, P.W., 1976, The role of fluids in the formation of platinum-group minerals, Duluth Complex, Minnesota: Mineralogic, textural, and chemical evidence: Econ. Geol. 86:15061518. Morey, G.B., and Nathan, H.D., 1978, Geologic map of the Gunflint Lake quadrangle, Cook County, Minnesota: Minnesota Geological Survey Misc. Map M-42, scale 1:24,000. Morey, G.B., Weiblen, P.W., Papike, J.J., and Anderson, D.H., 1981, Geologic map of the Long Island Lake quadrangle, Cook County, Minnesota: Minnesota Geological Survey Misc. Map M-46, scale 1:24,000. Mouginis-Mark, P.J., and Garbeil, H., 2007, Crater geometry and ejecta thickness of the Martian impact crater Tooting: Meteoritics and Planetary Science, v. 42, p. 1615-1625. 135 �Osinski, G., 2006, Effect of volatiles and target lithology on the generation and emplacement of impact crater fill and ejecta deposits on Mars: Meteoritics and Planetary Science, v. 41, p. 1571-1586. Peterson, D.P., Gallup, C., Jirsa, M.A., and Davis, D.W., 2001, Correlation of Archean assemblages across the U.S.Canadian border: Phase I geochronology, (abstract): Institute on Lake Superior Geology Proceedings, 47th Annual Meeting, Madison, WI, 2001, Part 1, p.77-78. Pratt, B.R., 2001, Syneresis cracks: subaqueous shrinkage in argillaceous sediments caused by earthquake-induced dewatering: Sedimentary Geol., 117:1-10. Pufahl, P.K., and Fralick, P.W., 2000, Depositional environments of the Paleoproterozoic Gunflint Formation: in Institute on Lake Superior Geology Proceedings, 46th Annual Meeting, Thunder Bay, Ontario, Part 2 Field Trip Guidebook, v. 51. Pufahl, P.K., and Fralick, P.W., 2004, Depositional controls on Paleoproterozoic iron formation accumulation, Gogebic Range, Lake Superior region, USA: Sedimentology 51:791-808. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, G.J., and Edwards, C.T., 2007, Physical and chemical evidence of the 1850 Ma Sudbury impact event in the Baraga Group, Michigan: Geology 35:827-830. Vervoort, J.D., 1987, Petrology and geochemistry of the Archean of the JAP Lake area, northeastern Minnesota: M.S. Thesis, University of Minnesota-Duluth, 193 p. Weiblen, P.W., Morey, G.B., and Mudrey, M.G., 1971, Guide to the Precambrian rocks of northwestern Cook County as exposed along the Gunflint Trail: in Davidson, D.M., Darby, D.G., Green, J.C., and Grant, J.A., eds., Technical Sessions, Abstracts and Field Guides, 17th Annual Institute on Lake Superior Geology, Annual meeting, Duluth, Minnesota, p.97-127. Wolff, J.F., 1917, Recent geologic developments on the Mesabi Iron Range, Minnesota: American Institute of Mining and Metallurgical Engineers Transactions, v. 56 p. 229-257. 136 �FIELD TRIP 8 Saturday, May 7, 2016 KEWEENAWAN GEOLOGY OF THE HOVLAND AREA Terry Boerboom (Minnesota Geological Survey John Green (University of Minnesota-Duluth Emeritus) INTRODUCTION This field trip will show several examples of the varied types of volcanic rocks within the Northeast sequence of the North Shore Volcanic Group (NSVG), beginning in the upper part of the lower, reversely-polarized sequence and working up the stratigraphic section into the upper, normally-polarized sequence (table 1). It also includes some stops in the Hovland sill, a layered intrusion that is subconformable to the host volcanic rocks, and some small hybrid intrusions. The trip will start NE of Hovland, and work back southwest toward and past Grand Marais (Figure 1). Although brief driving directions are given, the UTM coordinates should provide the most accurate locations. Some of the easternmost UTM coordinates (for stops 1-7) are given using NAD 83, Zone 16N, the rest (stops 7-21) use NAD 83 Zone 15N. It is likely that more stops are described here than can reasonably be covered in a single day. Stops missed during this excursion can be visited by individuals using the guide, with the caveat that some stops may require permission from land owners. Figure 1. General stop locations with respect to the towns of Grand Marais and Hovland. Almost all the stops are along Highway 61, which parallels the shore. The North Shore Volcanic Group (Figure 2), part of the Mesoproterozoic Midcontinent Rift System, is well described in a multitude of publications. Some of these include the Geological Society of America Special Paper 312 (Ojakangas, Dickas, and Green, editors, 1997); Minnesota Geological Survey Report of Investigations 58 (Green, 2002); and field trip number 7 in the Geological Society of America Field Guide 24 (Green and others, 2011). These are just a few examples, and within those publications numerous references to other publications on the topic are listed. Given the widespread background descriptions already available, the interested user is referred to those publications, and to the references therein. A series of detailed geologic maps, based on 1:24,000-scale quadrangles, are available for almost all of the quadrangles that intersect the shoreline of Lake Superior. These maps as well as all Minnesota 137 �Geological Survey publications are available for free download at the Minnesota Geological Survey website (www.mngs.umn.edu) under the ‘search or browse’ link. The geologic maps pertinent to the stops for this field trip are listed below. Published Minnesota Geological Survey bedrock geology maps (all 1:24,000 scale) pertaining to this field trip: Stops 1-10, MGS Map M-195, Marr Island and Hovland quadrangles (Boerboom and Green, 2013) Stops 11-15, MGS Map M-190, Kadunce River (Boerboom and Green, 2011) Stop 16, MGS Map M-189, Grand Marais (Boerboom and Green, 2010) Stops 17-21, MGS Map M-179, Deer Yard Lake – Good Harbor Bay (Boerboom and Green, 2008) Figure 2. Generalized geology of the Mesoproterozoic rocks of northeastern Minnesota showing the major subdivisions of the North Shore Volcanic Group. The black bar denotes the general traverse of this field trip. 138 �Table1. Generalized stratigraphy of the northeast limb of the North Shore Volcanic Group showing U/Pb ages (Davis and Green, 1997; Green and others, 2001; Boerboom and others, 2014). Positions of intrusions denote approximate stratigraphic level affected and not age of emplacement, except rocks of the Beaver Bay Complex affect multiple stratigraphic levels. Thickness(m) Lithostratigraphic units Lithologic character U/Pb ages 7359 325 Total section Schroeder–Lutsen sequence (normal polarity) Lutsen basalts 3998 olivine tholeiite; includes Indian Camp sandstone and thin conglomerate. angular unconformity Upper northeast sequence (normal polarity) 130 100 131 Terrace Point basalt (within Good Harbor Bay andesites) (Stop 21) Cut Face Creek sandstone (Stop 21) Good Harbor Bay andesites (Stop 20) ophitic, olivine tholeiite basalt red, laminated, ripple-marked sandstone brown, porphyritic basaltic andesites Beaver Bay Complex – Beaver River diabase, Leveaux ferrodiorite, etc. (Stops 2, 3, 5 & 6) 122 Breakwater basalt (Stop 17, 19) 348 Grand Marais felsites (Stop 18) 335 Croftville basalts (Stop 16) 250 70 Devil Track rhyolite (Stop 15) Maple Hill rhyolite (Stop 14) 274 Red Cliff basalts (Stop 13) 366 Kimball Creek felsites (rhyolite and icelandite) (Stops 11 & 12) 539 Marr Island lavas (Stops 9 & 10) 198 235 900 Naniboujou basalts Devil’s Kettle rhyolite Brule River lavas (Stop 4) brown, columnar-jointed basalt flow pink to gray porphyritic rhyolite and felsite intergranular basalt and andesite flows, thick interflow sandstone aphyric, intergranular rhyolite flow porphyritic rhyolite flow ophitic olivine tholeiite flows, some plagioclase-phyric Porphyritic Kimball Creek rhyolite flow; Kadunce icelandite. mixed basalt, tholeiitic andesite, and icelandite flows intergranular basalt flows porphyritic ash-flow tuff interbedded basalt and rhyolite flows Brule Lake gabbro, Hovland sill, (of the Beaver Bay Complex) (Stops 7 & 8) 3036 1097.7±1.7 1100.2±1.9 1095.94±0.62 (Hovland sill) Lower northeast sequence (reversed polarity) 1932 Hovland lavas (Stop 1) 67 92 945 Red Rock rhyolite Deronda Bay andesite Grand Portage lavas 60 Puckwunge Sandstone mostly plagioclase-phyric basalt flows, some rhyolite and andesite red/tan, porphyritic rhyolite tan/brown, porphyritic andesite basalt lava flows, pillowed at base BASE 60 1097.26±0.67 Cross-bedded quartz sandstone slight angular unconformity Rove Formation (Paleoproterozoic) 139 1107.7±1.9 1107.9±1.8 �FIELD TRIP STOPS STOP 1 – Porphyritic andesitic lavas – lower reverse sequence Location: UTM (NAD 83, Zone 16T): 0284488E/5305983N, Hovland quadrangle, Milepost 133.2. Park along Highway 61 just beyond the Axtell mailbox. Note: This is private property except for the roadside outcrops. Index map: rrd – Reservation River diabase; hyf – hybrid dike (stop 2); pqm; ferromonzonitic dike (stop 3); nnu – NSVG undivided, normally-polarized; nrh – Hovland lavas rhyolite; nhp – Hovland lavas porphyritic andesite (stop 1). Relationship between hyf and pqm modified from M-195. The gray areas on this and subsequent index maps denote mapped outcrops. Description: Strongly porphyritic basaltic andesite lava flow with as much as 35% large plagioclase phenocrysts (An46) up to 5 cm in length which show a subparallel flow alignment. There are also a few small phenocrysts of magnetite and altered olivine. Upper amygdaloidal portions of the flows contain amygdules filled with epidote and chlorite. The lava flows at this stop are part of the Hovland lavas (mainly basalts and basaltic andesites), which are near the top of the reversely-polarized Lower Northeast sequence of the North Shore Volcanic Group. UPb ages for rhyolites within the Lower Northeast sequence include the 1107.7 ±1.9 Ma Tom Lake rhyolite (located inland to the northwest), and the 1107.9 ±1.8 Ma Red Rock rhyolite, which is located further northeast up the shore (Davis and Green, 1997). --------------------------------------------DIRECTIONS: Cross the highway and proceed to the lake shore in the small cove by the boat house. Note: This is private property and permission from the house across the road is needed to access. ---------------------------------------------STOP 2 –Hybrid ferromonzonite to ferrodiorite dike Location: UTM (NAD 83, Zone 16T): 0284647E/5306009N, Hovland quadrangle, Milepost 133.3 (see inset map under stop 1 for location). This is just across the highway and slightly east of the last stop, outcrops in a small cove on the lake shore. Description: Pink to gray, variably porphyritic, intermingled rhyolitic (finely granophyric) to basaltic rock types that exhibit mutually cross-cutting relationships. Contacts between the phases vary from sharp to gradational. Plagioclase phenocrysts vary from few to abundant; also present are small phenocrysts of apatite, pyroxene, oxides, and possible altered olivine. Viewed from the proper perspective this dike appears to be curvilinear in form. This dike can be traced to the west-southwest, subparallel to the shore, for nearly two miles. Other similar small hybrid dikes and intrusions have been mapped to the southwest in the Marr Island, Kadunce River, and Devil Track Lake quadrangles. These dikes commonly contain resorbed quartz and feldspar xenocrysts, implying that they may be the product of mafic magmas having melted porphyritic rhyolite flows at depth, and comingling with the felsic melts as they were emplaced. The hybrid dike at this stop was shown to be reversely-polarized (K.G. Books, unpub. Data); it is considered to be part of the reversely-polarized Grand Portage dike swarm (Green and others, 1987). On the published bedrock geologic map (M-195) this information was overlooked, and this hybrid dike was mistakenly portrayed as cross-cutting the larger ferromonzodiorite dike that we will visit at the next stop. --------------------------------------------- 140 �DIRECTIONS: From the driveway near Stop 1, proceed back southwest approximately 0.5 mile to an outcrop along the north side of the highway and park on the edge of the road --------------------------------------------STOP 3 –Ferromonzodiorite dike Location: UTM (NAD 83, Zone 16T): 0283964E/5305764N, Hovland quadrangle (see inset map under stop 1 for location). Description: Medium-coarse grained pyroxene-quartz ferromonzodiorite dike – gray, but weathers to a pinkish color, due to granophyre in the mesostasis, and cm-sized clots of granophyre are commonly visible on weathered faces. Contains an average of 49% plagioclase, 20% micrographic felsic mesostasis, 22% blocky to prismatic augite, 6% Fe-Ti oxides, up to 3% altered pigeonite, and trace amounts of apatite and possible altered olivine. This north–south vertical dike averages 328 feet (100 meters) in width and forms a prominent topographic ridge that makes a point out in the lake and extends inland about one mile from the shore. Its full extent, or how it relates to other hypabyssal diabase intrusions to the north, is not known as that area is incompletely mapped. ---------------------------------------------DIRECTIONS: Go back to Highway 61 and drive southwest approximately 2.5 miles to small road leading to a gravel pit, directly across from Big Bay Point road. Veer left at the first intersection on this road and proceed into an old gravel pit which has outcrops in the pit floor. ---------------------------------------------- STOP 4 –Porphyritic rhyolite (Big Bay rhyolite; 1,100.2± 2 Ma) Location: UTM (NAD 83, Zone 16T): 0280325E/5304581N, Hovland quadrangle. Index map: hfd – Hybrid ferromonzonite (stop 5); hfm – Hybrid ferromonzonite and remobilized rhyolite; hod– ophitic diabase; cgf – augite ferromonzodiorite; hba – coarse-grained amygdaloidal basalt; nbf – porphyritic rhyolite (stop 4). Description: Maroonish-pink, feldspar-phyric rhyolite that is quite vesicular (drusy) at the western-most outcrops, but going east passes through a less vesiculated, spherulitic zone and farthest east into a dense, grayish-pink more massive zone; speculatively passing from the upper to lower part of a flow. The western-most outcrops contain possible relict fiamme features, and the eastern/speculatively lower part of the flow may be a welded tuff. Small garnets are locally present, presumably due to contact metamorphism by the surrounding diabases. A sample from here gave a U-Pb zircon age of 1,100.2 ± 2 Ma (the Big Bay rhyolite of Davis and Green, 1997). This rhyolite was determined by Val Chandler (Minnesota Geological Survey, pers. comm.) to have normal magnetic polarity, and thus must be very near the base of the Upper Northeast sequence. Based on the phenocryst assemblage and flow characteristics, this rhyolite is grouped with other disparate occurrences of similar rhyolite in the area, but which are separated by one of the many intrusions. ---------------------------------------------DIRECTIONS: Go back to Highway 61 and drive southwest approximately 0.4 miles to roadcut on both sides of the highway, just past a right bend in the road. NOTE: There is a small pull-off on the north side of the highway just past/west of the outcrop roadcut where one could park a car. The highway is narrow and dangerous – please be careful! ---------------------------------------------- 141 �STOP 5 – Hybrid ferromonzonite phase within ophitic diabase Location: UTM (NAD 83, Zone 16T): 02799459E/5304014N, Hovland quadrangle (see inset map for stop 4 for location). Description: This outcrop shows a gradation from ophitic olivine diabase (the Horseshoe Bay ophitic diabase – Beaver Bay Complex) into a plug-like body of prismatic pyroxene-quartz ferromonzonite. The western-most outcrop is spheroidally-weathered diabase with normal cm-sized pyroxene oikocrysts; going east the oikocrysts transition into bronzy clotted ophites, then to prismatic pyroxene grains; concurrently the diabase texture grades from ophitic, to intergranular and weakly granophyric, into increasingly coarse-grained, granophyric, and prismatic ferromonzodiorite, and ultimately into very coarse-grained ferromonzonite with large curved-prismatic clinopyroxene and plagioclase laths greater than 1cm in size. At the east end within the monzonite the pyroxene prisms and plagioclase laths are aligned into a vertical to steeply east-dipping flow structure. One petrographic sample of the ferromonzonite contains 40 percent strongly zoned plagioclase, 20 percent variably uralitized prismatic augite, 8 percent Fe-Ti oxide minerals, 20 percent felsic mesostasis that is dominated by micrographic quartz/alkali feldspar but also includes independent quartz and sanidine, 10 percent red-brown secondary clay-type minerals, 1 percent hornblende, and nearly 1 percent apatite. This ferromonzonite body is one of several similar bodies that occur within or marginal to the Horseshoe Bay ophitic diabase (unit hod). Some of the marginal bodies may have formed as partial melt segregations from the underlying rhyolite. ---------------------------------------------DIRECTIONS: Continue southwest on Highway 61 for approximately 1.1 miles, to the Flute Reed River in the town of Hovland. Park near the river at a safe place along the highway or on the street parallel to the river and find your way down to the river. The water level must be sufficiently low to access the outcrops. ---------------------------------------------- STOP 6 – Chicago Bay ophitic olivine diabase Location: UTM (NAD 83, Zone 16T): 0278355E/5303179N, Hovland quadrangle. Index map: cbd – ophitic diabase (stop 6); htd – troctolitic ophitic diabase; ndk– Devil’s Kettle rhyolite; nbo – ophitic basalt; nbf – feldspar-phyric rhyolite; nb – sparsely porphyritic basalt. Description: This ophitic olivine diabase (part of the Beaver Bay Complex) exhibits strong sheet joints that dip more or less 10 degrees toward Lake Superior. It locally verges on augite troctolite; and in general contains 60 to 65% plagioclase (dominantly labradorite but includes andesine and bytownite; average An66Ab33Or1), 11 to 20% ophitic augite (Wo38En43Fs19, Mg# 70), 2 to 3% Fe-Ti oxides, 7 to 20% partially iddingsite-altered olivine (average Fo60), trace to 5% felsic mesostasis and quartz, up to 2% hypersthene, up to 1.5% fine-grained chlorite and/or clay mesostasis, and trace amounts of pigeonite, apatite, hornblende, and biotite. This typical ophitic olivine diabase is thought to be a sill-like body that underlies the Hovland sill, but it is not clear whether or not they are related. The extent of this unit is well established in the Hovland quadrangle by outcrops, water well cuttings, and topography; but the extension to the west in the Marr Island quadrangle below the Hovland sill is based on only one set of water well cuttings. This unit is of normal polarity (H.C. Palmer, unpub. data, 1972). 142 �---------------------------------------------DIRECTIONS: From the Flute Reed River continue southwest on Highway 61approximately 1.3 miles to a set of outcrops on the uphill side of the road next to a small pull-off. ---------------------------------------------- Overview of the Hovland Sill, Stops 7 and 8 (actual stops below) The Hovland sill, previously mapped in part by Jones (1963), is a gently-dipping (approximately 15° SSE), subcordant body composed of a basal massive ferrogabbro (stop 7), a middle zone of cumulatefoliated granophyric ferrogabbro to ferromonzodiorite (stop 7), and an upper coarse-grained felsic cap (stop 8). Overall the sill is estimated to be at least 984 feet (300 meters) thick. This sill is very similar to a less well exposed unit a few miles north that has informally been named the Lookout sill, which dips approximately 15° S-SE, and like the Hovland sill, has a cumulate portion with abundant coarse ilmenite plates and an upper felsic cap that contains fayalitic olivine. Both of these evolved intrusions are broadly similar to the ‘Silver Bay ferrogabbro’ type of zoned intrusions that are late intrusions associated with the Beaver Bay Complex (e.g. Miller and Green, 2002). A sample (MH047A-AD) of coarse prismatic olivine-pyroxene ferromonzonite from the upper felsic phase (Stop 8) was submitted for age dating to Dr. Mark Schmitz at the Boise State Isotope Laboratory. No zircon was separated from the sample, however relatively abundant, although small (approximately 100 microns in long dimension), flattened, light brown baddeleyite crystals were recovered. Six baddeleyite crystals selected for dissolution were all variably discordant, but gave equivalent 207 Pb/206Pb dates with a weighted mean of 1095.94±0.62 (n=6; MSWD 0.37; Figure 3; Boerboom and others, 2014). This age falls within the range of published ages for various Figure 3. Isochron diagram for sample MH047A-AD. other units of the Beaver Bay Complex, including the Wilson Lake ferrogabbro (1095.75±0.92; Hoaglund and others, 2010), Sonju Lake intrusion (1096.1±0.8; Paces and Miller, 1993), Silver Bay ferrogabbro (1095.8±1.2; Paces and Miller, 1993), Pine Mountain granophyre (1095.3±3.8; Vervoort and others, 2007), as well as others. The sample was collected from a roadcut on Highway 6, 1.3 miles northeast of the Brule River near Hovland. (UTM zone 15T 722714E, 5301024N). Figures 4 and 5 demonstrate various aspects of chemical differentiation trends within the Hovland sill, and Figure 6 shows examples of the varied textures between the phases. 143 �Figure 4. Variation in Mg# of olivine (A) and An content of plagioclase (B) within the Hovland sill. The olivine Mg plot shows the number of points analyzed for each individual sample in parentheses. The An diagram is from many samples which are not differentiated. Semi-quantitative SEM analyses were provided by Jeff Thole, Macalester College. Figure 5. Whole-rock compositional variations through the Hovland sill. Note the increase in TiO2 near the transition from the lower to middle zone, which is reflected by abundant cumulate ilmenite plates near the bottom of the cumulate zone. Analyses provided by Karl Wirth, Macalester College. 144 �A B C Figure 6. Photographs of thin sections (plane-polarized on left, cross-polarized on right) of phases of the Hovland sill. A. Upper felsic cap (Stop 8) B. Middle cumulate phase (stop 7) C. Lower massive phase STOP 7 – Strongly foliated cumulate ferromonzodiorite of the Hovland sill Location: UTM (NAD 83, Zone 16T): 0276907E/5301831N, Hovland quadrangle. Index map: fdd – small ferrodiorite dikes; hcg – Hovland sill cumulate ferromonzodiorite (stop 7); hgc– Hovland sill ferrogabbro); cbd – ophitic diabase. Description: This stop is within the middle cumulate zone of the Hovland sill. The cumulate ferromonzodiorite is strongly foliated, and granophyric, with abundant magnetite and plates of ilmenite plates (Figure 6B). The cumulate phases consist of plagioclase, augite, Fe-Ti oxides, olivine, and minor apatite. Prismatic augite crystals up to 2 centimeters in length are randomly oriented within the foliation plane. Olivine content (mostly altered) is generally low, around 2-4% in most of the intrusion. Pigeonite (Wo13En49Fs38, Mg# 49) occurs as thin discontinuous rims on augite and as small post-cumulate grains within the felsic mesostasis. Samples from this unit examined petrographically contain 45 to 55% strongly zoned plagioclase, 16 to 33% augite, 0 to 3% pigeonite, 5 to 11% Fe-Ti oxides, 2 to 10% mostly 145 �altered olivine, 8 to 15% felsic mesostasis, minor apatite, and near the top, rare hornblende. The felsic mesostasis is composed of a mixture of quartz paramorphs of tridymite, euhedral sanidine, micrographic to granophyric quartz and alkali feldspar, and abundant secondary iddingsite and clay minerals. Based on limited SEM semiquantitative analyses, average Fe/Mg ratios of augite increase from the base (augite; Wo35En38Fs27, average Mg# 57) to the top (ferroaugite; Wo37En22Fs41, average Mg# 35). Mg numbers for olivine (Figure 4A) range from Fo29 near the base, to Fo17 near the top; olivine is typically altered to reddish-brown iddingsite and/or green bowlingite (saponite). Limited feldspar analyses (Figure 4B) indicate that plagioclase becomes increasingly sodic, ranging from labradorite to mainly andesine at the base (average An51Ab47Or2), and andesine to oligoclase near the top (average An41Ab56Or3). Sanidine is common within the felsic mesostasis (average An2Ab45Or53). The basal ferrogabbro (Figure 6C) is poorly exposed along the highway to the east and will not be visited by this trip. It is dark greenish-gray with a rusty-weathered surface, medium- to coarse-grained, non- to weakly-foliated, and typically contains evenly distributed 3- to 4-millimeter, reddish-brown altered olivine spots. Based on several point counts this lower unit contains 48 to 55% plagioclase (labradorite), 25 to 35% granular augite (Wo37En41Fs22, Mg# 59 at the base and Wo32En37Fs32, Mg# 54 at the top), up to 2% pigeonite (Wo10En37Fs53, Mg# 41), up to 5% altered olivine clots, 5 to 8% Fe- Ti oxides, 6 to 8% felsic mesostasis, 1% reddish-brown clay-like material, and traces of apatite. ---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 2.2 miles; pull over near the auto repair shop and go to outcrop on lake side of highway under powerline, just east of driveway marked by 4300. Be careful if you walk across highway! NOTE: Entering UTM zone 15 (NAD 83). ---------------------------------------------STOP 8 – Upper felsic phase of the Hovland sill Location: UTM (NAD 83, Zone 15T): 722453E, 5300915N; Marr Island quadrangle. Index map: hfg – upper felsic cap of Hovland sill (stop 8); hcg – Hovland sill cumulate ferromonzodiorite (stop 7) Description: Very coarse-grained, prismatic ferromonzonitic upper felsic portion of the Hovland sill; location of age date sample MH047A.AD (Figure 3). Note prismatic to trellis-shaped pyroxene, incomplete fayalitic olivine trellises, and brownish granophyric matrix. Overall description of this unit taken from published geologic bedrock map (M-195) below: Figure 6A shows the typical texture of this unit. Prismatic olivine-pyroxene ferromonzodiorite to ferrogranite— Rusty reddish-brown where weathered, dark brownish- to greenish-gray where fresh, coarse-to very coarse-grained, granophyre-rich, prismatic. Contains 30 to 50 percent strongly zoned andesine (average An39Ab58Or3) to oligoclase, 8 to 15 percent variably prismatic ferroaugite (Wo22En41Fs37, Mg# 35), 10 to 15 percent fayalitic olivine (Fo10) that is mostly altered to reddish-brown iddingsite and varies from irregular coarse prismatic grains and clots to acicular trellises up to30 centimeters in length, 2 to 8 percent Fe-Ti oxides, 30 to 40 percent felsic mesostasis (combinations of micrographic quartz and alkali feldspar, crystalline quartz, and sanidine crystals), 1 to 2 percent apatite, trace amounts of rutile within quartz, and rare fine-grained bornite. The felsic mesostasis also contains abundant reddish-brown iddingsite-like needles interpreted to be former fine-grained masses and prisms of Fe-olivine. Outcrops at the border zone along the Brule River are generally darker in color, slightly more fine-grained, and locally contain small round chlorite amygdules. This unit is of normal polarity (K.G. Books, unpub. data, 1972). 146 �---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 2.5 miles to where the beach nearly touches the highway. Pull over and park along the small pull off near the beach. Walk back east along the beach to low outcrops, or park farther east along the edge of the highway. ---------------------------------------------STOP 9 – Icelandite of the Marr Island lavas Location: UTM (NAD 83, Zone 15T): 719901E, 5299658N; Marr Island quadrangle. Index map: hfg – upper felsic cap of Hovland sill (stop 8); nic – icelandite (stop 9) na – pigeonitic andesite; nob – ophitic basalt; nmr – aphyric rhyolite; nmb – strongly amygdaloidal basalt; ndk –Devil’s Kettle rhyolite. All volcanic units are part of the Marr Island lavas. The volcanic units shown here are also known as the Naniboujou basalts; this unit has been incorporated into the larger Marr Island lava package on M-195 (Boerboom and Green, 2013). Description: Low outcrops along the beach are fine-grained, sparsely porphyritic icelandite, with phenocrysts mainly of plagioclase but also some phenocrysts of magnetite, pyroxene, and apatite in a matrix of fine felty plagioclase and brownishweathered alkali feldspar mesostasis. As you work east along beach there are zones that are variably amygdaloidal, but it is difficult to demarcate flow contacts. Icelandite is a felsic rock characterized by 62-66% SiO2, high FeO (~7%), and Na2O + K2O (6.5-9%) (Carmichael, 1964, as summarized in Green and Fitz, 1993). Icelandites characteristically contain a few percent of small rectangular plagioclase phenocrysts that bleach white on an otherwise pinkish-brown weathered surface. Icelandite can be hard to differentiate from plagioclase-phyric rhyolite; however icelandite typically is brownish in color, has a fine felty texture, is weakly magnetic, and has small apatite phenocrysts compared to rhyolite which is more pink, saccharoidal in texture, non-magnetic, and lacking in apatite phenocrysts. Icelandite in some exposures has a strong flaggy parting which forms slabs about 6-10 cm thick (which would make ideal paving stones!). ---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 2 miles to County road 14, and directly across from it, turn left towards the lake onto an old section of the highway (Fire number 3500) and park. Walk west on the old highway and cut down to the beach to outcrops of ophitic basalt. ---------------------------------------------STOP 10 – Ophitic basalt of the Marr Island lavas, and another small hybrid intrusion. Index map: nhd – hybrid ferromonzonite (stop 10); na – pigeonitic andesite; nmo – ophitic basalt (stop 10);nba –basaltic andesite; npa – strongly porphyritic andesite. All volcanic units are part of the Marr Island lavas. Dashed red line is path of newer Highway 61. Location: UTM (NAD 83, Zone 15T): 716788E, 5297938N; Marr Island quadrangle. Description: Typical ophitic basalt, not only of the Marr Island lavas but of the North Shore Volcanic Group in general. Examine old road cuts and beach outcrops. Note scattered plagioclase phenocrysts. 147 �---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 1 mile to the intersection with Kelly’s Hill Road. Outcrop at the northwest corner of the intersection. ---------------------------------------------STOP 11 – Rangeline icelandite of the Kadunce icelandites. Location: UTM (NAD 83, Zone 15T): 715255E, 5297698N; Kadunce River quadrangle. Index map: nhd – hybrid ferromonzonite; nki – Kadunce icelandite (stop 11); nkq – porphyritic rhyolite; na – pigeonitic andesite; nmo – ophitic basalt. Description: The roadcuts along the north side of the highway expose the Rangeline icelandite, which is typical of the icelandites of the NSVG. It is brownish, with ~15% phenocrysts of mainly plagioclase but also altered Fe-olivine, Fe-augite (En15Wo42Fs43), magnetite, and apatite in a fine-grained groundmass of mainly plagioclase, alkali feldspar, and quartz (Figure 7). Figure 7. Photomicrographs (plane-polarized light) of the Rangeline icelandite showing phenocryst assemblage. The glomerophenocryst in the center of photo on left includes pale green-altered pyroxene; the brownish matrix is alkali feldspar. Pf – plagioclase; Ex-Ol – altered olivine; Px – ferroaugite; Ap – apatite. ---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 2.3 miles and park just beyond mile marker 118 at base of trail going up the hill on north side of the highway. Note: this is private property so please use discretion and stay near the road. The loose pieces of rhyolite at the base of the cliff are identical to those on the cliff face, which is dangerous. ---------------------------------------------- 148 �STOP 12 – Kimball Creek Rhyolite Rheoignimbrite Location: UTM (NAD 83, Zone 15T): 711750E, 5296730N; Kadunce River quadrangle. Index map: mld – Monker Lake diabase; nkr – Kimball Creek rhyolite (stop 12). Description: The Kimball Creek rhyolite is thought to be the second largest felsic flow in the NSVG. It is ~350 m thick and extends at least 20 miles/32 km to the west (Green and Fitz, 1993). The composition of this rhyolite is midway between typical rhyolites and typical icelandites. It contains 5-10% small phenocrysts, mostly plagioclase but also magnetite, zircon, apatite, quartz, and altered Feaugite. The fine-grained groundmass is composed mainly of alkali feldspar and variably poikilitic quartz, much of which occurs as paramorphs after tridymite tablets. As in the Devil Track rhyolite, the size of these ‘ex-tridymite’ tablets increases toward the flow center from the top and the base, implying emplacement as a single cooling unit. At both the top and bottom of this flow, outcrops show pyroclastic texture, with flattened fiamme and shards in a dense, probably originally vitric-ash groundmass. Near the base, these stretched fiamme are involved in the flow-folds (Green an Fitz, 1993). These observations imply that this flow was emplaced as a high-temperature ignimbrite that consolidated and underwent bulk flow. ---------------------------------------------DIRECTIONS: Continue toward Grand Marais for about 1.4 miles to mailbox #2524. Walk down the driveway toward the lake to outcrop on the shore below house. NOTE: This is private property – please obtain permission of the owner before proceeding to the shore! NO HAMMERS! Alternate outcrops of this unit can be viewed at roadcuts along the highway, but they are not as strongly porphyritic. ---------------------------------------------STOP 13 – Porphyritic basalt flow of Red Cliff basalts Location: UTM (NAD 83, Zone 15T): 710132E, 5295236N; Kadunce River quadrangle. Index map: nrb – Red Cliff basalts (stop 13); nkr – Kimball Creek rhyolite (stop 12). Description: The Red Cliff basalts are a series of olivine tholeiite flows sandwiched between large rhyolite flows in this section of the Upper Northeast sequence. This group of basalt flows is approximately 300 m thick, and can be traced inland to the west for at least 18 mi / 30 km. The thick ophitic to subophitic flow at this stop is remarkable for its concentration of large plagioclase phenocrysts (~An70) near its top; the phenocrysts apparently floated in the lava after eruption. Locally it also contains dm-sized, angular inclusions of coarse-grained anorthosite. It appears as though some of the phenocrysts may have floated away from disaggregating anorthosite inclusions (Figure 8) 149 �Figure 8. Photograph of Red Cliff basalt flow with coarse-grained anorthosite inclusion and plagioclase phenocrysts that appear to have disaggregated and floated away from it. Hammer is 40 cm long. ---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 1.9 miles and park along the edge of the highway near outcrops on the uphill side of the road. ---------------------------------------------STOP 14 – Maple Hill rhyolite Location: UTM (NAD 83, Zone 15T): 707109E, 5295074N; Kadunce River quadrangle. Index map: nhd – hybrid ferromonzonite; ndr – Devil Track rhyolite (stop 15); nwo – Woods Creek basalt; nhr – Maple Hill rhyolite (stop 14); nrb – Red Cliff basalts. Description: The Maple Hill rhyolite varies from 260 to nearly 400 feet in thickness, and extends west for at least 12 mi / 20 km and possibly as far as 25 mi / 40 km. Structures such as lineated vesicles and folded vesicle trains, coupled with the lack of pyroclastic textures, indicate that this rhyolite erupted as a lava flow rather than a rheoignimbrite (Green and Fitz, 1993). The upper part of the flow has quartz-lined stretched vesicles and amygdules of quartz and calcite, and local veinlets of purple fluorite. Below the upper vesiculated zone the rhyolite commonly shows tightly folded flow layering, and locally contains abundant spherulites and lithophysae as large as 3 centimeters. In general, this unit contains 4 to 6 percent alkali feldspar phenocrysts, 2 to 4 percent quartz phenocrysts, and rare microphenocrysts of zircon in a groundmass of fine-grained quartz and feldspar with minor fluorite (Fitz, 1988), as well as phenocrysts of plagioclase, altered Fe-olivine and Fe-augite, and Fe oxides (Green and Fitz, 1993). This rhyolite, in contrast to the aphyric Devil Track rhyolite (stop 15), contains abundant phenocrysts. 150 �An enigmatic thin flow of ophitic basalt (Woods Creek basalt; now on inset map), which is exposed a couple miles inland to the west and also intersected in water wells, appears to have erupted at the same time as the Maple Hill rhyolite, or more likely between the Maple Hill and Devil Track rhyolites. Remarkably, the same stratigraphic relationships are noted over 40 km to the west in the Lutsen quadrangle (Boerboom and others, 2007), where a thin basalt flow overlies the western extension of the Maple Hill Rhyolite and in turn is overlain by the Devil Track rhyolite. ---------------------------------------------DIRECTIONS: Continue towards Grand Marais for about 0.25 miles and park along the edge of the highway near outcrops at the abandoned wave-cut cliff on the uphill side of the road. ---------------------------------------------STOP 15 – Devil Track rhyolite Location: UTM (NAD 83, Zone 15T): 706718E, 5294928N; Kadunce River quadrangle (see inset map for stop 14 above for location). Description: The Devil Track rhyolite is the largest known flow of felsic volcanic rocks in the North Shore Volcanic Group and is inferred to have been either a hot superliquidus lava flow or possibly a thick, hot rheoignimbrite that flowed and underwent complete crystallization after deposition (Green and Fitz, 1993). Basal outcrops just northeast of here show strong lamination containing a marked flow lineation. This rhyolite varies from 750 to over 950 feet / 230-300 meters) in thickness, and extends west for at least 25 mi / 40 km from the mouth of the Devil Track River (and an unknown distance to the east, beneath Lake Superior). This rhyolite is light pink to grayish-pink, fine-grained and saccharoidal textured, essentially aphyric, and contains abundant small, tabular paramorphs of quartz after primary tridymite. Grain size increases toward the center of the flow (Green and Fitz, 1993). Flaggy parting is typical, as well as a planar flow layering that is gently warped and is generally not parallel to the flaggy parting; neither parting nor flow layering provide consistent measured structural orientations. A well that penetrates the top of this rhyolite to the southwest of here shows that the upper part of the flow is perlitic, and is overlain by a thin, discontinuous sandstone. ---------------------------------------------DIRECTIONS: Drive to Grand Marais, and at the east edge of town, follow signs to go up the Gunflint Trail (Cook County Highway 12). Drive up the Gunflint Trail for approximately 2.5 miles, and turn right on the road to Pincushion Mountain ski trails. Drive this road to the parking lot. ---------------------------------------------STOP 16 – Andesitic Croftville lavas – Pincushion Mountain overlook Location: UTM (NAD 83, Zone 15T): 701012E, 5294305N; Grand Marais quadrangle. Index map: nco – ophitic to intergranular pigeonitic basalt; nca – andesite (stop 16). Both are part of the Croftville lavas. Description: Artists’ point and the breakwater that forms the enclosure around the Grand Marais Harbor visible below are formed by the Breakwater basalt flow, which will be the next stop. This landform, with an island tied to the mainland by a gravel bar, is a classical tombolo. The rubbly outcrops here below this overlook are fine-grained, sparsely porphyritic andesite which is part of the 151 �Croftville lava sequence. Exposures on the dipslope and in creek valleys below the overlook show the andesite flows contain thick rubbly Aa-type flow tops, typical of lava flows of this composition. ---------------------------------------------DIRECTIONS: Drive back down the Gunflint Trail to Grand Marais, turn right (west) on Highway 61, and proceed to Broadway Avenue. Turn left on Broadway and drive to the parking lot next to the harbor just before the Coast Guard station, and continue walking toward the lake to the breakwater. ---------------------------------------------STOP 17 – Breakwater basalt – Artist’s Point Location: UTM (NAD 83, Zone 15T): 699945E, 5291458N; Good Harbor Bay quadrangle. Index map: mmd – Murphy Mountain diabase; nbb – Breakwater basalt (stop 17); nba – Amygdaloidal porphyritic basalt; ngp –Grand Marais porphyritic rhyolite (stop 18); ngr – Grand Marais aphyric rhyolite. Description: The broad ledges here that help form the Grand Marais harbor are made of a thick (>100 m) flow of transitional basalt called the Breakwater basalt. It has distinctive texture and columnar jointing, and forms some of the ridges of the ‘Sawtooth Range’ visible to the west from the Breakwater. The western end of the breakwater has wellpreserved glacial striations and well-developed joint-plucking. This is the tombolo seen from the last stop. The basalt is massive, gray to maroon, and fine- to medium- grained, with abundant small clustered plagioclase phenocrysts and minor augite, altered olivine, and magnetite in a felty-intergranular groundmass. The Breakwater basalt can be traced at least 7.5 mi/12 km to the west, where it apparently pinches out. It is not present in the Cascade River, which is approximately 10.5 mi / 16 km to the west. ---------------------------------------------DIRECTIONS: Drive back to Highway 61, turn left (west) and go 1 mile up the hill to small roadcut on the right (north) side of the road opposite the lake shore. It might be best to pull into the Harbor Light parking lot and walk back to outcrop. ---------------------------------------------STOP 18 – Porphyritic Grand Marais rhyolite (1097.26±0.67 Ma) Location: UTM (NAD 83, Zone 15T): 698364E, 5291847N; Good Harbor Bay quadrangle. Index map units same as stop 17. Description: This outcrop (underneath large billboard) is the location of age date sample DG073-AD, porphyritic rhyolite (207Pb/206Pb age of 1097.26±0.67). This strongly porphyritic rhyolite contains in general 2-4% each of quartz and feldspar phenocrysts, as well as minor magnetite and altered mafic silicate phenocrysts. The rhyolite varies from massive to flow-banded, and is commonly strongly blocky-fragmental, a texture well-exhibited in the small creek just east of this outcrop where angular, flow-banded rhyolite blocks up to 1.5 m in size are evident; mapping in the vicinity indicates that loose blocks of rhyolite were overrun by the Breakwater basalt (see next stop for more description). 152 �Age dating on rhyolite from this outcrop was performed by Dr. Mark Schmitz at the Boise State Isochron lab. Six zircon crystals were selected for CA-TIMS (Chemical Abrasion Thermal Ionization Mass Spectrometry) analysis, from which five grains produced concordant isotopic ratios, with a weighted mean 206Pb/238U date of 1095.00±0.33 (MSWD (Mean Square Weighted Deviation) = 0.07) and a weighted mean 207Pb/206Pb age of 1097.26±0.67 (n=5; MSWD 1.47). Using the 207Pb/206Pb weighted mean date, this age is only slightly younger than the Devil’s Kettle rhyolite (1097.7±1.7; Davis and Green, 1997), which lies roughly 8,000 feet stratigraphically below and is separated by several thick mafic to felsic volcanic units. The nearly identical ages for these two units indicates rapid and voluminous volcanic activity in the upper part of the northeast limb of the North Shore Volcanic Group. ---------------------------------------------DIRECTIONS: Continue southwest on Highway 61 approximately 1.7 miles to the Fall River. Park on wide spot at edge of Highway near river. First cross highway and follow trail along east side of river to the shore, then cross the river (if possible) to outcrops on the west. ---------------------------------------------- STOP 19 – Breakwater basalt and Grand Marais rhyolite, Fall River Location: UTM (NAD 83, Zone 15T): 695809E, 5290906N; Good Harbor Bay quadrangle. Index map: nga – Good Harbor Bay andesites; nbb – Breakwater basalt; nba – amygdaloidal Breakwater Basalt; ngp – Grand Marais porphyritic rhyolite; nbd – inclusion-rich basalt sill or dike; nmi – icelandite (outcrop not shown). Description: Outcrops of porphyritic rhyolite cross-cut and/or overrun by the Breakwater basalt. Dikelets of basalt intruded into fractures in the rhyolite contain abundant small chips and slivers of rhyolite, and larger inclusions of rhyolite or possibly more andesitic rocks with stretched vesicles are contained in the basalt. Although the rhyolite here may be as inclusions in the basalt, farther up Fall Creek and also along the shoreline there are several ‘windows’ through the Breakwater basalt where the underlying rhyolite is exposed. In all cases, it appears as though the rhyolite was a ‘loose breccia’ that was overrun by the Breakwater basalt. Figure 9 is a photograph taken along the shore towards Grand Marais, which shows a block of rhyolite that is draped by the Breakwater basalt; evidence for a loose, blocky rhyolite surface having been overrun by a basalt flow. 153 �Figure 9. Block of porphyritic rhyolite overrun by the base of the Breakwater basalt flow, lakeshore west of Grand Marais (see map for stop 18, small unit labeled ‘ngp’ along shoreline due south of stop 18. Chilled, finely amygdaloidal margin of the basalt drapes over the block of rhyolite. ----Hike back up to the highway and climb down into the river just upstream of the highway----Exposed here (on the east side of the river) is another contact between the Breakwater basalt and rhyolite, which in this case has only feldspar phenocrysts. The basalt in general becomes increasingly amygdaloidal near the rhyolite, but the relationships are somewhat ambiguous because there are also inclusions of identical-appearing amygdaloidal basalt within the rhyolite. Another example of this can be found farther upstream, where the rhyolite is extremely brecciated, with fragments from 1 m or more to as small as 1 cm, and the fragments are intruded by the Breakwater basalt, or a possibly a feeder dike to the basalt. ---------------------------------------------DIRECTIONS: Continue southwest on Highway 61 for about 2.3 miles and park along the road edge adjacent to a long roadcut. 154 �STOP 20 – Good Harbor Bay andesites Location: UTM (NAD 83, Zone 15T): 692225E, 5289747N); Good Harbor Bay quadrangle. Index map: ngs – Cut Face Creek sandstone (stop 21); nga – Good Harbor Bay andesites (stop 20); nbb – Breakwater basalt. Description: The Good Harbor Bay andesites extend west from this location for nearly 20 mi\32 km, to beyond the Onion River, where they are terminated by the Leveaux Porphyry. The unit is at least 200 ft / 60 m thick. It overlies the Breakwater basalt and is overlain by the Cut Face Creek Sandstone, which in turn is overlain by the Terrace Point basalt flow (Stop 21). However, approximately 22 km to the west of here the sandstone pinches out, and the Good Harbor Bay andesites are in direct contact with the Terrace Point basalt. This locality is typical of the Good Harbor Bay andesites – fine-grained, fresh, sparsely porphyritic, and moderately to strongly magnetic. This roadcut exposes a flow contact where a lower rubbly amygdaloidal Aa flow-top is overlain by a massive flow base. A discontinuous, meter thick bed of sandstone locally overlies the rubbly flow-top breccia. The flow-top breccia can be recognized by blocks of strongly amygdaloidal/vesicular andesite infilled by sandstone, and the base of the overlying flow contains abundant amygdules that are highly stretched parallel to the flow contact. The sandstone between the two flows can be identified by its red-spotted appearance (oxidation spots). Just east of this outcrop is a small creek that crosses the highway. If the water is low enough to traverse up the creek, there are excellent fresh exposures of the Good Harbor Bay andesites. Approximately 200 m upstream is a small waterfall formed by an approximately N20°E, 20° north dipping, 30 cm wide sharply bounded brittle fault that has pink zeolite minerals infilled around the fault breccia clasts. Other small, flat faults, some with slickensides, may be visible in the stream bed. ---------------------------------------------DIRECTIONS: Continue southwest on Highway 61 about ¾ of a mile to the scenic overlook across from the high road cut. ---------------------------------------------- STOP 21A – Terrace Point basalt flow and Cut Face Creek sandstone (Cut Face Creek Road cut) Location: UTM (NAD 83, Zone 15T): 691780E, 5289170N Good Harbor Bay quadrangle. Index map: ngt – Terrace Point basalt flow; ngx – basaltic breccia; ngs – Cut Face Creek sandstone; nga – Good Harbor Bay andesites. Description: Thick interflow sandstone with ripple marks, deformation features in sandstone at base of flow, shale rip-up chips, and desiccation cracks. At the west end is a fragmental/scoriaceous phase of the Terrace Point basalt intruded and overrun by the main basalt flow. In high roadcut on northwest side of highway is an obvious contact between the Cut Face Creek Sandstone and the overlying Terrace Point basalt flow. The sandstone overlies the Good Harbor Bay andesites (stop 20). 155 �Basalt – The Terrace Point basalt, which overlies the sandstone, is a major ridge-forming unit from here to the southwest, forming ‘sawtooth mountains’. It is a distinctive flow characterized by a dark green color, white thomsonite amygdules, uniform 3-4mm ophitic texture, and scattered small glassy plagioclase phenocrysts. In general the contact with the sandstone is sharp and straight, but in places the sandstone has been slightly deformed, though not appreciably metamorphosed, by the basalt flow. Near the south end of the road cut is a unit of scoriaceous, fragmental basalt that is intruded and overrun by the Terrace Point flow. Similar rock types and relationships have been observed to the southwest, also near the flow base. The breccia contains 1-100 cm angular basalt fragments that are both massive and amygdaloidal, and scattered large blocks of massive basalt. At this locality and others, sub-volcanic dikes of Terrace Point basalt that intrude the fragmental basalt are slightly chilled, and contain small amygdules stretched parallel to, and columnar joints perpendicular to, curvilinear dike margins. The fragmental unit is interpreted as a cinder cone or lahar-type deposit that may have formed by interaction between a basalt feeder and water-saturated sediment. Sandstone –Jirsa (1984) measured approximately 73 meters of sandstone and 3 meters of shale in this section of the Cut Face Creek sandstone, but reported overall that nearly 30 percent of it is composed of thinly bedded, graded layers of fine-grained sand, silt, and clay. He reports both symmetrical and asymmetrical ripple marks, and bimodal paleocurrent distribution, and concluded that the Cut Face Creek sandstone was deposited in a fluvial-lacustrine environment. Unlike nearly all the other exposed interflow sandstones within the NSVG, planar cross-bedding is predominant, but some trough cross beds are present near the top of the section. Other features that may be visible in the sandstone are desiccation cracks filled with sandstone or coarse pink zeolite minerals, and rip-up textures. Compositionally it is predominantly a lithic arkose composed mostly of plagioclase feldspar and mafic rock fragments, with lesser amounts of quartz, altered clinopyroxene, and opaque grains; cemented by calcite and zeolite. The base of the sandstone is exposed to the north in the Cut Face Creek valley, where it overlies the Good Harbor Bay andesite (see stop 21B). The Cut Face Creek Sandstone was earlier considered to represent clastic deposition during a significant hiatus in volcanism prior to eruption of the Terrace Point/Schroeder-Lutsen basalts. It was also considered to be somewhat unique in that it was one of the few thick sandstone units in the northeast limb of the North Shore Volcanic Group, along with the 68 meter thick Indian Camp sandstone (which is within the Schroeder-Lutsen basalts). However, recent remapping has shown that there are likely at least three more substantially thick sandstone units that occur within the lower series of lava flows to the north of the Good Harbor Bay lavas. Thus, it is now recognized that the Cut Face Creek and Indian Camp sandstones, exposed because of more active erosion near the Lake Superior coast, are only one part of a larger set of thick sandstone units (Figure 10), some of which are located over 5 miles inland and at least 500 and possibly more than 1,000 meters downsection from here. The recognition of thick interflow sandstones throughout several different series of lava flows at different stratigraphic levels implies active sedimentation during a prolonged period of volcanism. Fragmental flow tops in the Good Harbor Bay Lavas contain abundant sandstone infillings, and thin, discontinuous, layers and crack fillings (clastic dikes) of sandstone are common in the Schroeder-Lutsen basalts. Elevated levels of clastic deposition during active volcanism may account for the relative abundance of sand at the tops of the lava flows, and the thick interflow sandstones may have formed during periods of relative volcanic quiescence, but continued basin subsidence. Available data indicate that the thicker sandstones may vary in thickness along strike, consistent with deposition onto an irregular lava surface. 156 �Figure 10. Simplified geologic map showing the distribution of sandstone units in the northeast limb of the NSVG. The black arrows at the lower left indicate thin sandstone units (one of which is the southwestern extension of the Cut Face Creek Sandstone at stop 21). ---------------------------------------------DIRECTIONS: Drive back towards Grand Marais a short distance to the wayside rest parking area on the lake side of the highway. ---------------------------------------------STOP 21B – Traverse up Cut Face Creek; Good Harbor Bay andesites, Cut Face Creek sandstone Location: UTM (NAD 83, Zone 15T): 691953E, 5289526N Good Harbor Bay quadrangle. See index map for stop 21A. Description: Small outcrop directly across highway from parking area and to northeast, and outcrops at the mouth of Cut Face Creek are fine-grained, brownish-gray, sparsely porphyritic Good Harbor Bay andesite. The andesite here is overlain by the Cut Face Creek sandstone, and the contact is exposed in numerous locations in the meandering Cut Face Creek valley. The basal contact appears to be conformable with the flat-surfaced, uneroded, amygdaloidal andesite. Small pebbles of massive to amygdaloidal andesite are common in the lower 3 meters of the sandstone, typically in 3-25 cm thick planar cross-bedded pebbly beds. The upper part of the andesite typically contains stretched amygdules, and has cracks filled with sandstone (Figure 11). 157 �Figure 11. Photograph of base of Cut Face Creek sandstone (red arrow and above), showing clastic dike (black arrow) filling a crack in underlying andesite. Hammer in ellipse is 45 cm long. REFERENCES Boerboom, T.J., Wirth, K., 2014, and Evers, J.F., 2014, Five newly acquired high-precision U-Pb ages in Minnesota, and their geological implications [abs]: Institute on Lake Superior Geology, May 14-17, 2014, Proceedings Volume 60, Part 1 – Programs and Abstracts, p.13-14. Boerboom, T.J., and Green, J.C., 2013, Bedrock geology of the Marr Island and Hovland quadrangles, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series Map M-195, scale 1:24,000. Boerboom, T.J., and Green, J.C., 2011, Bedrock geology of the Kadunce River quadrangle, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series Map M-190, scale 1:24,000. Boerboom, T.J., and Green, J.C., 2010, Bedrock geology of the Grand Marais quadrangle, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series Map M-189, scale 1:24,000. Boerboom, T.J., and Green, J.C., 2008, Bedrock geologic map of the Deer Yard Lake and Good Harbor Bay quadrangles, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous Map Series Map M-179, scale 1:24,000. Boerboom, T.J., Green, J.C., and Albers, P.B., 2007, Bedrock geology of the Lutsen quadrangle, Cook County, Minnesota: Minnesota Geological Survey Miscellaneous map M-174, scale 1:24,000. Carmichael, I.S.E, 1964, The petrology of Thingmuli, a Tertiary volcano in eastern Iceland: Journal of Petrology 5, p. 435-460. Davis, D.W., and Green, J.C., 1997, Geochronology of the North American Midcontinent Rift in western Lake Superior and implications for its geodynamic evolution: Canadian Journal of Earth Sciences, v. 34, no. 4, p. 476488 Fitz, T.J., III, 1988, Large felsic flows in the Keweenawan North Shore Volcanic Group in Cook County, Minnesota: unpubl. M.S. thesis, University of Minnesota-Duluth, Duluth, MN; 165 p. 158 �Green, J.C., Boerboom, T.J., Schmidt, S.Th., and Fitz, T.J., 2011, The North Shore Volcanic Group: Mesoproterozoic plateau volcanic rocks of the Midcontinent Rift System in northeastern Minnesota: Geological Society of America Field Guide 24, Trip number 7, p. 121-146; Miller, J.D. Jr. Houdak, G.J., Wittop, C, and McLughlin, P.I, eds. Green, J.C., 2002, Volcanic and sedimentary rocks of the Keweenawan Supergroup in northeastern Minnesota, in Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., and Wahl, T.E., Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations 58, p. 94-102. Green, J. C., Davis, D.W., and Schmitz, M.D., 2001, Three new zircon dates for the Midcontinent Rift, North Shore, Minnesota: More data, more questions: Institute on Lake Superior Geology, 47th Annual Meeting, Madison, Wis., Proceedings, pt. 1, Programs and Abstracts, p. 29. Green, J.C., Bornhorst, T.J., Chandler, V.W., Mudrey, M.G., Jr., Myers, P.E., Pesonen, L.J., and Wilband, J.T., 1987, Keweenawan dykes of the Lake Superior region: Evidence for evolution of the Middle Proterozoic Midcontinent Rift of North America, in Halls, H.C., and Fahrig, W.F., eds., Geological Association of Canada special paper 34, Mafic Dyke Swarms, Publication No. 0120 of the International Lithosphere Program, p. 289 – 302. Green, J.C., and Fitz, T.J. III, 1993, Extensive felsic lavas and rheoignimbrites in the Keweenawan Midcontinent Rift plateau volcanics, Minnesota: petrographic and field recognition: Journal of Volcanology and Geothermal Research, v. 54, p. 177-196. Hoaglund, S.A., Miller, J.D., Jr., Crowley, J.L., and Schmitz, M.D., U-Pb zircon geochron9llgy of the Duluth Complex and related hypabyssal intrusions: Investigating the emplacement history of a large multiphase intrusive complex related to the 1.1 Ga Midcontinent rift. Jirsa, M.A., 1984, Interflow Sedimentary Rocks in the Keweenawan North Shore Volcanic Group, Northeastern Minnesota; Minnesota Geological Survey Report of Investigation 30, 20 p. Jones, N.W., 1963, The relationship between the Duluth Gabbro and the dikes and sills in the vicinity of Hovland: Minneapolis, University of Minnesota, M.S. thesis, 90 p. Miller, J.D., Jr., and Green, J.C., 2002, Geology of the Beaver Bay Complex and related hypabyssal intrusions, in Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minnesota Geological Survey Report of Investigations 58, p. 144-163. Ojakangas, RW., Dickas, A.B., and Green, J.C., eds., 1997, Middle Proterozoic to Cambrian rifting, central North America: Geological of Society Special Paper 312, 322 p. Paces J.B., and Miller, J.D., Jr., 1993, Precise U-Pb ages of Duluth Complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagmatic processes associated with the 1.1 Ga Midcontinent Rift System: Journal of Geophysical Research, V. 98, No. B8, p. 13,997-14,013. Vervoort, J.D., Wirth, K., and Kennedy, B., 2007, The magmatic evolution of the Midcontinent rift: New geochronologic and geochemical evidence from felsic magmatism: Precambrian Research, vol. 157, no. 1-4, p. 235-268. 159 �FIELD TRIP 9 Saturday, May 7, 2016 DULUTH HARBOR GEOLOGIC HISTORY BOAT CRUISE - PLEISTOCENE TO ANTHROPOCENE Dr. Andy Breckenridge (University of Wisconsin-Superior; Natural Sciences Department) Todd Kremmin (University of Minnesota-Duluth; Dept. of Earth & Environmental Sciences) Eric Dott, P.G. and Irvin Mossberger, P.G. (Barr Engineering, Duluth, MN) “there is nothing -- absolutely nothing -- half so much worth doing as simply messing about in boats” -Wind in the Willows by Kenneth Grahame Introduction The St. Louis River flows into western Lake Superior to create one of the most remarkable coastal systems on the Great Lakes - the St. Louis estuary and harbor (Figure 1). The system is a ria, a coastal inlet formed by submergence of a former river valley. Post-glacial rebound drives the modern submergence of the St. Louis River valley, which probably began around 1000 years ago, but this is only the most recent of at least three periods of submergence, the first two of which were succeeded by lake level lowering and downcutting of the former lake plain. Morning stops highlight a few recent efforts to improve our understanding of the coastal morphology and geologic history of the western basin. These include Minnesota Point, the longest freshwater baymouth bar in the United States, and a tour of the RV Blue Heron, the only UNOLS research vessel on the Great Lakes. In the afternoon we will board the Vista Queen, a commercial tourist vessel, to draw attention to modern environmental challenges within the estuary. The harbor is the largest and busiest on the Great Lakes, and the long history of industry led to the St. Louis River being one of many federally designated “Great Lakes Areas of Concern,” which are targeted for remediation and restoration due to severe environmental degradation. At the forefront of these restoration efforts are geoscientists. We hope that a cruise along the river will reveal that despite development, the St. Louis estuary is an urban wilderness waterway, with exceptional recreational opportunities and a natural beauty that rivals the Great Lake. Background The Pleistocene The Superior Lobe of the Laurentide Ice Sheet filled the western Lake Superior basin during the last glacial maximum. Superior Lobe sediments are distinguished by a striking red color created by erosion of redbed sandstones and mafic igneous rock that underlie the lake. Ice retreat took thousands of years, and was punctuated by multiple readvances. Determining the extent of retreat prior to each readvance is challenging because readvances bury and often erode older landforms and sediments. Multiple Superior Lobe till sheets have been mapped. Younger tills are enriched in silt and clay upsection, suggesting that successive readvances overrode clay-rich glaciolacustrine sediments from pro-glacial lakes that formed after the ice front retreated to positions within the Lake Superior drainage basin. 160 �Figure 1. Field Trip 9 Overview: St. Louis River, Park Point, and Field Trip Stops. In the western Lake Superior basin, most of the glaciogenic sediment has been mapped as the Barnum Formation (Hobbs, 2003; Knaeble and Hobbs, 2009; Johnson et al., 2016). Within the Barnum Formation are three texturally distinct tills (i.e. members). The youngest tills are the Moose Lake and Knife River members. Between these two tills are deltaic and lacustrine sediments mapped as the Wrenshall Member, which were deposited in glacial Lake Wrenshall (Figure 2A; Wright et al., 1970; Johnson et al., 2016). Glacial Lake Wrenshall (also referred to as an early phase of glacial Lake Duluth) was a pro-glacial lake fronted by Superior Lobe ice within the Lake Superior basin (Breckenridge, 2013). Overflow of glacial Lake Wrenshall routed initially through the Moose Lake (or Portage) outlet into the Kettle River, but ice margin retreat opened the slightly lower Brule outlet into the St. Croix River (Figures 2A, 3). At this time, lake levels within the St. Louis estuary were at around 325 m (~1070 ft) above sea level, probably at the Epi-Duluth level (Figure 4). The ancestral St. Louis River entered the lake just downstream of the area which is now Jay Cooke State Park and deposited deltaic sediment that has not been mapped in detail. Because the ancestral St. Louis River drained meltwater and glacial lakes associated with the Rainy Lobe to the north (Figure 2A), the discharge must have been far greater than for the modern river. Finer-grained lacustrine sediments were deposited distal to the river mouth, including glacial varves (Wright et al., 1970; Knaeble and Hobbs, 2009; Hobbs and Breckenridge, 2010). Clay associated with these lacustrine sediments was quarried for bricks, near the town of Wrenshall during the late 19th and early 20th centuries (Grout, 1919). The timing of glacial Lake Wrenshall is poorly constrained. Basal radiocarbon dates from kettle lakes on the Nickerson moraine, (which precedes Lake Wrenshall), suggest the lake is at least as old as 12,800 cal 161 �yr BP (~10.8 14C ka), but these are minimum ages and likely lag ice retreat (Florin and Wright, 1969). Spruce logs have been dated to as old as 11,700 cal yr BP (~10.1 14C ka) on the former Lake Figure 2. Selected paleogeographic reconstructions of the Great Lakes region through time (adapted from Breckenridge et al., 2010). Outlets for glacial lakes are provided by white arrows. Lake Superior outlets varied through time, and include the Brule (BR), Au train-Whitefish (AWH), St. Mary’s River (SMR), and North Bay (NB). 162 �Wrenshall lake plain in Wisconsin and Michigan, which necessitates lake levels lowered below the Brule outlet by this time (Black, 1976). This requires that the Superior Lobe retreated far enough to the northeast to open lower elevation routes into the Lake Michigan basin. This would have resulted in the first period of downcutting of the former lake plain by the St. Louis River within the reach that is now the St. Louis estuary. Note that a disconformity from this period of low lake levels has not been found within the western basin, but the area lacks detailed stratigraphic research. The radiocarbon dated logs on the former Lake Wrenshall plain were buried by red clayey till, therefore the Superior Lobe re-advanced, blocked routes into the Lake Michigan basin, and raised lake levels back to the Brule outlet. This advance, known as the Marquette Advance in the eastern Lake Superior basin, created glacial Lake Duluth. The name Lake Duluth originates from Leverett (1928) after a series of prominent gravel beaches in the city of Duluth described by early geologists (Taylor, 1894). The Duluth level follows in a general way Skyline Parkway, but development and construction within the city of Duluth have obscured evidence of the strandline. Notable locations along Skyline Parkway that are at the lake Duluth level include a prominent terrace along the Spirit Mountain ski slope, the 1st United Methodist Coppertop church, and Heller Hall on the University of Minnesota Duluth campus where the Department of Geological Sciences is housed (Hobbs and Breckenridge, 2013). At the Lake Duluth level, overflow routed through the Brule outlet into the St. Croix River for a second time. Eventually ice retreat once again opened eastern outlets to the Lake Michigan basin causing lake levels to fall. Unlike the prior phase of lower levels, which is only inferred by wood on the former lake plain, the geomorphologic record of these lowered levels is preserved by a series of strandlines that are clearly visible on high resolution lidar DEMs (Figure 4; Breckenridge, 2013). Every one of these lower lake levels are named, typically after a town in which a strandline associated with the lake level is found (e.g. Highbridge, Washburn, and Beaver Bay) (Farrand, 1960; Farrand and Drexler, 1985). These strandlines are particularly useful for understanding the nature of glacial isostatic adjustment (GIA) because they can be traced around the basin. Glacial isostatic adjustment (or post-glacial rebound) is the rise of the lithosphere following deglaciation. Rebound rates are highest where the ice was thickest. In the Lake Superior basin, geomorphologic features provide a record of former lake levels (waterplanes) that rise in elevation to the northeast, where ice was thickest (Figure 3). Older waterplanes have undergone a longer period of GIA, and therefore rise more steeply. The end of glacial Lake Duluth occurred when ice retreat allowed a union with glacial Lake Minong, a lake that initially was limited in extent to Whitefish Bay in southeastern Lake Superior and fronted to the northwest by ice from the Marquette advance (Figure 2B). The merger between glacial lakes Duluth and Minong created a lake that retains the name glacial Lake Minong (Farrand, 1960) (Figure 2C). Glacial Lake Minong drained over a drift-covered bedrock sill, called the Nodaway barrier, into the St. Mary’s River (SMR). Progressive downcutting of the drift-covered sill at the St. Mary’s River caused falling lake levels in Lake Minong, which include the Minong through post-Minong levels (Figure 2C; Breckenridge, 2013). 163 �Figure 3. Map of Lake Superior with inlets and outlets (arrows), and contours for glacial isostatic adjustment used in Figure 4. Figure 4. Former waterplanes of Lake Superior (adapted from Breckenridge, 2013). 164 �The Holocene Not until ~2500 years after the onset of the Holocene ~9100 cal yr BP (8100 14C yr BP), did glacial meltwater cease to discharge into the glacial Lake Minong (Breckenridge et al., 2004; Hyodo and Longstaffe, 2011). The earliest post-glacial (i.e. post-Minong) lake level is called the Houghton, which was established when the Nadoway barrier was cut down to bedrock at the St. Mary’s River (SMR) (Figure 2D). Glacial isostatic adjustment (GIA) depressed the bedrock sill relative to the western basin, which resulted in the lowest lake levels ever in the Twin Ports region, probably 60-m below the modern level. In the Michigan and Huron basins, in situ stumps have been found that establish the immediate, post-glacial lake levels, and they appear to be ~20-m lower that the outlets for each lake (Lewis et al., 2008). This suggests a drier climate created closed-basins. No data has been found from the Superior basin to determine whether or not Lake Superior was a closed-basin lake at this time. By 8300 cal yrs BP, with the onset of a wetter climate, lake levels appear to have risen enough to overflow from Lake Huron into the North Bay outlet, which discharged to the Ottawa River (Figure 2D). During the early to mid-Holocene, differential GIA caused the outlet for Lake Huron (North Bay) to rise above the elevation of Lake Superior’s St. Mary’s River (SMR) outlet. Rising levels in Lake Huron drowned the SMR and created a shared waterplane between the Superior, Michigan, and Huron basins known as Lake Nipissing (Figure 2E; Larsen, 1985; Baedke and Thompson, 2000). Lake Nipissing levels peaked in the Huron and Michigan basins at around 4500 cal yr BP (Thompson et al., 2011). At this time, rising lake levels likely breached the basin drainage divide between the Huron and Erie basins, or perhaps the Michigan basin and Mississippi River via the Chicago River (Thompson et al., 2011; Johnston et al., 2012). This resulted in a major shift in the drainage pathway for the upper Great Lakes, from a route into the Ottawa River via North Bay (Figure 2E), to a southern outlet (Figure 2F). Subsequent lake levels dropped due to a combination of sill incision of the new outlet and perhaps climate change (Thompson et al., 2011). The Lake Nipissing level is commonly referred to as the Nipissing highstand, and strandlines formed at this level are readily apparent across much of the upper Great Lakes. At Sault Ste Marie (SSM), the Nipissing strandline is at 198 meters asl, 16-m above the modern level (Cowan, 1985), but the Nipissing elevation decreases to the west, converging towards modern lake levels due to GIA (Figure 4). Prior studies have suggested that Connor’s and Rice’s Point were a former baymouth bar formed by lake level rise to the Nipissing level (Loy, 1963; Barlaz, 1983). Longshore drift of sand eroded primarily from the southern shore, combined with sediment sourced from the Nemadji and St. Louis Rivers, likely built a spit across the head of the lake, but there have been no sediment or geomorphic studies to test this hypothesis. Former lake levels since the Nipissing highstand have been established for Lake Superior by coring and dating foreshore sand deposits from multiple strandplains in the eastern Lake Superior basin (Figure 5; Johnston et al., 2012). The work is an impressive undertaking, and is currently the most detailed paleohydrograph anywhere on the Great Lakes. The data indicate a lake-level drop of almost 4-m shortly after the peak Nipissing, followed by steady lake level lowering until around 1000 cal yr BP; thereafter lake levels have been stable near the Lake Superior outlet. The steady drop in lake levels from 4000 to 1000 cal yr BP was probably the result of GIA; an outlet on the southern side of a basin would have caused lake levels to fall everywhere north of the outlet. Stabilization of lake levels at 1000 cal yr BP is attributed to the separation of Lake Superior from Lake Huron (Johnston et al., 2012). Lake levels have been constant at SSM since this time, but this paleohydrograph must be corrected for GIA to understand former lake levels within the estuary and Twin Ports region. GIA is causing SSM to rise relative to the St. Louis River estuary. For example, gauge data suggest that lake levels in the estuary have risen around 25cm over the last 100 years due to differential GIA (Mainville and Craymer, 2005). 165 �For this field guide, an empirically derived model of GIA by Lewis et al (2005) has been adapted to correct the Sault Ste Marie paleohydrograph using the isobases of Breckenridge (2013) (Figure 6). The model applies an exponential decay function to estimate the rate of uplift necessary to result in warped strandlines in the Great Lakes of known age (Figure 6B). The modeled GIA correction is poorly constrained and could be in error, but the underlying processes that affected lake levels in the Twin Ports are generally understood. The resultant hydrograph (Figure 6C) suggests rapid drawdown from the Nipissing highstand around 4000 cal yr BP was too fast to be countered by relatively slow rates of GIA, resulting in a rapid lake level drop. Steadily falling lake levels at Sault Ste Marie from 4000 to 1000 cal yr BP were likely countered by GIA in the Twin Ports, which may have resulted in relatively stable lake levels. When lake levels stabilized at Sault Ste Marie at 1000 cal yr BP, lake levels would have risen in the Twin Ports. This abrupt change to rising lake levels is most likely responsible for drowning the St. Louis River to create the estuary. In addition, the rising lake levels probably initiated formation of Minnesota and Wisconsin Points. Ground penetrating radar surveys of Minnesota and Wisconsin Points suggest that the baymouth bar system is prograding lakeward (Morrison et al., 2015), presumably in response to lake level rise and increased sediment supply. One possibility is that the baymouth bar system is accreting vertically and lakeward on a former beach ridge that is just one of many that comprise a drowned strandplain now buried in the St. Louis Harbor. Evidence for this strandplain exists on the incredibly detailed bathymetric survey of the harbor and estuary completed by William Hearding in 1861 (Figure 7). Testing this model will necessitate detailed sediment and stratigraphic analysis of the harbor and baymouth bar sediments, combined with robust age dating. Figure 5. Lake levels at Sault Ste Marie since the Nipissing highstand, adapted from Johnston et al. (2012). 166 �Figure 6. Relative lake level curves for the Twin Ports (unpublished). A) Lake levels since 12,000 cal yr BP include a lowstand that exposes the glacial Lake Wrenshall plain, which is reflooded, presumably due to ice re-advance around 11,500 cal yr BP. Subsequent ice retreat lowers lake level in abrupt drop as new outlets open. The lowest level is the Houghton, which is also the end of glacial Lake Minong. Lake levels rise to the Nipissing due to GIA of the North Bay and St. Mary’s River outlets. B) Lake levels since the Nipissing highstand have been constrained Johnston et al., 2012 (see also Figure 5). Differential GIA caused levels to rise relative to Sault Ste Marie (SSM). C) Modeled GIA has been subtracted from the SSM hydrograph to estimate relative lake levels in the Twin Ports. Figure 7. Bathymetric data from Hearding (1861) converted to raster data and overlayed with modern LIDAR DEM. Water depths in the harbor were generally between 6 and 9 feet, except for the deep river channel. 167 �The Anthropocene Various Native American peoples inhabited the western Lake Superior area throughout the Holocene. After the glaciers receded, Paleo-Indian cultures were the first to inhabit the land, succeeded by Eastern Archaic peoples, until about 1,000 B.C (Dierckins, 2006). The Eastern Archaic peoples gave way to the Woodland cultures, until roughly 1600 A.D. Next were Dakota tribes, who were soon pushed west by the Ojibwe, near the same time as the arrival of the first explorers and gun/fur traders, including Daniel Greysolon, Sieur du Lhut. The 1842 and 1854 Treaties of La Pointe ceded rights of ownership near Lake Superior in areas of Wisconsin and Minnesota respectively, to European settlers in the region, ushering in development of the Industrial Age. The first iteration of modern locks at Sault Ste. Marie was completed in May 1855. In the Duluth/Superior Harbor, breakwaters were built, a shipbuilding industry began, and commercial fishing was established. Jay Cooke brought the Lake Superior & Mississippi Railroad to Duluth from St. Paul in the 1860’s, spurring logging throughout the region, with lumber mills appearing from Rice’s Point (Figure 1) to West Duluth (Dierckins, 2006). Grain elevators and railroad docks soon followed, connecting the waterfront to the railway. Other railroads began working their way to Duluth. In 1870 Duluth incorporated as a city. At this time the Superior Entry was the only waterway connecting Lake Superior to the Duluth/Superior Harbor. Soon after, the Duluth Shipping Canal was built from 1870 to 1877. This new connection to the lake changed the currents and hydrodynamics of the harbor. In the 1880’s, with the development of the Mesabi, Cuyuna, and Vermillion Iron Ranges, iron ore shipping began, along with the subsequent construction of docks to handle the ore. In 1907 Duluth surpassed New York City in shipping tonnage (Dierckins, 2006). The United States Steel Corporation began construction of a steel plant at Spirit Lake around this time. Shipbuilding operations were founded in Superior and later Riverside around the time of the two World Wars. The uppermost reach of the St. Louis River estuary became constrained by the Fond du Lac Dam in Jay Cooke State Park, when its construction was completed in 1924. In 1959 the St. Lawrence Seaway was created, allowing large ocean vessels into the Port. To accommodate ever-larger ships, navigational dredging deepened the harbor in shipping channels, and the dredge spoils were used variously as fill for the port, to create man-made islands (e.g., Barker’s Island, Figure 1), and the 90-acre Erie Pier dredge disposal facility. Recently, dredge spoils have begun to be beneficially reused to restore habitat at the 21st Ave. W. (Stop 3), 40th Ave. W., and Grassy Point sites (Figure 1). The Western Lake Superior Sanitary District (WLSSD) treatment plant began operating in September 1978, consolidating 17 inadequately treated wastewater discharges. Water quality in the St. Louis River rapidly improved from essentially a sewer condition to becoming suitable again for fishing and recreation. The St. Louis River Great Lakes Area of Concern (AOC) was established in 1987 by the EPA. Work continues today to remove several beneficial use impairments from the AOC. For example, river sediment cleanup projects have been initiated at Stryker Bay (Stop 4) and the Former US Steel plant upstream at Spirit Lake. A revitalization plan for West Duluth neighborhoods was begun in 2012 to capitalize on the neighborhood’s unique location along the St. Louis River corridor. Today, the harbor still handles many commodities ranging from coal, iron ore, grain, and limestone to cement, salt, wood pulp, wind turbine components and other heavy equipment. 168 �Field Trip Stops Vehicle Tour (meet at Stop 1 at 9:00am) Stop 1: Blue Heron Research Vessel – Lakehead Boat Basin (Breckenridge) Location: UTM Zone 15, 569377E 5180363N The RV Blue Heron is part of the fleet of University National Oceanographic System (UNOLS) vessels and is operated by the Large Lakes Observatory at the University of Minnesota Duluth as a charter vessel for research scientists. The vessel was built in 1985 for fishing the Grand Banks, but was sold to the University of Minnesota and converted for research in 1998. Over the last 18 years many researchers have used a wide array of equipment aboard the Blue Heron to extend our knowledge about the geologic and modern history of Lake Superior. Notable equipment includes acoustic seismic profilers that operate on a range of frequencies for imaging both shallow and deep sediments, side-scan sonar for imaging the sea floor surface and composition, and a piston corer capable of recovering sediment cores up to 9 meters in length. At this stop, we can tour the vessel and examine geophysical data collected aboard the Blue Heron from the harbor and greater lake. Examples include side scan sonar images of the harbor and the Thomas Wilson, a whaleback freighter that sank just outside the Duluth entrance in 1902. Selected sections from Lake Superior sediment cores BH02-5P, taken aboard the Blue Heron, will also be available for inspection. The core sections are on loan from the National Lacustrine Core Repository in Minneapolis where they are permanently archived in cold storage. Core BH02-5P (Figure 8) is from the deep Caribou sub-basin of Lake Superior, east of the Keeweenaw peninsula (Figure 3). BH02-5P has been correlated with photographs from S62-8, a long gravity core that penetrated to bedrock. By combining these records, 1406 varves (annual couplets) have been measured that overlie a red till and the Cambrian Jacobsville sandstone. The entire record dates to between 9.3 and 8.1 14C ka BP (11,500-9,100 cal yr BP). The most intriguing aspect of this record is a series of ~36 anomalously thick varves that correlate with those from the northern Lake Superior sub-basins. These varves are at the very top of the glaciolacustrine record, when the ice margin was most distal, and must have resulted from anomalously high sediment and water fluxes into the Superior basin (Breckenridge et al., 2004; Breckenridge, 2007). In the Isle Royale trough, these individual varves can be up to 14-cm thick. Presumably this 36-year event of anomalously high sediment flux was caused by abnormally high discharge by the ice sheet, or by the influx of anomalously great overflow from glacial Lake Agassiz which spilled into the basin (Figure 2C). These great floods of water may have caused short term rises in lake levels in the Michigan, Huron, and Superior basins; presumably the flux of water was too great to be accommodated by outlet channels. Evidence for rapid, short term increases in water level at around this time, perhaps 12-m or more, is found in small lakes that appear to be flooded by sediments from Lake Superior, including Fenton Lake (Breckenridge et al., 2010) and Beaver Lake (Fisher and Whitman, 1999). 169 Figure 8. Glacial varves (annual couplets) from BH02-5P (see Fig.3 for core location). Winter (W) and summer (S) sediment laminae are noted. �Stop 2: – Formation of Minnesota/Wisconsin Point and the St. Louis River Estuary (Kremmin/Breckenridge) Location: UTM Zone 15, 572438.5E 5175762.3N. PLS: T.49N., R.14W., S.13, NE1/4. Figure 9. View of Lake Superior, Duluth Harbor, Superior Bay, Minnesota Point, and Wisconsin Point from Enger Tower, Duluth, Minnesota (October 3rd, 2015). Photographer – Todd Kremmin. Introduction – Situated at the southwestern tip of Lake Superior, Minnesota Point and Wisconsin Point form a baymouth bar stretching northwest-southeast between Duluth, Minnesota and Superior, Wisconsin establishing the outer breakwater for the largest and farthest inland freshwater seaport in North America (Figure 9 and 10) (Duluth Seaway Port Authority, 2015). In combination, the baymouth bar of Minnesota and Wisconsin Points is approximately 16 kilometers in length and considered one of the longest freshwater bars in the world (Loy, 1962 & 1963; Bernard and Davidson, 1969). The material necessary for the formation of this sandy baymouth bar is attributed largely to littoral drift of sediments from the south shore of Lake Superior in Wisconsin, as well as secondary sources of sediment derived from outflows of the St. Louis and Nemadji Rivers. Research pertaining to the geomorphic development of the system has remained relatively surficial in nature: examining present day geomorphic structures, surveying bathymetry and topography, collecting surface sediment samples, extrapolating offset boreholes, and comparing analogous marine system bar developments (Merrill, 1939; Loy, 1962 & 1963; Kemp et al., 1978; Thomas and Dell, 1978; Sydor et al., 1979; Barlaz, 1983; Rasid and Hufferd, 1989; Rasid, et al., 1992; Albrecht, 2005; Hobbs, 2009). Moreover, research regarding the development of this baymouth bar in the context of post glacial isostatic rebound of the Lake Superior basin along with lake level variation has not been investigated. 170 �DULUTH Lake Superior Duluth Harbor Minnesota Point St. Louis River Park Point SUPERIOR Superior Bay Wisconsin Point Nemadji River Allouez Bay 0 1 Miles Figure 10. Map view of Lake Superior, Duluth Harbor, Superior Bay, Allouez Bay, Minnesota Point, Park Point, Wisconsin Point, City of Duluth, City of Superior, St. Louis River, and Nemadji River (Google Earth, 2016). The objective of this thesis research is to enhance previous observations of geomorphic development and continue further investigation into the subsurface of the baymouth bar system using Ground-Penetrating Radar (GPR). 40 kilometers of GPR data using both 250MHz and 100MHz antennae were obtained on Minnesota Point near the Park Point Recreational Area (Figure 11) for this project. To supplement and ground truth the GPR data, 8 vibracores (11.24 meters total) were drilled within the expanse of the acquired geophysical data (Figure 11). All sediment cores underwent Loss On Ignition (LOI) analysis and select portions of the cores were radiocarbon dated to provide chronology to the subsurface system and ultimately extrapolate findings across the entire baymouth bar. Background – Ground-Penetrating Radar GPR is a geophysical method (similar to seismic reflection) which enables indirect insight into the subsurface using radar energy. This geophysical method allows researchers to view large expanses of the subsurface without altering the landscape, uncovering the architectural radar stratigraphy of shallow well sorted fluvial, aeolian, or coastal sediments (Stoker et al., 1997; Jol and Bristow, 2003). Radar signals are “pinged” into the ground and reflected back to a receiver, recording travel times (a proxy for depth) and the variable signatures of properties and materials encountered. Within the region of Minnesota Point and 171 �specifically the Park Point Recreation Area, water saturation and organic content of the subsurface play an important role in the variations of reflections received. A variety of antennas can be used in GPR surveys, higher frequency antennas provide high resolution records, but limit subsurface penetration depths, lower frequency antennas provide deeper penetration depths, but limit resolution. When collected in tight grid spacing (≤1m), a spatially large (3D) volume of GPR data can be observed in planar dimensional slices with depth, permitting views on how systems have geomorphically changed over time at larger scales than traditional core observations. Lake Superior Superior Bay 100 0 Meters Figure 11. Map view of Park Point Recreation Area on Minnesota Point. GPR grids – (Grid_Orig, Grid_01, Grid_02, Grid_03) and lines are shown along with core locations - (1A, 1B, 1C, 2A, 2B, 4A, 5A, 6A). (Google Earth, 2016). Historical Context and Anthropogenic Influences The baymouth bar made up of Minnesota and Wisconsin Points’ remained in a natural state up until the 1860’s (Figure 12) (Hearding, 1863; Demeter, 1993; Zenith City, 2016). During this period Duluth had recently paved a railroad under the direction of Jay Cooke. This railroad provided a shaky commercial industry, which began to rival the neighboring city of Superior, Wisconsin. The City of Superior had been commercially well off prior to the increase in population within Duluth simply due to the natural opening 172 �of the baymouth bar. This natural opening, known as the Superior entry, allowed ships to enter from Lake Superior and easily harbor at the shores near the City of Superior. Around 1869 Duluth leaders discussed the idea of developing their own entry, but instead the Superior entry was securely modified with concrete N 0 1 Miles Figure 12. Historical Survey of the Northern and North Western Lakes made in obedience to acts of Congress and under the direction of The Bureau of Topographical Engineers of the War Department. Insets highlight the Duluth Canal prior to excavation (upper inset), along with the Park Point Recreation Area prior to dredge dumping (lower inset). (Hearding, 1863). breakwaters in 1869 by the U.S. Army Corps of Engineers (Demeter, 1993). Outraged, Minnesota leaders lobbied for their own ship canal, and, in 1870, began dredging. Businessmen of Superior filed a federal suit against these actions, fearing a new canal would diminish the City of Superior’s commercial industry. Shortly before the courier could arrive with the U.S. Supreme Courts orders to desist from excavating the canal, legend has it that all available citizens of Duluth had picked up their shovels and hand dug the remainder of the canal. This legend may be exaggerated, as other accounts claim the dredging operations handled most of the excavation. Regardless, a new unnatural entry had been built and marked the beginnings of the anthropogenic influences of the baymouth bar system (Zenith City, 2016). Following the establishment of the “Twin Ports,” dredging operations have been ongoing and extensive to keep channels open for shipping vessels. In the mid 1930’s Duluth Parks Superintendent Rodney Paine submitted a proposal to the federal government for $338,000 to develop a new recreation center on Minnesota Point (Figure 13). Unfortunately, this proposal was submitted in the midst of a financial crisis during the depression era, 173 �ending the idea before it even gained attention. President Franklin D. Roosevelt’s Works Progress Administration (WPA) program revived the Park Point Recreation Area project, and in 1936 workers brought in approximately 150,000 yards³ of fill to create new ‘land’ just south of 43rd street (Paine, 1936). A comparison of natural and man-made land is available to view in Figure 14. This historical context has relevance to this thesis project because a considerable amount of GPR data was collected on this new land made up of dredge material. These dredge spoils consist mainly of silty-sand and pebbles and can be readily differentiated between the natural underlying materials in both GPR and core observations. 100 0 Meters Figure 13. Preliminary drawings of proposed Park Point Recreation Area Project (Zenith City Online, 2016). N Natural System 1863 0 100 Meters Figure 14. LiDAR map showing existing Park Point Recreation area and Minnesota Point system (grayscale). Natural Minnesota Point system (yellow) (Hearding, 1863; DNR LiDAR Data, 2016). 174 �Results The most prominent grid - (Grid_02) (Figure 15) is 85m x 85m, with 1.0m spacing between lines, 0.05m step size, 0.25m antennae separation, and was collected in a half day using the 250 MHz Noggin Smartcart™ from Sensors & Software, Inc. Select planar slices were chosen showing amplitude variation of the evolving geomorphic system in 0.1m intervals (Figures 16 and 17). The lower right corner of each grid slice is North. Cross sections of GPR data are presented in Figures 18 and 19. Lake Superior A’ GRID_02 B’ B A Superior Bay 0 100 Meters Figure 15. Map view of Park Point Recreation Area on Minnesota Point. GPR Grid_02 is shown with cross section lines A-A’ and B-B’ - please reference Figures 18 and 19 to view cross sections (Google Earth, 2016). 175 �B’ A’ A A B C D B Figure 16. Grid_02 planar depth slices showing amplitude variation (red-high amplitude, blue-low amplitude). Variations in amplitude are believed to be derived from water and organic content variability in the subsurface (i.e. red = changing water/organic content with depth, blue = no change in water/organic content with depth). The lower right corner of each grid slice is North. Crosshatched region contains no data. Depths slices are indicated as the following intervals: A) 0.7-0.8m, B) 0.8-0.9m, C) 0.9-1.0m, D) 1.0-1.1m. A-A’ and B-B’ cross sections – refer to Figure 18 and 19. (Please refer to original color version for precise amplitude variation analysis in upcoming master’s thesis – Todd Kremmin UMD ‘16). High amplitude coherent variations of radar energy are visible (red) with depth beginning from 0.7m1.5m. These coherent packages seem to migrate North-Northwest over time and extend from Superior Bay towards Lake Superior/Duluth (Figures 16 and 17). Medium to lower amplitude variations of radar energy (yellow-blue) do not have as coherent of patterns in these depth ranges. Figure 16 (C) and (D) and Figure 17 (A) and (B) show blotchy medium-low amplitude variation in the right half of the depth slices. 176 �B’ A A’ A B C D B Figure 17. Grid_02 planar depth slices showing amplitude variation (red-high amplitude, blue-low amplitude). Variations in amplitude are believed to be derived from water and organic content variability in the subsurface (i.e. red = changing water/organic content with depth, blue = no change in water/organic content with depth). The lower right corner of each grid slice is North. Crosshatched region contains no data. Depths slices are indicated as the following intervals: A) 1.1-1.2m, B) 1.2-1.3m, C) 1.3-1.4m, D) 1.4-1.5m. A-A’ and B-B’ cross sections – refer to Figure 18 and 19. (Please refer to original color version for precise amplitude variation analysis in upcoming masters thesis – Todd Kremmin UMD ‘16). 177 �A A’ Figure 18. Cross Section GPR line A-A’ (Grid_02 Line X65). Depth in meters on left, time in nanoseconds on right, and position lines in meters at top. Red window shows planar depth slice from 1.3-1.4m. 0.0-0.7m is believed to be dredge material. 0.7-3.0m is believed to be dune clinoforms. 3.0-5.0m is thought to be glacio-lacustrine material. B B’ Figure 19. Cross Section GPR line B-B’ (Grid_02 Line Y15). Depth in meters on left, time in nanoseconds on right, and position lines in meters at top. Red window shows planar depth slice from 1.3-1.4m. 0.0-0.7m is believed to be dredge material. 0.7-3.0m is believed to be dune clinoforms. 3.0-5.0m is thought to be glacio-lacustrine material Conclusions Subsurface (0.7-3.0m) natural baymouth bar system development in a back bar setting indicates progradational/aggradational migration towards Lake Superior/Duluth. Above this interval (0.7-3.0m), incoherent radar amplitude reflectors differentiate the natural system with dredge material dumped here around 1936 (0.0-0.7m) (Paine, 1936). Core observations and radiocarbon ages support this hypothesis*. GPR cross section lines A-A’ show clinoformal migration towards Lake Superior/Duluth at depths (0.73.0m). On occasion, GPR cross sections happen to cut clinoformal structures along-strike or obliquely (AA’ position lines 20-35m), (B-B’ position lines 54-70m), negating the true stratigraphic architecture and dip of the reflectors. Loss of signal with depth indicates radar energy is attenuated (which happens when radar energy interacts with fine-grained materials). It is believed the baymouth bar system has developed over glacial/lacustrine material. The western Lake Superior Basin has seen increases in lake levels throughout the last ~2000 years due to differential isostatic rebound and changing outlets (Farrand and Drexler, 1985; Mainville and Craymer, 2005). This baymouth bar has responded to the increasing lake 178 �level in a progradational and aggradational fashion. Smaller climatic driven lake level variations are observed in detail from core observations. Although this baymouth bar is a young, non-marine system, reconstructing its geomorphic evolution in response to lake level change may become a useful analogue for similar, larger systems involved with base level change. In addition, stratigraphic findings of how such a system’s architecture is configured may yield insightful clues towards vintage conventional exploration reservoirs. Finally, a stronger understanding of how such a system geomorphically evolved in the context of the Lake Superior region post glaciation may aid in reshaping knowledge of how other geomorphic features and processes have developed throughout the region, perhaps providing tangible framework for future engineering and environmental management undertakings. *Sediment cores, radiocarbon ages, and additional GPR data will be available to view at Stop 2 during the field trip. Lunch at Barr Engineering 5th floor lunchroom (see handout map) Location: UTM Zone 15, 568970E 5181519N Vista Queen Tour (boarding at 1:00pm, departing 1:30pm) Stop 3: 21st Ave. West/Miller Creek restoration (Dott/Mossberger) Location: UTM Zone 15, 567314E 5178430N The St. Louis estuary is the largest estuary on Lake Superior, and is an important source of biological productivity, and wetland and aquatic habitat types for a wide variety of fish and wildlife communities. The 21st Avenue West Habitat Complex (Figure 20) is one of several habitat restoration projects that are incorporating beneficial reuse of navigational dredge material (Host, et al., 2013). The site is located near the WLSSD treatment plant, which discharges an average of 43 million gallons per day of treated wastewater. The intent of the restoration is to improve habitat by implementing remedial activities to address sediment contamination while complementing the desired ecological vision of stakeholder teams. The US Army Corps of Engineers started a pilot project for placement of material in the 21st Ave. W. Complex in 2013. Barr Engineering performed turbidity monitoring to help evaluate how suspended sediment generated during material placement would be transported, and provided preliminary cut and fill estimates for material quantities needed to implement the habitat plan. The MPCA received an approved Work in Public Waters permit in spring 2015 from the MNDNR for the design shown in Figure 20, so the final stages of the 21st Ave. W. Project may begin. It is anticipated that the base material and features will be constructed by 2017. 179 �Figure 20. 21st Ave. W. Placement Plans – Islands (modified from USACE 2016). Stop 4: SLIDRT, aka Stryker Bay (Dott/Mossberger) Location: UTM Zone 15, 563072E 5174671N The 255-acre St. Louis River/Interlake/Duluth Tar (SLRIDT) Superfund site is the largest sediment remediation project in Minnesota’s history. Seven decades of industrial use had left polycyclic aromatic hydrocarbons (PAHs) and other contaminants in the St. Louis River estuary. Barr’s work at the site included development of one of the nation’s first “hybrid” sediment remedies, which combined dredging and capping at the site; surcharge capping to maintain wetland habitat; use of an activated carbon mat to prevent re-contamination – the first commercial application of this technology; and integration of remediation, mitigation and restoration. Costing $90 million less than an all-dredging approach, the project restored 106 acres of aquatic and riparian habitat for fish, wildlife, and the Duluth community. Part of the investigation phase involved researching the history of the development of Stryker Bay and associated docks and slips. Figure 21 (SERVICE, 2002) shows a schematic of development of the site throughout the 98 years from 1903-2001. Fill (stippled areas in Figure 21) was brought into the protoStryker Bay in various phases, extending the land southward to accommodate industrial uses. Interestingly, the dock walls became confining barriers to shallow groundwater flow from the uplands to the river, causing artesian conditions to exist on the docks. An industrial water supply well was drilled near the southern tip on Dock 7, which was an artesian flowing well. After the site was investigated with numerous soil borings and sediment cores, the remedial project began with capping of impacted material in Slip 7 (Figure 22). In subsequent years, construction of a dike across Slip 6 created a Confined Aquatic Disposal facility (CAD). Impacted material was dredged from Stryker Bay and the shipping channel, hydraulically pumped to the CAD, and evenly spread throughout 180 �the CAD using a spreader barge. Part of Stryker Bay was separated from the rest with a sheet pile wall and was covered with an activated carbon mat and a surcharge to compress impacted material and cap it in place. The impacted material in the CAD was covered with an activated carbon mat, clean cap sand, and “environmental media” – material dredged and hydraulically pumped from a restoration project at Tallas Island approximately 2 miles upstream. The dike across the CAD was later removed, reconnecting it with the estuary. The site is currently in a long-term monitoring and maintenance phase to confirm that the constructed caps are properly containing the contaminants of concern and that aquatic plant and benthic communities are recovering consistently with other areas of the St. Louis River estuary. Figure 21. Industrial development of the SLRIDT site since 1903 (Modified from SERVICE, 2002). 181 �Figure 22. SLRIDT Caps and Covers Map (Modified from Hard Hat Services, 2013). Stop 5: St. Louis River and Glacial Lake Duluth Geomorphology: Pokegama Bay, Strandlines (Breckenridge/Mossberger) Location: UTM Zone 15, 562881E 5173280N The Pokegama River is a tributary of the St. Louis River, originating near Jay Cooke State Park on the Minnesota/Wisconsin border (Figure 23). It flows through the Superior Municipal Forest, where it empties into Pokegama Bay and the St. Louis River. The Pokegama-Carnegie wetlands are identified by the WDNR Bureau of Endangered Resources as a Lake Superior Basin Priority Site due to high quality wetland and occurrence of rare plant populations. The river is an important spawning area for walleye, northern pike, and other fish species. The Pokegama and its tributaries are deeply incised into red clay of the former lake bed, often forming steep stream banks with exposed clay. The exposed clay is susceptible to slumping and accelerated erosion (Mossberger, 2010). Slumping of stream banks, such as those in the Pokegama River, increases erosion and downstream sedimentation by supplying freshly exposed sediment to the stream. Slumping and erosion can also threaten the stability of nearby structures. If a slump intersects a confined aquifer, water under high hydraulic head in the confined aquifer can then seep to the surface, contribute baseflow, and increase the sediment load to the streams. Sedimentation from the St. Louis River deposits about half of the sediment yielded to the Duluth-Superior harbor (NRCS, 1996), and causes various economic and environmental problems. For example, a large plume of suspended sediment from the Pokegama River is often visible in the St. Louis River after rainstorms (Figure 24). 182 �Figure 23. LiDar image of Clough Island and Dwight’s Point at Pokegama Bay (DNR LiDAR Data, (2016). Figure 24. Aerial image of Clough Island and Dwight’s Point at Pokegama Bay (Google Earth, 2016). 183 �One economic impact of excessive turbidity is the cost of dredging. Dredging is necessary to maintain adequate draft for ships that use the 17 miles of harbor shipping channels. Insufficient draft requires ships to reduce their cargo load, leading to increased transportation costs. The United States Army Corps of Engineers (USACE) has estimated that approximately 33,000 tons of sediment each year is dredged from the Duluth-Superior harbor (NRCS, 1998). The St. Louis River’s contribution to the dredged sediment is approximately 14,000 tons, or 1,000 dump-truck loads, of sediment to the harbor annually, resulting in a burden to taxpayers for dredging and disposal of the river’s sedimentation. The USACE‘s current dredge disposal facility for the harbor, Erie Pier needed to be expanded and reengineered because it was expected to run out of capacity in 2017 due to the accumulation of excess fine-grained material. In addition to economic impacts, sediments can cause environmental problems. Suspended sediments are considered non-point pollution when they occur in high enough concentrations in designated surface waters. Industrial pollutants such as mercury, dioxins, and PCBs can attach to sediment particles, trapping toxins or oxygen-demanding materials in the harbor (Bridges, 2008). Dredging agitates and re-suspends settled pollutants, elevating toxin levels in the water and biota (MPCA, 1992). The harbor became one of the 43 Areas of Concern (AOC) under the Great Lakes Water Quality Agreement (WQA) in 1972. In 1987, Remedial Action Plans (RAPs) were developed to improve the health of the Nemadji and St. Louis Rivers. The harbor was designated as impaired for five uses, including fish consumption advisories, degradation of benthos, restrictions on dredging, degradation of aesthetics, and loss of fish and wildlife habitat (NRCS, 1998). In Pokegama Bay, there is a seasonal anoxic zone. The NOAA Sentinel Site Program is designed to create a national network of long-term research sites that measure the effects of climate change on our estuaries. The Lake Superior National Estuarine Research Reserve (LSNERR) has set up a Sentinel Site in Pokegama Bay. The goal will be to measure changes in climate and the associated effects on water quality, erosion, decomposition, marsh morphology, vegetation, and wildlife. A weather station, a water quality station, and 12 surface elevation tables (SETs) used to measure sediment accretion and subsidence were installed in 2011-2014. Acknowledgements Thanks to the following people and sources of information that helped shape this guidebook: Mehgan Blair and Pete Kero, Barr Engineering. Dan Breneman, Minnesota Pollution Control Agency. http://zenithcity.com/ http://wlssd.com/about-us/history/ http://www.lre.usace.army.mil/Missions/Recreation/SooLocksVisitorCenter/SooLocksHistory.aspx http://www.duluthport.com/port-history.php 184 �REFERENCES Albrecht, D.R. (2005). Geochemical Fingerprinting and the Role of the Nemadji River in Sediment Transport within Western Lake Superior. [Master’s Thesis]: University of Minnesota Duluth, Duluth, Minnesota. Barlaz, D.B. (1983). Sedimentation in the Duluth-Superior Harbor, Lake Superior. [M.S. Thesis]: University of Minnesota Duluth, Duluth, Minnesota. Bernard, J.M., and Davidson, D.W. (1969). A Floristic Resurvey of a Landfill Area 32 Years after Deposition: The Oatka Beach Addition, Minnesota Point, Minnesota. American Midland Naturalist, 82(2), 559-563. Black, R.F., 1976. Quaternary geology of Wisconsin and contiguous Upper Michigan. In: Mahaney, W.C. (Ed.), Quaternary Stratigraphy of North America. Dowden, Hutchinson, and Ross, Stroudesburg, Pennsylvania, pp. 93–117. Breckenridge, A., 2007. Lake Superior varve stratigraphy and implications for eastern Lake Agassiz outflow from 10,700 to 9,040 cal BP (9.5-8.1 14C ka). Palaeogeography, Palaeoclimatology, Palaeoecology 246, 45-61. Breckenridge, A. 2013, An analysis of the late glacial lake levels within the western Lake Superior basin based on digital elevation models: Quaternary Research v. 80, p. 383-395. Breckenridge, A.J., Johnson, T.C., Beske-Diehl, S., Mothersill, J.S., 2004. The timing of regional late glacial events and post-glacial sedimentation rates from Lake Superior. Quaternary Science Reviews 23, 2355-2367. Breckenridge, A., Lowell, T.V., Fisher, T.G., Yu, S., 2010. A late Minong transgression in the Lake Superior basin as documented by sediments from Fenton Lake, Ontario. Journal of Paleolimnology 47, 313–326. Bridges, T.S., et al., 2008. The Four Rs of Environmental Dredging: Resuspension, Release, Residual, and Risk. U.S. Army Corps of Engineers Dredging Operations and Environmental Research Program, ERDC/EL TR08-4. Cowan, W.R., 1985. Deglacial Great Lakes shorelines at Sault Ste Marie, Ontario. In: Karrow, P.K., Calkin, P.E. (Eds.), Quaternary Evolution of the Great Lakes. Geological Association of Canada Special Paper 30, 3337. Demeter, S.C. (1993). U.S. Army Corps of Engineers Historic American Engineering Record No. MN-10: Duluth Ship Canal. Commonwealth Cultural Resources Group, Inc. Dierckins, Tony. Zenith: A Postcard Perspective of Historic Duluth. Zenith City Press (X-Communication), Duluth, Minnesota: 2006. DNR LiDAR Data, (2016). Retrieved from: http://arcgis.dnr.state.mn.us/maps/mntopo/ Duluth Seaway Port Authority, (2015). Retrieved from: http://www.duluthport.com/port.php Farrand, W.R., 1960. Former Shorelines in Western and Northern Lake Superior Basin. PhD dissertation, University of Michigan, Ann Arbor, MI, 266 pp. Farrand, W.R., & Drexler, C.W. (1985). Late Wisconsinan and Holocene History of the Lake Superior Basin. In: Karrow, P.K., and Calkin, P.E. (Eds.) Quaternary Evolution of the Great Lakes: Geological Association of Canada Special Paper, 30, 17-32. Fisher, T.G., Whitman, R.L., 1999. Deglacial and Lake Level Fluctuation History Recorded in Cores, Beaver Lake, Upper Peninsula, Michigan. Journal of Great Lakes Research 25, 263-274. Florin, M., and Wright, H.E., Jr., 1969, Diatom evidence for the persistence of stagnant glacial ice in Minnesota: GSA Bulletin, v. 80, p. 695-704. Grout, F., 1919. Clays and Shales of Minnesota, United States Geological Survey, Bulletin 678, Washington, Government Printing Office. Hard Hat Services, 2013. Long Term Monitoring & Maintenance Plan. St. Louis River/Interlake/Duluth Tar Site, Sediment Operable Unit Remediation. Prepared March 14, 20013 for XIK Corp. 185 �Hearding, W.H. (1863). Survey of the Northern and North Western Lakes: West End of Fond Du Lac of Lake Superior Map. The Bureau of Topographical Engineers of the War Department. Hobbs, H.C., 2003, Surficial geology of the Knife River quadrangle, St. Louis and Lake Counties, Minnesota; Minnesota Geological Survey Miscellaneous Map M-137, scale 1:24,000. Hobbs, H.C. (2009). Surficial Geology of the Duluth Quadrangle, St. Louis County, Minnesota. Minnesota Geological Survey. Hobbs, H.C., Breckenridge, A., 2011. Ice advances and retreats, inlets and outlets, sediments and strandlines of the western Lake Superior basin. In: Miller, J.D., Hudak, G.J., C. Wittkop, C., McLaughlin, P.I. (Eds.), Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America, GSA Field Guide 24, 299-315. Hyodo, A., Longstaffe, F.J., 2011. The chronostratigraphy of Holocene sediments from four Lake Superior subbasins. Canadian Journal of Earth Sciences 48, 1581–1599 . Host, James, Buttermore, et al., An Ecological Design for the 21st Avenue West Remediation-to-Restoration Project. Natural Resources Research Institute, University of Minnesota Duluth Large Lakes Observatory, U.S. Fish and Wildlife Service. USFWS Cooperative Agreement F11AC00517June 2013 NRRI/TR2013/24. Johnson, M. D., Adams, R. S., Gowan, A. S., Harris, K. L.; Hobbs, H. C., Jennings, C. E., Knaeble, A.R.; Lusardi, B. A.; Meyer, G. N.. (2016). RI-68 Quaternary Lithostratigraphic Units of Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, http://hdl.handle.net/11299/177675. Johnston, J.W., Argyilan, E.P., Thompson, T.A., Baedke, S.J., Lepper, K., Forman, S.L., Wilcox, D.A., 2012. A Sault-outlet-referenced mid- to late-Holocene paleohydrograph for Lake Superior constructed from strandplains of beach ridges. Canadian Journal of Earth Sciences 49, 1364-1371. Jol, H.M., & Bristow, C.S. (2003). GPR in Sediments: Advice on Data Collection, Basic Processing and Interpretation, a Good Practice Guide. Geological Society London, Special Publications, 211, 9-27. Kemp et al., (1978). Sedimentation rates and a sediment budget for Lake Superior. Journal of Great Lakes Research, 4, 276-287. Knaeble, A.R., and Hobbs, H.C., 2009, surficial geology, Plate 3, and Quaternary stratigraphy, Plate 4 of Boerboom, T.J., project manager, geologic atlas of Carlton County, Minnesota: Minnesota Geological Survey County Atlas Series C-19, Part A, 6 plates, scale 1:100,000. Larsen, C.E., 1985. Lake Level, uplift, and outlet incision, the Nipissing and Algoma Great Lakes. In: Karrow, P.K., Calkin, P.E. (Eds), Quaternary Evolution of the Great Lakes: Geological Association of Canada Special Paper 30, pp. 63-77. Leverett, F. 1928. Moraines and shorelines of the Lake Superior basin. USGS paper 154-A, p. 1-72. Lewis, CFM, et al., 2008. Dry climate disconnected the Laurentian Great Lakes, EOS 89(52), 541-542. Loy, W.G. (1962). The Coastal Geomorphology of Western Lake Superior. [Master’s Thesis]: University of Chicago, Chicago, Illinois. Loy, W.G. (1963). Formation of the Duluth-Superior Harbor. The Minnesota Academy of Science Proceedings, 31, 28-35. Loy, W.G. (1963). The Evolution of Bay-Head Bars in Western Lake Superior. Journal of Great Lakes Research, 10, 150-157. Mainville, A., & Craymer, M.R. (2005). Present-Day Tilting of the Great Lakes Region Based on Water Level Gauges. Geological Society of America Bulletin, 117(7), 1070. Minnesota Pollution Control Agency and Wisconsin Department of Natural Resources, 1992. St. Louis River System Remedial Action Plan, Stage One. Merrill, J.A. (1939). The Wonderland of Lake Superior. Minneapolis, Minnesota: Burgess Publishing Company. 186 �Morrison, S. M., Jol, H. M, Alger, R. Using radar stratigraphic analysis to identify erosion and deposition in the Duluth bay barrier, Lake Superior, NC-GSA poster presentation, Madison, WI. Mossberger, I. G. (2010). Potential For Slumps, Sediment Volcanoes, And Excess Turbidity In The Nemadji River Basin (Masters Thesis, University of Minnesota-Duluth). Natural Resources Conservation Service, 1996. Memorandum. Natural Resources Conservation Service and U.S. Forest Service, 1998. Executive summary report: Erosion and sedimentation in the Nemadji River Basin: Nemadji River Basin Project, 16 p. Paine, R.F (1936). City of Duluth Park Department – Annual Report 1936. Duluth Public Library, Duluth, Minnesota. Call 352.7. No. D88a. Vol.1936. Rasid, H., & Hufferd, J. (1989). Hazards of Living on the Edge of Water: The Case of Minnesota Point, Duluth, Minnesota. Human Ecology, 17, 85-100. Rasid et al., (1992). Coping with Great Lakes Flood and Erosion Hazards; Long Point, Lake Erie, vs. Minnesota Point, Lake Superior. Journal of Great Lakes Research, 18(1), 29-42. SERVICE Engineering Group, 2002. Geology-Hydrogeology (of the) St. Louis River/Interlake/Duluth Tar Site, Duluth, MN. Appendix GH of the Data Gap Report. Stoker, M.S., Pheasant, J.B., and Josenhans, H., 1997, Seismic Methods and Interpretations, in Davies, T.A., et al., eds., 1997, Glaciated Continental Margins: An Atlas of Acoustic Images; London, Chapman & Hall, 315 p. Sydor et al., (1979). Red Clay Turbidity and its Transport in Lake Superior. Environmental Protection Agency Document. Grant No. R005175-01. Taylor, F. B., 1894. A reconnaissance of the abandoned shore lines of the south coast of Lake Superior. American Geologist, 365-383. Thomas, R.L. & Dell, C.I. (1978). Sediments of Lake Superior. Journal of Great Lakes Research, 4(3-4), 264-275. Thompson, T.A., Lepper, K., Endres, A.L., Johnston, J.W., Baedke, S.J., Argyilan, E.P., Booth, R.K., Wilcox, D.A. 2011. Mid Holocene lake level and shoreline behavior during the Nipissing phase of the upper Great Lakes at Alpena, Michigan, USA. Journal of Great Lakes Research 37, 567–576. U. S. Army Corps of Engineers, 2016. 21st Ave. West Restoration Placement Plan – Islands. Sheet GI001. Unpublished. Wright, H.E., Jr., Mattson, L.A., Thomas, J.A., 1970. Geology of the Cloquet quadrangle, Carlton County, Minnesota: Minnesota Geological Survey Geologic Map Series GM-3. Zenith City Historical Archive, (2016). Retrieved from: http://zenithcity.com/zenith-city-history-archiv 187 ��1 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2016 Description An account of the resource Institute on Lake Superior Geology. Duluth, Minnesota. May 4-8, 2016. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2016 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/8cc74af26712cd59ea3d647b21961546.pdf 465d7b12c828300cfef32450ada060e7 https://digitalcollections.lakeheadu.ca/files/original/9d120470c1c05757ed646d2dee949394.pdf 28759b6a6e8f43b02466eb45a44d0338 PDF Text Text 63rd Annual Meeting Institute on Lake Superior Geology Wawa, Ontario May 8 - 12, 2017 Wawa Wild Goose or Land of the Big Goose in Ojibway Proceedings Volume 63 Part 2: Field Trip Guidebook �Institute on Lake Superior Geology 63rd ANNUAL MEETING May 8-12, 2017 Wawa, Ontario SPONSORED BY: ONTARIO GEOLOGICAL SURVEY MINISTRY OF NORTHERN DEVELOPMENT AND MINES AND A. E. SEAMAN MINERAL MUSEUM MICHIGAN TECHOLOGICAL UNIVERSITY Meeting Co-Chairs Anthony Pace, Ann Wilson, and Theodore J. Bornhorst Proceedings Volume 63 Part 2: Field Trip Guidebook Compiled by Theodore J. Bornhorst and Margaret J. Hanson Cover Photos: Top left, Calcite crystals on pyrite from George W. MacLeod Mine, Wawa, collection of the A. E. Seaman Mineral Museum donate, photograph by George Robinson; iconic Wawa goose at junction of Trans-Canada Highway and Highway 101, image from http://www.northernontario.travel/algoma-country/fun-facts-about-the-famous-wawa-goose; Sir James Dunn open pit iron mine on the Eleanor Iron Range, photograph by Anthony Pace. �63rd INSTITUTE ON LAKE SUPERIOR GEOLOGY VOLUME 63 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD TRIP GUIDEBOOK TRIP 1: ARCHEAN AND PROTEROZOIC GEOLOGY OF THE MARATHON-HEMLO AREA Day 1: Geology of the Coldwell Alkaline Complex Part 1: Transect through the Coldwell alkaline complex Part 2: Marathon Cu-PGM deposit Day 2: Geology of the eastern Schreiber-Hemlo Greenstone Belt in the vicinity of Heron Bay and Hemlo TRIP 2: MORE UNUSUAL DIAMOND-BEARING ROCKS OF THE WAWA AREA TRIP 3: GEOLOGY OF THE WAWA GOLD PROJECT TRIP 4: GEOLOGY OF THE ISLAND GOLD MINE TRIP 5: GEOLOGY OF THE RENABIE AREA TRIP 6: KAPUSKASING STRUCTURAL ZONE AND BORDEN LAKE GOLD DEPOSIT Reference to material in Part 1 should follow the example below: Authors, 2017, abstract title, 63rd Institute on Lake Superior Geology Proceedings, v. 63, Part 1, Field Trip Guidebook, p. xx. Proceedings Volume 63, Part 1: Program and Abstracts, and Part 2: Field Trip Guidebook are published by the 63rd Institute on Lake Superior Geology and distributed by the Institute Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca Some figures in this volume were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the volume when it is available on-line at: http://www.lakesuperiorgeology.org ISSN 1042-9964 i �Part 1: Field Trip Guidebook Table of Contents Trip 1: Archean and Proterozoic geology of the Marathon-Hemlo area 1 Trip 1, Day 1: Geology of the Coldwell Alkaline Complex 1 Trip 1, Day 2: Geology of the eastern Schreiber-Hemlo Greenstone Belt 44 Trip 2: More unusual diamond-bearing rocks of the Wawa area 79 Trip 3: Geology of the Wawa gold Pproject 109 Trip 4: Geology of the Island Gold Mine 148 Trip 5: Geology of the Renabie area 164 Trip 6: Kapuskasing structural zone and Borden Lake gold deposit 187 ii �Field Trip 1 Archean and Proterozoic geology of the Marathon-Hemlo area Day 1 Geology of the Coldwell Alkaline Complex The Coldwell Alkaline Complex portion (Day 1) of Field Trip 1 consists of two parts: 1) a west to east transect through southern part of the complex and 2) a visit to the Marathon Cu-PGE Deposit. 1 �Part 1: Transect Through the Coldwell Alkaline Complex Allan MacTavish1 and Mark Smyk2 Panoramic PGMs (Canada) Limited, Thunder Bay, ON; 2 Resident Geologist Program, Ontario Geological Survey, Thunder Bay, ON 1 A variety of Mesoproterozoic, Midcontinent Rift-related alkalic and carbonatitic rocks occur within several intrusive complexes on or near the northern shore of Lake Superior (see Figures 1 and 2). They include the Coldwell and Killala Lake alkaline complexes, the Prairie Lake and Chipman Lake carbonatites, and numerous diatremes and related dikes in the vicinity of Dead Horse Creek (Sage 1982, 1985, 1987) (see Figure 2). These complexes are spatially localized and structurally controlled by the Trans-Superior Tectonic Zone (TSTZ), a north-northeast-trending structure that extends for over 600km and includes the Thiel Fault in Lake Superior (Klasner et al. 1982). Alkaline magmatism related to Midcontinent rifting occurred along the TSTZ from approximately 1.2 to 1.0 Ga, as is shown in Table 1 below: Table 1 Lithologic Unit/Complex Age(s) (Method) Reference Coldwell Alkaline Complex 1108 ± 1 Ma (U/Pb) Heaman and Michado (1987) Be-Zr Zone crosscutting Dead Horse Creek diatreme Prairie Lake Carbonatite Lamprophyre Dyke, McKellar Harbour Gabbro (biotite), Killala Lake Complex 1112.7 ± 4 Ma (U/Pb)1 1128.7 ± 6 Ma (U/Pb)2 1130 ± 10Ma (Rb/Sr) 1145 + 15/10 Ma (U/Pb) 1185 ± 90 Ma (K/Ar) Krogh and Wilkinson (M. Smyk pers. Comm. 1995) Pollock (1987) Queen et al. (1996) Coats (1970) Syenite, Killala Lake Complex 1050 ± 35 Ma (Rb/Sr) Bell and Blenkinsop (1980) (1-2.49% discordant; 2 1.82% discordant) It has been postulated that the TSTZ may represent part of a failed arm of a Keweenawan-age triple junction (Weiblen 1982; Mitchell and Platt 1982b) or the intersection of a late fracture system with the rift (Mitchell et al. 1983). Local alkalic and carbonatite complexes have been emplaced at inflections in the trends of major structural zones, or at sites of cross-faulting (Sage 1991). The Coldwell and Killala Lake alkaline complexes are both thought by some to have formed as the result of ring fracturing and caldera collapse. The abundance of observed xenolithic blocks and roof pendants suggests that these complexes are presently exposed at relatively high structural levels. Similar ages for numerous mafic intrusions in the Nipigon Embayment (cf. Heaman et al. 2007) and the alkalic rocks of the Coldwell Complex (1108 ± 1 Ma; Heaman and Machado 1992) indicate the contemporaneous production of tholeiitic and alkalic magmas during Midcontinent rifting. The oldest magnetization, found in the gabbros and augite syenites on the eastern side of the complex, records a concordant pole position with reversed polarity at about 1109 ± 5 Ma on the Keweenawan segment of Precambrian apparent polar wander path (Lewchuk and Symons 1990). The localization of the alkalic magmatism off-axis, dominantly northeast of the central rift, prompted Heaman and Machado (1992) to suggest that this may have been a region of maximum lithospheric extension during rifting. U/Pb data (Heaman and Machado 1992) demonstrate that most rock units in the Coldwell Complex were emplaced within a relatively short time span (<3 million years) ca. 1108 Ma, and support the contention that the complex experienced relatively rapid cooling from initial emplacement temperatures to at least ~500º C. Strontium-, neodymium- and lead-isotopic compositions of selected minerals from different phases of the complex (Heaman and Machado 1992) display considerable scatter, suggesting that their magmas had different isotopic compositions. The initial strontium- and neodymium-isotopic compositions of clinopyroxene and plagioclase from one of the earliest gabbroic phases are identical to data derived from primitive olivine tholeiites from the Midcontinent Rift and indicate that the majority of magmas, both 2 �tholeiitic and alkaline, have a uniform, nearly chondritic isotopic composition (ibid). Samariumneodymium data, supported by oxygen-isotopic and whole-rock geochemical data, indicate that crustal contamination played a small, varied role in the generation of the Coldwell magmas (Bohay 1997). In addition to small, variable amounts of assimilation of upper and lower crust, the parental plume magmas also interacted with the lithospheric upper mantle to a small degree (ibid). Local alkalic and carbonatitic intrusive rocks host a variety of characteristic base, precious, titaniferous, phosphate, and rare metal occurrences (cf. Smyk and Sage 1995). They include the following: 1. 2. 3. 4. 5. Magmatic Cu-Ni-PGE (± Au, Ag) in gabbros of the Killala Lake and Coldwell complexes; Magmatic Ti-V±apatite deposits in the Eastern Border Gabbros of the Coldwell Complex; Magmatic U, Nb (+ wollastonite, apatite) in the Prairie Lake carbonatite (Sage 1987); Late-stage magmatic Nb-Y-F-family rare earth elements in syenite pegmatites (Alexander 2007); A Be-Zr-U-Th-Y mineralized zone crosscutting the Dead Horse Creek diatreme (Smyk et al. 1993; Potter, 2004); and 6. Pb-Zn-Ag-mineralized quartz-carbonate veins (Kissin and McCuaig 1988). The Coldwell Alkaline Complex (see Figure 3) covers an area of ~580km2, making it one of the largest alkalic complexes in the world and the largest in North America. It was emplaced during the early stages of the Midcontinent rift system, which includes: early large and small mafic to ultramafic intrusions (i.e. Seagull Lake Complex, Thunder Bay North Complex); Keweenawan flood basalts, the Duluth Complex, the Nipigon and Logan sills, and a variety of non-diabase mafic to ultramafic dyke-rocks. The Coldwell Complex was mapped by Kerr (1910a, 1910b), Puskas (1967), and Walker et al. (1993b, 1993c), and comprises 3, superimposed ring sub-complexes or magmatic centers (Mitchell and Platt 1978) that young progressively (Centers 1 to 3) to the southwest (see Figure 4). Walker et al. (1993) and Barrie et al. (2002) dispute the series of ring dykes or sheeted cones interpretation and suggest that the complex is a composite lopolith or sill. The intrusive centres can be generally described as follows: 1. Center 1: Generally silica-saturated rocks with oversaturated residue; chiefly consisting of the Eastern and Western border gabbros (the oldest rocks within the complex) and later iron-rich augite syenite and syenite-syenodiorite (Mitchell and Platt, 1978, 1982; Mulja, 1989); 2. Center 2: Generally silica-undersaturated alkalic rocks with oversaturated residue; consisting of locally nepheline- and hastingsite-bearing miaskitic nepheline syenite, and numerous volumetrically minor alkaline lamprophyre and analcime tinguaite dykes (Mitchell and Platt, 1978, 1982; Laderoute, 1989; and Mulja, 1989); and 3. Center 3: Silica-oversaturated alkalic rocks with oversaturated residue; consisting of magnesio-hornblende syenites, quartz syenites, and minor granites (Mitchell and Platt, 1994; Lukosius-Sanders, 1988). The mineralogy of the main lithologic units is listed in Table 2. The superimposition of the three intrusive centres and a complex, protracted magmatic history has produced a myriad of hybrid rocks, igneous breccias, and ambiguous crosscutting relationships. The wide variety of lamprophyric and other dyke rocks occurring within the complex (as described by Mitchell and Platt, 1994) include (in order of emplacement): 1. 2. 3. 4. 5. 6. Mafic ocellar lamprophyre (camptonitic variety) Quartz-bearing, mafic lamprophyres (camptonitic variety) Sannaite-type lamprophyres Monchiquitic-type lamprophyres Feldspar glomeroporphyritic and alkali basalt dikes Analcime tinguaite (heronite) 3 �Figure 1. Mid-continent Rift geology and the locations of mafic/ultramafic intrusions (After Miller et al. 1995). 4 �Figure 2. Regional geology (Sage 1991) in the vicinity of the Trans-Superior Tectonic Zone (TSTZ), extension of the Thiel Fault (B). Key to numbering: 30 – Chipman Lake fenites / carbonatite dikes; 31 – Killala Lake alkaline complex; 32 – Prairie Lake Carbonatite; 33 – Coldwell alkaline complex; 36 – Slate Islands; 47 – Dead Horse Creek diatremes; 48 – McKellar Creek diatreme; 49 – Gold Range Diatreme; 50 – Neys Diatreme; A – Michipicoten Fault; C – Killala Lake Deformation Zone. 5 �Figure 3. Generalized geology of the Coldwell Alkaline Complex (after Walker et al. 1993) with field trip stops. 6 �Figure 4. Coldwell Alkaline Complex magmatic centres: CI (Centre 1), CII (Centre 2), and CIII (Centre 3). Generalized geology after Mitchell and Platt, 1994. 7 �Abundant large rafts and/or roof pendants of mafic volcanic rocks are mapped throughout the Coldwell Complex and in places exhibit horizontal extensional cooling cracks on a ten to hundreds of metres scale that are thought consistent by some workers with sub-horizontal bedding. For the most part the roof pendants may be the lowermost portions of the Keweenawan flood basalt sequences suggesting that the complex is barely unroofed and is exposed at a very shallow crustal level (Mitchell and Platt 1994; Sage and Watkinson 1995; Barrie et al. 2002). It is also highly probably that some, or most of the mafic rafts observed within the complex that could not be roof pendants are detached portions of chilled complex roof or wall rocks (fine-grained gabbros). The mafic intrusive rocks occurring within Centers 1 and 2 are tabulated, with their associated mineralized zones, below: Table 2 Intrusion (Centre) Eastern Gabbro (1) Lithologic Units Reference(s) Shaw (1994, 1997); Lum (1973); Barrie et al. (2002) Malpas Lake (1) Layered gabbro cumulates (olivine gabbro, gabbro, troctolite, anorthositic (leuco-) gabbro); Fe-Ti oxide ± apatite cumulates Massive and layered series gabbro; olivine-bearing Gabbronorite, olivine gabbronorite, olivine-bearing gabbro, leucogabbro Hornblende gabbro to monzodiorite; olivine ferrogabbro to ferrodiorite; olivine gabbro to diorite Amphibole-bearing olivine gabbro Geordie Lake (2?) Troctolite, olivine gabbro Alkalic gabbro (2) Biotite gabbro Biotite- and olivine-gabbro Mulja (1989); MacTavish et al. (1987) Mitchell and Platt (1982b) Walker et al. (1993a) Western Gabbro (1) Two Duck Lake (1) Penczak (1992); Wilkinson (1983) Shaw (1994, 1997) Dahl et al. (1987) Shaw (1994, 1997) Magmatic, gabbro-hosted Cu-Ni-PGE deposits in the Coldwell Complex have been the focus of much exploration and research for the past 60 years. Mineralized zones occur within the border gabbro at the eastern (Marathon deposit; Skipper Lake Zone) and western (Middleton occurrences) margins of the complex, and within its interior at Geordie Lake. The Geordie Lake mineralized zones, hosted by a younger (?) gabbro, are enriched in tellurium and silver and have higher Pd:Pt ratios (~19) (Mulja and Mitchell 1991) than the border gabbro-hosted deposits (~4) (Smyk 2001). Geochemical variations in mineralized zones in the Coldwell Complex are shown in Figure 5. A table of selected Coldwell Complex deposits and mineralized zones is shown below (see Table 3). 8 �Table 3 Mineralized Zone Grade / Significant Assays Marathon Measured and Indicated In-Pit Resources : 114.8 Mt @ 0.775 g/t Pd, 0.228 g/t Pt, 0.083 g/t Au, 0.241% Cu, 1.567 g/t Ag; Proven and Probable In-Pit Reserves: 91.447 Mt @ 0.832 g/t Pd, 0.237 g/t Pt, 0.085 g/t Au, 0.247% Cu, 1.440 g/t Ag (January 2010) Chalcopyrite Marathon PGM > cubanite  pyrite ; s Corporation hollingworthite, atokite-zvyaginstevite, Ohnenstetter et al. sperrylite, Bi-kotulskite, michenerite, (1991); Watkinson merenskyite, monceite, stibiopalladinite, and Ohnenstetter paolovite, merteite II, palladoarsenide, (1992); Good and unnamed (Pd5As2), nickeline, majakite, Crocket (1994a, argentian gold 1994b) Measured and Indicated Resources (above $13.00/t cut-off): 32.42 Mt @ 0.61 g/t Pd, 0.04 g/t Pt, 0.05 g/t Au, 0.37% Cu, 2.93 g/t Ag Chalcopyrite, bornite, pyrite, millerite, siegenite, pentlandite, galena, chalcocite, melonite, hessite, unnamed (Ag3Te2), altaite, kotulskite, merenskyite, michenerite, sopcheite, Pd-bismuthotelluride, paolovite, Pd-arsenide, guanglinite, Pd-antimonide, sperrylite, electrum, Pd1.6As1.5Ni, AgSb4 Chalcopyrite, pyrrhotite, pentlandite, sphalerite, pyrite Chalcopyrite, bornite, pentlandite, cobaltite, galena, chalcocite; telargpalite, polarite, kotulskite, taimyrite, merteite, zvyagintsevite, plumbopalladinite, majakite, tetraferroplatinum n/a Geordie Lake Middleton Skipper Lake Area 41 average grade of 1.05 g/t Pd+Pt+Au over 12 m 0.48 g/t Pt+Pd+Au over 202 m, incl. 1.23 g/t Pt+Pd+Au over 61 m Ore Mineralogy Reference(s) news release, Marathon PGM Corporation, May 04, 2010 Mulja (1989); Mulja and Mitchell (1990, 1991) Penczak (1992) MacTavish (2000) Benton Resources Corp. Marked similarities exist between the mineralization style, geochemistry, and host rocks of Coldwell Complex-, Duluth Complex-, and the Crystal Lake gabbro-hosted deposits near Thunder Bay. Similarities include mineral textures, abundance and compositions, crystallization paths for the host gabbros, silicate-sulphide associations, trace-element trends and chalcophile element fractionation trends (Good and Crockett 1994a). Research by Watkinson and Ohnenstetter (1992) and Good and Crockett (1994a, 1994b) produced debate between the relative importance of magmatic and hydrothermal processes in local copper-nickel-PGE mineralization processes. Watkinson and Ohnenstetter (1992) presented field, petrographic and mineral-chemical data that support the interaction of magmatic sulphide mineral assemblages with a chlorine-rich mixture of magmatic (deuteric) fluid and volatile species generated by the breakdown of assimilated xenoliths at low temperatures. However, Good and Crockett (1994a, 1994b) contended that element migration took place over only very short distances and that the original, bulk sulphides were not enriched in copper and PGE by later fluids. The information within this field trip guide was taken from a variety of sources, including guidebooks from previous field trips to the Coldwell Complex: Puskas (1970); Loubat (1972); Mitchell and Platt (1977, 1982a, 1994); Smyk and Sage (1995), Smyk (2001), Smyk (2010), and unpublished field observations and mapping completed by A. MacTavish (1992). All UTM co-ordinates listed are NAD83 Zone 16. 9 �Figure 5. Discrimination plot for PGE-mineralized samples for Coldwell and other Midcontinent Rift-related intrusions. Data from Good (1992), MacTavish (unpublished data, Resident Geologist’s Files, Thunder Bay), Watkinson et al. (1983), Wilkinson (1983), and unpublished data, Resident Geologist’s Files, Thunder Bay. Duluth Complex composite data from Hauck et al. (1997). 10 �Coldwell Alkaline Complex Transect - Field Trip Stops STOP C1: Natrolite-Bearing Syenite and Massive Fe-Ti-oxides Location: Lat: 48o47'45" N, Long: 86o39'15"W; UTM 525528E, 5405511N 29.4 to 29.9km west of the Highway 626 and Highway 17 junction Description: This exposure displays natrolite-bearing, pegmatitic syenite (see Photo 1). Reddish orange natrolite (an acicular or prismatic zeolite mineral replacing nepheline) patches up to 15cm in diameter, crystals of perthitic feldspar up to 30cm in length, and crystals or black amphibole up to 25cm in length comprise the bulk of this syenite (see Photo 2). Mitchell and Platt (1994) reported accessory pleochroic clinopyroxene, zircon, titanite, and biotite. Natrolite has locally been ascribed to the hydrothermal alteration of primary nepheline and has also been referred to as “hydronepheline” by local workers. The syenite is intruded by a camptonite lamprophyre dike (Mitchell and Platt 1994) and also hosts large, medium-grained gabbro xenoliths (see Photo 3), up to 1m in thickness and sometimes up to 5m in length (west-side of the highway), that exhibit 1 to 2cm wide, dark reaction rims adjacent to the enclosing syenite. To the east, the pegmatitic syenite gives way to finer grained nepheline syenite in which chalky-weathering nepheline may be observed. Rare natrolite grains are also present. Farther east, a variety of equigranular and pegmatitic syenites are exposed. Near the eastern end of the outcrop (UTM 525625E, 5404825N), a large xenolith of gabbro-hosted, massive titaniferous magnetite has been exposed. Minor clinopyroxene, plagioclase and apatite occur within the massive oxide unit. Analyses completed in 1951 and reported by Hinz and Landry (1994) indicated total iron and titanium values ranging between 33 and 45%, and 4.5 and 13.5%, respectively; phosphorus contents ranged up to 0.371%. Photo 1. Pink, natrolite-bearing, pegmatitic augite syenite. Photo credit D. Campbell. 11 �Photo 2. Pegmatitic syenite containing reddish orange patches of natrolite, light pinkish perthitic feldspar, and black amphibole. Photo credit A. MacTavish. Photo 3. Large gabbro xenolith located on the west side of the highway. Please note that the xenolith has been crosscut by fine-grained syenite veins and that the syenite below the xenolith is varitextured to pegmatitic in texture, whereas the syenite above is medium- to coarse-grained. Photo credit A. MacTavish. 12 �STOP C2: Little Pic River Breccia Zone Location: Lat: 48o47'45"N, Long: 86o37'30"W; UTM 527478E, 5405531N 27.3 km west of the Highway 626 and Highway 17 junction Description: These road cuts, particularly along the south side of the highway, expose spectacular intrusive breccias within the youngest rocks of the complex, along the east side of the fault zone that the Little Pic River occupies. The breccias often occur as semi-continuous, fragmented, elongate rafts (western end of southern rock cut, see Photo 4) that consist of angular to rounded blocks of fine- to medium-grained, equigranular, mafic (gabbroic?) rocks within a groundmass of pink, medium-grained, quartz syenite. In some cases blocks can exhibit both angular and lobate to cuspate (amoeboid) margins (see Photo 5). The mafic rocks comprising the blocks were interpreted as oligoclase-bearing basalt by Mitchell and Platt (1982a). Subsequent discussion and study has led to the suggestion of perhaps 2 texturally discernable types of basic xenoliths, those with: (1) sharp, angular margins, and (2) those with lobate to cuspate margins. In this model, the angular xenoliths represent synplutonic basalts which are now preserved elsewhere as megaxenoliths in younger intrusions. The cuspate-margined xenoliths may represent the effects of mixing between 2 contemporaneous gabbroic/basaltic and syenite magmas (i.e., magma mixing or co-mingling). Cuspate, possible chilled margins with quench-textured clinopyroxene, plagioclase and skeletal olivine have been noted in similar xenoliths to the south on the Coldwell Peninsula by G. Shore (personal communication with M. Smyk, 1995) and suggest the quenching of the basic magma against the cooler, syenitic magma. These are reasonable hypotheses and there are definitely at least 2 types and textures of xenoliths; however, they do not completely explain the presence of blocks exhibiting both margin types as observed by the senior author of this guide and shown in Photo 5. Texturally there also seems to be 3 different types of xenoliths: the most abundant are dark grey to black, very fine-grained xenoliths (Photo 4); medium-grained, greyish pink xenoliths with somewhat less distinct, but still relatively sharp margins (Photo 5), and several unusual zones where there is are subvertical zones of rounded, dark grey, amphibole-phyric xenoliths within a pinkish, mafic groundmass. Are these some sort of breccia dykes or just hybridized zones of xenoliths (see Photo 6; what do you think?)? Although isolated xenoliths are common, there are many areas within these outcrops where incipient or in-situ brecciation characterized by syenite dykes and “jig-saw puzzle/jig-saw fit” breccias are observed, where brecciated fragments can be fitted back together. Miarolitic cavities, up to several centimetres in width, contain euhedral quartz, feldspar, and calcite crystals. The breccia zone persists to the east, towards the scenic lookout located 800m to the east. The south side of the highway is underlain by oligoclase gabbro and quartz syenite, while various, xenolithic-bearing syenitic rocks are exposed on the north side. These pyroxene- and amphibole-(ferro-edenite) bearing syenites contain xenoliths of alkali gabbro, alkali diorite and other, equigranular to porphyritic syenites. Near the lookout turnoff, gray, nepheline-bearing syenite intrudes the mafic rocks and contains orange natrolite. Sannaite and ocellar, camptonitic lamprophyre dikes have been reported near this site by Mitchell and Platt (1994) who proposed the following order of local emplacement: Mg-hornblende syenite  contaminated Fe-edenite syenite  Fe-edenite syenite  quartz syenite (earliest  latest) Lukosius-Sanders (1988) classified the local rocks as miaskitic, metaluminous syenites enriched in U, Th, REE and Zr. These syenites have affinities to A-type granites and have been interpreted to be the result of fractional crystallization of mantle-derived, basaltic magma (Lukosius-Sanders 1988; Mitchell et al. 1993). 13 �Photo 4. Syenite outcrop on the south side of the highway containing blocks of mafic xenoliths often occurring in elongated, semi-continuous rafts. Photo credit A. MacTavish. Photo 5. Elongated, jig-saw-fit xenolithic block exhibiting both angular and lobate/cuspate (amoeboid) margins. Please note the lighter-coloured, pinkish-grey elongated xenolithic block with apparently sharp margins located below the dark grey block. This lower block appears to be coarser-grained and may possibly be different in composition. Photo credit A. MacTavish. 14 �Photo 6. Rounded dark grey, amphibole-phyric xenoliths/inclusions within a fine- to medium-grained, pinkish mafic groundmass (breccia dyke?). Photo credit M. Puumala. Stops C3: Prisoners Cove, Neys Provincial Park (Sample Collecting Prohibited!). The descriptions (unpublished mapping/field descriptions, MacTavish 1992) herein are for 14 sub-stops along the shoreline south of Prisoner’s Cove; however, due to time constraints only the northernmost stops will be visited. Location: Lat: 48o46'30"N, Long: 86o37'10"W; UTM 527984E, 5402537N; 23.5km west of the Hwy 626 and Hwy 17 junction; 2.8km south of Hwy 17 to the park headquarters and then south along the shoreline trail General Description: The wave-washed, glacially polished outcrops along the shoreline of Lake Superior at Prisoner Cove and south for over a kilometre along the western side of the Coldwell Peninsula exhibit a variety of lithologic, textural, and crosscutting features that characterize much of the Center 2 magmatism in the Coldwell Complex. In its simplest sense, this composite stop displays the contact between alkalic biotite gabbro and amphibole-nepheline syenite, but the enigmatic effects of assimilation and hybridization have severely complicated and obscured many of the primary features. In all cases within the nepheline syenitic rocks exposed along the shoreline at this stop the nepheline has been completely altered to the zeolite mineral natrolite which weathers to orange-coloured pits. Medium- to coarse-grained, olivine- and enclave-bearing, biotite gabbro comprises much of the eastern portion of the outcrops. Gabbro xenoliths occur within the syenite and within hybrid phases along their mutual contact, which trends roughly north-south, parallel to the shoreline. The outcrops often exhibit a pitted surface resulting from the preferential weathering of mafic enclaves consisting of biotite-olivine gabbro to biotite-clinopyroxene gabbro or leucogabbro (Walker et al. 1992) within a more syenitic groundmass. The syenitic groundmass consists of fine- to coarse-grained nepheline (altered to natrolite) syenite with minor acicular amphibole and poikilitic biotite. Mitchell and Platt (1994) have identified the amphibole as hastingsite. 15 �Distinct to diffuse layering, a nebulous to locally distinct igneous foliation, and localized soft-sediment style magmatic deformation exists within the amphibole-nepheline syenite. Identifiable, undisturbed layering may be oriented parallel to the apparent syenite/gabbro contact and dips from very steeply to vertically in the north to shallow to moderate to the east (where measurable) in the south. Observed soft-sediment deformation features consist of flame structures, fluid-escape features, slump folds, and isolated well-layered syenite blocks surrounded by obvious fluid escape textures. Much past discussion has focused on whether the observed structures have resulted simply from igneous process, syn- or post-intrusion shearing, or a combination of these processes. The present author’s strongly favour igneous processes since the observed fracturing is very localized, is late and brittle, and does not appear to have affected the foliation or layering within the surrounding rocks in any observable way. It is highly probable that the crystallizing magma chamber was often shocked by MCR tectonic activity. These earthquakes then caused the slumping of unstable crystallizing layers along chamber walls; allowed the isolation of broken, but relatively intact layered blocks; and allowed trapped deuteric fluids formed during the fractionation process to escape upwards through the broken layers. Upon close examination the fracturing presently observed in outcrop obviously took place after the chamber was completely crystallized and was able to deform in a brittle manner. Sub-Stop C3a (527980E, 5402571N): This area, located on the point to the north and west of the old flat-bottomed boats, mainly consists of foliated, wispy, hybridized amphibole-nepheline syenite with diffuse discontinuous “layers” (see Photo 7). The core of this outcrop is flanked to the northeast by a heterogeneous zone containing large numbers of rounded to angular, variably assimilated (metasomatized?) and disaggregated inclusions/xenoliths of biotite gabbro. Reaction rims around these inclusions are readily visible. Also the inclusion-rich zone, as a whole, seems to be enclosed within a diffuse reaction zone when compared to the hybrid syenites adjacent to the west. The western margin of the exposure is a medium- to coarse-grained hybridized syenite with numerous very coarse-grained to pegmatitic inclusions of amphibole-nepheline syenite. At the northwestern tip of the outcrop is an elongate, diffuse zone of apparently non-hybridized, non-foliated syenite (possibly the original parent syenite?). Sub-Stop C3b (527950E, 5402535N): This stop, located 30m west-southwest of the old boats near the shoreline, consists of a 4 to 5m wide, west-northwest-striking, brittle fracture zone hosting a 70 to 100cm thick, dark greenish-grey, ocellar lamprophyre dyke at its northern margin near the water’s edge. The ocellae present within the dyke are composed of reddish, recessive-weathering carbonate (±zeolites?) which are elongated parallel to dyke margins (elongated by flow?). The lamprophyre dyke is also enveloped by a brick-red alteration halo that is not completely within the fracture zone and also extends into the unfractured hybrid syenites to the north for up to 5m. This red halo could be due to either hematization or K-alteration. Similar, subparallel fracture zones can also be observed 10m and 23m to the south. Sub-Stop C3c (527966E, 5402475N): This stop is located ~50m east-southeast of Sub-stop C3b, and consists of a zone of large blocks (?) of coarse-grained, natrolite-bearing, biotite gabbro to biotite melagabbro that are surrounded by fine- to medium-grained amphibole-nepheline syenite containing diffuse gabbro xenolith ghosts. It is distinctly possible that this is not a zone of xenoliths/inclusions at all, but the exposed upper contact of an underlying biotite gabbro body that is part of the biotite gabbro body located about 40m to the southeast (see Sub-Stop C3e, below) where the syenite is observed to overly the gabbro. These blocks (?) are cross-cut by narrow horizontal and subvertical syenite veins and dykes (see Photo 8). 16 �Photo 7. Wispy, relatively mafic in appearance, hybrid amphibole-nepheline syenite exhibiting diffuse discontinuous layers. Photo credit A. MacTavish. Photo 8. Large biotite gabbro xenolith/inclusion (?) crosscut by veins and dykes of amphibole-nepheline syenite. Photo credit A. MacTavish. 17 �Sub-Stop C3d (528004E, 5402440N): This location (45m southeast of Sub-stop C3c) consists of an irregular zone of gabbro xenoliths/inclusions, surrounded by fine- to medium-grained, weakly foliated syenite (see Photos 9 and 10). The xenoliths are in the process of being broken down in stages from originally angular, cohesive blocks to diffuse groupings of amoeboid to wispy mafic remnants within a mafic mineral-rich, hybrid syenite. This process is probably not assimilation in the strictest sense, but is more likely a process of chemical (rather than thermal) invasion through metasomatism that over time breaks down the xenoliths and then eventually disaggregates the mineral constituents of the blocks to the point where they are then assimilated into the syenite melt. The hybridized (?) syenite surrounding the xenoliths exhibit aligned amphibole grains that may indicate flow (?) around and between fragments. There are also places where there are noticeable (up to 15cm thick) halos surrounding zones of xenoliths that consist of aligned amhibole grains that are somewhat separated into diffuse bands. Sub-Stop C3e (528014E, 5402433N): Located only 12m southeast of Sub-Stop 3d and consists of coarse-grained, knobby-weathering, biotite gabbro that has been cross-cut by numerous hair thin to 5cm thick, very fine- to fine-grained syenite stringers and veins and the occasional, larger, fine-grained to pegmatitic syenite vein (pegmatite is in centre of these veins) (see Photo 11). There are numerous leucocratic clots (oikocrysts?) of plagioclase (see Photo 12) throughout. Photo 9. Biotite gabbro xenoliths/inclusions separating from and original larger block and beginning to assimilate (?) into the surrounding syenite melt through a process of metasomatism and disaggregation. Photo credit A. MacTavish. 18 �Photo 10. Disaggregating biotite gabbro xenoliths exhibiting subparallel reaction haloes. Photo credit M. Puumala. Photo 11. Biotite gabbro crosscut by fine-grained to pegmatitic syenite dyke (centre) and thinner syenite veins (centre left). Photo credit A. MacTavish. 19 �Photo 12. Leucocratic clots of plagioclase (oikocrysts) within biotite gabbro. Photo credit A. MacTavish. Sub-Stop C3f (527982E, 5402324N): This sub-stop (~115m west-southwest of Sub-stop C3e) consists of an irregular, variably assimilated zone of mafic (gabbroic?) xenoliths of highly variable size ranges. Many blocks are in the last stages of assimilation where the original xenoliths are now merely ghosts infilled with isolated mafic remnants and considerable numbers of hornblende grains. Sub-Stop C3g (527978E, 5402292N): This location (~30m south of Sub-stop C3f) consists of locally well-developed modal layering within medium- to locally coarse-grained amphibole-nepheline syenite. The magmatic layering dips shallowly to the west and west-southwest at between 20 and 26o and there is a possible weak alignment of K-feldspar laths parallel to layering. The bases of the undulating modal layers are defined by a sharp increase in amphibole content. The best defined layering is near the lake with layering becoming increasingly more diffuse, disrupted, folded (slumping?), and contorted to the east until it becomes unrecognizable. Sub-Stop C3h (527971E, 5402265N): At this location (~27m south of Sub-stop C3g) are 2, sub-parallel, aphanitic to fine-grained, ocellar lamprophyre dykes (see Photo 13) occupying a narrow southeast-striking fracture zone. The dykes dip to the northeast between 54o and subvertical. The ocellae (immiscible liquid droplets) are usually centralized within the dykes away from the strongly chilled dyke margins and are infilled with several minerals including apple-green and greyish minerals (zeolites?), and possibly white calcite. 20 �Photo 13. Ocellar lamprophyre dyke in narrow fracture zone. Photo credit A. MacTavish Sub-Stop C3i (527987E, 5402198N): At this location (~70m south of Sub-stop C3h) is a zone of leopard mottles in moderately mafic, often grain-size-layered (?) amphibole nepheline syenite. Sub-Stop C3j (527971E, 5402265N): This location (~80m south of Sub-stop C3i) is, for lack of a better name, a “Layer Breccia Zone” where there has been strong disruption, localized rotation, and folding of original magmatic syenite layers. Finer-grained syenite containing acicular amphibole grains has flowed around the layer blocks and alignment of those amphibole grains mirrors flow directions. The zone is surrounded by a highly disturbed hybrid mixtite with few measurable features. Thinner blocks are composed of a series of thin modal layers of highly variable textures. The thicker layers are usually the coarsest, are sometimes size-graded, and contain glomerocrysts of K-feldspar (with include amphibole and natrolite after nepheline) up to 1.5cm in diameter surrounded by acicular amphibole grains and recessive-weathering altered nepheline. Sub-Stop C3k (528000E, 5402039N): Located 77m south of Sub-stop C3j. This sub-stop consists of a well-layered block of amphibole-nepheline syenite (~6.5m by 3.5m in dimensions) that is surrounded by a highly distorted zone of fine- to very-coarse-grained (varitextured) syenitic material that appears to have flowed around the block. This isolated block is composed of a sequence of 3 thick layers where amphibole and K-feldspar are aligned subparallel to layer bases. The base of each layer is undulatory on the scale of a single very coarse feldspar crystal. Sub-Stop C3l (528004E, 5402016N): A further 23m south of Sub-stop C3k is a zone characterized by well-developed syenite layering (see Photo 14), some possible magmatic channel scours, some localized soft-sediment-style deformation, and a few zones of intense, localized layer disruption. Most of the layers within the southern part of the zone are quite flat lying (18oW dip). 21 �Photo 14. Well-developed, relatively flat-lying magmatic layering within amphibole-nepheline syenite. Photo credit A. MacTavish Sub-Stop C3m (528011E, 5402006N): This sub-stop is located 12m southeast of Sub-stop C3l and is directly adjacent to it. It consists of a zone of disturbed and distorted syenite layering similar to that observed north of the isolated block observed at Sub-stop C3k. Contorted and convoluted layering is common and folding is observed locally with the disturbance increasing in intensity to the south. Most noticeable in this area is a deformed, crosscutting, texturally variable, vein-like body (see Photo 15) composed of mobile “flow-banded” material. The “vein” margins are often irregular, possibly due to volatile fluid seepage (?) and it is often cored by coarse-grained to pegmatitic veinlets and pods. It is possible that this structure has erupted from the nose of a slump fold. Sub-Stop C3n (528020E, 5401977N): This final sub-stop is located 30m east-southeast of Sub-stop C3m and consists of a large slump-fold composed of medium- to very coarse-grained, modally- and normally grain-size graded syenite layers (see Photo 16). This was interpreted as slump folding due to the presence of at least 3 axial planar directions present within 3 separate folds all in close proximity to each other. Unfortunately since mapping was completed in 1992 this exposure has become partially obscured by the growth of lichen. 22 �Photo 15. Crosscutting, vein-like body possibly resulting from the movement of volatile-rich magmatic fluids. Photo credit A. MacTavish. Photo 16. Lichen-obscured slump fold within layered amphibole-nepheline syenite. Photo credit A. MacTavish. 23 �STOP C4: Hornfelsed Basaltic Roof Pendants, Wolf Camp Lake Location: Lat: 48o47'40"N, Long: 86o26'05"W; UTM 541775E, 5404189N 8.6km west of the Highway 626 and Highway 17 junction Description: Hornfelsed basaltic rocks overlying the complex were recognized early in its mapping by Tuominen (1967) and Puskas (1970) and likely represent a volcanic edifice that has been subsequently eroded (Sage 1986). Mitchell and Platt (1994) and Nicol (1990) have considered these basalts to have a tholeiitic lineage, contemporaneous with the Coldwell Complex. Fresh, metasomatized and hornfelsed, andesine-oligoclase basalt flows are estimated to attain a thickness of 5km (Mitchell and Platt 1994; Nicol 1990). Assimilation and brecciation of the flows by subsequent gabbroic to syenitic magmatism has resulted in the widespread development of basaltic xenoliths ranging from 1m to over 1km in size, comprising a roof pendant in the central part of the complex (Walker et al. 1992). Walker et al. (1992) subdivided these basaltic rocks into 3 main units: 1. Aphanitic to fine-grained, massive, locally amygdaloidal (?) / ocellar basalt; 2. Medium-grained, diabasic (ophitic) basalt; and 3. Aphanitic to medium-grained, feldspar-phyric, diabasic (ophitic) basalt. At Wolf Camp Lake, aphanitic basalts contain round to amoeboid, epidote- and quartz-filled structures up to 2cm in diameter that have been interpreted as amygdules (see Photo 17). Well-defined, amygdule-bearing zones dip 8o to the southwest in this vicinity (Walker et al. 1992). The basaltic roof pendant is locally underlain and enveloped by feldspar-phyric amphibole syenite and Fe-rich augite syenite. Photo 17. Amygdules within the basaltic roof pendant located near Wolf Camp Lake. Photo credit D. Campbell. 24 �STOP C5: Layered Fe-rich Augite Syenite (Alternate stop if time allows) Location: Lat: 48o44'20"N, Long: 86o23'25"W; UTM 544782E, 5398443N 680m west along the shoreline of Lake Superior from the end of the James River industrial road along the waterfront in Marathon; OR 150m south of Carden Cove road, 0.3km past CPR tracks (park at 544864E, 5398750N) Description: Broad expanses of glacially polished and wave-washed, massive Fe-rich augite syenite occur all along this part of the Lake Superior shoreline near Marathon. Fresh surfaces vary from dark green-brown to black, despite a buff to white weathered surface. Small dimension stone quarries were developed and produced in this area during the 1930’s. Much of the stone was shipped to larger centres in the American mid-west and Toronto. Fe-rich augite syenite (formerly referred to as ferroaugite syenite) comprises a large portion of the exposure in the eastern half of the Coldwell Complex. It appears to be a sheet-like intrusion that dips approximately 15o toward the center of the complex, sandwiched between the underlying Eastern Border Gabbro and an overlying, recrystallized amphibole-quartz syenite; it also intrudes the basaltic roof pendants (Walker et al. 1992; 1993a). Crystallization of the syenite inwards from its upper and lower contacts produced mineralogical and compositional variations across it (Walker et al. 1993a). Constituent minerals include iridescent, lathlike, cryptoperthitic feldspar (up to 30% interstitial), and variable amounts of fayalite, amphibole, aenigmatite, and rare quartz. Coarse-grained to pegmatitic portions of the syenite host a variety of REE-bearing fluoro-carbonates, quartz, chalcedony, and molybdenite. Iridescent feldspar, known locally as “spectrolite”, was recently (2010) commercially extracted on a very small-scale from pegmatite at Shack Lake near Marathon. Although this unit is typically massive, rhythmic to chaotic layering is locally developed and where observed commonly dips shallowly towards the centre of the complex. At this site, layering strikes at 070o and dips 60o north. The layering is unusual in that it is defined by an intercumulus mineral (augite) rather that by cumulus phases (feldspar). STOP C6: Layered Eastern (Border) Gabbro Location: Lat: 48o44'00"N, Long: 86o20'00"W; UTM 549199E, 5398010N 1.7 km east of the Highway 626 and Highway 17 junction Description: Layering in the Eastern Border Gabbro shows distinct variations in style, is usually parallel to the eastern contact of the gabbro, and dips 20o to 60o toward the center of the complex (Shaw 1994; 1997). At this stop, layering strikes approximately north and dips west towards the rest of the complex at ~45o. This thickly layered sequence is underlain by massive gabbro near the contact with the Archean country rocks. The macrorhythmic layering is laterally discontinuous, pinching out over distances of 5 to 10m and contacts are sharp and conformable (Shaw 1994; 1997). Rhythmic layering is modal and has been related to variation in the respective proportions of plagioclase (An60-35), augite (Fo67-43), minor orthopyroxene (En55-66), and Fe-Ti-oxides by Lum (1973). Modal plagioclase varies from approximately 60 to 80% in the leucocratic layers and 20 to 35% in the meso- to melanocratic layers (Shaw 1994). A second band of layered gabbro, separated from the first by massive gabbro, is exposed on top of the long rock cut (see Photo 18). Here, the macrorhythmic layering (see Photo 19) produces relatively thin (1 to 5cm) to medium thick (5 to 100cm) layers that can be traced for over 35m along strike. Layer contacts are sharp, locally scalloped and conformable. Trough cross-bedding has been noted on vertical faces by Shaw (1994). This stop is also close to the contact between the Eastern Border Gabbro and the Fe-rich augite syenite to the west. Pegmatitic syenite dykes intrude the gabbro at this locality and contain miarolitic cavities. McLaughlin (1990) has reported the presence of a variety of REE-bearing fluorocarbonates (bastnaesite, parisite, synchisite), Nb-bearing phases, and zircon in pegmatitic syenite with quartz, feldspar, and sodic amphibole. 25 �Photo 18. Macrorhythmic layering within the Eastern Border Gabbro. Photo credit M. Smyk. Photo 19. Macrorhythmic modal layering within the Eastern Border Gabbro. Photo credit A. MacTavish. 26 �STOP C7 (Alternate Stop): Eastern Contact of the Coldwell Complex Location: Lat: 48o43'10"N, Long: 86o19'30"W; UTM 549656E, 5396238N 3.3 to 3.6km east of the Highway 626 and Highway 17 junction Description: A number of highway rock cuts and outcrops expose the eastern contact of the Coldwell Complex with the enclosing Archean greenstone belt country rocks. Center 1 gabbros, the oldest rocks of the complex, form a ring dyke that forms the eastern and northern margins of the complex where it is in contact with Archean supracrustal and granitoid rocks. The reverse magnetization of these gabbros (Lilley 1964) produces prominent magnetic “lows” on aeromagnetic maps. The most recent and comprehensive study of the Eastern Border Gabbro was conducted by Shaw (1994, 1997) who noted that more than 90% of the unit consists of layered gabbro. At this location, varitextured, unlayered Eastern Border Gabbro is in contact with, and contains numerous xenoliths of, Archean metasedimentary rocks. This has produced hybrid and contaminated phases and rheomorphic breccia. Crosscutting Center 1 syenite dikes are commonly pegmatitic. Amethystine quartz, calcite, and molybdenite occur in vugs within this chaotic contact zone. Disseminated iron- and copper-sulphides occur in biotite-rich, varitextured gabbro (Dunlop Occurrence), which has experienced sporadic exploration since the discovery of copper in the early 1950’s. It was last drilled in 1992 by Noranda Inc. with the best assay intervals grading 0.35% Cu/6.0m and 0.42% Cu/4.0m, respectively (Resident Geologist’s Files, Thunder Bay). A grab sample of rusty-weathering, moderately magnetic, fine- to medium-grained gabbro with coarse biotite and blebby chalcopyrite graded 5090ppm Cu, 494ppm Ni, 241ppm Zn, 8ppb Pd, 2ppb Pt and 22ppb Au (ibid). Overgrown pits are located just inside the tree line, west of the highway (UTM 549575E, 5396290N). Shaw (1994; 1997), Walker et al. (1993a, 1993b, 1993c), Currie (1980), and Tucker (1995) have documented a number of occurrences of rheomorphic breccia associated with the Eastern Border Gabbro along its intrusive, basal contact with the Archean supracrustal country rocks. Breccia units are characterized by chaotic flow fabrics that surround flow-oriented clasts situated in a medium-grained, granitic matrix. This unit has been somewhat enigmatic, having been alternatively described by earlier workers as conglomerate and ignimbrite (Resident Geologist’s Files, Thunder Bay). Similar exposures of this map unit also occur along the western contact of the complex, north of Middleton (cf. Wilkinson 1983). Locally, pods of breccia vary from 20 to 75m in width and are up to 250m long. The breccia exposed along Highway 17 at this site contains mainly hornfelsed Archean clastic metasedimentary and metavolcanic rocks and massive vein quartz. In the vicinity of Two Duck Lake, the breccia contains fine-grained gabbro clasts (Tucker 1995). The breccia varies from clast- to matrix-supported; the matrix consists of equigranular quartz, feldspar, and minor biotite, clino- and orthopyroxene, and opaque minerals; and tourmaline and prehnite overgrowths have been noted (Tucker 1995). Rounded to angular clasts range in size from 0.5 to over 100cm and locally have developed 1 to 2cm wide, chlorite-rich reaction rims that are thickest where they are matrix-supported (Shaw 1994). Magnetite and quartzfeldspar-tourmaline veins cut both matrix and clasts. Quartzo-feldspathic rinds and crosscutting veinlets have been interpreted to be the result of partial melting of the felsic material during assimilation. The close association between rheomorphic breccia and the Eastern Border Gabbro suggests that the intrusion of the gabbro led to the brecciation and partial melting of the country rocks (Shaw 1994, 1997; Tucker 1995). 27 �COLDWELL ALKALINE COMPLEX REFERENCES Alexander, M. 2007. The mineralogy of NYF pegmatites from the Coldwell Alkaline Complex, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario. Barrie, C. Tucker, MacTavish, A.D., Walford, P.C., Chataway, R., and Middaugh, R., 2002. Contact-type and magnetite reef-type Pd-Cu mineralization in ferroan olivine gabbros of the Coldwell Complex, Ontario; in The Geology, Geochemistry, Mineralogy and Mineral Beneficiation of Platinum-Group Elements. Edited by L.J. Cabri; Canadian Institute of Mining and Metallurgy and Petroleum, Special Volume 54, p.321-337. 87 86 Bell, K. and Blenkinsop, J. 1980. Grant 42: Ages and initial Sr- Sr ratios from alkaline complexes of Ontario; in Geoscience Research Grant Program, Summary of Research, 1974-1980, Ontario Geological Survey, Miscellaneous Paper 93, p.16-23. Bohay, T.J. 1997. The Coldwell alkaline complex, Ontario: Magmatic affinity as determined by an isotopic and geochemical study; unpublished MSc thesis, McMaster University, Hamilton, Ontario, 135p. Coates, M.E. 1970. Geology of the Killala–Vein lakes area, Ontario; Ontario Department of Mines, Geological Report 81, 35p. Good, D.J. 1992. Genesis of copper-precious metal sulfide deposits in the Port Coldwell alkalic complex, Ontario; unpublished PhD thesis, McMaster University, Hamilton, Ontario, 203p. Good, D.J. and Crockett, J.H. 1994a. Genesis of the Marathon Cu-platinum-group element deposit, Port Coldwell alkaline complex, Ontario: A Midcontinent Rift-related magmatic sulfide deposit; Economic Geology, v.89, p.131-149. ———. 1994b. Origin of albite pods in the Geordie Lake gabbro, Port Coldwell alkaline complex, northwestern Ontario: Evidence for late-stage hydrothermal Cu-Pd mineralization; The Canadian Mineralogist, v.32, p.681701. Hauck, S.A., Severson, M.J, Zanko, L., Barnes, S.-J., Morton, P., Alminas, H., Foord, E.E. and Dahlberg, E.H. 1997. An overview of the geology and oxide, sulfide and platinum-group element mineralization along the western and northern contacts of the Duluth Complex; Geological Society of America, Special Paper 312, p.137-185. Heaman, L.M., Easton, R.M., Hart, T.R., Hollings, P., MacDonald, C.A. and Smyk, M. 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon Region, Ontario; Canadian Journal of Earth Sciences, v.44, no.8, p.1055-1086. Heaman, L.M. and Machado, N. 1987. Isotope geochemistry of the Coldwell alkaline complex: 1. U-Pb studies on accessory minerals; Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Saskatoon, Saskatchewan, Program with abstracts, p.54. ——— 1992. Timing and origin of the Midcontinent Rift alkaline magmatism, North America: Evidence from the Coldwell complex; Contributions to Mineralogy and Petrology, v.110, p.289-303. Hinz, P. and Landry, R. 1994. Industrial mineral occurrences and deposits in northwestern Ontario; Ontario Geological Survey, Open File Report 5889, 145p. Kerr, H.L. 1910a. Geological map of part of the north shore of Lake Superior, District of Thunder Bay; Ontario Bureau of Mines, Annual Report Map 19B, scale 1:63 360. Kerr, H.L. 1910b. Nepheline syenites of Port Coldwell; Ontario Bureau of Mines, Annual Report, v.19, p.194-232. Kissin, S.A. and McCuaig, T.C. 1988. The genesis of silver vein deposits in the Thunder Bay area, northwestern Ontario: Geoscience Research Grant Program, Summary of Research, 1987-1988; Ontario Geological Survey, Miscellaneous Paper 140, p.146-156. Klasner, J.S., Cannon, W.F. and Van Schmus, E.R. 1982. The Pre-Keweenawan tectonic history of the southern Canadian Shield and its influence on the formation of the Midcontinent Rift; in Geology and Tectonics of the Lake Superior Basin, Geological Society of America, Memoir 156, p.27-46. Lewchuk, M.T. and Symons, D.T.A. 1990. Paleomagnetism of the late Precambrian Coldwell complex, Ontario, Canada; Tectonophysics, v.184, p.73-86. Laderoute, D.G. 1987. The petrology, geochemistry, and petrogenesis of alkaline dyke rocks from the Coldwell Alkaline complex; unpublished M.Sc. Thesis, Lakehead University, Thunder Bay, Ontario, 89p. Tectonophysics, v.184, p.73-86. 28 �Lilley, F.E.M. 1964. An analysis of the magnetic features of the Port Coldwell intrusion; unpublished BSc thesis, University of Western Ontario, London, Ontario, 89p. Lukosius-Sanders, J. 1988. Petrology of the syenites from Center III of the Coldwell alkaline complex, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 141p. Lum, H.K. 1973. Petrology of the eastern gabbro and associated sulphide mineralization of the Coldwell alkaline complex, Ontario; unpublished BSc thesis, Carleton University, Ottawa, Ontario, 68p. MacTavish, A. 2000. A new style of PGE mineralization within the Coldwell alkaline complex, northwestern Ontario; Ontario Exploration and Geoscience Symposium, Toronto, December 11-12, 2000, Speaker Abstracts, p.3. MacTavish, A., Lukosius-Sanders, J. and Jowett, R. 1987. Geological report of the Joa Option (Geordie Lake property), St. Joe Canada Inc.; unpublished report, Resident Geologist’s Files, Thunder Bay, 7p. McLaughlin, R.M. 1990. Accessory rare metal mineralization in the Coldwell alkaline complex, northwest Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 123p. Miller, J.D., Jr., Nicholson, S.W., and Cannon, W.F. 1995. The Midcontinent rift in the Lake Superior region, in Miller, J.D., Jr., ed., Field trip guidebook for the geology of ore deposits of the Midcontinent rift in the Lake Superior region; Minnesota Geological Survey Guidebook Series, no. 20, p.1-22. Mitchell, R.H. and Platt, G. R. 1978. Mafic mineralogy of ferroaugite syenite from the Coldwell alkaline complex, Ontario, Canada; Journal of Petrology, v.19, p.627-651. ——— 1982a. The Coldwell alkaline complex; in Field Trip Guidebook, Proterozoic geology of the northern Lake Superior area, Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Winnipeg, Manitoba, p.42-61. ——— 1982b. Mineralogy and petrology of nepheline syenites from the Coldwell alkaline complex, Ontario, Canada; Journal of Petrology, v.23, p.186-214. ——— 1994. Aspects of the geology of the Coldwell alkaline complex: Field trip A2, Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Waterloo, Ontario, 36p. Mitchell, R.H., Platt, G.R., Lukosius-Sanders, J., Artist-Downey, M. and Moogk-Pickard, S. 1993. Petrology of syenites from Center III of the Coldwell alkaline complex, northwestern Ontario, Canada; Canadian Journal of Earth Sciences, v.30, p.145-158. Mitchell, R.H., Platt, R.G. and Cheadle, S.P. 1983. A gravity study of the Coldwell complex, northwestern Ontario, and its petrological significance; Canadian Journal of Earth Sciences, v.20, p.1631-1638. Mulja, T. 1989. Petrology, geochemistry, sulphide- and platinum-group element mineralization of the Geordie Lake intrusion; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 234p. Mulja, T. and Mitchell, R.H. 1990. Platinum-group minerals and tellurides from the Geordie Lake intrusion, Coldwell complex, northwestern Ontario; Canadian Mineralogist, v.28, p.489-501. ——— 1991. The Geordie Lake intrusion, Coldwell Complex, Ontario: Palladium- and tellurium-rich disseminated sulfide occurrence derived from an evolved tholeiitic magma; Economic Geology, v.86, p.1050-1069. Nicol, D.N. 1990. Assimilation of basic xenoliths with Center 3 syenites of the Coldwell Complex, Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 59p. Ohnenstetter, D., Watkinson, D.H. and Dahl, R. 1991. Zoned hollingworthite from the Two Duck Lake intrusion, Coldwell complex, Ontario; American Mineralogist, v.76, p.1694-1700. Penczak, R.S. 1992. Petrology and mineral chemistry of the Middleton copper occurrence of the Western gabbro, Coldwell alkaline complex, Ontario; unpublished BSc thesis, University of Waterloo, Ontario. Pollock, S.J. 1987. The isotopic geochemistry of the Prairie Lake carbonatite complex; unpublished MSc thesis, Carleton University, Ottawa, Ontario, 71p. Potter, E.G. 2004. The rare and exotic mineralogy of the western subcomplex of the Deadhorse Creek diatreme, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario. 29 �Puskas, F.W. 1967. Port Coldwell area; Ontario Department of Mines, Preliminary Map P.114, scale 1:31 680. th ——— 1970. The Port Coldwell alkali complex; in Proceedings, 16 Institute on Lake Superior Geology, Thunder Bay, Ontario, p.87-100. 40 39 Queen, M., Heaman, L.M., Hanes, J.A., Archibald, D.A. and Farrar, E. 1996. Ar/ Ar phlogopite and U-Pb perovskite dating of lamprophyre dykes from the eastern Lake Superior region: Evidence for a 1.14 Ga magmatic precursor to Midcontinent Rift volcanism; Canadian Journal of Earth Sciences, v.33, p.958-965. Sage, R.P. 1982. Mineralization in diatreme structures north of Lake Superior; Ontario Geological Survey, Study 27, 79p. ——— 1985. Geology of carbonatite-alkaline rock complexes in Ontario: Chipman Lake area; Ontario Geological Survey, Study 44, 40p. ——— 1986. Alkalic rock complexes – carbonatites of northern Ontario and their economic potential; unpublished PhD thesis, Carleton University, Ottawa, Ontario, 335p. ——— 1987. Geology of carbonatite-alkaline rock complexes in Ontario: Prairie Lake carbonatite complex, District of Thunder Bay; Ontario Geological Survey, Study 46, 91p. ——— 1991. Alkaline rock, carbonatite and kimberlite complexes of Ontario, Superior Province; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p. 683-709. Sage, R.P. and Watkinson, D.H. 1995. Alkalic rocks of the Midcontinent rift; Institute on Lake Superior Geology, Marathon, ON, Proceedings Volume 41:2A, 79p. Shaw, C.S.J. 1994. Petrogenesis of the eastern gabbro, Coldwell alkaline complex, Ontario; unpublished PhD thesis, University of Western Ontario, London, Ontario, 292p. ——— 1997. The petrology of the layered gabbro intrusion, eastern gabbro, Coldwell alkaline complex, northwestern Ontario, Canada: Evidence for multiple phases of intrusion in a ring dyke; Lithos, v.40. p.243-259. Smyk, M.C., Taylor, R.P., Jones, P.C. and Kingston, D.M. 1993. Geology and geochemistry of the West Dead Horse Creek rare-metal occurrence, northwestern Ontario; Exploration and Mining Geology, v.2, no.3, p.245-251. Smyk, M.C. and Sage, R.P. 1995. Geology and mineralization of intrusive complexes of the Marathon, Ontario area; in Field Trip Guidebook for the Geology and Ore Deposits of the Midcontinent Rift in the Lake Superior region, International Geological Correlation Program, Project 336, Field Conference and Symposium, Duluth, Minnesota, August 19 to September 1, 1995, p.182-193. Tucker, C. 1995. Origin of breccia associated with the Eastern Gabbro, Coldwell alkaline complex, northwestern Ontario; unpublished BSc thesis, University of Western Ontario, London, 57p. Tuominen, H.V. 1967. Port Coldwell area; Ontario Department of Mines, Map P.114, scale 1:15 840. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T. and Penczak, R.S. 1992. Geology of the Port Coldwell alkaline complex; in Summary of Field Work, 1992, Ontario Geological Survey, Miscellaneous Paper 160, p.108-119. ——— 1993a. Precambrian geology of the Coldwell Alkaline Complex; Ontario Geological Survey, Open File Report 5868, 30p. ——— 1993b. Precambrian geology, Port Coldwell complex, west half; Ontario Geological Survey, Preliminary Map P.3232, scale 1:20 000. ——— 1993c. Precambrian geology, Port Coldwell complex, east half; Ontario Geological Survey, Preliminary Map P.3233, scale 1:20 000. Watkinson, D.H., Whittaker, P.J. and Jones, P.L. 1983. Platinum group elements in the eastern gabbro, Coldwell complex, northwestern Ontario; Ontario Geological Survey, Miscellaneous Paper 113, p.183-191. Weiblen, P.W. 1982. Keweenawan intrusive rocks; Geological Society of America Memoir 156, p.57-82. Wilkinson, S.J. 1983. Geology and sulphide mineralization of the marginal phases of the Coldwell complex, northwestern Ontario; unpublished MSc thesis, Carleton University, Ottawa, Ontario, 129p. 30 �PART 2: MARATHON CU-PGM DEPOSIT 3 David Good1 and John McBride2 Earth Sciences Dept., Western University, London, ON 2 Stillwater Canada Inc., Marathon, ON Introduction to the Marathon Deposit: The Cu-PGM sulphide mineralization of the Marathon deposit is hosted by the Two Duck Lake Gabbro, the latest mafic intrusive event and consequently the most continuous gabbroic body within the Eastern Gabbro Suite at the Marathon deposit. The Eastern Gabbro Suite, located around the eastern and northern margin of the Coldwell, was composed initially of a thick sequence of tholeiitic basalt that was subsequently intruded by a much larger volume of leucocratic to ultramafic intrusions that caused contact metamorphism of the basalt to pyroxene-hornfels grade (Good et al., 2015). All of these units are represented at the Marathon deposit (Figure 1). Figure 1: Geology of the Marathon deposit (after Good et al. 2015) highlighting location of field trip stops. Stops are marked with red dot and labelled as stop 1a, etc. Note two normal faults that correspond to strong surface lineaments (dashed lines) 31 �The topography of the Coldwell is characterized by deep valleys and steep cliffs that form strong surface lineaments. Two lineaments at the Marathon deposit correspond to north dipping normal faults (north side down) with displacement of approximately 50 metres. Two Duck Lake Intrusion: The Two Duck Lake intrusion is irregular in shape and elongated north-south (Figure 2). The dip at the east contact is variable from nearly flat (at the south end) to vertical and locally over turned where the footwall overhangs the intrusion. The intrusion is composed of coarse-grained to pegmatitic olivine gabbro and troctolite. Modal layering is rare. The TDL gabbro was interpreted to have formed by intrusion of a nearly homogeneous plagioclase crystal mush by Good and Crocket (1994). But recent work suggests the intrusion formed by accumulation of several pulses of magma in a conduit setting (Good 2010; Ruthart, 2012; Good et al., 2015; and Shahabi Far, 2017). Multiple feeder channels were inferred by Good et al. (2015) to occur in the vicinity of several coincident features, including: deep V- or U-shaped channels in the footwall contact; topographic lineaments; very thick mineralized intervals; and irregular-shaped intrusions of olivine-magnetite-clinopyroxene-apatite. Figure 2: 3d isometric view of the Two Duck Lake intrusion (from Good et al. 2015). Three coloured portions indicate blocks that were offset by normal faults with north side down by up to 60 metres. Note that numerous intrusions of mineralized Mt-Ol-Cpx-Ap rock (yellow) occur in the vicinity of the 6300 feeder but are not shown. Age relationships Evidence suggests that all units in the Coldwell were emplaced within a short time interval between about 1108 Ma and 1105 Ma (Heaman and Machado 1992; Good et al. in preparation). Age relationships, based on cross-cutting contacts and U-Pb age dating for the various units are summarized in Figure 3. The metabasalt is interpreted to correlate with Mamainse Point Volcanic Group 1 (Figure 4) which was emplaced at approximately 1108 Ma (Keays and Lightfoot, 2015). 32 �The metabasalt was subsequently intruded by the following units, listed in order from oldest to youngest, layered troctolite sill of the Marathon Series, gabbroic anorthosite and olivine gabbro of the Layered Series, Two Duck Lake gabbro and various ultramafic units composed of magnetite +/- olivine +/- apatite +/clinopyroxene of the Marathon Series, Malpa Lake intrusion, and syenite. Figure 3: Relative timing of mafic metavolcanic and intrusive events (age dates after Good et. al, in preparation) in the Eastern Gabbro Suite of the Coldwell Alkaline Complex. Figure 4: Correlation diagram showing range of ages for the Coldwell units compared to volcanic and intrusive units in the Midcontinent Rift (after Keays and Lightfoot, 2015). 33 �Mineralization: Disseminated sulfide mineralization is hosted by the Two Duck Lake gabbro and associated breccia (Figure 5) and occurs within several thick and continuous shallow-dipping lenses that parallel the footwall contact. The lenses are referred to as the Footwall, Main, and Hangingwall zones and the W Horizon. Sulfides in the Footwall, Main, and Hanging-wall zones consist predominantly of chalcopyrite and pyrrhotite with minor amounts of cubanite, bornite, pentlandite, cobaltite, and pyrite. Sulfides occur interstitial to primary silicates and also in association with hydrous silicates such as amphibole, chlorite, and minor serpentine (Watkinson and Ohnenstetter, 1992; Samson et al., 2008). Chalcopyrite occurs as separate grains or as rims on pyrrhotite grains. Some chalcopyrite is intergrown with highly calcic plagioclase (An70–An80) in replacement zones at the margins of plagioclase crystals (Good and Crocket, 1994; Shahabifar, 2016). The W horizon is characterised by extreme PGE enrichment relative to Cu with several 2m thick drill hole intersections having 20 to 70 ppm Pd and Cu/Pd as low as 3. The best intersection contains 34 ppm Pd and 9.6 ppm Pt over 10 m. Mass balance considerations, assuming initial magma contained 10 ppb Pd, would require a magma column on the order of 34 km to generate the 34 ppm Pd in this interval. The W Horizon is commonly difficult to identify in drill core because it typically contains only trace sulfides, but if sulfides are present, they consist of chalcopyrite and bornite with minor pyrrhotite and trace amounts of pentlandite, cobaltite, and pyrite (Ruthart, 2012). Figure 5: Stratigraphic section through the Main zone and overlying troctolite sill. Note the saw tooth pattern for Cu, Pd and Cu/Pd indicating individual pulses of sulphide-bearing crystal slurry. Unit 2d, breccia of metabasalt blocks and Two Duck Lake gabbro; unit 3bd, coarse grained ophitic and pegmatitic Two Duck Lake gabbro; unit 4a, breccia of footwall blocks and Two Duck Lake gabbro (from Good et al., 2015). 34 �Figure 6: Three versions of top view for the Marathon deposit showing 3d topography (green surface) and contoured footwall surface models. Note the troughs and ridges (left hand image) correspond to surface lineaments. Note the higher grade assays for Cu (>0.5%) and Pd (>3 ppm) are aligned within zones that parallel troughs within the 3d footwall surface model. Deposit Model: Figure 7: Step 3 of schematic illustration for magmatic plumbing system (after Barnes et al, 2016) Evidence for magma conduit setting at Marathon include: • • • • • • • • 35 association with volcanic rocks fault control brecciation and assimilation accumulation in trough setting flow through PGE upgrading tube shaped intrusion gravity driven back flow high heat flux �Marathon Cu-PGM Deposit - Field Trip Stops Figure 8: Map at south end of Marathon deposit showing location of stops 1 and 2. 36 �Stop 1a: South end of Troctolite Sill (Figures 8 and 9b): Trench exposure with coarse grained mottled augite troctolite shows large fresh oikocrysts of olivine (brown), clinopyroxene (black) and magnetite (black) and subhedral plagioclase (white). The layered troctolite sill is an important marker horizon because it occurs just above the top of the Main Mineralized zone and is an indicator of the relative fault offset that occurred along E-W–trending normal faults at 5404500 and 5404900 North (Figure 1). The sill dips moderately west at the north end, but flattens out in the south to sub-horizontal (Figure 10b). Note layering is approximately east west at stop 1a, consistent with rotation of sill from west dipping in the north to sub horizontal or north dipping in the south. Figure 9: 3d image (iso view) of geology at south end of Marathon deposit showing location of stops 1 and 2 on trenched outcrops (black polygons): (a) shows orientation of footwall surface troughs approximately perpendicular to contact, and the red centre line of the W horizon at surface on the splat trench; (b) top (light blue) and bottom surfaces (dark blue) of the troctolite sill. Gap in the troctolite sill surfaces represents location where W Horizon and TDL gabbro cuts the through the sill; (c) surface model of W Horizon (yellow). 37 �Stop 1b: South end of Marathon deposit (Figures 9, 11 and 12): Trench outcrop exposure of Two Duck Lake gabbro shows shallow dipping W Horizon and Main zone type mineralization. Figure 10: Plan map of trench at the south end of the Marathon deposit. Red circle marks location of historic trench (ca. mid 1960’s) with high copper mineralization. The unit was not assayed for Pd until 2005. The channel sample located just north of the red circle returned assays of 3.37 ppm Pd+Pt+Au, and 0.35% Cu over 18.6 m East-west layering in TDL gabbro is visible just south of trench. Note textural evidence for cross cutting intrusions. East-west channel samples 44 to 56 (just north of red dot at stop 1b). Interval average: 3.37 ppm Pd+Pt+Au, and 0.35% Cu over 18.6 m. 38 �Stop 2a: Splat trench and the W Horizon Figure 11: Figure 12: PGE rich samples from the W Horizon contain fresh clinopyroxene, olivine and plagioclase. Top photo sample with 107 ppm Pd+Pt+Au and 203 ppm Cu. Bottom photo sample with 70 ppm Pd+Pt+Au and 0.86 % Cu. 39 �Stop 2b: Splat trench and the W Horizon Figure 12: Stop 2b: Splat Trench - NW branch 40 �Stop 3: Ultramafic magnetite-olivine-clinopyroxene-apatite intrusions Numerous pod shaped ultramafic apatite intrusions occur within the metavolcanic pile to the west of and stratigraphically above the Two Duck Lake intrusion. Similar intrusions occur within the TDL intrusion. These units were previously referred to as reef type accumulations, but trenched exposures show they are discontinuous intrusive bodies that cut the metavolcanic and Layered Gabbro Series. The Four Dams and Willie Lake Cu-Pd occurrences are larger examples of this type of mineralization. At the Marathon deposit, the ultramafic units occur as discontinuous pods that range in size from about 20 to 100 m along strike and less than about 10-20 m in thickness. The best intersections in surface outcrops include 9.3m at 1.6 g/t PGM+Au and 0.13% Cu and 4.94 m at 2.15 g/t PGM+Au and 0.29% Cu. The ultramafic pods are concentrated in the vicinity above trough structures in the footwall and are potential indicators of underlying magma conduits. These intrusions are proposed to have formed by backflow of dense minerals in a conduit setting. 41 �Stop 4: Main Zone at Two Duck Lake 42 �MARATHON Cu-PGE DEPOSIT REFERENCES Barnes SJ, Cruden AR, Arndt N, Saumur BM, 2016. The mineral system approach applied to magmatic Ni–Cu–PGE sulphide deposits, Ore Geology Reviews 76, 296-316. Good DJ, 1993. Genesis of copper-precious metal sulfide deposits in the Port Coldwell Alkalic Complex, Ontario Geoscience Research Grant Program, Grant No. 341, Ontario Geological Survey, Open File Report 5839, 23. Good DJ and Crocket JH, 1994. Genesis of the Marathon Cu-Platinum-Group Element Deposit, Port Coldwell Alkalic Complex, Ontario: A Midcontinent Rift-Related Magmatic Sulfide Deposit, Economic Geology, v. 89, p. 131-149. Good, DJ, Epstein, R, McLean, K, Linnen, RL & Samson, IM, 2015. Evolution of the Main Zone at the Marathon Cu-PGE Sulfide Deposit, Midcontinent Rift, Canada: Spatial Relationships in a Magma Conduit Setting, Econ Geol v.110, p.983-1008. Keays RR, and Lightfoot PC, 2015. Geochemical Stratigraphy of the Keweenawan Midcontinent Rift Volcanic Rocks with Regional Implications for the Genesis of Associated Ni, Cu, Co, and Platinum Group Element Sulfide Mineralization, Econ Geol, 110, 1235-1267. Ruthart R, 2012. Characterization of high-PGE, low-sulphur mineralization at the Marathon PGE-Cu deposit, Ontario: M.Sc. thesis, Waterloo, ON, University of Waterloo, 145 p. Samson, IM, Fryer, BJ, and Gagnon, JE, 2008. The Marathon Cu-PGE deposit, Ontario: Insights from sulphide chemistry and textures, in Goldschmidt conference, p. 820. Shahabi Far, M, 2016. The magmatic and volatile evolution of gabbros hosting the Marathon PGE-Cu deposit: evolution of a conduit system, PhD thesis, University of Windsor, Ontario. Walker EC, Sutcliff RH, Shaw CSJ, Shore GT and Penczak RS, 1993. Precambrian geology of the Coldwell Alkaline Complex, Ontario Geological Survey, Open File Report, v. 5868, 30p. Watkinson DH and Ohnenstetter D, 1992. Hydrothermal origin of platinum-group mineralization in the Two Duck Lake intrusion, Coldwell Complex, Northwestern Ontario: Canadian Mineralogist, v. 30, p. 121-136. 43 �Field Trip 1 Archean and Proterozoic geology of the Marathon-Hemlo area Day 2 Geology of the eastern Schreiber-Hemlo Greenstone Belt in the vicinity of Heron Bay and Hemlo Mark Puumala and Mark Smyk Resident Geologist Program, Ontario Geological Survey, Thunder Bay, Ontario, Canada and Tom Muir Ontario Geological Survey (retired) INTRODUCTION Day 2 of the Marathon-Hemlo field trip will begin in the Heron Bay area, just to the east of the contact between the Mesoproterozoic Coldwell alkaline complex and the eastern half of the Neoarchean Schreiber-Hemlo greenstone belt. The first three stops will highlight the lowermetamorphic grade (predominantly greenschist facies) supracrustal rocks and variable deformation that are characteristic of this portion of the belt. The trip will then proceed eastward to the world-renowned Hemlo gold camp (production to date of close to 22 Moz Au) where the bulk of the trip will be spent examining the strongly deformed and higher-metamorphic grade (amphibolite facies) supracrustal rocks that are exposed along Highway 17 adjacent to Barrick Gold Corporation’s Williams Mine property. These roadside exposures provide an overview of the bedrock types, geological structures, mineralization and alteration that are present in the vicinity of the Hemlo gold deposit, and include outcrops of the up-dip projections of two goldmineralized zones. REGIONAL GEOLOGY The Schreiber-Hemlo greenstone belt is located in the Neoarchean Wawa subprovince. The belt extends from Schreiber in the west, almost to White River in the east, and is bisected by the Mesoproterozoic Coldwell alkaline complex between the Little Pic River and Marathon. Consequently, the two halves of the greenstone belt are sometimes referred to separately as the Schreiber greenstone belt and the Hemlo greenstone belt (see Figure 1). 44 �Figure 1. Regional geological setting of the Schreiber-Hemlo greenstone belt (from Jackson et al. 1998). The western half of the Schreiber-Hemlo greenstone belt consists of a mixture of tholeiitic mafic metavolcanic rocks, calc-alkalic mafic to felsic volcanic rocks and sedimentary rocks whose stratigraphic and structural relationships have not previously been resolved (Williams et al. 1991). An Ontario Geological Survey (OGS) mapping project that is currently being undertaken in the western portion of the belt (Magnus and Walker 2015; Magnus and Arnold 2016) is designed to address this knowledge gap. The eastern half of the Schreiber-Hemlo greenstone belt (i.e. east of the Coldwell complex) is much better understood (Muir 2000) and consists of three main units. The lowermost unit is dominated by tholeiitic metavolcanic rocks that are intercalated with lesser amounts of mafic to ultramafic intrusions and flows, and is older than 2697 Ma. This is overlain by calc-alkalic metavolcanic rocks (flows and pyroclastic rocks) and related intrusions that are between 2695 and 2688 Ma. This metavolcanic rock sequence is intercalated with metasedimentary rocks and is eventually superseded by an overlying metasedimentary rock dominated sequence (Muir et al. 1999). The greenstone belt is bounded to the north and south by earlier (ca. 2720 Ma), gneissic granitoid rocks of the Black-Pic and Pukaskwa batholiths (see Figure 2), and has been intruded internally and along its margins by younger granitoid plutons ranging in age from 2697 to 2677 Ma. These include the ca. 2697 Ma Dotted Lake pluton, the ca. 2690 to 2684 Ma Cedar Lake pluton, Cedar Lake stock and Heron Bay pluton and the ca. 2679 to 2677 Ma Gowan Lake, Musher Lake and Bremner plutons (Beakhouse 2001). 45 �Figure 2. Generalized geology of the eastern half of the Schreiber-Hemlo greenstone belt (from Muir and Smyk 2006). The eastern Schreiber-Hemlo greenstone belt has the morphology of an open synclinorium with complex internal structural patterns. It is interpreted to have undergone 2 main stages of structural deformation (Muir et al. 1999; Jackson 1998). Each deformation stage occurred over an extended period of time and may have consisted of multiple events (i.e., as documented by Muir (1997, 2003) in the immediate vicinity of the Hemlo gold deposit). The first phase of deformation (D1R) affected rocks younger than 2693 Ma and resulted in the development of penetrative foliation that is generally subparallel to stratigraphy and intrusion boundaries. The second regional deformation event (D2R) affected rocks as young as 2675 Ma and is considered to have been responsible for the development of map-scale folds with upright axial planar fabrics (Jackson 1998). The eastern Schreiber-Hemlo greenstone belt is dominated by westward-plunging structural elements that suggest that the portion of the belt adjacent to the Coldwell complex exposes a shallower structural level of the belt than the eastern end (Muir et al. 1999; Jackson 1998). This 46 �interpretation is consistent with the general eastward increase in metamorphic grade from greenschist facies in the Heron Bay area, to amphibolite facies in the Hemlo area (Thompson 2006). Numerous high strain zones have been documented in the vicinity of the Hemlo gold deposit, with the most prominent of these being the Hemlo fault zone (Muir 1997). This high-strain zone may represent a portion of a regional-scale boundary fault that extends along the entire length of the greenstone belt between Heron Bay and Hemlo (Williams et al. 1991; Muir et al. 1999). GEOLOGY OF THE HEMLO GOLD DEPOSIT The Hemlo gold deposit is located 35 km east of Marathon in the south-central portion of the eastern Schreiber-Hemlo greenstone belt (see Figure 2). The following overview of the deposit geology is excerpted from Muir et al. (1995). The Hemlo deposit is situated within supracrustal rocks in a southern bifurcated segment of the eastern part of the Schreiber-Hemlo greenstone belt. Page (1947a, 1947b; 1948; 1949) was the first geologist who recognized the Hemlo Fault as an important structural feature and identified the close co-planar relationship the fault had with the Lake Superior Shear Zone. He also recognized that all the gold discoveries at that time were hosted by the Lake Superior Shear Zone, and stated that the zone had been traced on surface for over 12 kilometres. Page also noted that the gold mineralization was associated with felsic porphyritic bodies and that the emplacement of these porphyries was related to a major structure which he termed the "Heron Bay-Hemlo Break". The Hemlo deposit largely lies at or near the contact between felsic to intermediate quartz-feldspar-phyric rocks (pyroclastic and subvolcanic(?) varieties) and metasedimentary rocks. Here, the rocks generally strike at 290o to 295o and dip between 60o and 70o to the northeast. The Hemlo deposit is presently interpreted by the authors as being hosted within 290ostriking, highly strained, transposed, and juxtaposed, lithotectonic supracrustal segments, which lie in a generally east-striking greenstone belt. The deposit cannot be demonstrated to be stratiform or stratabound. Sporadically distributed, anomalous gold mineralization has been noted, about several kilometres southeast and east-southeast of the Hemlo deposit on the Lac Minerals Limited, White River property, as being spatially associated with sericitic and pyritic rocks within what is interpreted as a brittle-ductile shear zone (Pan and Fleet 1988, 1989, 1990; Pan 1990). Underground mapping and drilling have demonstrated the existence of parallel mineralized zones within both the metavolcanic and metasedimentary rocks, as well as mineralized zones which transect the metavolcanic-metasedimentary contact. The Hemlo deposit orebodies, collectively, extend for a strike length of about 3.7 km, a depth of 1.35 47 �km, and an approximate down-plunge distance of 2.5 km. The main mineralized zone extends for a strike length of about 2.9 km, and a down-dip distance of 2.5 km (Harris 1989). The thickness of the main mineralized zone ranges from about 2 m in the David Bell Mine (Burk et al. 1986) to 50 m in the Williams Mine (Walford et al. 1986). Several types of ore are delineated in each of the 3 mines, based largely on the predominant mineral(s) and/or textures present. Commonly, because of extensive metasomatism and deformation, the mineralized zones comprise rocks of equivocal protolith(s). Alteration, collectively, is in the form of widely various degrees of microclinization, sericitization, biotitization, silicification, carbonatization, albitization, pyritization, and tourmalinization. Significant amounts of barite of equivocal origin are locally present. Bright green vanadian muscovite (Harris 1989) is commonly present in the altered rocks, as is molybdenite. At least 2 ages of quartz veins can be found within the ore zones: some veins display considerable folding, attenuation, boudinage, and dismemberment, whereas others display minimal deformation. In some cases, outside the ore zone, there are numerous quartz veins which tend to display a lower degree of deformation. Collectively, the ores are enriched in Au, Mo, Sb, Hg, As, Tl, V, and Ba. Gold is commonly disseminated along with molybdenite. Native gold grains are mercury rich and occur along quartz-feldspar and pyrite grain boundaries and fractures, as well as inclusions in, or rimmed with, several varieties of sulphide minerals including, rarely, pyrite and molybdenite (Harris 1989). Visible gold is not common overall, but does occur within quartz veins in feldspathized, molybdenite-bearing rocks, along molybdenite-green-mica-bearing fractures, in stibnite- and cinnabar-bearing quartz pods, and rarely in fractures in some of the plagioclase-porphyritic dikes. Molybdenite is the second most abundant sulphide, after pyrite, and occurs as fine- to very fine-grained, foliation-parallel blades, euhedral crystals, and platy masses mostly in association with silicate minerals, chiefly feldspar and quartz (Harris 1989). Previous field guides for the Hemlo area have largely focussed on the Hemlo “Main Zone,” which continues to contribute to the production of the Williams Mine, and is known as the “B Zone” in current mine terminology. However, much of the current mine production is obtained from a distinct ore body known as the “C Zone.” The following description of the C Zone is excerpted from Langlais and Barber (2015). The C Zone represents multiple sub-parallel lenses of irregular, generally narrow, gold mineralization. C Zone ore is stratigraphically different from the main zone and occurs in two broad geological domains, the porphyritic felsic metavolcanics and the intermediate to felsic volcaniclastic sediment unit. The open pit is located within the C Zone. 48 �The general stratigraphy from south to north is Lower Metasedimentary rocks, Porphyritic Felsic Metavolcanics (Moose Lake Porphyry), Quartz Eye Muscovite Schist, Intermediate to Felsic Volcaniclastic Sediments (fragmental unit) and the Upper Metasedimentary rock sequence. Lower and upper denote the relative structural positions of the metasedimentary rock units as the younging directions are unclear. All of the major rock units are highly deformed with multiple events of deformation. Structural geology is complex. Rocks in the deposit area exhibit high strain. At the deposit scale, rocks in the area are tightly isoclinally folded. Most of the ore bodies occur on one or more limbs of these folds. Local drag folding can be seen in the ore. Occasional transverse faults offset ore and wall rock units up to a few meters, and there is some shearing along major contacts. Regional metamorphism is up to amphibolite grade. The deposit has also been cut by a number of north-south trending diabase and lamprophyre dikes which post-date mineralization. Although much research has been carried out in the Hemlo area since gold production commenced in 1985, it is interesting to note that considerable debate continues around the application of a genetic model to the Hemlo gold deposit (especially for the Main Zone). Epithermal, intrusion-related (i.e., porphyry) and shear zone-hosted mesothermal models have all been suggested to explain the origins of this enigmatic deposit, with none having achieved widespread acceptance Muir (2002). Irrespective of the model used to explain its origins, it is clear that the Hemlo deposit’s current morphology is largely related to deformation associated with the Lake Superior shear zone. Hemlo Gold deposit Area Exploration and Development History The chronology of events leading up to the discovery and development of the Hemlo mines is condensed from Muir et al. (1995). The subsequent development summary is largely based on Langlais and Barber (2015), with additional information obtained from the Thunder Bay South District Resident Geologist’s files. 1869: Gold was discovered by Moses Pee-Kong-Gay near the present town of Heron Bay. 1920's: J. LeCours sank test pits on a mineralized shear zone near Hemlo station, 6 km southwest of the Hemlo deposit. Assays of up to 4.16 ounces gold per ton were reported. At about the same time, a group of claims was staked on a small quartz vein that returned low gold assay values just to the north of railway mile post 38 (measured from White River). 1930-31: J.E. Thomson mapped the area for the Ontario Department of Mines. 1937: Bowhill Mines shipped a 500 lb. (227 kg) test sample from the Heron Bay area that returned 0.30 ounce gold per ton and 1.53 ounces silver per ton. 49 �1944-46: Prospector Peter Moses discovered a siliceous, mineralized shear zone approximately one-half mile (0.8 km) north of mile post 37 on the railway with samples returning assays up to 0.415 ounce gold per ton. Harry Ollmann and Dr. Jack K. Williams staked 11 claims in the discovery area (Ollmann-Williams property). Gold values were encountered in a large shear zone and stripping, trenching, and diamond drilling were completed. The claims were subsequently patented in 1947. After Ollmann died in December 1947, the claims were placed in Williams’ name in trust by mutual consent. 1946-50: Adjoining claims were staked by consulting geologist Trevor Page. These claims, together with others staked by associates including Moses Fisher (J.E. Thomson's guide), subsequently became part of the Lake Superior Mining Corporation Limited property. Samples collected by Page assayed up to 0.13 ounce gold per ton. Lake Superior Mining Corporation carried out mapping, trenching, chip and channel sampling, and X-Ray diamond drilling on the Lake Superior shear zone. By 1950, a mineralized zone containing 31 543 tons to a depth of 300 feet (91 m) with a calculated width of 8.8 feet (2.7 m) and a cut grade of 0.22 ounce gold per ton was outlined. 1951: The Lake Superior property was optioned by Teck-Hughes Gold Mining Limited. 6 diamond drill holes totalling 2733 feet (833 m) were completed to add to the over 6000 feet of drilling completed to that time by Lake Superior Mining Corporation Limited. The size of zone No. 1 (formerly the 'A' zone) was increased to 76 653 tons at a grade of 0.27 ounce gold per ton. 1957: Teck Exploration Company Ltd. drilled 7 “packsack” drill holes totalling 289 feet (88 m). 1957: Bartley and Page wrote a geological report on the Hemlo area for the Canadian Pacific Railway. 1958: Cusco Mines Ltd. optioned the Lake Superior claims and carried out diamond drilling to test the main mineralized zone. 1960s: The Lake Superior property was staked intermittently by prospectors during the 1960's. 1973: The Lake Superior property was staked by J.E. Halonen for Ardel Explorations Ltd. Three holes totalling 789.9 feet (241 m) were drilled and the deposit tonnage was increased to 150 000 tons grading 0.21 ounce gold per ton above 60 feet (18 m) depth. 1976-77: Claims were staked by R.G. Newman to the west of the Williams property. The claims were investigated by Copper Lake Explorations Ltd. (soil and bedrock geochemical sampling). 50 �1977-78: T.L. Muir of the Ontario Geological Survey mapped the Heron Bay and Hemlo areas, in 1977 and 1978, respectively (Muir 1982a, 1982b). Muir reported an occurrence (0.32 oz/ton Au), presently referred to as the “Highway Zone,” in altered felsic metavolcanic rocks several kilometres west of the main Hemlo deposit. This occurrence is probably in the vicinity of the mile post 38 discovery of the 1920's. Claim stakers and explorationists would later base much of their land acquisition during the staking rush on Muir's maps. 1979: Prospectors Donald McKinnon and John Larche staked the claims surrounding the 11 patented Williams claims. 1980: Corona Resources Ltd. (later International Corona Resources Limited) optioned the McKinnon and Larche claims and completed preliminary linecutting and geophysical surveys. 1981: Corona began a $600 000 diamond drilling program in January. In March, R. Hughes and F. Lang optioned 156 claims which lie to the east and west of the Williams and Corona properties. These claims were put into the holding of their companies, Golden Sceptre Resources Ltd. and Goliath Gold Mines Ltd., who subsequently relinquished controlling interest to Noranda Exploration Company Ltd. In May, representatives of LAC Minerals Ltd. visited Corona's drill site and exchanged information pursuant to a possible joint-venture agreement. While negotiations with Williams' widow in Maryland for the Williams property were ongoing, diamond drilling was stepped back to the east of the outlined deposit. Drill hole 76 intersected a 10.5-foot (3.2 m) section grading 0.209 ounce gold per ton at a depth of 336.5 feet (102.5 m). This new, separate zone was the main Hemlo orebody. By August, 120 drill holes totalling 43 000 feet (13 106 m) had delineated 750 000 tons of rock grading 0.10 ounce gold per ton in the 'West' zone and had begun to indicate the much larger reserves of the 'East' or main zone. The Hemlo gold rush, ultimately involving 180 companies, ensued. Both Corona and LAC had been actively negotiating for the Williams property. In July, Mrs. Williams accepted LAC’s offer. Corona, citing a breach of a fiduciary agreement, sued LAC for ownership of the Williams claims. Teck Corporation subsequently entered into a joint venture agreement with Corona in December to develop what would become the David Bell Mine. 1982: LAC announced the discovery of the deposit on their property, which would become known as the Page-Williams Mine. Drilling by Goliath Gold Mines intersected the northward-dipping extension of the ore zone. The Goliath part of the deposit became, after a joint venture with Noranda Exploration Company Limited, the Golden Giant Mine. 51 �1985: The David Bell, Golden Giant and Page-Williams mines commenced gold production. 1986-87: The Supreme Court of Ontario awarded the Page-Williams Mine to International Corona Resources Ltd. LAC appealed the decision to the Ontario Court of Appeal but continued to operate the mine under conditions imposed by the court. The Ontario Court of Appeal upheld the earlier decision in October, 1987. The Supreme Court of Canada later granted LAC the right to appeal the provincial court ruling. 1987: Golden Sceptre Resources Ltd., Goliath Gold Mines Ltd., and Noranda Minerals Inc. amalgamated their holdings and formed Hemlo Gold Mines Inc. 1989: The Supreme Court of Canada awarded the Page-Williams Mine, Canada's largest gold producer, to Corona who subsequently re-named it the Williams Mine (the mine operates under the name Williams Operating Corporation). 1991: Homestake Mining Corporation purchased the assets of International Corona Resources, including their interest in the Williams and David Bell Mines. 1992: Noranda Minerals Inc. transferred ownership of the Golden Giant Mine to Hemlo Gold. 1996: Battle Mountain Canada Ltd. acquired the Golden Giant Mine. 1998-99: Williams Operating Corporation acquired the surface and mineral rights to the Sceptre claims from Battle Mountain Canada to the 9450 elevation of the Williams Mine grid in 1998. In 1999, Williams also acquired the surface and mining rights on the Horizon claims from Battle Mountain Canada to the 10150 elevation of the Williams Mine grid. These acquisitions would permit pit expansion to the west, and allow evaluation of underground mining of the down dip extension of the C-Zone pit. 1999: Homestake’s interest in the Williams and David Bell Mines was purchased by Barrick Gold Inc. in 1999. Milling operations ceased at the David Bell Mine when ore processing was transferred to the Williams mill. 2001: Ownership of Golden Giant changed to Newmont Canada Ltd. 2002: Williams acquired the surface and mineral rights from surface to the 10150 level on lease 273 and the remainder of lease 274 from Newmont Canada Ltd., providing an area for barren waste stockpiles from the expanded pit. 2006: Williams acquired the surface and mineral rights on lease 106623 from Newmont Canada Ltd. This acquisition allowed Williams to mine C Zone mineralization above the 9450 level as well as the down dip extension of the C Zone mineralization on the Interlake property. 52 �2006: The Golden Giant Mine became the first Hemlo mining operation to close. The mine produced a total of 6,780,373 ounces of gold. 2008: Newmont and Williams entered into an agreement to allow Williams to extend its underground mining operations on the Williams property through a 60 m restricted area (Boundary Pillar). 2009: Barrick Gold acquires Teck’s interests in the Williams and David Bell mines, giving Barrick sole ownership. 2010: Barrick acquires claims from Newmont that include the Golden Giant Mine workings. 2014: David Bell Mine closed. 2015: Barrick acquired additional claims from Newmont that are located immediately to the north and west of the Williams Mine. These acquisitions have opened up new opportunities for ore body expansion and exploration. Barrick’s land holdings as of August 2016 are shown on Figure 3. 2017: The Williams Mine is currently owned by Barrick Gold Corporation and continues to produce from open pit and underground operations. 2016 production was 235 000 ounces Au from 3 408 000 tonnes milled (Barrick Gold Corporation, Q4 and Year-End Mine Statistics, February 15, 2017). 53 �Figure 3. Barrick Gold Corporation’s Hemlo Mine area land holdings (from Williams Mine Closure Plan Amendment, August 2016). Hemlo Gold Deposit Production, Reserves and Resources Cumulative gold production from the Hemlo gold camp to the end of 2016 has been 21 667 271 ounces Au (worth $26 billion in $US at today’s gold price of $1200/oz) from 108 983 846 tonnes of milled ore. The following graph (Figure 4) illustrates annual production from the three mines between 1985 and 2016. Production peaked at approximately 1.35 Moz in 1990. Proven and Probable Reserve figures for the Williams Mine as of December 31, 2016, totalled 25 782 000 tonnes at a grade of 1.92 g/t Au for a total of 1 588 000 ounces Au. Measured and Indicated Resources currently stand at 58 897 000 tonnes at a grade of 0.908 g/ton Au for a total of 1 720 000 ounces Au (Barrick Gold Corporation, 2016 Mineral Reserves and Mineral Resources, February 15, 2017). 54 �Figure 4. Graph illustrating annual gold production from the Hemlo camp to the end of 2016. Note that between 2009 and 2014 only the combined production statistics for the Williams and David Bell mines were published. These combined statistics were assigned to the Williams mine for the purposes of this graphic. Figure 5. Long section view of the Hemlo deposit illustrating the location of the orebody (mined-out and reserves), mineral resources and areas of exploration potential along strike and down-plunge toward the west (from Barrick Gold Corporation, investor day presentation, February 22, 2016). 55 �Field Trip Stops Most of the field trip stop descriptions contained in this guide are mildly edited excerpts from an unpublished Ontario Geological Survey field trip guide that was prepared by Muir and Smyk (2006). Many of the coloured maps (unless otherwise noted) are also excerpted from that field guide. Field trip locations for the Heron Bay area are shown on Figure 6, while Hemlo area stops are shown on Figure 9. Figure 6. Geological map illustrating field trip stop locations (white triangles) in the Heron Bay area (geology from Ontario Geological Survey 2011) 56 �Stop H1: Pillowed and massive mafic metavolcanic rocks UTM Zone 16, 552024E, 5393482N From Marathon, travel 6.8 km southeast on Highway 17 from the corner of Peninsula Road to the intersection with Highway 627. Then travel south for 1.7 km along highway 627 to an outcrop exposure on the west side of the highway, at a power line crossing. Figure 7. Pillowed mafic metavolcanic rocks at stop H1. At this stop are several small outcrops that display sections of a steeply dipping, thick mafic, tholeiitic flow. The more northerly ones consist of "massive," medium- to fine-grained basalt. One outcrop displays what appears to be a dike with a highly irregular orientation. Although large amphibole porphyoblasts give the impression that the rock is somewhat gabbroic, grain size decreases upwards through the flow to its aphanitic top, where small pillows and autoclastic breccias are developed. Although the pillows are somewhat flattened (see Figure 7), they have well-preserved selvages that indicate that the top direction is toward the south. The most southerly outcrop exposes a section that displays possible flow banding and what appears to be a flow top breccia or pillow breccia. Minor quartz + carbonate veins are present, particularly in the main outcrop. 57 �Stop H2: Felsic pyroclastic rocks and “Heronite” dike Continue 5.6 km south along Highway 627 (and through the community of Heron Bay) to an outcrop located on the west side of the highway. This stop is located approximately 300 m past the intersection with a dirt road that heads east from the highway. UTM Zone 16, 553640E, 5388442N Figure 8. Analcite tinguaite (heronite) dike cross-cutting felsic pyroclastic rocks at stop H2. This stop provides an opportunity to view some of the felsic to intermediate pyroclastic rocks that occur in the immediate vicinity of Heron Bay. The metavolcanic rocks in this outcrop exposure have been cross-cut by an approximately 1 m wide Mesoproterozoic alkalic mafic dike that is likely to be related to the alkalic rocks of the Coldwell complex (see Figure 8). The Heron Bay area pyroclastic rocks are quartz-plagioclase-phyric, and include pyroclastic breccia, tuff-breccia, lapilli-tuff, tuff and crystal tuff (Muir 1982). These rocks are mostly calcalkalic dacite (intermediate), with some rhyolite breccias occurring near the lakeshore of Heron Bay). Overall, the fragments are subrounded to subangular, heterolithic in texture and composition, commonly more felsic than the matrix, and quartz-feldspar phyric. Some of the 58 �fragments are more mafic than the matrix and include intermediate, feldspar ± quartz-phyric rocks, and mafic, aphyric rocks. Bluish quartz phenocrysts are locally common. The matrix consists of feldspar, quartz, sericite, ± chlorite. These rocks were dated at 2695 ± 2 Ma by Corfu and Muir (1989). The alkalic mafic dike exposed in this outcrop has been classified as analcite tinguaite (heronite). A number of dikes with this composition occur in the Heron Bay area and were first described by Coleman (1899; 1900). Interesting features include the presence of ocelli within the dike and fluorite mineralization in the metavolcanic rocks near the dike contact. Stop H3: Deformed and altered felsic rocks along the shoreline of Heron Bay Turn around and travel north on Highway 627 for 1.3 km back to the community of Heron Bay and turn west onto a gravel road just before the railway crossing. Then travel 1.7 km to a boat launching site on the shore of Heron Bay, Lake Superior. UTM Zone 16, 551238E, 5388997N The felsic metavolcanic rocks exposed along the shoreline at this stop provide an example of some of the deformation and alteration effects that are associated with the Heron Bay deformation zone. The following description of the Heron Bay deformation zone is excerpted from MacTavish and Osmani (1996). The 'Heron Bay Deformation Zone' (HDZ) is a strong, locally intense, northeasterly trending, roughly 900 m wide zone, consisting of numerous, discrete, anastomosing shears, faults, and lineaments. Shearing has deformed the host rocks into a highly variable assemblage of quartz-sericite-carbonate±chlorite schist. Hematization, silicification and iron carbonatization are locally common. Intermediate to felsic hypabyssal intrusive rocks (eg. porphyries) are often emplaced along the margins of the zone in a region of more brittle-ductile deformation. These dykes/sills are shallow dipping, and with the exception of the margins of a few bodies, are undeformed. This suggests that they were emplaced during the later stages of the regional deformation events. A number of gold occurrences are found in association with the Heron Bay deformation zone. (e.g. Heron Bay Mine (Peekongay) and Bowhill Mines occurrences; Patterson 1984) One of those occurrences, known as the Screamer Zone is located approximately 200 m north-northwest of this stop. Gold mineralization at the Screamer Zone occurs in two variably brecciated quartzcarbonate veins. These veins occur within sheared and iron-carbonatized mafic metavolcanic host rocks at the contact with a feldspar porphyry dike. The dike itself intruded along the contact between mafic and felsic to intermediate metavolcanic rocks (MacTavish and Osmani 1996). 59 �60 Figure 9. Geological map of the Hemlo gold deposit area showing field trip stop locations (white triangles). Stops H4 and H15 are approximately 2.25 km east and west of the map limits respectively. Geology from Muir (2002). �Stop H4: Cedar Lake pluton Return to Highway 17 and then proceed east. After travelling for approximately 33 km along Highway 17, and just after passing the Williams Mine, you will see the intersection with Highway 614. Continue eastward along Highway 17 past this intersection for another 1.4 km to a location with road cuts on both sides of the highway. UTM Zone 16, 585227E, 5395748N The Cedar Lake pluton consists of massive to very weakly foliated medium-grained, microclinemegacrystic, hornblende-biotite granodiorite that contains mafic/ultramafic clots or xenoliths and diorite to monzodiorite inclusions and dikes (varies with location in pluton). The diorite and monzodiorite inclusions and dikes display a variety of relationships with the granodiorite including back-veining by the granodiorite. These relationships suggest a co-magmatic relationship (Beakhouse, 2001). The mafic-ultramafic and dioritic inclusions are commonly preferentially oriented, generally parallel or subparallel to the contact of the pluton (in a broad sense) and to the foliation in the granodiorite where it is discernible. The foliation locally deflects around the inclusions. The foliation may be related to original flow foliation; synchronous D2 deformation (i.e., syntectonic pluton); and/or a later D3 deformation. The pluton here has been intruded by a few intermediate-composition Archean dikes and a Proterozoic biotite lamprophyre dike. The Archean dikes are locally sheared with an apparent dextral sense. The granodiorite at Stop H4 was initially dated at 2688 +3 Ma (Corfu and Muir 1989). A more recent sample provided a somewhat younger age of 2680 ±1 Ma (Beakhouse and Davis, 2005). Stop H5: Cedar Lake Pluton western contact with metasedimentary rocks Turn around vehicle (there is a gravel road on the south side of the highway that can be used as a turn-around site just east of stop H4) and travel west along Highway 17 for 2.6 to the Yellow Brick Road sign. UTM Zone 16, 582891E, 5394626N At this stop, the Cedar Lake pluton consists of a relatively magnetite-rich, fine- to mediumgrained, biotite-hornblende granodiorite, which is massive except for a very weak foliation within several metres of the contact. A plagioclase-phyric, hornblende-biotite sheet exists to the southwest (Figure 10). The granodiorite at Stop H5 is intruded by many aplite and pegmatite dikes, some of which are composite: generally, pegmatite tends to crosscut aplite. The dikes typically terminate against the country rock schist, although some display ductile deformation within the contact granodiorite. The strike of the dikes within the granodiorite tends to “rotate” clockwise toward the contact, consistent with a dextral component of displacement at the contact. To the north-northwest of here, dextral shear along the contact is related to D2, even though D2 is an overall sinistral event. 61 �The country rocks within about 9 m of the contact consist of D3-crenulated mafic schist containing several locally dismembered, aplitic and granodioritic dikelets and stringers, which display refolded folds. This refolded nature of country rock layering and/or dikes in the immediate contact aureole is common with at least the Cedar Lake pluton and the Pukaskwa Batholith. For approximately the next 150 m westward, the country rocks consist mostly of various sets of metawacke and metasiltstone “packages”, with locally isoclinally folded layering. There are numerous mafic to felsic dikes and several swarms of dikes, generally granitic. Boudinage in dikes is locally present. The granitoid dikes consist mostly of weakly foliated granodiorite, which is similar in grain size and composition to the marginal phase of the Cedar Lake Pluton, as well as some plagioclase-porphyritic dikes. Within this sequence of turbiditic sedimentary rocks is a thick, composite sheet (Figure 10) of foliated, medium-grained, plagioclase-porphyritic, hornblende-biotite granodiorite (2687 ± 3 Ma; Corfu and Muir 1989), and foliated, finer-grained, biotite-hornblende granodiorite, both of which were intruded by aplitic dikes and subsequently by quartz veins. At the structurally lower (western) contact of this sheet, apophyses of the porphyritic granodiorite are folded about a less steeply dipping, S2-like foliation, with attendant attenuation and boudinage which has taken place in 2 dimensions. The folds have been interpreted by others as F2 folds (i.e., dikes predated F2), but it is more likely that they represent a local contact strain phenomenon related to a progressive and complex D2 event. Figure 10. Sketch map of outcrops at the Cedar Lake pluton western contact (from Muir et al. 1995). Note that the number labels shown on this figure denote the 1995 field trip stop numbers. 62 �Stop H6: Deformed metaconglomerate Travel 1.5 km to an outcrop located on the north side of the highway approximately 150 m past the access road to the former David Bell Mine site. UTM Zone 16, 581605E, 5393852N The main outcrop in this set of 3 small outcrops consists mostly of 2 metaconglomerate “layers” (one with predominantly cobbles-boulders; the other with predominantly pebbles), separated by a medium- to coarse-grained wacke locally with entrained pebble- to cobble-sized clasts. This represents a weak, remnant bedding (S0/S1). Flattened clasts tend to be aligned parallel to S2, and oriented lengthwise slightly clockwise with respect to the crude layering, consistent with this unit being on the northeast limb of the northwest-closing Williams fold (see Figure 9). This fold is part of the large-scale S-shaped Williams-Teck fold pair, formed during the D2 sinistral shear event and best delineated by the Moose Lake volcanic complex (yellow unit; Figure 9). Note that some of the clasts have been deformed by F3 folds (see Figure 11). Figure 11. Deformed conglomerate (F3 fold) at Stop H6. 63 �Stop H7: Cedar Creek fault zone area Travel 300 m west along Highway 17 to a series of outcrops on the north side of the highway. UTM Zone 16, 581337E, 5393743N 7A 7C 7B Figure 12. Sketch map of outcrops at stop H7 (from Muir et al. 1995). Outcrop H7A: Folded, sheared, porphyroblastic metasedimentary units Compositional layering in this outcrop is generally well defined and likely reflects modified original bedding. Folded layering defines F2 folds and retrograded porphyroblasts (cordierite?) are aligned parallel to S2. Other porphyroblast species found include garnet, and anthophyllite / cummingtonite. Evidence of superposed D3 dextral shear consists of: spaced shear bands and microfabrics with the apparent appropriate geometry and sense of fabric deflection. Some small Z-shaped folds are ambiguous F2 or F3 folds. Note that the orientation of the axial plane of the main F2 fold here is about 270o, whereas the overall orientation of the Williams fold axial plane is about 290o, possibly consistent with some back-rotation during D3. Some of the amphibolerich (“calc-silicate”) layers reveal evidence of alteration, commonly on both sides of the layers, and this evidence increases toward the Barren Sulphide zone (Outcrop 7B). This suggests that all of the rocks may have been altered. Toward this zone, there is a general change to more thinly laminated pelitic(?) sedimentary rocks, which also show more evidence of alteration (pyrite, muscovite). An increase in the degree of strain toward the zone is evident from the presence of tighter F2 folds and boudinage in calc-silicate layers and a mafic dike (Figure 13). Here the tight F2 folds display S-shaped asymmetry. 64 �Figure 13. Geological map of outcrop area H7A (after Muir 1990). Figure 14. F2 fold closure at Outcrop H7A. 65 �Outcrop H7B: Barren Sulphide Zone (Cedar Creek fault zone) This zone (at least 15 m thick) was also known in the earlier days as the Sucker zone. This is because it had all of the appearances of being a good candidate for a mineralized “horizon,” but in fact is essentially barren. The significant degree of oxidation precludes many unambiguous observations, but it appears to comprise sheared, pyritiferous, quartz-feldspar-sericite schist with small amounts of green mica and irregularly distributed, trace amounts of gold. The rock is possibly derived from the hanging wall pelitic/wacke metasediments (Outcrop H7A). Shearing is quite likely due to D3 based on fabric deflection, geometry of lozenges and the presence of inferred shear bands. Earlier involvement of D2 is possible. The Barren Sulphide zone is inferred to coincide with the Cedar Creek fault zone, one of several high-strain zones along this stretch of highway. Outcrop H7C: Reworked felsic volcaniclastic rocks These rocks consist of somewhat rusty weathering, layered, feldspathic, quartz-crystal-bearing rocks, interpreted to be derived from reworked felsic volcanic detritus of the Moose Lake volcanic complex (MLVC; Figure 9) which, in part, is spatially associated with the Hemlo gold deposit. The rocks do not appear to be the protolith to most if any of the muscovitic schists in the Barren Sulphide zone. Minor disseminated pyrite and rare green (vanadian?) muscovite are locally present in the structurally upper parts of this unit, indicating some alteration has taken place. Stop H8: Bedded felsic fragmental rocks Cross to the south side of the highway (while carefully watching for traffic) and walk a short distance to a low outcrop that is adjacent to the highway. UTM Zone 16, 581210E, 5393696N This is an outcrop of layered, quartz-plagioclase-phyric, heterolithic, felsic fragmental rocks that display a variety of clast sizes. The clasts are flattened parallel to S2 and are oriented counterclockwise with respect to layering (consistent with being on the central limb of the WilliamsTeck fold pair). Stop H9: Felsic fragmental and pelitic sedimentary rocks Carefully re-cross the highway and walk 100 m west to an outcrop on the north side of the highway. UTM Zone 16, 581099E, 5393719N The structurally lower part of Stop H9 consists of an enigmatic unit inferred to be altered, felsic fragmental rocks of volcanic origin. Alteration here is uncharacteristic, showing parts that are greenish, charcoal grey, and pink. Lenses are locally discernible, in part defined by the 66 �distribution of the coloured minerals. Staining for K-feldspar shows microcline tends to be locally distributed around some of the mafic lenses and along some of the dominant cleavage. Note the high degree of strain in this outcrop (likely another, currently unnamed, high-strain zone). The structurally upper part of this outcrop, which displays a sharp, locally slightly discordant contact with the felsic rocks, consists of hard, thinly laminated pelitic rocks that contain abundant staurolite, garnet and sillimanite (fibrolite). A very tight east-closing F2 fold in the layering is evident, similar in style to that in the pelitic rocks of Outcrop H7A. A possible westclosing fold (south part of unit) is enigmatic but, if present, would make this a Z-shaped fold pair. Some staurolite crystals have overgrown and incorporated the laminations. Clockwise rotation of many staurolite crystals, on both limbs of the F2 fold, as well as the orientation of fibrolite, locally about staurolite and counterclockwise to laminations, are inferred as D3-related features. Reddish alteration along undeformed fractures may be related to common alteration (hematitization and/or epidotization) found near diabase dikes, one of which lies at the west end of the outcrop. Stop H10: Main Mineralized Zone Travel 450 m west along Highway 17 to a low outcrop in the ditch on the north side of the highway. UTM Zone 16, 580652E, 5393727N This outcrop exposes the gold-mineralized rocks of what had been previously termed the West or Lake Superior zone. Although not directly connected with the main Hemlo orebody (aka “B” zone) it represents both its on-strike and up-dip projection that was the focus of the vast majority of exploration efforts prior to the discovery of the main orebody in May, 1981. Figure 15 attempts to indicate various features that were visible several years ago. Close examination near the westernmost channel sample (outcrop “B”) reveals highly strained pyritiferous fragmental rocks with, among other things, sparse green (vanadian) muscovite (gMs) lenses and a general grey to bluish grey colour in parts of the matrix due to finely disseminated molybdenite, which is generally a “pathfinder” to gold in this deposit. Grab and channel sampling has returned up to 6 g/t Au. The outcrop contains appreciable barite crystals (Davis and Lin 2003). The high degree of strain in this fragmental rock reflects that these rocks lie within the Lake Superior Shear zone, one of several high-strain zones in the Hemlo deposit area. This high-strain zone is up to 50 m thick. Complex fabric relationships with a possible chronologic interpretation are indicated in the inset at the upper part of Figure 15 pertaining to outcrop “B”. F3 folds (“flanking structures”) can be seen in the southeast part of outcrop “B”. 67 �About 50 m west along the highway ditch is outcrop “A”, which displays the more “massive” quartz-plagioclase porphyry (QPP) that structurally underlies the Main Mineralized zone. Also evident is an example of one of many (in the area) swarms of plagioclase-phyric dikes. Note that the dikes are deformed: extension followed by shortening, reflecting boudinage related to D2 predated shortening related to D3. Also note that in these outcrops and all other ones already seen and to be seen later today, that quartz veins are generally few and small; there typically being as many, if not more, small pods or lenses of quartz as there are veins. Full interpretation of kinetics is commonly enigmatic. This attests to the complex history of strain. Figure 15. Main Mineralized Zone (also known as West Zone) geology (after Muir and Smyk 2006). 68 �Stop H11: Inter-ore zones metasedimentary rocks. Travel 180 m west along Highway 17 to a pair of outcrops on the north side of the highway that are bisected by a road that formerly accessed the Williams A Zone open pit. UTM Zone 16, 580471E, 5393729N The outcrop to the east of the A Zone pit road (see Figure 16) displays relatively feldspathic, turbiditic sedimentary layers, with well-preserved grain-size gradation indicating overturned, south-southwest facing units. The well-preserved nature of this outcrop, contrasts with the adjacent rocks (hanging wall and footwall) and suggests that it may represent a low-strain lithon. To the west of the A Zone road, the sediments are darker grey and locally contain magnetitebearing layers and discrete folds. At the western end of the outcrop, the layering is interpreted as transposed. The increase in degree of strain over an inferred 5 to 10 m across strike between outcrops is considerable and is inferred to reflect, in part, the Moose Lake fault zone. At least 2 types of feldspar porphyry dikes are exposed in this outcrop. Out of interest, the structurally lower one, locally called '‘popcorn” porphyry because of its coarser plagioclase phenocrysts, was used in the early stages of exploration as a cutoff marker unit for drilling collared to the north. Note that, ironically, the Lower Mineralized zone occurs structurally below this dike, and is exposed about 110 m further west along the highway. Figure 16. Sketch map of outcrops at Stop H11 (from Muir et al. 1995). The number labels shown on this figure denote the 1995 field trip stop numbers. The A Zone pit can be seen through the fence located just to the north of the outcrops at this stop. Figure 17 illustrates the pre-mining geology of the surface expression of the A Zone. The ore zone occurs at the contact between metasedimentary and fragmental rocks. 69 �Figure 17. Pre-mining geological map of the A Zone. The approximate outline of the ore zone is shown in red (from Muir 2002; inset showing ore outline derived from Walford et al. 1986). Stop H12: Up-dip projection of the Lower Mineralized Zone. Travel 110 m west to another outcrop on the north side of the highway. UTM Zone 16, 580311E, 5393727N This outcrop is the up-dip projected zone of alteration associated with the actual Lower Mineralized Zone (which occurs at about 900 m depth here). However, the rocks at surface are altered, and locally contain low gold values, minor amounts of green (vanadian?) muscovite, pyrite, medium to coarse-grained tourmaline, muscovite, and rarely microcline porphyroblasts. Barite veinlets and layer-parallel seams have been reported (Schnieders, pers. com. 1994). Magnetite-bearing wacke, which is at the north part of the outcrop (see Figure 18), becomes muscovitic and possibly silicified adjacent to the quartz-plagioclase porphyry (QPP of the Moose Lake Volcanic Complex). The contact between the altered sedimentary rocks and the QPP is locally at an uncharacteristically high angle to the layering in part of the outcrop (folded? crosscutting?). Granodioritic, feldspar porphyritic, and intermediate to mafic dikes occur here. 70 �Figure 18. Sketch map of the up-dip projection of the Lower Mineralized Zone (from Muir et al. 1995). The number label shown on this figure denotes the 1995 field trip stop number. Stop H13: Hemlo Fault zone. Travel 1.5 km west along highway 17 (past the Williams Mine entrance road and tailings line overpass) to a steep road cut on the north side of the highway. UTM Zone 16, 578843E, 5393791N This exposure shows part of the fault zone (which is probably up to tens of metres thick), and features the structurally lower amphibolite/gneissic amphibolite (inferred to be derived from pillowed and massive mafic volcanic units, based on outcrops to the west) and the structurally overlying feldspathic sedimentary rocks. Mafic-ultramafic dikes occur largely as chlorite ± actinolite ± talc phyllonites, mostly within the amphibolitic rocks, and provide much of the evidence for relatively late D3 dextral shear, as can be seen by fabric relationships. Also occurring here are a few intermediate dikes (presently biotite schists) and a plagioclase-phyric dike swarm. Very-coarse-grained tourmaline crystals and crystal “clots” are associated with a feldspar ± quartz dike or vein. The outcrop, featured in Figure 19, was modified in 1989 (postmapping by Smart 1988) by roadway blasting. 71 �Figure 19. Geology of the Hemlo fault zone at Stop H13 (modified from Smart 1988). Stop H14: Highway Zone Travel 1.3 km west along Highway 17 to a large outcrop area on the north side of the highway. UTM Zone 16, 577540E, 5393717N This outcrop (see geological map, Figure 20) demonstrates the heterogeneous strain that occurs even away from the Hemlo gold deposit. It presumably represents the structurally upper part of the tens-of-metres thick Hemlo fault zone. The north part of the outcrop displays disrupted layering in feldspathic siltstone and arenite with S-shaped folds. As one approaches the highway (and hence the Hemlo fault zone), the units change to: 1) quartz-plagioclase-phyric, heterolithic and monolithic, felsic fragmental rocks, displaying contacts tightly interfolded(?) with wacke The felsic fragmental rocks are pyritiferous, erratically auriferous and locally contain green (vanadian?) muscovite; and then 2) wacke-siltstone which displays well-developed laminations that are thinly spaced and inferred to be derived, in part, from a combination of primary layering and compositional differentiation along the main cleavage. Very tight folds (south of the diabase dike) with S-shaped asymmetry are evident. One interpretation is that there is a composite S1/S2 fabric that is locally preserved and oriented 72 �slightly clockwise to the dominant Sm (mylonitic) fabric. The S3 flattening fabric (micaceous) can be locally seen within the Sm cleavage, counter-clockwise to it. The contact with the amphibolitic and mafic-ultramafic rocks seen in Stop H13 presumably lies to the south of the highway. Figure 20. Geological map of the Highway Zone at Stop H14 (modified from Muir et al. 1995). 73 �Stop H15: Homestake F3 fold. Travel 3 km west along Highway 17 to a high road cut on the south side of the highway. UTM Zone 16, 574529E, 5393489N This outcrop (see geological map, Figure 21) is possibly north of the Hemlo fault zone, but ambiguity abounds. Several features in this otherwise over-lichened outcrop are of note: 1) Rare, calc-alkalic or “alkalic” pillows with feldspar phenocrysts and hornblende crystals, locally with amygdaloidal/vesicular textures; 2) Abrupt strain partitioning from slightly strained pillows (near the highway) to moderately strained pillows (D2) over a distance of less than 1 m southward; 3) A related volcanic breccia or volcaniclastic unit (fragments with similar composition and textures); 4) Adjacent feldspathic arenite and/or altered wacke; 5) A short, thin gossan, which returned 103 ppm Au and 100 ppm Mo (Resident Geologist’s files, Thunder Bay South District) – the highest molybdenite results outside of the Hemlo gold deposit proper and similar to some feldspathized tholeiitic pillows “along strike” to the west of the deposit; 6) Varieties of variably deformed metawacke and metasiltstone, locally with garnet ± magnetite ± staurolite porphyroblasts; 7) Examples of S2 and F2 structural elements overprinted by S3 and F3 structural elements (e.g., folded S2); 8) Numerous variably deformed (folded, transposed) dikes such as mafic schists; mafic biotite lamprophyre schists ± mafic-ultramafic xenoliths; a mafic net-veined or backveined dike (mixture of previous 2 types?); and 2 types of intermediate dikes; 9) The xenolith-rich parts of the lamprophyre dikes are very similar to some of the diamondiferous intrusions found in the Wawa greenstone belt. 74 �Figure 21. Detailed geology of the Homestake F3 fold exposure (modified from Muir et al. 1995). 75 �References Barrick-Hemlo Williams Operating Corporation and Amec Foster Wheeler Environment and Infrastructure 2016. Williams Mine closure plan amendment; Ministry of Northern Development and Mines, Thunder Bay Mineral Development and Lands Branch Office, Mine Closure Plan Files, 207 p. Beakhouse, G.P. 2001. Nature, timing and significance of intermediate to felsic intrusive rocks associated with the Hemlo greenstone belt and implications for the regional geological setting of the Hemlo gold deposit; Ontario Geological Survey, Open File Report 6020, 248p. Beakhouse, G.P. and Davis, D.W. 2005. Evolution and tectonic significance of intermediate to felsic plutonism associated with the Hemlo greenstone belt, Superior Province, Canada; Precambrian Research, v.137, p.61-92. Burk, R., Hodgson, C.J., and Quartermain, R.A. 1986. The geological setting of the Teck-Corona Au-MoBa deposit, Hemlo, Ontario, Canada; in Proceedings of gold '86, an international symposium on the geology of gold, Toronto, Ontario, 1986, p.311-326. Coleman, A.P. 1899. Dyke rocks near Heron Bay; Ontario Bureau of Mines, Vol.8, Pt.2, p. 172-174. Coleman, A.P. 1900. Heronite or analcite tinguaite; Ontario Bureau of Mines, Vol.9, p. 186-191. Corfu, F. and Muir, T.L. 1989. The Hemlo–Heron Bay greenstone belt and Hemlo Au-Mo deposit, Superior Province, Canada, 1. Sequence of igneous activity determined by zircon U-Pb geochronology. Chemical Geology (Isotope Geoscience Section), 79: 183-200. Davis, D.W. 1998. U-Pb zircon and titanite geochronology; Part 3, in Geological setting of the Hemlo gold deposit; an Interim Progress Report, Jackson, S.L., Beakhouse, G.P. and Davis, D.W. (eds), Ontario Geological Survey, Open File Report 5977, p.1-10. Davis, D.W. and Lin, S. 2003. Unraveling the geologic history of the Hemlo Archean gold deposit, Superior Province, Canada: a U–Pb geochronological study; Economic Geology, v.98, p.51-67. Harris, D.C. 1989. The mineralogy and geochemistry of the Hemlo gold deposit, Ontario; Geological Survey of Canada, Economic Geology Report 38, 88p. Jackson, S.L. 1998. Stratigraphy, structure and metamorphism; Part 1, in Geological setting of the Hemlo gold deposit; an Interim Progress Report, Jackson, S.L., Beakhouse, G.P. and Davis, D.W. (eds), Ontario Geological Survey, Open File Report 5977, p.1-71. Langlais, A. and Barber, R. 2015. Barrick Gold Inc. work assessment report, Bomby Township, Thunder Bay Mining District; Thunder Bay South Resident Geologist District, Assessment Files, AFRO number 2.56670, 75p. MacTavish, A. and Osmani, I. 1996. Geological report for the Toothpick West and East properties, Heron Bay project, Pic Township, northern Ontario; Thunder Bay South Resident Geologist District, Assessment Files, AFRO number 2.16854, 22p. 76 �Magnus, S.J. and Arnold, K.A. 2016. Geology and mineral potential of the western Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2016, Ontario Geological Survey, Open File Report 6323, p.11-1 to 11-17. Magnus, S.J. and Walker, J. 2015. Geology and mineral potential of Walsh, Tuuri and Syine Townships, Schreiber-Hemlo greenstone belt; in Summary of Field Work and Other Activities 2015, Ontario Geological Survey, Open File Report 6313, p.14-1 to 14-12. Muir, T.L. 1997. Precambrian geology, Hemlo gold deposit area; Ontario Geological Survey, Report 289, 219p. Muir, T.L. 2000. Geological compilation of the eastern half of the Schreiber–Hemlo greenstone belt; Ontario Geological Survey, Map 2614, scale 1:50 000. Muir, T.L. 2002. The Hemlo gold deposit, Ontario, Canada: principal deposit characteristics and constraints on mineralization; Ore Geology Reviews, v.21, p.1-66. Muir, T.L. 2003. Structural evolution of the Hemlo greenstone belt in the vicinity of the world-class Hemlo gold deposit; Canadian Journal of Earth Sciences, v.40, p.395-430. Muir, T.L., Jackson, S.L. and Beakhouse, G.P. 1999. The regional framework of the Hemlo Gold Deposit; in Summary of Field Work and Other Activities 1999, Ontario Geological Survey, Open File Report 6000, p.15-1 to 15-6. Muir, T.L., Schnieders, B.R. and Smyk, M.C. 1995. Geology and gold deposits of the Hemlo area revised edition; Institute on Lake Superior Geology, 41st Annual Meeting, Marathon, ON, 1995, v.41, part 2d, 120p. Muir, T.L. and Smyk, M.C. 2006. Hemlo greenstone belt: Hemlo gold deposit area and Heron Bay area; unpublished OGS Precambrian Geoscience Section field trip guide, May 2-3, 2006, 28p. Ontario Geological Survey 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release—Data 126–Revision 1. Page, T.W. 1947a. A report on the Ollmann-Williams group of claims, Hemlo, Ontario; assessment files, Resident Geologist's office, Schreiber--Hemlo District, Thunder Bay, 8p. Page, T.W. 1947b. Report on the property of the Lake Superior Mining Corporation Ltd.; assessment files, Resident Geologist's office, Schreiber--Hemlo District, Thunder Bay, 1p. Page, T.W. 1948. A report on the property of Lake Superior Mining Corporation Ltd., Hemlo Area, Ontario; assessment files, Resident Geologist's office, Schreiber--Hemlo District, Thunder Bay, 23p. Page, T.W. 1949. Report on the properties of Lake Superior Mining Corporation Limited; assessment files, Resident Geologist's office, Schreiber--Hemlo District, Thunder Bay, 4p. Pan, Y. 1990. Metamorphic petrology and gold mineralization of the White River gold prospect, Hemlo area, Ontario; unpublished PhD thesis, University of Western Ontario, London, Ontario, 256p. 77 �Pan, Y. and Fleet, M.E. 1988. Metamorphic petrology of the White River gold prospect, Hemlo area, Ontario; in Geoscience Research Grant Program, Summary of Research 1987-1988, Ontario Geological Survey, Miscellaneous Paper, p.164-176. Pan, Y. and Fleet, M.E. 1989. Metamorphic petrology and gold mineralization of the White River gold prospect, Hemlo area; in Geoscience Research Grant Program, Summary of Research 1988-1989, Ontario Geological Survey, Miscellaneous Paper 143, p.42-52. Pan, Y. and Fleet, M.E. 1990. Metamorphic petrology and gold mineralization of the White River gold prospect, Hemlo area; in Geoscience Research Grant Program, Summary of Research 1989-1990, Ontario Geological Survey, Miscellaneous Paper 150, p.13-26. Patterson, G.C. 1984: Field Trip Guidebook to the Hemlo Area; Ontario Geological Survey, Miscellaneous Paper 118, 33p. Smart, P. 1988. A lithological study along the Hemlo fault; unpublished Hon. BSc thesis, Queen's University, Kingston, Ontario, 43p. Thompson, P.H. 2006. A new metamorphic framework for the Hemlo greenstone belt: Implications for deformation, plutonism, alteration and gold mineralization; Ontario Geological Survey, Open File Report 6190, 80p. Walford, P., Stephens, J., Skrecky, G., and Barnett, R. 1986. The geology of the 'A' Zone, Page-Williams Mine, Hemlo, Ontario, Canada; in proceedings of Gold '86, an international symposium on the geology of gold, Toronto, Ontario, 1986, p.362-378. Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L. and Sage, R.P. 1991. Wawa Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.485-541. 78 �Field Trip 2 Unusual Archean Diamond-bearing rocks of the Wawa Area by Ann C. Wilson Ministry of Northern Development and Mines, Resident Geologist’s Program, Ontario Geological Survey, Timmins, Ontario On the cover (clockwise from top): Giant lower crustal to upper mantle xenoliths, Enigma Property, Menzies Township, Oasis Diamond Corporation Inc.; Diamonds from the Festival Property (photo courtesy of Pele Mountain Resources Inc.); Sandor Diamond Occurrence, Highway 17, Spider Resources Inc. & KWG Resources Inc.; Heterolithic diamond-bearing breccia, Engagement Zone, Northern Sierra Minerals Corporation 79 �Unusual Diamond-bearing rocks of the Wawa Area Introduction In 1991, local prospector C. “Mickey” Clement recovered at least three alluvial diamonds from the Michipicoten River. Two of the stones were sent to the Department of Mineralogy, Royal Ontario Museum, where they were identified as industrial-grade diamonds with weights of 1.05 and 1.13 carats. Both stones were described as frosted and graphite-inclusion riddled. In 1995, prospector Sandor Surmacz and geologist Marcelle Hauseux recovered diamonds from a bedrock occurrence on the Trans Canada Highway, approximately 20 km north of the town of Wawa. An 18.1-kg bulk sample of a xenolith-rich lamprophyre yielded 1 macrodiamond and 5 microdiamonds. All but one was gem quality. Since then, more than 50 occurrences of diamondiferous bedrock have been reported in an area of approximately 30 km2 in size, centred approximately 20 km north of the town of Wawa. The occurrences are hosted within a sequence of unusual, Archean-aged, heterolithic breccias. Historically, this sequence of rocks has received little exploration interest and was considered to have little economic significance. This field trip will focus on exposures of a series of foliated lamprophyre dikes and associated heterolithic breccias outcropping in Lalibert, Leclaire, Menzies and Musquash townships. This field trip guide represents a summary of information available at the time of writing and should not be considered the final analysis of these rock types. Much more research is required on these rocks. Active exploration and research is still underway on the properties included in this field guide. Given the limitations of time, the field trip will visit only some of the more accessible properties. Bear in mind that when visiting active exploration or mine properties, permission must be granted by the property owner. Current ownership information can be obtained from the Resident Geologist’s Office in Timmins, or the District Geologist’s Office in Sault Ste Marie, Ontario. Geological Overview of the Michipicoten Greenstone Belt The Wawa region lies within the Wawa subprovince of the Canadian Shield. The Michipicoten greenstone belt extends inland for approximately 150 km from the Lake Superior shore and has an average width of 38 km. The greenstone belt consists of supracrustal rocks of Archean age. Younger Archean granitic rocks surround the greenstone belt. Figure1 shows a generalized geological map of the Michipicoten greenstone belt. The oldest volcanic cycle is approximately 2900 Ma and is of limited distribution. This cycle is best developed in Esquega Township along the southern flank of the supracrustal assemblage. Portions of this metavolcanic cycle extend into eastern McMurray and western Lastheels townships. The base of this volcanic cycle consists of massive to pillowed komatiitic flows intruded by mafic sills. The ultramafic rocks are overlain by intermediate to felsic tuffs capped by thinly bedded chert-magnetite-sulphide iron formation (Judith-Kathleen Iron Formation). Intermediate to felsic metavolcanic tuffs below the Judith iron formation have been dated at 2889 ± 9 Ma (Turek et al. 1992). Overlying the 2900 Ma cycle is a 2750 Ma volcanic cycle. This volcanic cycle is predominately composed of intermediate to felsic tuffs, breccias and flows. Porphyritic and spherulitic flows are not common and most of the intermediate to felsic metavolcanic rock is fragmental. The base of this cycle consists of high magnesium and high iron tholeiitic massive and pillowed flows. It lies conformably atop the JudithKathleen iron formation at the east end of Wawa Lake, but the basal unit is poorly exposed elsewhere. Overlying the mafic metavolcanic rocks is a sequence of heterolithic, intermediate to mafic breccia that has been traced for a distance of over 13 kilometers. This unit is in turn, overlain by a thick section of intermediate to felsic tuffs, breccias and massive flows that reaches a maximum thickness of approximately 2000 m below the Helen Iron Formation. The upper part of the intermediate to felsic metavolcanic rocks has been dated at 2749 ± 2 Ma (Turek et al. 1992). The Michipicoten (Helen) iron formation caps this 80 �81 Figure 1. Generalized geological map of the Michipicoten greenstone belt showing some of the diamond occurrences (modified after Stott et al. 2002). �this volcanic cycle. It commonly exceeds 100 m in thickness and, in the vicinity of the past-producing Helen Mine; it has been tectonically thickened to over 300 m. This iron formation was the source of virtually all commercial iron ore production in the Wawa area from 1898 to 1998. The youngest metavolcanic rocks in the area are those of the 2700 Ma volcanic cycle. These rocks underlie approximately fifty-percent of the Michipicoten greenstone belt and are found in the central and northern parts of the belt. The basal unit is composed of massive and pillowed mafic amygdaloidal flows that attain a maximum thickness of 5.5 km in the Goudreau area. This unit is overlain by intermediate to felsic metavolcanic rocks, or their stratigraphic equivalent, the Doré metasedimentary rocks. These metavolcanic rocks are typically composed of tuffs and monolithic and heterolithic breccias. Quartz + feldspar crystal tuff and an intermediate tuff from this volcanic cycle returned a U-Pb zircon age of 2701 ± 8 Ma (Turek et al. 1992) and 2701.4 ± 2.1 Ma (Ayer et al. 2003). The intermediate to felsic tuffs interdigitate with clastic metasedimentary sequences that include cross-bedded sandstone and a tonalite cobble conglomerate (Doré conglomerate). Corfu and Sage (1987, 1992) reported an age of 2698 ± 2 Ma for a tonalite clast in the Doré conglomerate and maximum ages of 2680 ± 3 and 2682 ± 3 Ma for sedimentary sequences in northern and central parts of the Michipicoten greenstone belt. Geochronological and structural evidence indicates that sedimentation continued after cycle 3 volcanism and predated a major folding and faulting event. Arias (1996) noted that the rocks comprising cycle 3 in the central part of the Michipicoten greenstone belt are upside down and represent the overturned limb of a belt-scale recumbent nappe fold. This inverted limb has been refolded and imbricated by subsequent southverging thrust faults, which caused local repetition of the stratigraphic sequence (Wilson 2004). Felsic plutonism occurred synchronous with all of the major volcanic cycles and continued after volcanism ceased at Wawa. Plutonic rocks associated with cycle 1 volcanism include the Murray-Algoma porphyry (2881± 6 Ma) and the Regnery biotite granite of the Hawk Lake granitic complex (2888 ± 2 Ma). Both intrusions are situated in Esquega Township. The Jubilee granitic stock, located in McMurray Township, was dated at 2745 ± 3 Ma and is coeval with cycle 2 volcanism. Plutons associated with cycle 3 volcanism range in composition from tonalite through granodiorite and granite and have ages ranging from 2698 to 2693 Ma. These plutons are located south and west of the Michipicoten greenstone belt (Stone and Semenyna 2004). The Kapuskasing Structural Zone extends east from the shore of Lake Superior, northeast through Kapuskasing and into the Hudson Bay Lowland. Local features interpreted to be associated with it include northeast-striking Proterozoic lamprophyre dikes (Sage 1994; Morris 1999). Lamprophyre dikes of middle Proterozoic age are common in the region south of the Wawa-HawkManitowik Lake Fault and rare to non-existent north of the fault. They are carbonate, biotite, and sometimes olivine-rich and usually less than 1.0 m in width. The dikes generally strike northeast. Dikes exposed in McMurray Township commonly have blue to blue-green sodic amphibole developed in their wall rocks. This mineral has been interpreted to be a product of fenitization. These dikes are likely spatially and temporally related to the emplacement of the Keewenawan-age Firesand River Carbonatite (Sage 1994). North of the Wawa-Hawk-Manitowik Lake Fault, in the area extending west from the former Magpie Mine (Leclaire Township), to the east side of the Dickenson Lake Stock (Lalibert Township), there are a series of what appear to be narrow, biotite-amphibole-rich dikes that have been interpreted as Archean lamprophyres. These dikes commonly have large, rounded inclusions (lower crust to upper mantle-derived xenoliths) up to 3.0 m in size, the centres of which are often completely altered to radiating clusters of actinolite crystals. Titanite from the matrix of one of these dikes returned an age of 2703 ± 42 Ma (Sage 2000). The date is interpreted to be a minimum age of intrusion. Subsequent dating of a zircon from a gneissic xenolith from the same dike returned an age of 2684.9 ± 1.4 Ma (Ketchum, Kamo and Davis 2003). 82 �Several narrow (<10 m) kimberlite dikes also intrude the Michipicoten greenstone belt. These dikes are largely restricted to the eastern part of the greenstone belt and are spatially related to the northeast-trending Kapuskasing Structural Zone. Several of these dikes have been dated. Intrusions K115 and K121, located in Isaac Township, returned an average 87Rb/86Sr age of 1097 ± 7 Ma (Kaminsky et al. 2002). A kimberlite dike intersected in Pele Mountain Resources Inc.’s drill hole 97-34, drilled in Riggs Township, returned a 207 Pb/206Pb age of 1197 ± 24 Ma (R. P. Sage, Ontario Geological Survey, unpublished, 2000; Wilson 2004). Quaternary Geology All Quaternary deposits within the Wawa area were deposited during the Late Wisconsin by the Labrador sector of the Laurentide Ice Sheet. Peat, recovered from a bog located within the surface of a terrace associated with the highest glacial lake in the Lake Superior basin, was radiocarbon dated at 9759 ± 170 BP. A caribou antler, recovered from near the surface of the delta of glacial lake Minong III, yielded a radiocarbon date of 8820 ± 80 BP (Morris 2001). Bedrock striae indicate that there were two prominent ice flow directions. The oldest and most pervasive ice flow was south to southwest (159° - 240°). A later, weaker ice flow was to the southwest and west (220° - 290°). The younger set of striae was formed during the latter stages of glaciation as the ice sheet began to thin and bedrock topography began to influence the direction of ice flow. Much of the overburden consists of a thin (< 1 m), discontinuous till veneer draped over the bedrock. Most of the area's thicker surficial deposits are located within bedrock controlled valleys where glaciolacustrine waters from the Lake Superior basin covered the area. In several of these valleys, the glaciofluvial material reaches thicknesses up to 32 m (Morris 2001). Structural Geology of the Diamond-bearing Rocks The greatest concentration of diamond-bearing rocks in the Michipicoten greenstone belt is constrained to a roughly 30 km2 block of land coinciding with a D1 recumbent nappe identified by Arias (1996). The nappe lies immediately north and adjacent to the Kapuskasing Structural Zone and has affected the two youngest rock sequences in the greenstone belt. In cross section, a generalized structural section across the greenstone belt would be upside down and represents the overturned limb of a belt-scale recumbent fold. This inverted limb of the nappe fold has been refolded and imbricated by subsequent south-verging thrust faults. The result is that the geometry of the rocks hosting the diamondiferous bedrock is that of an inverted anticline (Arias and Helmsteadt 1990). The effect is to create a tectonic repetition of the diamond-bearing sequence of rocks. In simplified terms, the rocks hosting the diamonds in the Wawa area can be considered as a single, overturned fold limb that has been faulted along a series of thrust zones. As a result, the diamond-bearing breccias, the lamprophyres associated with them and the surface onto which these metavolcanic rocks were deposited has been repeated at least four times (Walker 2002). The western limit of the nappe is likely the Dickenson Lake fault that passes along the west side of the Dickenson Lake stock. Extrusive lithologies similar to those hosting diamonds were recognized to the east of the stock by Sage (1993) during reconnaissance mapping. To date, no diamondiferous occurrences have been discovered to the west of the Dickenson lake fault. The eastern limit of the nappe is the Marsden Lake fault. Prior to 2002, rocks favourable to hosting diamonds had not been observed east of this structure (Wilson 2004). However, in 2002, Oasis Diamond Exploration Inc. made a discovery of diamonds on the east shore of the Magpie River and in 2004, diamondiferous occurrences of bedrock were found on the west shore of the Magpie River in Chabanel Township. These discoveries suggest that the potential for these host rocks extends farther eastward than previously believed. Figure 2 shows a graphic representation of the nappe structure. 83 �Figure 2. Composite structural section through the central part of the Michipicoten greenstone belt. Section X-Y from south-central Corbiere Tp. (Josephine Iron Range) to Andre Lake central Corbiere Tp. Section Y-Z is a schematic section from west central Lalibert Tp. The sketch in the lower left explains the present configuration of the belt as a regional nappe fold (F1) refolded about F2. Imbricate thrusts are considered related to Fs (Arias 1996). Description of the Diamond-bearing Rocks The diamond-bearing breccia and associated lamprophyre is broadly distributed throughout Lalibert, Leclaire, Menzies and Musquash townships. On-going mapping by Pele Mountain Resources Inc. on the Festival property and by Nathalie Lefebvre on the GQ Property has helped to refine the classification of these rocks. Systematic exploration and sampling suggest that the individual diamond occurrences are part of a much larger suite of rocks and that diamonds occur primarily within discrete layers at the base of diamond-bearing zones. North-northwest-trending diamond-bearing zones of breccia and lamprophyre are up to 1500 m in length and up to 800 m in width (Pele Mountain Resources Inc., press release, January 18, 2005). The breccia forms thick units (maximum true thickness is approximately 110 m) dipping to the northeast 30°. The lateral extent and thickness of the breccia unit is not well constrained, owing to the large-scale regional folding and thrusting (Lefebvre 2004). Figure 3 provides a detailed map of the southwest corner of the Festival Property showing a recent interpretation of these diamond-bearing zones. The diamond-bearing rocks can be visually subdivided into two classes, lamprophyre (dikes and bodies of indeterminate morphology) and heterolithic or polymict breccias. It is often difficult to differentiate between the two classes since the lamprophyre dikes frequently contain an assortment of inclusions that 84 �85 Figure 3. Detailed geology of the southwestern corner of the Festival Property (modified from Pele Mountain Resources, Inc. press release, January 18, 2005). The units shown in green consist of mafic to intermediate massive and pillowed flows, breccias and tuffs (Vaillancourt 2005b, Sage et al. 1982) �give them the appearance of breccia. The lamprophyre dikes cut the breccia units. Both lithologies have been metamorphosed to upper greenschist facies. The breccia primarily consists of angular, pebble-sized, lithic fragments, mainly of volcanic composition, contained within a green to grey fine-grained matrix. The matrix grain size ranges from < 2 mm to 1 mm. At least eleven distinctive types of lithic fragments have been observed in the breccia and the fragments are irregularly distributed throughout the breccia. The clast population is primarily derived from rocks with which the breccia is intercalated. Most typically these clasts are mafic and felsic metavolcanic rocks and intermediate to mafic intrusive rocks. Other clast types include fragments of clast-supported breccia within matrix supported breccia, fragments of earlier matrix-supported breccia with fewer than 5% fragments and coated lithic fragments (Lefebvre 2004). The breccia is characteristically massive, unstratified and poorly sorted with clast size ranging from sand to boulders up to 9 m. Primary sedimentary structures such as bedding and crude grading are rare (Lefebvre 2004). Petrographic work by Lefebvre (2004) on the breccias identified a typical fragmental texture within a mineralogically variable matrix. Typically the groundmass is dominated by actinolite, but chlorite and biotite dominated groundmass are locally predominant. Juvenile magmatic material also was observed as discrete fragments and rims on other clasts in the breccia. Its petrography is distinct from the breccia groundmass. The juvenile magmatic material contains more abundant actinolite grains; fewer epidote and fine-grained plagioclase grains and more fine-grained oligoclase and muscovite; and more oscillatoryzoned hornblende. Most exploration efforts over the past few years have concentrated on the breccias because they are considered to have the best potential for hosting commercial diamond deposits. Over the course of the last few years, explorationists have subdivided the diamond-bearing breccias into three separate facies. These facies are volcanic (pyroclastic), subvolcanic/intrusive breccia and hypabyssal facies. A variable, but distinctive proportion and composition of fragments and/or xenoliths characterizes each facies (Wilson 2004). The volcanic facies contains breccia, lapilli- and ash sized fragments and consists of medium to thickly bedded pyroclastic air-fall deposits. They are characterized by angular to sub-angular Archean supracrustal fragments, some hypabyssal fragments and rare lower crustal to upper mantle xenoliths. The subvolcanic/intrusive breccia facies are the most variable in texture and appearance. The rock is characterized by observed intrusive relationships, a high proportion of fragments and a close proximity to the volcanic facies. The fragment characterization is variable. The facies can contain all or some of the following fragment types: supracrustal fragments, crustal fragments and lower crustal to upper mantle xenoliths. This facies was once included with the lamprophyres but is now interpreted by various exploration companies as a debris flow. The hypabyssal facies hosts variable proportions (<25%), of sub-rounded to rounded mantle xenoliths, as well as minor proportions of gneiss and/or trondhjemite fragments. This facies also was originally classified as lamprophyre but is now considered to be part of the debris flows (Lefebvre et al. 2003). The brecciated unit(s) could also be a single or multiple diatreme(s) localized within Lalibert, Leclaire, Menzies and Musquash townships. A tectonic repetition of the diamond-bearing diatreme(s) has been achieved through the regional deformation events described previously. This regional deformation may help to explain the layered appearance observed at some of the diamond occurrences (J. Ayer, Ontario Geological Survey, personal communication 2006). Lamprophyre occurs as narrow dike-like intrusions or in bodies of indeterminate morphology cross-cutting or intercalated with local country rocks. Dikes range in width from 50 cm to 2 m and may display 2-3 cm offshoots in some locations. Contacts with the host rocks vary from sharp and straight to highly irregular. No variation in grain size within the dikes has been observed and the no variation in colour or mineralogy 86 �has been observed within country rock adjacent to the dikes. The lamprophyre is fine-grained, grey in colour and contains approximately 5-10% subrounded to subangular fragments. The fragment population is dominated by actinolite-rich monomineralic rocks, or by biotite-rich greenstone and hornblende-rich ultramafic rocks (Lefebvre 2004). Lamprophyre is differentiated from the breccia by: a lower clast content; a predominance of highly altered, coarse-grained actinolite fragments; scarcity of wall rock fragments; the rounded shape of the xenoliths; and the presence of a weaker fabric (Lefebvre 2004). The lamprophyre dikes post-date all other lithologies. The lamprophyre with indeterminate morphology predates most of the host lithologies since fragments of each are found within these lamprophyres. However, in some locations, lamprophyre fragments have been observed within intermediate and mafic intrusive rocks near the contact with lamprophyre (Lefebvre 2004). Fragments within lamprophyre dikes commonly have a biotite-rich rim enclosing the xenoliths and the fragments found in the lamprophyres of indeterminate morphology do not. The lamprophyres of indeterminate morphology also show positive relief of less severely weathered xenoliths more so than the lamprophyre dikes. Lastly, the lamprophyre dikes have a less variable fragment lithology (Lefebvre 2004). The lamprophyre is petrographically distinct from the breccia. It contains a lower abundance of clasts and fewer clast types. Unlike the breccia, the lamprophyre contains no juvenile magmatic material and oscillatory-zoned hornblende grains are rare. Detailed petrographic descriptions of the lamprophyre can be found in Lefebvre (2004). Relationship between diamond content and lithology Microdiamonds have been recovered from a wide array of breccias and lamprophyres in the Wawa area. Between 2003 and 2004, the Ontario Geological Survey (OGS) investigated a number of diamond occurrences during a mapping program conducted in Menzies and Musquash townships. Closer look at locations where bulk samples had been collected suggested that the bulk samples probably included more than one rock type. Diamonds were recovered from each of these bulk sample sites, but it would be difficult to establish from which of the rock types the diamonds were recovered. The OGS undertook a limited sampling program to further investigate the diamond content of specific lithologies (Vaillancourt et al. 2005a). The OGS collected three small samples (95.21 kg total weight) from three different lithological units at the Cristal and Genesis diamond occurrences. These occurrences are on the Festival Diamond Property owned by Pele Mountain Resources Inc. Results from this limited sampling program reinforce the observation that microdiamonds are not restricted to a single unit. Microdiamonds were recovered from heterolithic breccia, both with and without ultramafic magma pockets, and from a fragment-free ultramafic dike. There is the possibility, however, that the diamonds recovered from the ultramafic dike are xenocrysts derived from diamond-bearing host breccia (Vaillancourt et al. 2005a). The OGS concluded that the results from only three samples are not sufficient to draw irrefutable conclusions regarding the location of the microdiamonds. Collection and analysis of more, well constrained samples is necessary to further refine the diamond potential of specific host rocks. In 2005, Spider Resources Inc. and KWG Resources Inc. conducted a similar bedrock sampling program to investigate the diamond-bearing potential of specific lithologies. Separate representative samples (approximate weight 16 kg each) of the matrix, xenoliths and run of mill (ROM) portions of the bedrock were collected from the Wawa Diamond Project and were sent for caustic dissolution. The results are as follows: matrix sample returned 67 diamonds (0.008 total ct), xenolith sample returned 244 diamonds (0.051 total ct) and the ROM sample returned 86 diamonds (0.006 total ct) (Spider Resources Inc., press release, February 20, 2006). 87 �Geochronology In conjunction with recent geological mapping in the area, the Ontario Geological Survey has been conducting geochronological work to help understand the nature and timing of the diamondiferous units and their host rocks within the Michipicoten greenstone belt. A felsic volcanic horizon hosting diamondiferous units returned a 207Pb/206Pb age of 2701.4 ± 2.1 Ma. Maximum 207Pb/206Pb ages of 2685.1 ± 1.0 Ma and 2684.9 ± 1.4 Ma have been returned for diamondiferous lamprophyre dikes cutting the Catfish assemblage (2.7 Ga) intermediate to felsic metavolcanic rocks in Lalibert and Menzies townships (Ayer et al. 2003). A second sample of felsic lapilli tuff, part of the Catfish assemblage, adjacent to the Moet Occurrence contains zircons that returned a 207Pb/206Pb age of 2698.7 ± 1.1 Ma (Vaillancourt et al. 2004). A sample was collected from the diamondiferous breccia at the Moet Occurrence in order to determine the age of brecciation. Five zircons were analyzed. The three oldest ages are 2687 ± 2 Ma, 2683 ± 2 Ma and 2681 ± 2 Ma. The two youngest zircons cave results that precisely overlapped one another at 2679.2 ± 2.1 Ma (Vaillancourt et al. 2004). Since this is the youngest zircon age obtained from the breccia, it represents either the time of cystallization or emplacement of the body if the zircons are magmatic, or a maximum time of emplacement if the zircons are xenocrystic (Vaillancourt et al. 2005a). If the zircons are xenocrystic, this age must still be close to the time of breccia emplacement since a lamprophyre dike from the GQ property returned a date of 2673 ± 8 Ma from titanite (R. P. Sage, Ontario Geological Survey, unpublished 2000). The data indicate that the felsic metavolcanic rocks hosting the diamondiferous breccias are part of the Catfish assemblage. The maximum age for the diamondiferous breccias and the associated dikes is less than 2680 Ma. These absolute age constraints indicate that the breccias are not volcaniclastic units belonging to the Catfish assemblage (Vaillancourt et al. 2005a). Zircons from a sample of felsic lapilli tuff from northwestern Menzies township returned an age of 2736.0 ± 0.8 Ma which is taken to represent the age of eruption and crystallization of the tuff. This age clearly indicates that the volcanic package underlying the iron formation in the western part of Menzies Township is part of the Wawa assemblage (2.75 Ga) and brackets the uppermost part of the assemblage at 2736 Ma (Vaillancourt et al. 2005a). Geochemistry of the Diamond-bearing Rocks Both Williams (2002) and Lefebvre (2004) conclude that the whole rock major element geochemistry is consistent with a calc-alkaline classification for both the lamprophyres and the associated breccia. Both authors also noted that the compositions of chromite in the Wawa metavolcanic rocks are in the range typical for lamprophyres and dissimilar to those in kimberlites and lamproites. Whole rock geochemistry for the diamond-bearing rocks is tabulated in Sage (2000), Williams (2002), Lefebvre (2004), Stone and Semenyna (2004) and Vaillancourt et al. (2005c). Sage and Williams’ work is specific to the diamond-bearing and non diamond-bearing lamprophyres. Work by the other authors relates to both the breccia and the lamprophyres. Whole rock geochemistry for kimberlites of the Wawa area can be found in Kaminsky et al. (2002). Diamond characteristics Lefebvre (2004) undertook a study examining a parcel of 80 macrodiamonds recovered from the volcaniclastic breccia on the GQ Property. Results from this work are summarized below and in De Stefano et al. (2006). Additional work on the morphology of the Wawa diamonds can be found in Stone and Semenyna (2004). Stachel et al. (2004) summarize results of analysis conducted on diamonds from the Genesis and Cristal diamond occurrences presently held by Pele Mountain Resources Inc. 88 �Shape: The diamonds display a highly variable primary growth form. Most of the diamonds are either octahedral aggregates (44% of the population) or single octahedral crystals (26%). Single cubic and cubicoctahedral crystals and their aggregates, as well as macles form the remainder of the population. Fortyeight percent of the diamond population is single crystals and only 28 diamonds could be evaluated for crystal regularity. The majority of the diamonds also display some degree of distortion. Colour and transparency: The diamonds included in this study are classified into colourless, brown, grey, black, yellow and white. No pink, green, violet or blue diamonds were observed. The colour distribution within the population is: colourless (48%), heterogeneous (24%), yellow (11%), black (3%), brown (10%) and grey (3%). The heterogeneity in colour is observed only in aggregates. The diamond population consists of 48% transparent crystals, 25% translucent crystals, 14% opaque crystals and 14% is a combination of opaque and translucent crystals. Transparent crystals are typically colourless and also comprise a few yellow octahedral single crystal and coarse aggregates as well as macles. Translucent crystals comprise all possible primary crystal forms and colours. Opaque crystals are mostly fine-grained aggregates which have black body colouring. Resorption: Generally speaking, the diamonds have experienced low degrees of resorption. Only 21% of the diamond population displays extensive resorption. Some crystals (14%) exhibit non-uniform resorption where one part of the crystal is more strongly resorbed than another. Inclusions: Mineral inclusions were identified in 58% of the diamonds. Both primary and secondary inclusions were observed. The mineralogy of the recovered primary inclusions is listed in descending order of abundance: olivine (Fo92 and Fo89), clinopyroxene (omphacite), plagioclase (albite and An-rich), orthopyroxene (En93) and Fe-Ni sulphide (pentlandite). Cathodoluminescence: The relative abundances of cathodoluminscence (CL) colours for the Wawa diamonds are: orange-red (46%), yellow (28%), orange-green (10%), green (6%), and other non-uniform colours (10%). None of the 69 diamond examined displayed the more common blue CL. Impurities: Fournier Transform Infrared (FTR) spectrometry was used to investigate the nitrogen and aggregation states for 41 diamonds. The majority of the diamonds have low nitrogen contents, < 300 ppm. The diamonds show two modes of nitrogen aggregation suggesting mantle storage at 1100 - 1170° C. Diamonds from the Genesis occurrence are almost exclusively cubes including some fragmented, twinned and moderately resorbed cubes. Most of the crystals contain clouds. Fully transparent stones are dominantly brown although colourless stones also are common; one diamond was yellow in colour. Nitrogen concentrations range from below detection (<10 ppm) to 600 atomic ppm. Nitrogen aggregation is very low (Stachel et al. 2004). Diamonds from the Cristal occurrence range from un-resorbed octahedra to highly resorbed dodecahedra. Octahedral and weakly resorbed octahedral stones dominate the population. About 25% of the population are irregular crystals, macles are common (15%) and about 5% of the diamonds show cubo-octahedral growth. The stones fall into two dominant colour classifications, colourless and a range of brown colouration. Nitrogen contents range from <10 – 560 ppm, but with only one exception nitrogen is ≤170 ppm. Nitrogen aggregation varies between 0 and 97% B-centre. Olivine is the most common mineral inclusion, followed by pyrope garnet and Mg-chromite (Stachel et al. 2004). Origin of the Diamond Deposits Based on published data on the diamond-bearing rocks at Wawa and Cobalt, Wyman et al. (in press) suggest that the tectonic setting of the deposits and nature of the host rocks indicate that the diamonds may be derived from the asthenospheric wedge and subducted slab at shallow depths (100 – 160 km) rather than the deep keels of Archean cratons associated with traditional diamond deposit types. Models of lowtemperature Phanerozoic diamond formation in active subduction zones, or rapid uplift and emplacement of peridotite massif occurrences, can be adapted to the Archean deposits. The stability field of diamonds in 89 �most Phanerozoic scenarios may be too deep to be accessed by the lamprophyric magmas. Shallow subduction, as proposed for these occurrences of adakitic-type rocks in the Wawa subprovince, could generate two different diamond stability windows at sufficiently shallow depths to account for their presence in lamprophyric magmas. Wyman et al. (in press) states that any tectonic model for these Archean diamond occurrences must address several requirements. These requirements include 1. a deep source for oxidized metasomatic fluids that is activated prior to lamprophyre emplacement 2. a mechanism to isolate this isotopically aged and depleted source for tens or hundreds of millions of years until it is heated in the mantle during orogeny 3. a hybridized mantle source for primitive, hydrous, shoshonitic lamprophyres 4. sustained cold finer effect in the mantle to establish a shallow-mantle diamond stability window Two theories of diamond origin are postulated by De Stefano et al. (2006). Both a cratonic and orogenic model of diamond formation are discussed in an effort to rationalize the observed diamond characteristics. The authors conclude that neither model fully explains all of the observed characteristics. 90 �Field Trip Stops Field Trip Road Log Stop 1 2 3 4 5 6 7 8 Locality Intersection of Hwy 101 and Hwy 17 Take Highway 17 north km 0 Catfish Road forestry road – turn east 19.2 GQ Diamond Discovery Northern Sierra Minerals Corporation Area B Northern Sierra Minerals Corporation Engagement Zone 23.3 23.9 29.1 Return to Highway 17 reset odometer Drive north to access road – turn east Park Walk eastward along trail Moet Occurrence, Festival Property 0 6.2 6.6 7.6 7.6 Return to Highway 17 reset odometer Drive north, park on shoulder of highway Sandor Diamond Occurrence Continue north on Highway 17 to intersection of Highway 519, turn right 0 4.3 4.3 12.6 Safely turn in parking area and return to Intersection of Hwy 17 and 519 reset odometer Drive south on Highway 17, park on shoulder of road Dubreuilville Dike Continue south on Highway 17 Turn left into access road to gravel pit, park Walk south approximately 150 m Monchiquite Dike Drive north on Hwy 17 to Wawa Wawa Motor Hotel 0 3.1 3.1 44.9 45 53.2 STOP 1 - GQ Diamond Discovery Site Northern Sierra Minerals Corporation Area A UTM co-ordinates – 0665570E 5333291N NAD83 Several outcrops of diamondiferous breccia outcrop on the west side of a forestry road. This exposure is an example of the hypabyssal facies of the three identified diamond-bearing units. The rock cut displays the apparently conformable nature of these “lamprophyre” dikes. The most notable features of these outcrops are the actinolite-rich nature of the matrix and the presence of biotite-rich reaction rims around the xenoliths. It is frequently difficult to distinguish these dikes from the mafic to intermediate agglomeratic and tuffaceous host rocks. In the vicinity of the discovery area, located at the south end of Area A, the diamondiferous breccias are arrayed linearly along the logging road where the topography indicates a 5 – 10 m thick, northwest-trending dike (Cavey 2002). A compilation of the geology of the GQ Property is shown in figure 4. Figure 5 provides a compilation of the diamond occurrences on the GQ Property. 91 �Figure 4. Geological compilation map of the GQ Property, Musquash Township, Northern Sierra Minerals Corporation (Cavey 2004). 92 �Figure 5. Occurrences of diamondiferous bedrock on the GQ Property, Musquash Township, Northern Sierra Minerals Corporation (Cavey 2004). 93 �A relatively precise 207Pb/206Pb age of 2674 ± 8 Ma was returned from a sample collected in 2000 by R. P. Sage (R. P. Sage, Ontario Geological Survey, unpublished 2000). The sample analysed was titanite. The titanite grains are considered to be a primary mineral and not xenocrystic. In thin section, the rock is characterized by a green, medium-grained, granoblastic to decussate groundmass of actinolite, biotite, chlorite, plagioclase and accessory minerals. Amphibole or biotite macrocrysts up to 1 mm are common (Stone and Semenyna 2004). Local prospectors T. Nicholson, J. Robert and M. Tremblay made the discovery in the fall of 1999. The first two bedrock samples (63.4 kg and 70.5 kg) collected were processed by Kennecott Canada Exploration Inc. in their Thunder Bay laboratory. According to Kennecott’s report, the 63.4 kg sample yielded 45 diamonds, of which 10 were macro diamonds and 35 were microdiamonds. One of the macro diamonds measured 1.01 mm in one dimension. The 70.5 kg sample yielded 9 microdiamonds. All stones were white in colour and transparent in clarity. Duplicate samples were collected by Band-Ore Resources Ltd. in early 2000 and were processed at SGS Lakefield Research Limited. A 54.6 kg sample yielded 98 microdiamonds. A confirmation sample from the same area yielded 98 microdiamonds from a 54.6 kg sample. In 2000, Band-Ore Resources drilled 3 short holes (75 m total) at the discovery site. Table 1 details the diamond recovery results from the drill program. Only partial intervals from drill hole DDH GQ-00-3 were submitted for microdiamond recovery since portions of the core were used for thin sections, microprobe analysis and display purposes. In total 5 diamonds were recovered from DDH GQ-00-3, including one champagne coloured macro diamond and one white microdiamond from a sample weighing 7.5 kg. Drill sections for DD GQ-00-01 through 03 are shown on figure 6. Table 1. Diamond recovery results from 2000 Band-Ore Resources Ltd. diamond-drilling program Drill Hole DDH GQ-00-1 DDH GQ-00-2 Sample No. Sample Size (kg) No. Macro Diamonds No. Micro Diamonds Sample 1A 63.35 1 Sample 1B 30.17 10 Sample 1C 28.51 6 Sample 2A 37.52 1 30 Sample 2B 30.97 434 Sample 2C 28.51 30 results compiled from Band-Ore Resources Ltd. press releases 2000 To date, the discovery site has yielded 746 diamonds, including 15 macro diamonds, from sample material weighing 785 kg. The largest diamond recovered exceeds 1.0 mm in size and the majority of the stones are gem quality, white, clear and transparent. STOP 2 - Northern Sierra Minerals Corporation Area B UTM co-ordinates – 0665425E 5334748N NAD83 The exposure on the east side of the forestry road provides an excellent exposure of the subvolcanic/intrusive breccia facies. Subrounded to rounded xenoliths dominate the vertical exposure (Figure 7). Field relationships between the intrusive breccia and other heterolithic breccias can be observed in several outcrops along the road. Texturally, the subvolcanic/intrusive facies may resemble both the hypabyssal facies and the intrusive heterolithic breccias. The facies consists of mica and amphibole phenocrysts (<2mm) in a groundmass of mica, actinolitic amphibole and lesser albite, carbonate, sphene and oxides. Alteration includes variably chloritized mica while the other phenocrysts have been extensively altered to varying proportions of mica, albite and actinolitic amphibole. 94 �Figure 6. Drill hole sections GQ-00-01, GQ-00-02 and GQ-00-03, GQ Property, Musquash Township, Northern Sierra Minerals Corporation (Cavey 2002). 95 �Figure 6 cont’d. Drill hole sections GQ-00-01, GQ-00-02 and GQ-00-03, GQ Property, Musquash Township, Northern Sierra Minerals Corporation (Cavey 2002). 96 �Figure 7. Stop 2 - Northern Sierra Minerals Corporation Area B – rounded lower crustal to upper mantle xenoliths Stone and Semenyna (2004) completed a petrographic examination of one of the ultramafic xenoliths from this site. The sample was dominated by a coarse, radiating to decussate, clear to pale green amphibole of tremolitic to magnesium-rich actinolite composition. Carbonate occurs locally and biotite is concentrated at the rims of the xenolith. Band-Ore Resources Ltd. discovered area B in 2000. Thirty-three (33) reconnaissance samples were collected from this area and a total of 273 diamonds was recovered from 352 kg of material (Cavey 2002). One 24 kg sample returned 126 microdiamonds (1.37 mg total weight). No macro diamonds were recovered from Area B. Band-Ore Resources Ltd. completed only a reconnaissance sampling program, minor stripping and trenching in this area. The Barnett Zone lies approximately 1.6 km to the northwest. It was discovered in the fall of 2001 by Kennecott Canada Exploration Inc. who completed a limited program of mechanical stripping, channel sampling and washing of outcrops over the area. A total of 27 outcrop channel samples (270 kg) were collected. The channel samples returned 330 microdiamonds and 3 macro diamonds. A single 24 kg sample of heterolithic breccia from the Barnett Zone returned 3 macrodiamonds and 123 microdiamonds. A total of 273 diamonds (261 microdiamonds and 12 macrodiamonds) were recovered from 34 samples (352 kg) collected between September 2000 and July 2001 by Kennecott Canada Exploration Inc. in Area B (including the Barnett Zone). 97 �STOP 3 - Northern Sierra Minerals Corporation Engagement Zone Bulk Sample Site UTM co-ordinates – 0667760E 5336073N NAD83 The main outcrop exposure consists of medium to thickly bedded pyroclastic air-fall deposits and debris flows described as heterolithic tuff-breccias that grade upwards to lapilli-tuff and tuff. The groundmass of the extrusive phase is fine-grained with variable proportions of relatively small (<2mm) altered phenocrysts, including chloritized mica. The groundmass consists primarily of actinolitic amphibole with rare to minor mica or granular albite. The pyroclastic rocks are mineralogically and compositionally similar to intrusive varieties, but have a significantly higher proportion of mica phenocrysts. Both diamond-bearing intrusive and extrusive rocks host significant proportions of fragments derived from the local country rock. Rare to minor, deep crustal and upper mantle xenoliths, such as banded gneiss and extensively altered talcose ultramafic xenoliths are present. Fresh mantle rocks, such as lhertzolite, harzburgite and eclogite have not been identified. The matrix material is typically fine-grained, green, weakly foliated actinolite schist. Albite is present, although it is less abundant (<10%) than in the lamprophyre dikes and implies a more ultramafic composition for the breccia matrix than for the lamprophyre dikes. Titanite is fairly abundant and calcite, epidote, apatite and sulphide minerals occur locally. Macrocrysts of actinolite are commonly observed. Rare macrocrysts of amphibole also are observed and are frequently altered to actinolite. The actinolite macocrysts are probably metamorphic in origin, whereas the amphibole macrocrysts may represent accidental or cognate crystals (Stone and Semenyna 2004). The Engagement Zone has a minimum strike length of 335 m and a horizontal width in excess of 75m. The zone strikes northwesterly and has a shallow dip to the northeast. This zone may represent the southeast extension of Pele Mountain Resources Inc.’s Cristal diamond occurrence located approximately 2 km to the northwest. In 2003, a 0.72 carat macrodiamond was recovered from a bulk sample collected from the Cristal. The geology of the Engagement Zone and sample locations are shown in Figure 8. A simplified cross section of the Engagement Zone is shown in Figure 9. Band-Ore Resources Ltd. discovered this zone in January 2001. A 16 kg sample from a single angular boulder of diatreme breccia returned 128 microdiamonds. Four subsequent samples (96 kg total weight) returned 5045 microdiamonds and 65 macrodiamonds. In 2001 a mini-bulk sample weighing 12.5 tonnes was collected under the supervision of Kennecott Canada Exploration Inc. and shipped to the Saskatchewan Research Council. The largest diamond recovered from this sample was a 0.254 carat, broken, white octahedral stone. Two additional bulk samples were collected in 2003. A 22 tonne sample tested an area where 6 channel samples (63 kg total) recovered 1752 stones. A 20 tonne sample tested an area where 5 channel samples weighing 41.8 kg returned 552 stones. The results from these bulk samples are shown in Table 2. Table 2. Diamond results from Engagement Zone bulk samples (2001) Occurrence Sample Weight Sieve +1mm Sieve +2mm Sieve +3mm Sieve +5mm Sieve +6mm Engagement Zone East 22.1 tonnes 1 4 4 Engagement Zone West 20.4 tonnes 2 3 3 2 1 Total Diamonds Total Carat Weight 12 0.375 8 0.155 results compiled from Band-Ore Resources Ltd. press releases 2004 To date, exploration on the Engagement Zone has included an orientation geochemical survey, geological mapping, channel sampling, trenching, bulk sampling and a 9-hole (1775 m) diamond drilling program designed to test the strike continuation of the zone. The drilling program demonstrated that thick diamondiferous breccia deposits can feather out and thin to a few centimeters thickness (Cavey 2003). 98 �Figure 8. Stop 5 – Geology and sample locations at the Engagement Zone, GQ Property, Northern Sierra Minerals Corporation (Cavey 2002) 99 �Figure 9. Simplified cross section through the Engagement Zone looking northwest (Cavey 2004). STOP 4: - Moet Occurrence, Festival Property Pele Mountain Resources Inc. UTM co-ordinates – 0662709E 5338009N NAD83 The Moet occurrence is a large stripped outcrop that extends in a north-south orientation across the forestry road. The outcrop displays all three facies of diamond-bearing bedrock exposed over an area 500 m by 300 m. It is hosted within fine- to medium-grained mafic metavolcanic rocks with closely associated intermediate to felsic metavolcanic rocks and metasediments. The volcanic facies was found concentrated within a series of outcrop exposures along the west side of a 5-8 m north-trending ridge and the subvolcanic/intrusive breccias and hypabyssal rocks are present in several outcrops east of the volcanic facies. The subvolcanic/intrusive breccias are hosted in the metavolcanics and the volcanic facies appears to overlie these rocks, and are in turn overlain by intermediate to felsic metavolcanic rocks. The fragments 100 �within the volcanic facies consist of country rock fragments with lesser crustal fragments and lower crust to upper mantle xenoliths. The hypabssyal facies displays primarily mantle xenoliths (Walker 2003). A detailed geological map of the Moet locality is shown in Figure 10. Figure 10. Detailed geological map of the Moet Occurrence, Festival Property, Pele Mountain Resources Inc. (Kjarsgaard et al. 2003). The access trail to the outcrop passes through a sequence of intermediate to felsic tuffs and tuffaceous breccias. The composition of the fragments and that of the matrix are highly variable from felsic to intermediate and a combination of felsic matrix with intermediate fragments and vice versa is not uncommon. At a distance of approximately 1km, the trail passes through a sequence of mafic to intermediate tuffs and lapilli tuffs (Vaillancourt et al. 2005b). An age date of 2698 ±1 Ma was returned from sample of the felsic metavolcanic rocks at the occurrence. A sample of the breccia returned an age date of 2680 Ma (J. Ayer, Ontario Geological Survey, personal communication, 2004). Discovered by Pele Mountain Resources Inc. in 2001, the occurrence initially gained interest because the breccia has a size distribution of diamonds that includes coarser sized diamonds from relatively small samples. For example, an 8 kg sample collected in 2001 recovered a total of 9 diamonds, 4 of which were in the +600 mesh fraction. Diamonds are consistently recovered from both the volcanic and subvolcanic facies at this showing. A summary of the results of the 2001 and 2002 sampling of the occurrence by Pele Mountain Resources Inc. is found in Table 3. 101 �Table 3. Summary of the diamond results from the 2001-02 sampling of the Moet Occurrence Facies Weight (kg) <425 mesh >425 mesh volcanic volcanic hypabyssal hypabyssal volcanic volcanic hypabyssal subvolcanic breccia 117 16 24 8.6 9.3 14.1 9.3 32 48 2 0 0 7 12 0 30 0 1 0 0 0 5 0 1 total diamonds 48 3 0 0 7 17 0 31 results compiled from Pele Mountain Resources Inc. press releases 2002 Further exploration was conducted at the occurrence in 2003 when DeBeers Canada Exploration Inc. completed a detailed airborne geophysical survey, stripping, mapping and sampling program. A 47.8tonne bulk sample from the site returned 5 diamonds with a total carat weight of 0.13. All diamonds were recovered from the +1 to +3 sieve class screens (Pele Mountain Resources Inc., press release, March 17, 2004). STOP 5: - Sandor Diamond Occurrence Spider Resources Inc. & KWG Resources Inc. UTM co-ordinates – 0659805E 5342191N NAD83 The Sandor occurrence is the first confirmed occurrence of diamonds in bedrock in the Wawa area. The occurrence is located in a 4 m high road cut on the east side of the Trans Canada Highway (Highway 17). The dike is approximately 5 m wide, steeply dipping and strikes roughly parallel to the regional schistosity at 120°. The dark, greenish-grey rock weathers olive grey, is highly fractured, moderately carbonatized and is non-magnetic. It is composed of up to 40% actinolite replaced mantle xenoliths and supracrustal xenoliths. Towards the margins of the dike xenoliths are less common and the rock grades into an adjacent micaceous dike. Only remnants of the dike remain in situ. The dike is hosted by gabbros and intermediate to felsic crystal tuffs. A (308.6 kg) sample of the dike, collected by Spider Resources Inc. in 1997, returned a total of 97 diamonds comprising 1 commercial stone, 13 macrodiamonds and 83 microdiamonds. A short walk into the forest from the top of the outcrop leads to a larger stripped area where field relationships between the host gabbro and the dike can be observed. A second, xenolith-bearing dike (occurrence LAL-3) is located at the north end of the outcrop. This dike is 2 m wide and closely resembles the Sandor occurrence. A 34.6 kg sample of this dike contained 1 microdiamond. A detailed geological map of the Sandor occurrence is shown in Figure 11. A compilation map of the geology of the Spider Resources Inc. and KWG Resources Inc. property is shown in Figure 12. Using normative mineralogy, Sage (2000) concluded that this dike should be classified as a spessartite. A spessartite is defined as a lamprophyre composed of phenocrysts of green hornblende or clinopyroxene in a groundmass of sodic plagioclase with accessory olivine, biotite, apatite and opaque oxides. Titanite and rutile from the matrix of the Sandor dike returned an age of 2703 ± 42 Ma (Sage 2000). The date is interpreted to be a minimum age of intrusion. Subsequent dating of a zircon from a gneissic xenolith from the Sandor dike returned a 207Pb/206Pb age of 2684.9 ± 1.4 Ma (Ketchum, Kamo and Davis 2003). Spider Resources Inc. and joint venture partner KWG Resources Inc. have taken 3 mini-bulk samples along the 1 km strike extent of the dike since the fall of 2001. The three bulk samples had a combined weight of 7.61 tonnes and returned 11 commercial stones and 9 macro diamonds. These samples were tested only for macro diamond and commercial diamond content (Spider Resources Inc., press releases, February and March 2002). 102 �Figure 11. Detailed geology of the Sandor Occurrence, Spider Resources Inc. & KWG Resources Inc. (Kjarsgaard et al. 2003). Figure 12. Generalized geology of Wawa Project, Spider Resources Inc. and KWG Resources Inc. (Spider Resources Inc. and KWG Resources Inc., CD-ROM presentation, update February 24, 2004) 103 �STOP 6: - Dubreuilville Dike - Xenolith-rich lamprophyre Spider Resources Inc. & KWG Resources Inc. UTM co-ordinates – 0657251E 5347023N NAD83 The dike outcrops in a 1.5 m high exposure on the west side of the Trans-Canada Highway (Highway 17). In the area between the former Magpie Iron Mine in Leclaire Township to approximately the eastern contact of the Dickenson Lake Stock, there are a number of exposures of unusual dike rocks. These rocks are characterized by prominent round, to elliptical inclusions of actinolite or actinolite plus talc. Xenoliths are altered to fine-grained actinolite with or without talc, and some display zoning from talc core to an actinolite rim. The actinolite inclusions may consist of prismatic green crystals as large as 8 cm in length, which may be randomly oriented or radiating inward towards the core. The inclusions containing talc consist of a talc core with the prismatic to acicular actinolite projecting radially inward towards the core. The xenoliths are believed to represent at least two original mafic compositions which are likely to be originally of lower crust origin or deeper. A weakly developed regional schistosity crosses the dike implying an Archean age (Sage 1993, 2000). An example of one of the lower crustal to upper mantle xenoliths is shown in Figure 13. Figure 13. Lower crustal to upper mantle xenolith, Dubreuilville Dike Stop 6, Highway 17 North. This dike may be the one described by Higgins (1986). He reported that the dike consists of 60% euhedral amphibole, 20% biotite replacing amphibole and 15%plagioclase. Minor sphene and opaque minerals are present and chromite is reported from the core of the talc-bearing clasts. The bulk composition of the nodules is reportedly pyroxenite, but Higgins did not indicate whether their source was mantle or crust. The rounded outline of the clasts may reflect magmatic erosion during transport and the present mineralogy is the product of regional metamorphism (Higgins 1986). 104 �STOP 7: - Monchiquite Dike Michipicoten Post Provincial Park UTM co-ordinates – 0663600E 5309223N NAD83 On the west side of Trans Canada Highway (Highway 17) two, 1-metre wide and one, 3-metre wide monchiquite dikes outcrop at the base of a 3 m high outcrop. The dikes are parallel, strike 060° and crosscut a portion of the Mission Stock. A 4 m wide, east-trending microsyenite dike also can be observed cutting the granodiorite and the monchiquite at this location. The north-northwest-trending Trembley Fault cuts across the north end of the outcrop (Massey 1985). These lamprophyres are most commonly seen in road cuts along the highway and along the shore of Lake Superior. Typically, they are black when fresh and weather to an orange-brown colour. The dikes are usually narrow and are typically less than 1 m wide. Biotite-phyric and olivine-phyric varieties are most common, but pyroxene and feldspar phenocrysts also are observed. Often the dikes show evidence of multiple and composite intrusion, sometimes with tin screens of country rock trapped within the dikes. Lamprophyre dikes south of the Michipicoten River typically have a northeast trend (Massey 1985). Lamprophyre dikes are commonly found cross cutting all lithologies south of the Wawa – Hawk – Manitowik Lakes Fault. Because of their ease of weathering they are infrequently seen in outcrop. However, the dikes are frequently found in underground mine workings. These dikes have a number of macrocrystic resemblances to the dikes seen here. As early as 1927, geologists working in the area had noted a similarity between these lamprophyre dikes and kimberlite (Gledhill 1927). Recent petrographic and mineralogical studies on several dikes in both McMurray and Lendrum townships have suggested that some of the geochemistry falls within the classification of type II kimberlites (orangeites) (Barnett 2001). To date, no diamonds have been recovered from these dikes. These lamprophyre dikes are probably Proterozoic in age and are interpreted to represent Proterozoic alkalic magma emplacement into structures related to the Kapuskasing Structural Zone, perhaps consanguineous with the nearby Firesand Creek carbonatite complex. Rocks from the Firesand Creek carbonatite complex have a U-Pb date of 1078 ± 2.4 Ma (Sage 2000). STOP 8: - Diamondiferous Conglomerates north of Wawa This unit, referred to as the Leadbetter Conglomerate, is the main diamond-bearing unit on the property; it is comprised of clast supported plymictic conglomerate that is poorly sorted. Predominantly massive, the unit is rarely stratified and when it is, the bedding varies in thickness from 10-30 centimetres to over 1 meter with no obvious graded bedding. The sdeiments range in clast size from mud-sized particles to bould-sized clasts which are sub-rounded to rounded, flattened and highly stretched. The matrix is very fine-grained and chlorite rich with a minor amount of quartzofeldspathic material. The clasts predominantly consists of volcanics, with a minor amount of non-volcanics (siltstone and chert). The conglomerate appears to pinch out to the northwest in the project area, but thickens dramatically in the central andeaster parts of the project area where surface exposures give an indicated thickness of up to 300 m. Deformation of this unit has resulted in stretched or elongated clasts. The degree of stretching is variable suggesting that there are local zones of intense deformation. The diamondiferous conglomerate is interpreted to represent valley fill deposits in the proximal reaches of an alluvial fan comples (Wendland, 2009) with the matrix supported cobble and boulder conglomerates forming a debris flow sequence. An ultramafic component in the source for the matrix component of the of the conglomerate is indicated by immobile element geochemistry. The fact that a distinct suite of diamond indicator minerals consisting of chromite, olivine, and rare ilmenite, pyrope garnet and chrome diaopside, have been recovered from the conglomerate makes it unique incomparison to the sparsely diamondiferous Wawa volcaniclastics located to the north which contain a paucity of conventional indicator minerals (Verley 2009). 105 �STOP 9: - Contemplation of the rocks on the fireplace at the Wawa Motor Hotel Acknowledgments The author would like to thank Wayne O’Connor, Northern Sierra Minerals Corporation, Al Shefsky, President, Pele Mountain Resources Inc. and Neil Novak, Vice-President of Exploration, Spider Resources Inc. for permission to access properties described in the field guide. Editorial comments by R. P. Sage also were appreciated during the preparation of the field guide. My thanks also go to Glenn Seim for his technical assistance in the preparation of the guidebook. 106 �Bibliography Arias, Z. G. 1996. Structural evolution of the central Michipicoten greenstone belt, Superior Province, Wawa, Ontario, Canada; unpublished MSc thesis, Queen’s University, Kingston, Ontario, 140p. Arias, Z. G. and Helmstaedt, H. 1990. Structural evolution of the Michipicoten (Wawa) greenstone belt, Superior Province: Evidence for an Archean fold and thrust belt; in Geoscience Research Grant Program, Summary of Research 1989-1990, Ontario Geological Survey, Miscellaneous Paper 150, p. 107-114. Ayer, J. A., Conceição, R. V., Ketchum, J. W. F., Sage, R. P., Semenyna, L. and Wyman, D. A. 2003. The timing and petrogenesis of diamondiferous lamprophyres in the Michipicoten and Abitibi greenstone belts; in Summary of Field Work and Other Activities 2003, Ontario Geological Survey, Open File Report 6120, p. 10-1 to 10-9. Barnett, R. L. 2001. Fletch Kimberlite correspondence, Report of Work, Matchinameigus and Fletch Properties, Dolson & Echum Townships, Sault Ste Marie Mining Division, Northern Ontario; Timmins Resident Geologist’s Office, Dolson Township, assessment file WP–Dolson–4, unpaginated. Cavey, G. 2002. Geological report on the Wawa properties for Band-Ore Resources Ltd. Sault Ste Marie Mining Division, Ontario; unpublished NI 43-101 report, 50p. Cavey, G. 2003. Updated geological report on the Wawa property for Band-Ore Resources Ltd. Sault Ste Marie Mining Division, Ontario; unpublished NI 43-101 report, 35p. Cavey, G. 2004. Summary geological report on the GQ Property for Band-Ore Resources Ltd. Sault Ste Marie Mining Division, Ontario; unpublished NI 43-101 report, 39p. Corfu, F. and Sage R. 1987. A precise U-Pb zircon age for a trondhjemite clast in Doré conglomerate, Wawa, Ontario; in Proceedings and Abstracts, Institute on Lake Superior Geology Annual Meeting, v. 33, p. 18. Corfu, F. and Sage, R. 1992. U-Pb age constraints for deposition of clastic metasedimentary rocks and late-tectonic plutonism, Michipicoten belt, Superior Province; Canadian Journal of Earth Sciences, v. 29, p. 1640-1651. De Stefano, A., Lefebvre, N. and Kopylova, M. 2006. Enigmatic diamonds in Archean calc-alkaline lamprophyres of Wawa, southern Ontario, Canada; Contributions to Mineralogy and Petrology, on-line edition, January 5, 2006. Gledhill, T. L. 1927. Michipicoten Gold Area, District of Algoma; Ontario Department of Mines, Annual Report, 1927, v. 36, pt. 2, p. 1-49. Higgins, M. D. 1986. Nodule-bearing spessartite (lamprophyre) dike near Wawa, Northern Ontario; in Program with Abstracts GACMAC-GCU-AGC-AMC-UCG, vol. 11, p. 81. Kaminsky, F. V., Sablukov, S. M., Sablukova, L. I, Shchukin, V. S. and Canil, D. 2002. Kimberlites from the Wawa area; Canadian Journal of Earth Sciences, v. 39, p. 1819-1838. Ketchum, J., Kamo, S. and Davis, D. 2003. U-Pb ages from the Superior and Grenville Provinces of Ontario; unpublished report of the Jack Satterly Geochronology Laboratory, Toronto. Kjarsgaard, B. A., McClenaghan, M. B., Boucher, D. R. and Kivi, K. 2003. Kimberlites and ultrabasic rocks of the Wawa, Chapleau, Kirkland Lake and Lake Timiskaming Areas; in VIIIth International Kimberlite Conference, Northern Ontario Field Trip Guidebook, B. A. Kjarsgaard (ed) p. 1-37. Lefebvre, N. S. 2004. Petrology, volcanology, and diamonds of Archean calc-alkaline lamprophyres, Wawa, Ontario, Canada; unpublished MSc thesis, The University of British Columbia, Vancouver, British Columbia, 265 p. Lefebvre, N., Kopylova, M., Kivi, K. and Barnett, R. 2003. Diamondiferous volcaniclastic debris flows of Wawa, Ontario Canada; long abstract prepared for VIIIth International Kimberlite Conference, Victoria, British Columbia, Canada. Massey, N. W. D.1985. Geology of the Mishewawa Lake Area, District of Algoma; Ontario Geological Survey, Open File Report 5532, 167p. Morris, T. F. 1999. Overburden as a media for kimberlite, base metal and gold exploration, Wawa region, northeastern Ontario; Geological Association of Canada – Mineralogical Association of Canada, Joint Annual Meeting, GAC-MAC Sudbury 1999, Field Trip B6 guidebook, 67p. Morris, T. F. 2001. Quaternary geology of the Wawa Area, Northeastern Ontario; Ontario Geological Survey, Open File Report 6055, 67p. 107 �Sage, R. P. 1993. Geology of Killins, Knicely and Lalibert townships, District of Algoma; Ontario Geological Survey, Open File Report 5589, 141p. Sage, R. P. 1994. Geology of the Michipicoten greenstone belt; Ontario Geological Survey, Open File Report 5888, 592p. Sage, R. P. 2000. The “Sandor” diamond occurrence, Michipicoten greenstone belt, Wawa, Ontario: A preliminary study; Ontario Geological Survey, Open File Report 6016, 49p. Sage, R. P., England, D., Calvert, T., Oudkerk, G., Worona, R. and Kusciusko, R. 1982. Precambrian geology of Musquash Township, Algoma District; Ontario Geological Survey, Map P.2556, scale 1:15 840. Stachel, T., Blackburn, L., Kurszlaukis, S., Barton, E. and Walker, E. C. 2004. Diamonds from the Cristal and Genesis volcanics, Wawa area, Ontario; in Abstracts of Talks & Posters, 32nd Annual Yellowknife Geoscience Forum, 16-18 November 2004, p. 74-75. Stone, D. and Semenyna, L. 2004. Petrography, chemistry and diamond characteristics of heterolithic breccia and lamprophyre dikes at Wawa, Ontario; Ontario Geological Survey, Open File Report 6134, 39p. Stott, G. M., Ayer, J. A., Wilson A. C. and Grabowski, G. P. B. 2002. Are Neoarchean diamond-bearing breccias in the Wawa area related to late-orogenic alkalic and “sanukitoid” intrusions?; in Summary of Field Work and Other Activities 2002, Ontario Geological Survey, Open File Report 6100, p. 9-1 to 9-10. Turek, A., Sage, R. P. and Van Schmus, W. R. 1992. Advances in the U-Pb ziron geochronology of the Michipicoten greenstone belt near Wawa, Ontario; Canadian Journal of Earth Sciences, v. 27, p. 649-656. Vaillancourt, C., Ayer, J. A., Zubowski, S. M. and Kamo, S. L. 2004. Synthesis and timing of Archean geology and diamond-bearing rocks in the Michipicoten Greenstone Belt: Menzies and Musquash townships; in Summary of Field Work and Other Activities 2004, Ontario Geological Survey, Open File Report 6145, p. 6-1 to 6-9. Vaillancourt, C., Ayer, J. A. and Hamilton, M. A. 2005a. Synthesis of Archean geology and diamond-bearing rocks in the Michipicoten greenstone belt: Results from microdiamond extraction and geochronological analyses; in Summary of Field Work and Other Activities 2005, Ontario Geological Survey, Open File Report 6172, p. 8-1 to 8-11. Vaillancourt, C., Dessureau, G. R. and Zubowski, S. M. 2005b. Precambrian geology of Menzies Township; Ontario Geological Survey, Preliminary Map P.3366, scale 1:20 000 Vaillancourt, C., Zubowski, S. M. and Dessureau, G. R. 2005c. Lithogeochemical data and field photographs for the Wawa area: Menzies and Musquash Townships; Ontario Geological Survey, Miscellaneous Release – Data 151. Verley, C. G.. 2009. 2009 Update of Activities on the Leadbetter Diamond Project; National Instrument-43-101, for Dianor Resources Inc. Ontario, 78p. Walker, E. C. 2002. Diamond deposits of the Festival Property, Wawa, Ontario; Report prepared for Pele Mountain Resources Inc. under National Instrument 43-101, 41p. Walker, E. C. 2003. Diamond deposits of the Festival Property, Wawa, Ontario; Report prepared for Pele Mountain Resources Inc. under National Instrument 43-101, 39p. Wendland, C. 2010. Diamondiferous Mass-Flow and Traction Currenct Deposits in a Neoarchean Fan Delta, Wawa Area, Superior Province unpublished MSc thesis, Lakehead University, 102 p. Williams, F. 2002. Diamonds in late Archean calc-alkaline lamprophyres Ontario, Canada: Origins and implications; unpublished BSc thesis, University of Sydney, Sydney, Australia, 82 p. Wilson, A. C. 2004. Diamond exploration targets, Michipicoten greenstone belt; Canadian Institute of Mining Bulletin, v. 97, no. 1077, p. 41-46. Wyman, D. A., Ayer, J. A., Conceição, R. V., and Sage, R. P. in press. Mantle processes in an Archean orogen: evidence from 2.67 Ga diamond-bearing lamprophyres and xenoliths, Lithos. 108 �Field Trip 3 Geology of the Wawa gold project Jean-Francios Montreuil, Quentin Yarie, and Conrad Dix Red Pine Exploration Ltd, Toronto, Ontario 109 �1.0 Introduction The fieldtrip on the Wawa Gold project combines an overview of the exploration and mining history in the Wawa Gold Camp and of the geology of the various mineralized structures of the property. The field trip is centered on the historic Surluga Road, along which most of the historic gold mines of the Wawa Gold Camp were exploited. The Wawa Gold Project is located 2 km east of the town of Wawa in Ontario and has a long exploration history that begun in the 1860s and has been discontinuously explored and developed since. This long period of activity resulted in the exploitation of 9 gold mines with preserved records of production (Fig. 1-1 and Table 1-1; Rupert, 1997); as well as many other small-scale mining operations without production records and the digging of numerous pits and the sinking of many shafts (Abandoned Mines Information System Database, 2017). The main gold concentration defined so far on the Wawa Gold Project is the Surluga Deposit a NI 43-101 compliant inferred resource of 1,088,000 ounces @ 1.71 g/t gold (0.5 g/t gold cut-off; Ronacher et al., 2015; Table 1-2). The inferred Surluga Deposit resource is entirely hosted within a brittle-ductile shear zone named the Jubilee Shear Zone. Recently discovered gold zones not identified in the immediate vicinity of the inferred resource are not incorporated in the current resource model. Table 1-1 Historic gold mine and gold production that were active on the Wawa Gold Project Mine Tonnes Milled Gold Grade (g/t) Gold Recovered (Ounces) Mariposa Grace+Darwin Parkhill Van Sickle Cooper Jubilee Minto Surluga 8 41,302 114,096 8,372 4,435 107,930 57,335 86,082 72.99 13.27 14.81 6.34 11.42 4.29 12.56 3.12 19 17634 54298 1710 1627 Totals 419,560 9.04 120,093 Production period(s) 1902-1903 1900-1903, 1907-1910, 1934-1937 1929-1938 1933-1936 1897-1900, 1926-1939 36178 8626 1964-1969 Table 1-2 Mineral resource statement* of the Surluga Deposit effective May 26th, 2015 estimated by SRK consulting (Canada) Inc. Resource Category Cut-off Gold (g/t) Quantity (‘000 t) Grade Gold (g/t) Contained Metal Gold (‘000 oz) Inferred** Pit-Constrained 0.4 10,239 2.05 676 Outside Pit-constrained 0.4 8,630 1.07 298 Underground 2.5 955 3.73 114 Total 0.5 19,824 1.71 1,088 * Mineral resources are not mineral reserves and have not demonstrated economic viability. All figures are rounded to reflect the relative accuracy of the estimate. Composites have been capped where appropriate. ** Open pit mineral resources are reported at a cut-off grade of 0.40 g/t gold in relation with a conceptual pit shell constructed by SRK. Underground mineral resources include classified modelled blocks below the conceptual pit shell and above a cut-off grade of 2.50 g/t gold. Cut-off grades are based on a gold price of US$1,250 per ounce, a gold recovery of 95 percent and a $US:$CAD exchange rate of 0.95. 110 �Figure 1-1 - Map showing the main historic mines as well as some of the main shafts and pits of the Wawa Gold Project 2.0 Overview of the property history 2.1 Discovery period - 1897 to 1910 The Wawa area has been explored for gold since the 1860s and gold was first discovered by William Teddy's wife in 1897 at Mackay Point along the south shore of Wawa Lake (Rupert, 1997; Frey, 1987). A staking rush followed the discovery and benefited from the change in claim staking adopted by the Ontario Government to encourage staking in 1895 (MacMillan and Rupert, 1990). This early rush period resulted in multiple discoveries. Attempts to produce gold from bedrock started in 1897 with the sinking of many shafts and the digging of many test pits throughout the property. Between 1897 and 1898, gold was identified on the Jubilee Shear Zone and a 103-foot shaft was sunk by the Great Northern Mining Company Ltd. in the sericite schists 111 �characteristic of the structure west of Jubilee Lake (Sage, 1994). Gold values encountered in that shaft were described as very low and the shaft was abandoned. In 1897, the Minto Mine was discovered by S. Berailldt who sold it to D. Tisdale. A 130-foot inclined shaft was sunk and the Minto mine operated until 1900. The Hornblende Shear Zone, located 300 metres west of the Jubilee Shear Zone, was discovered in 1899 by Mr. Peter Nissen. Two inclined shafts of 22 and 32 feet were sunk in the structure and in 1900, a test mill was constructed near Hornblende Lake by the Hornblende Mining Company. In 1902 and 1903, the Mariposa Gold Company sunk the 208-foot Mariposa shaft inclined at 80°NE in the footwall of the Mariposa Vein with two drifted levels at 100 and 200 feet (Sage, 1994). Gold production in the discovery period started on a larger scale in 1900 following the discovery of the Grace vein. The Algoma Commercial Company sunk a 304-foot shaft and extracted 6,097 tons of ore between 1900 and 1903 after which, commercial gold production ceased (Sage, 1994). The Lepage Gold Mining Company resumed production in the Grace Mine between 1907 and 1910 and produced 4,260 tons of ore (Fig. 2-1). Figure 2-1 - Grace Mine in 1908 Source: http://www.michiwawa.ca/albums/wawa/goldfields/goldfield_one.htm 2.2 Peak of mining activity - 1925 - 1938 During the period between 1910 and 1925, the Wawa Gold Project went into an exploration and production hiatus, with many land transactions between different parties and brief periods of field activity (Sage, 1994). The period extending from the mid-late 1920s to the late 1930s then saw the peak of mining activity on the property, with several mines in operation. By the late 1930s, 15 mines produced gold in the Wawa area (Frey, 1987). On the Wawa Gold Project, production records exist for eight of the mines that produced during that period (Cooper, Minto, Jubilee, Parkhill, Grace+Darwin, Mariposa and Van Sickle; Figure 1-1; Tables 1-1). The Cora vein, located in the Jubilee Shear Zone, was also briefly mined in the 112 �Cora shaft in 1927. The other larger mine from that period that produced gold from the Jubilee Shear Zone was the Jubilee Mine that produced 107,930 tonnes @ 4.29 g/t gold. The largest producer of that period was the Parkhill Mine that produced 54,298 ounces of gold from 114,096 tonnes @ 14.81 g/t gold between 1929 and 1938 (Table 1-1; Figure 2-2). Figure 2-2 - Mining town of Parkhill inhabited during the operation of the Parkhill Mine Source: http://www.michiwawa.ca/albums/wawa/goldfields/goldfield_one.htm 2.3 Surluga Mine discovery and first mining operation - 1960 to 1980 Exploration and development activity substantially slowed down in the 1940s and the 1950s, but resumed in 1960 when Tom Surluga and W.D. Sutherland drilled 25 holes between 1960 and 1961 that resulted in the discovery of the Surluga Deposit (Table 5-7; Sage, 1994). In 1962, Sutherland and company formed Surluga Gold Mines Limited to continue the development of the Surluga Deposit (Table 5-7). Between 1964 and 1968, Surluga Gold Mines Limited sank a 950-foot shaft with 7 levels spaced by 150 feet that formed the Surluga mine. In 1967, Surluga Gold Mines Limited constructed a 750 ton per day mill on the property and started an extensive underground drilling program (Amalgamated reports, 41N15NE0036). The mill was in operation in 1968 and 1969 but, inadequate definition of reserves in the mine before its construction precluded its optimal operation. The mill was shut down in 1969 and underground development continued between 1969 and 1971 with the objective of defining enough reserves to operate the mill (Amalgamated reports, 41N15NE0036). In 1973, under strong debt pressure, the Surluga Gold Mine company was re-organized as Pursides Gold Mines that conducted an extensive exploration program between 1973 and 1975 (Amalgamated reports, 41N15NE0036). Pursides Gold Mines ultimately went bankrupt and was re-organized as Citadel Gold Mines Inc. ("Citadel") in 1980. 2.5 Second development of the Surluga Mine by Citadel Gold Mines - 1986 to 1991 In 1986, before additional reserves were defined and the exploration model updated, the Surluga mine was dewatered, the Surluga mine shaft refurbished, the mill reconstructed, and a 3 year program of surface and underground drilling and geological mapping started on the Surluga Deposit (Rupert, 1997). One exploration success is the discovery in 1989 of the Old Tom zone in the southernmost part of the 113 �Surluga Deposit. During that period, Citadel also undertook an extensive exploration program of its property to find additional gold to feed the newly refurbished mill (Rupert, 1989a, b; Rupert and Leroy, 1989). This included diamond drilling of Root and Cooper-Ganley vein systems, stripping, trenching, channel sampling and geological mapping, as well as many airborne and ground geophysical surveys. In 1987, Citadel purchased from Duraine, the Parkhill and Grace-Darwin Mine properties (Rupert, 1997). Development and exploration in the Surluga Deposit however stopped again in 1989 because of the mill inefficiency, lack of defined reserves to feed the mill, the un-optimized design of the mine, including the difficulties of mechanizing production and problems with dilution control because of the cryptic boundaries of the high-grade zones (E. Hoffman, pers. comm.). Figure 2-3 Surluga mine in 2002 before site reclamation in the closure plan Source: http://www.michiwawa.ca/albums/wawa/goldfields/goldfield_one.htm 2.6 Change of the exploration model for the Surluga Deposit - 1990 to 1996 In 1996, while reviewing the available data on the Surluga Deposit, Bowdidge (1996) postulated that the Jubilee Shear Zone is a large-scale structure up to 150 feet thick and contains widespread low grade mineralization; and evaluated that a substantial resource of low-grade mineralization may exists in structure. This evaluation changed the perception of the property and forms the basis of the exploration model now used by Red Pine Exploration. 114 �3.0 Geology of the Wawa Gold Project 3.1 Regional Geology The Wawa Gold Project is in the southern part of the Michipicoten greenstone belt, one of two greenstone belts that form the Wawa Subprovince of the Superior Province (Fig. 2-1). The Michipicoten greenstone belt was formed by three cycles of mafic to felsic volcanism associated with concomitant subvolcanic intrusions (Sage, 1994). Zircon U-Pb ages date volcanic Cycle 1 to 2.9 Ga, volcanic Cycle 2 to 2.75 Ga, and volcanic Cycle 3 to 2.7 Ga (Turek et al., 1992; Sage, 1994). In the southern part of the Michipicoten greenstone belt, two large intrusive complexes were emplaced during cycles 1 (Hawk Lake Granitic Complex) and 2 (Jubilee Stock). These syn-volcanic intrusions have been interpreted to delineate the centres of calderas and to be the intrusive equivalent of their host felsic to intermediate volcanic rocks (Sage, 1994). Like other greenstone belts within the Superior Province, the Michipicoten greenstone belt contains basaltic to komatiitic volcanic rocks. Throughout most of the Michipicoten Greenstone Belt, the hiatus between volcanic Cycles 2 and 3 was marked by the deposition of large banded iron formations. In the Wawa area, from their discovery in 1898 to the closure of the last iron mine of the region in 1998, these iron formations were explored and mined over a 100 year period and were at one point the largest source of iron in Canada. All the rocks of the Michipicoten greenstone belt are metamorphosed at greenschist facies and its volcano-plutonic sequences have been repeatedly deformed and folded. Post-Archean magmatism includes diabase dikes and the emplacement of the Firesand River Carbonatite intruded along the Wawa-Hawk Lake-Manitowik Lake Fault System. The Wawa Gold Project is located within the southern part of the Michipicoten greenstone belt (Sherman, 2005). A prominent structure in the southern Michipicoten greenstone belt is the Wawa-Hawk Lake-Manitowik Lake Fault System, which defines the boundary between a lamprophyre-rich domain to the south and lamprophyre-free domain to the north (Figure 2-1). The emplacement of the Firesand River Carbonatite along the Wawa-Hawk Lake-Manitowik Lake Fault System suggests that the fault is deep-seated, whereas the location of the Jubilee Stock and Hawk Granite Complex along the fault indicate that it may follow an older structure active during the formation of the greenstone belt. 115 �Figure 3-1 - Regional Geology of the Michipicoten Greenstone Belt and location of the Wawa Gold Project within the belt Geology from Ontario Geological Survey (2011) open file MRD 126; Showing and deposit locations from Ontario Geological Survey (2014) Mineral Deposit Inventory 3.2 Property Geology The Wawa Gold Project occurs in a mafic to felsic volcano-plutonic sequence formed during the second cycle of volcanism in the Michipicoten Greenstone Belt (Sage, 1994). 3.2.1 Jubilee Stock The main intrusive complex of the property is named the Jubilee Stock that represents a key geological unit of the property for gold mineralization as all the largest historic mines (Darwin-Grace, Jubilee, Minto, Surluga, Parkhill) and most of the known gold showings are located within or at the margins of the Jubilee Stock (Fig. 3-2). The compositional and geometrical complexity of the Jubilee Stock, comprising many contacts zones between rocks of different rheology, are interpreted to be important controls on the geometry and distribution of the gold zones. 116 �Figure 3-2 - Geology map of the Wawa Gold Property from Ronacher et al. (2016) 117 �The Jubilee Stock is a composite calc-alkaline to transitional intrusive complex formed by many individual intrusions compositionally ranging from gabbroic-dioritic to granitic (Figs. 3-3a, b and 4a, b). Based on U-Pb ages on zircons (Sullivan et al., 1985), the coarser-grained intrusions of the Jubilee Stock was emplaced around 2,745 ± 3 Ma, whereas one porphyritic intrusions located at the margins of the stock was dated at 2,742 ± 6 Ma (Turek et al., 1992). Within error, these ages are contemporaneous to those of its host volcanic sequence dated at 2,746 ± 11Ma and 2,744 ± 10 Ma (Turek et al., 1992). The main intrusive facies of the Jubilee Stock include: medium- to coarse-grained "diorite" (key unit historically used to define the Jubilee Stock), porphyritic intrusions with variable phenocryst assemblages (biotite, feldspar-biotite, feldspar, quartz-feldspar) generally emplaced at the margins of the intrusive complex, fine- to medium-grained mafic intrusions (gabbroic), zones of magma mixing and mingling where mafic to intermediate-felsic magmas intrude each other; and broad zones of intrusive breccia developed at the contact between the intrusive complex and its host volcanic rocks, but also along the contact zones between individual intrusions of the Jubilee Stock. Mapping by Sage et al. (1982) shows that the core of the Jubilee Stock is curved-shaped into a sigmoid form, its long axis oriented at 20°, and has an approximate 6 x 1.3 kilometres surface expression. The actual extent of the Jubilee Stock remains however ill-defined; as the porphyritic intrusions typically forming the margins of the intrusive complex are texturally similar to some of the crystal-rich volcanic rocks of the volcanic sequence and their classification as intrusive or volcanic varied according to the generations of workers on the property. Recent re-mapping by Red Pine indicates that the surface extent of the Jubilee Stock is larger than the core zone currently illustrated on the geology maps of the property. Medium-grained dioritic to granitic intrusions (Figures 3-5 and 3-6) Medium-grained to coarse-grained dioritic to granitic intrusions form the diagnostic core zone of the Jubilee Stock. Historically to simplify the nomenclature of those units, all the operators of the property described the medium- to coarse-grained intrusions of the Jubilee Stock using diorite as a generic term. However, chemically based on the Pearce (1996) diagram, the coarser grained intrusions of the Jubilee Stock vary compositionally between diorite and granite/granodiorite (Fig. 3-3a). This is supported by the petrographic work from Sage (1994) that indicates some of the intrusions have modal 10-30% quartz, 4055% plagioclase and 10-20% biotite, which indicates granitic composition. Porphyritic intrusions (Figure 3-7) Many porphyritic intrusions surround the core of Jubilee Stock and were hypothesized by Sage (1994) to occupy the ring fracture of a large caldera centered on the Jubilee Stock. In the contact zones between different intrusions, the porphyritic intrusions are often intermixed together and with intrusions of the medium- to coarse-grained diorite. The main primary phenocrysts assemblages observed in the porphyritic units are: feldspar, biotite-feldspar, quartz-feldspar and quartz. A compositional continuum and visual gradation between the medium- to coarse-grained diorite and intrusions of the feldspar-phyric, biotite-feldspar-phyric and biotite-phyric units are commonly observed, indicating the likely coeval emplacement of those units. Because of the variability in the mapping and logging of the porphyritic units, the porphyritic units of the Jubilee Stock remain undivided. The porphyritic intrusions are compositionally similar to the Jubilee Stock (Fig. 3-4). 118 �Figure 3-3 Geochemical classification diagram of the intrusions typical of the core of the Jubilee Stock. A) Zr/Ti versus Nb/Y discrimination diagram from Pearce (1996); B) Th/Yb discrimination diagram for discrimination of magmatic affinities from Ross and Bédard (2009) 119 �Figure 3-4 Geochemical classification diagram of the porphyritic intrusions observed in the Jubilee Stock. A) Zr/Ti versus Nb/Y discrimination diagram from Pearce (1996); B) Th/Yb discrimination diagram for discrimination of magmatic affinities from Ross and Bédard (2009) 120 �Figure 3-5 Medium- to coarse-grained facies of the Jubilee Stock "diorite" near the contact with the volcanic units containing enclaves of volcanic rocks Figure 3-6 - Typical Jubilee Stock "diorite" in forming the core of the intrusive complex 121 �Figure 3-7 Feldspar-quartz porphyritic intrusion exposed near the Surluga Deposit Silicified/Albitized unit (Figure 3-8) This unit corresponds to rock altered by a strong to intense sodic-silica alteration and encompasses altered diorite, volcanic units and porphyritic intrusions. This unit prevails in certain zones of the Wawa Gold Corridor and may correspond to the hornfelsed units described by Sage (1994) as occurring along some of the contacts between the Jubilee Stock and the volcanic rocks. In zones of intense alteration, the primary textures of the host rocks are generally destroyed, the unit becomes quite homogeneous and protolith identification is difficult. In the transitional zones, strong alteration fronts are seen to replace the host units. The predominant precursor unit is most likely fine-grained volcanic units intruded by the Jubilee stock in which albitization was preferentially partitioned. 122 �Figure 3-8 Albitized unit formed near the contacts between the Jubilee Stock and the volcanic units Intrusive breccias (Figures 3-9 and 3-10) Many of the contact zones between the intrusions of the Jubilee Stock, but also between different intrusive facies of the stock, are characterized by the formation of intrusive breccia zones. The breccia cement is typically composed of the coarser-grained facies granitic intrusions, whereas the fragments, predominantly of volcanic origin, are fine- to very fine-grained, vary considerably in size, ranging from a few millimetres to tens of metres, and some are partially assimilated by the dioritic magma. A report by Sage (1994) and noted by Red Pine geologists, this is making the mapping of this unit, especially in drill cores, particularly challenging. Figure 3-9 Surface exposure of the intrusive breccia formed at the contact between the Jubilee Stock medium- to coarse-grained diorite and the volcanic units 123 �Figure 3-10 Intrusive breccia texture in drill hole and melanocratic feldspar-phyric unit in the contact zone between the Jubilee Stock coarse-grained diorite and the volcanic units 3.2.2 Gabbroic rocks Tholeiitic to calc-alkaline mafic-intermediate to mafic intrusions were documented by Red Pine in the Jubilee Stock (Fig. 3-11). Visually, the mafic intrusions were described and discriminated based on the grain size of the core of the intrusion and the absence or presence of a porphyritic texture with feldspars being the main phenocrysts. The mafic intrusions observed in the Jubilee Stock include: coarse-grained gabbros pertaining to the tholeiitic suite (Fig. 3-12), fine-grained gabbros pertaining to the tholeiitic suite (Figure 3-13), and feldspar-phyric very fine-grained gabbro to gabbro-diorite generally pertaining to the calc-alkaline suite (Figure 3-14). Both the tholeiitic and the calc-alkaline mafic intrusions are deformed in the gold-bearing structures of the Wawa Gold Corridor. Walker (2011) also recognized, based on the observation of magma mixing textures between felsic and a mafic magmas in the Jubilee Stock, that some of the mafic intrusions are comagmatic with the stock. This observation is supported by the occurrence of calc-alkaline mafic intrusions in the Jubilee Stock. The intersection between some intrusions of the tholeiitic suite and the gold-bearing structure was also observed to form zones of preferential gold enrichments. In some of the mafic intrusions of the calc-alkaline suite in the Surluga Deposit, Ni-Cu mineralization can occur as disseminated cluster of pyrrhotite-chalcopyrite, the pyrrhotite likely intermingled with pentlandite (Fig. 3-14). The largest mafic intrusion on the property is centered on Reed Lake and forms the Reed Lake maficultramafic complex which is composed of diorite, quartz-gabbro, leuco- to mela-gabbro and pyroxenite (Fig. 6-2). Sage (1994) inferred that the combined trends of the long axis of 315° for the Reed Lake and 20° for the Jubilee Stock may suggest there were emplaced in a conjugate fracture system and that they are possibly petrogenetically related. 124 �Fig. 3-11 Geochemical classification diagram of the mafic and mafic/intermediate intrusions observed in the Jubilee Stock. A) Zr/Ti versus Nb/Y discrimination diagram from Pearce (1996); B) Th/Yb discrimination diagram for discrimination of magmatic affinities from Ross and Bédard (2009) 125 �Figure 3-12 Coarse-grained tholeiitic gabbroic intrusion in the Jubilee Stock Figure 3-13 Fine-grained tholeiitic gabbro dyke in the Jubilee Stock 3.2.3 Volcanic units For most of the Wawa Gold Project where large surface mapping and drilling programs have been conducted, the descriptions of the volcanic units are constantly evolving depending of the geologist, exploration model, and time period. In many cases the sub-volcanic porphyritic intrusions, part of the Jubilee Stock, and the volcanic units, are confused and their classification inter-changed. To add to the difficulties of recognizing the volcanic units, in historic logs, in adjacent drill holes, some of the goldbearing structures are variably described as sedimentary rocks, fragmental volcanic units, volcanic flows and porphyritic intrusions. 126 �3.2.4 Lamprophyre dikes Lamprophyre dikes are pervasive throughout the Wawa Gold Project and at least two generations of lamprophyre exists. One generation is late-stage and cut all the gold mineralized zones of the property. Dikes of that generation are black, porphyritic, medium-grained and strongly magnetic with a blue amphibole alteration halo. A possible set of lamprophyre is likely older. Dikes of that set are generally smaller, their primary mineralogy is partially to completely replaced, which gives them a dark- to palegreenish color. One dike of this set is also possibly gold mineralized, indicating that some of the lamprophyre dykes could have been emplaced prior to the formation of the gold system. A few carbonatite dikes, likely related to the Firesand Carbonatite located a few hundred metres east of the northeastern corner of the property, were also observed in drill holes in the Surluga Deposit. 3.3 Late brittle faulting The main brittle fault of the Wawa Gold Project is the NW-oriented and subvertical Parkhill Fault. Following Sage (1994), the Parkhill Fault is the southeastern extension of the northwest-striking Black Trout Lake Fault and is seen to truncate the Wawa Gold Corridor, which likely resurfaces to the south in the Darwin Shear Zone. The age of the Parkhill Fault remains uncertain and its intrusions by gabbroic rocks interpreted to be Archean indicate that it is possibly a long lived-structure in the area, even possibly formed during the evolution of the gold system. The interpreted late movement along the Parkhill Fault is left-lateral with an unknown vertical component. 4.0 Gold zones of the Wawa Gold Project 4.1 Attributes of the gold zones of the Wawa Gold Project Gold mineralization is obvious throughout the Wawa Gold Project and is spatially related to the numerous shear zones, fractures and replacement zones of variable strike and dip. The Surluga Deposit forms the largest gold concentration currently defined on the Wawa Gold Project. The deposit is entirely hosted in the Jubilee Shear Zone that crosscut the various intrusive rocks of the Jubilee Stock. The Jubilee Shear Zone is a brittle-ductile structure consisting of a number of parallel, ~300–900 m long en-echelon segments that strikes northeast (018–034°) and dips (25–55°) to the southeast (Figs. 4 to 6). Its width ranges from 9 m to 75 m and extends from Wawa Lake to the northwest-trending Parkhill Fault (3.2 km). Helmstaedt (1988) interpreted that the gold-mineralized quartz veins predate some of the brittle-ductile shearing in the Jubilee Shear Zone and that the geometry of the high-grade zone of the deposit is controlled by a strong stretching lineation in the shear zone. Syn-deformation folding in the structure may also increase the thickness of the mineralized zone as parasitic folds, generally present in the thickest mineralized zones of the structure (Figure 6-19). Red Pine Exploration demonstrated that the Jubilee Shear Zone is part of a larger deformation corridor that includes many other gold-bearing structures and zones of gold mineralization. This larger deformation corridor, in which tectonic foliations oriented NNE to NE (20°–45°) and dipping between 30° and 80° and stretching lineation trending 160°-190° and plunging 20°-35° are systematically observed (Fig. 7), was named the Wawa Gold Corridor. In addition to the Jubilee Shear Zone, the main goldbearing structures of the Wawa Gold Corridor includes: the Darwin Shear Zone (interpreted southern extension Jubilee Shear Zone south of Parkhill Fault), the Grace Deformation Zone, the Hornblende shear zones (Hornblende Upper and Lower), the Parkhill Mine Shear Zone, the Minto shear zones (Minto A, B 127 �and C), the Surluga Road Shear Zone, and the William gold zones. From north to south, the Wawa Gold Corridor is continuously defined between Wawa Lake and the Darwin-Grace Mine. Laterally, the Wawa Gold Corridor is defined between the Hornblende Shear Zone and the Minto B Shear Zone north of the Parkhill Fault, and between the Darwin Shear Zone and the Darwin-Grace Mine, south of the Parkhill Fault (Figs. 1, 3, 8-10). Throughout the Wawa Gold Corridor, gold mineralization predominantly occurred in brittle-ductile shear zones, but also occurs in micro-breccias formed into albitized zones of the Jubilee Stock intrusions, and into network of quartz tension veins. Although the main orientation of the tectonic fabrics defining the Wawa Gold Corridor and some of its main gold zones is NNE (Hornblende, Jubilee, Minto B, Darwin), additional orientations exists for the corridor gold-bearing shear zones. This includes ENE-oriented shear zones (Minto C, Parkhill Mine Shear Zone, Mickelson Shear Zone, Root Vein structure) and WNW-oriented shear zones (Grace Deformation Zone, Minto A shear Zone). Domains of L-tectonite controlled by the stretching lineation trending 160°190° and plunging 20°-35° without strong planar fabric development are also relatively common. Figure 4 - Cora Shaft into the Jubilee Shear Zone 128 �Figure 5 - Brittle-ductile shearing and tight parasitic folding of auriferous quartz veins in the Jubilee Shear Zone Figure 6 - Zone of brittle deformation in the Jubilee Shear Zone 129 �Figure 7 - Characteristic stretching lineation of the Wawa Gold Corridor preferentially partitioned in a mafic dike (William Gold Zone) Figure 8 - Brittle-ductile deformation in the Minto B Shear Zone 130 �Figure 9 - Intersection of the Minto A Shear Zone, related to the Minto and Parkhill mines, and the Jubilee Shear Zone 131 �Figure 10 - Gold mineralization in the Grace Deformation Zone related to the historic Darwin-Grace Mine 132 �Figure 6-19 - Folded high-grade quartz shear vein in the Mickelson Shear Zone Channel sample contains 69.5 g/t gold over 0.7m In most of the gold zones on the property, key visual indicators of gold mineralization include the presence of sulfides, potassic alteration (sericitic or biotitic) and development of quartz veins with accessory sulfides (Fig. 11). The main style of gold mineralization in the Wawa Gold Corridor, in both the shear zones and the replacement zones, principally occurs as free gold associated with pyrite, accessory to minor pyrrhotite and minor to absent chalcopyrite deposited with silicification and sericitic alteration, and with or without quartz veining. Moderate to strong tourmalinization also occurred in some of those gold zones. The modal content of sulfides is generally around 0.2-3%, and locally up to 5-10% in the higher grade zones, which are almost systematically associated with quartz veins. Another style of gold mineralization is characterized by free gold associated with abundant arsenopyrite with accessory to minor pyrite and/or pyrrhotite deposited with chlorite+biotite alteration, weak to moderate-strong sericite alteration and with or without quartz veining. Arsenopyrite in those gold zones can occur in modal content exceeding 3% (Grace Deformation Zone, Minto C shear zone). In the Jubilee Shear, Hornblende and Minto A+B shear zones, arsenopyrite-bearing or pyrite-bearing gold zones occur in close spatial association. Many zones with a transitional metal signature between the As-rich and other metal assemblages also occur in those structures (Fig. 12). In the transitional zones, pyrite is the main sulfide, but these zones exhibit a variable arsenic content between 100-1000 ppm. In the Grace and Minto C shear zones, arsenopyrite-rich gold mineralization predominates and is associated with abundant quartz veining in the higher grade zones of the structure (Fig. 10). In the Minto C shear zone, quartz veining is also relatively minor to absent; and does not seem to be the main control on gold concentration as highgrade gold in the structure is not spatially associated with quartz veining like in the Grace Deformation Zone (Fig. 13). 133 �Figure 11 - Grey quartz vein with pyrite typical of the higher-grade zones of the pyritic gold zones of the Surluga Deposit Drilling intersection contains 18.62 g/t gold over 0.48 m A rarely observed sulfide paragenesis formed of pyrite with accessory sphalerite and galena deposited in quartz veins also occur in the Wawa Gold Corridor, and was so far only observed in the Surluga Deposit. However chemically Pb and Zn enrichments are characteristic of many of the property gold zones (Fig. 12). The main network of gold-bearing tension veins recognized by Red Pine in the Wawa Gold Corridor is preferentially formed in the medium- to coarse-grained diorite and is not associated with obvious shearing (Fig. 13). Individual tension veins of the network are moderately to steeply dipping (50-80°) to the WNW to NNE and typically contain accessory tourmaline, accessory pyrite and pyrrhotite, minor chalcopyrite and locally visible gold. The variable development of tectonic fabrics in the tension veins of the property indicate a pre- to syn-tectonic formation of the gold mineralized tension veins. 134 �Figure 12 - PC1 vs PC2 biplot of log-centered transformed concentration of metals to illustrate the diversified metal assemblages observed in the gold zones of the Wawa Gold Corridor Analyses from 326 drill core samples with Au > 1 g/t; because of the large number of analyses below detection, Ag, Se and Te were not used for the PCA. Figure 12 - High-grade core of the Minto C Shear Zone formed after a strongly sericitized and silicified quartz-feldspar porphyry Grab samples from this arsenopyrite-rich shear zone without ubiquitous quartz veins contains 17 g/t and 5.51 g/t gold 135 �Figure 13 - Quartz tension vein with visible gold in the footwall of the Surluga Deposit Drilling intersection contains 53.2 g/t gold over 1 m The William gold zones form a style of gold mineralization that contrasts with the shear-hosted gold zones of the Wawa Gold Corridor and were first identified during the fall 2015 drill program. In the William gold zones, tectonic fabrics (planar and linear) are absent to locally strong, hydrothermal alteration spatially related to gold deposition is weak to strong and mineralization is associated with pervasive micro-brecciation of the albitized host rock (Fig. 13). The main geometrical controls hypothesized for the William gold zone mineralized zone are: weak ENE-striking and shallowly dipping (20°-35°) shearing and the stretching lineation of the Wawa Gold Corridor that shallowly plunges to the ESE. An elevated arsenic background is diagnostic of these zones and likely relates to the finely disseminated pyrite interpreted to be related to gold mineralization. Because of the weak visual indicators associated with William-like mineralized zones, they were generally missed by all previous operators of the property. Figure 13 - William-like mineralization in the Jubilee Shear Zone hangingwall Drilling intersection contains 1.86 g/t gold over 1.15 metre Three main controls on the spatial distribution and geometry of the gold zones were defined in the Wawa Gold Corridor and include: • • • rheological contrasts between the various intrusive phases of the Jubilee Stock to partition and focus deformation and fluid circulation (Fig. 14); chemical contrast between units of the Jubilee Stock and more specifically the contact zones between gabbroic intrusions and the other intrusion types of the Jubilee Stock to partition and focus deformation and fluid circulation, and also to act as a chemical trap for gold; and the stretching lineation characteristic of the shear zones of the property. 136 �Figure 14 - Strain partitioning in the Minto B shear zone where shearing is preferentially partitioned in the mafic rocks whereas the dioritic domain remains weakly deformed 4.2 Hydrothermal alteration associated with gold mineralization Many mineralogical changes are associated with the formation of the gold-bearing shear zones. In mafic rocks, progressive chloritization and carbonatation (calcite and ankerite) with subsidiary tourmalinization and mariposite formation indicate increasing proximity to the gold-bearing shear zones is (Fig. 2). In the intermediate to felsic precursors, the mineralogical changes are progressive sericitization, silicification and carbonatation. In intermediate to mafic rocks, magnetite is also deposited in the shoulders of the goldbearing structures and can form haloes extending 25-30 metres away from the gold-bearing structure. In the Wawa Gold Corridor the hangingwall of the Jubilee Shear Zone the biotite alteration forms a distal halo extending tens of metres away from the structure that gets progressively chloritized with increasing proximity to the shear zone. Barren sericitization in sheared intermediate to felsic precursor rocks and barren chloritization and carbonatation in sheared gabbroic rocks also occurs in the gold-bearing structures of the property. A pervasive hydrothermal imprint related to the emplacement of swarms of lamprophyre dykes is present throughout the property. Mineralogically it leads to the formation of riebeckite in the contact zones of the lamprophyre dykes as well as networks of iron carbonate (pos. siderite) veins with K-feldspar selvages. 137 �5. Field Trip Stops Figure 5-1 Stops of the fieldtrip on the Wawa Gold Project Stop 1: Cora Shaft and Jubilee Shear Zone Coordinates: UTM 16 NAD83 667930mE, 5316243mN This stop provides an overview of the early mining history in the Jubilee Shear Zone, current host of the inferred resource estimated for the Wawa Gold Project, and what is now the Surluga Deposit. This stop is located just south of the small peninsula in Jubilee Lake where the historic Jubilee Mine was operated 138 �between 1926 and 1938 (Fig. 5-1), and where the Cora Shaft was sunk on a quartz vein in the Jubilee Shear Zone in 1927. The Cora Shaft outcrop provides a good overview of the structural complexity of the Jubilee Shear Zone, of the geometry of some of the parasitic folds in the shear zone, and on the higher grade quartz veins formed in the shear zone. Individual grab samples from the Cora Shaft quartz vein contain up to 50.8 g/t gold. Figure 5-1 Jubilee Mine during operation in the 1930s Source: http://www.michiwawa.ca/albums/wawa/goldfields/goldfield_one.htm 139 �Figure 4 Exposure of the Jubilee Shear Zone at the Cora Shaft Stop 2: Jubilee Shear Zone Southern Extension Coordinates: UTM 16 NAD83 667723 mE, 5315771 mN This stop illustrates the control of the zones rheological contrasts between different intrusions of the Jubilee Stock has on strain partitioning, fluid circulation and gold concentration in the Jubilee Shear Zone. The Jubilee Shear Zone on the outcrop traverses a contact zone between a mafic dyke and the Jubilee Stock "diorite". Brittle-ductile deformation is strongly partitioned into the mafic rocks in the shear zone, whereas quartz veining and gold concentration preferentially occur along the contact zones between the deformed mafic rocks and the "diorite". Channel sample from the contact zone contains 6.74 g/t gold over 1 metre. 140 �Figure 5 Exposure of the southern extension of the Jubilee Shear Zone Stops 3 to 7 - Wawa Gold Corridor transect This series of stops illustrates some of the main gold zones and gold mineralization types of the Wawa Gold Corridor in the northernmost extension of the Surluga Deposit. The transects begins in the Jubilee Shear Zone and also covers some the various intrusive facies of the Wawa Gold Corridor. Stop 3: Jubilee Shear Zone Northern Extension Coordinates: UTM 16 NAD83 668273 mE, 5317339 mN This stop provides an overview of the Jubilee Shear Zone in the northernmost extension of the Surluga Deposit. The trench also exposes the brittle to brittle-ductile deformation regime that were active in the Jubilee Shear Zone. Stop 4: Tholeiitic Gabbro and Jubilee Stock "diorite" Coordinates: UTM 16 NAD83 668231 mE, 5317362 mN This stop shows at surface, one of the tholeiitic gabbro dyke on the property. The intersection between that gabbro dyke and the Jubilee Shear Zone spatially correspond to one of the main high-grade zones defined in the Surluga Deposit. Gabbro dykes of the tholeiitic series are considered as having a critical role on the spatial distribution of the higher-grade zones in the gold-bearing structures of the property. 141 �Figure 6 Contact between the tholeiitic gabbro and an intermediate intrusion of the Jubilee Stock Stop 5: Surluga Road Shear Zone Coordinates: UTM 16 NAD83 668130 mE, 5317400 mN This stop shows one of the shear zone, parallel to the Jubilee Shear Zone, discovered by Red Pine in the Wawa Gold Corridor. Similar to the other structures on the property, the intersection between the structure and mafic/diorite contacts forms a preferential zone for gold enrichment and quartz veining. The shear is adjacent to the Surluga Road, but was not tested for gold before the exploration program of Red Pine during the summer of 2016. Channel sampling on the exposed outcrop indicate that the shear zone contains 1.59 g/t gold over 7 metres. 142 �Figure 7 Exposure of the Surluga Road Shear Zone next to the Surluga Road Stop 6: Hornblende Shear Zone Coordinates: UTM 16 NAD83 668054 mE, 5317474 mN The Hornblende Shear Zone is one of the large gold-bearing structures in the footwall of the Surluga Deposit and two contrasting styles of mineralization are visible on the exposure. The northern part of the exposure is characterized by a high-strain domain with low quartz vein density and pervasive dissemination of pyrite. The southern domain is characterized by weak strain intensity, more abundant quartz veining. Drilling in the northern domain hole HS-15-28 contains 1.25 g/t gold over 15.35 metres, whereas hole HS-15-27, drilled in the southern domain, contains 3.11 g/t gold over 8.1 metres. 143 �Figure 8 Hornblende Shear Zone exposure Stop 7: William Gold Zone Coordinates: UTM 16 NAD83 668093 mE, 5317233 mN This stop illustrates the cryptic style of gold mineralization associated with the William gold zones. The controls on the spatial distribution of the gold on the outcrop, and within the William gold zones in general, remain to be clearly established. Tectonic fabrics of the Wawa Gold Corridor are moderately to weakly developed on the outcrop and well-evident in the mafic dykes in which a strong stretching lineation is developed, but the relation between the tectonic fabric and gold distribution could not be unequivocally established. Stop 8: Root Vein Coordinates: UTM 16 NAD83 668775 mE, 5318459 mN This stop illustrates a gold-bearing quartz vein formed in a ENE-oriented shear zone. The overall style of gold mineralization characteristic of the Root Vein is similar to the structures mined in the Parkhill and Van Sickle mines. The outcrop also exposes at surface an example of a lamprophyre dyke that is part of the lamprophyre dyke swarm of the Wawa Gold Property; and provides good exposures of the porphyritic and coarser-grained intrusive facies of the Jubilee Stock, and of the tension veins formed in the Wawa Gold Corridor. 144 �Figure 9 Core of the Root vein Length of channel sample is 1.5 metres Stop 9 Darwin-Grace Mine site and exposure of Grace Deformation Zone UTM Coordinates: UTM 16 NAD83 668077 mE, 5313382 mN This stop illustrates the site of the historic Darwin-Grace mine and an exposure of the Grace Deformation Zone. The exposure of the Grace Deformation Zone illustrates the strong brittle-ductile deformation and arsenopyrite flooding with quartz veining characteristic of that structure. Grab samples collected from the Grace Deformation on the outcrop contain between 1.27 g/t and 18.4 g/t gold. Stop 10 Gold zones of the Wawa Gold Corridor (Red Pine Exploration core shack) This stop, at the Red Pine core shack, will present a suite of drill core intersections to illustrate the main gold zones discovered so far in the Wawa Gold Corridor. This will include drill intersections from: the Jubilee Shear Zone, the Hornblende Shear Zone, the William gold zones, the Minto shear zones and the Grace Deformation Zone, and some of the gold-bearing tension veins. 6. References Amalgamated Reports, Ontario: Ontario Ministry of Northern Development and Mines Assessment Report No. 41N15NE00036, 103 p. Bowdidge, C., 1996, Mineralization in the Jubilee Shear Zone - Re-appraisal as a large-tonnage, lowgrade bulk-mineable underground resource. Goldbrook Exploration Inc., 15 p. 145 �Frey, E.D., 1987, Geology of Wawa area gold mineralization: Institute of Lake Superior Geology Field Trip Guidebook, v. 33, Part 2, 31 p. Helmstaedt, H., 1988, Structural observations in the Surluga and Jubilee mines, Citadel Gold Mines Inc., Wawa, Ontario: Report for Citadel Gold Mines Inc., 29 p. MacMillan, D. and Rupert, R.J., 1990, Exploration Report -- Geological Mapping in the Vicinity of the Grace- Darwin, Parkhill and Minto Mines: Report for Citadel Gold Mines Inc., 61 p. Ontario Geological Survey 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release–Data 126 - Revision 1. Ontario Geological Survey 2014. Mineral Deposit Inventory—2014; Ontario Geological Survey. Pearce, J.A., 1996, A user guide to basalt discrimination diagrams, in Wyman, D.A., ed., Trace element geochemistry of volcanic rocks: Application for massive sulfide exploration: Geological Association of Canada, Short Course Note 12, p. 79–113. Ross, P.-S. and Bédard, J.H., 2009, Magmatic affinity of modern and ancient subalkaline volcanic rocks determined from trace element discriminant diagrams. Canadian Journal of Earth Sciences, v. 46, p. 823−839. Rupert, R.J., 1989a, Citadel Gold Mines Inc. Report on Magnetometer Survey Block B West of Firesand River, McMurray Township, Ontario: Ontario Ministry of Northern Development and Mines Assessment Report No. 41N15NE0023, 20 p. Rupert, R.J., 1989b, Citadel Gold Mines Inc. Report on Magnetometer Survey South Part of Block A at Deep Lake: Ontario Ministry of Northern Development and Mines Assessment Report No. 41N15NE0021, 14 p. Rupert, R.J., 1997, Exploration report on the Wawa area properties of Citadel Gold Mines Inc., Report for Citadel Gold Mines Inc., 51. Rupert, R.J. and Leroy, A., 1989, Citadel Gold Mines Inc., Technical Reports OMEP Project No. OM887- C-254: Ontario Ministry of Northern Development and Mines Assessment Report No. 42C02SE0220, 465 p. Sage, R.P., 1994, Geology of the Michipicoten greenstone belt: Ontario Geological Survey Open File Report 5888, 592 p. Sage, R.P., Sawitsky, E., Turner, J., Leeselleur, P., and Sagle, E., 1982, Precambrian Geology of McMurray Township, Wawa Area, Algoma District. Ontario Geological Survey, Preliminary Map Scale 1:15 840, P. 2441. Sherman, B., 2005, Illustrated Information to Accompany an Independent Assessment of the Mineral Exploration Potential of the Surluga Property of Citadel Gold Mines Inc., at Wawa, Ontario: Report for Citadel Gold Mines Inc., 48 p. 146 �Sullivan, R. W., Sage, R. P., and Card, K. D. 1985. U-Pb zircon age of the Jubilee stock in the Michipicoten greenstone belt near Wawa, Ontario. In Current research, part B. Geological Survey of Canada, Paper 85-1B, pp. 361-365. The Wawa History Album, 2009, http://www.michiwawa.ca/albums/wawa/goldfields/goldfield_one.htm, website consulted March 27-28th, 2017 Turek, A., Sage, R.P., and Van Schmus, W.R., 1992, Advances in the U-Pb zircon geochronology of the Michipicoten greenstone belt, Superior Province, Ontario. Canadian Journal of Earth Sciences, v. 29, p. 1154-1165. Walker, E.C., 2011, 2011 Prospecting, Geology, and Sampling Program: Re-Appraisal of an Old Gold Camp. Ministry of Northern Development and Mines, Assessment Report N. 20010231, 65 p. 147 �Field Trip 4 Geology of the Island Gold Mine Doug MacMillan, Simon Comtois-Urbain, Harold Tracanelli Richmont Mines Inc.- Island Gold Exploration Department Dubreuilville, Ontario 148 �Introduction This field trip will provide participants with an overview of the local property geology of the Island Gold Mine which is situated 83 kilometers north-west of Wawa. The trip will start at the Island Gold Mine coreshack where core will be displayed from ore intercepts within the mine horizon across a 2 kilometer wide strike length and then the visitation of outcrop exposures along a 4 kilometer segment of the immediate mine stratigraphy. Elements of the Goudreau Lake Deformation Zone brittle-ductile deformation which include shearing, folding and veining will be seen through many of the exposures in the area including the gabbro sill hosting the Kremzar Deposit, highly Z folded Michipicoten Iron formation at the Morisson No. 1, the exposed south east extension of the Island Gold mine structure and within the Webb Lake Sill and Stock which hosts the Magino Mine and Argonaut Gold open pit resource. The Goudreau-Lochalsh area is a historic gold camp with the first gold discovery in 1918 by D. J. McCarthy and Mr. W. J. Webb within the Webb Lake Stock and later in 1926 by Patrice Kremzar on what is now the Richmont Island Gold Mine Property. In the 1930’s the Cline Mine was in production 8 kilometers east of the Island Gold mine and eventually produced 63,000 ounces of gold. Regional Geology The Island Gold deposit located in the Wawa Sub-Province of the Superior Province of the Canadian Shield. The Wawa Subprovince is composed of numerous east-west trending sublinear supracrustal belts of greenstone enclosed by larger regions of granitoid rocks. Greenstone belts are dominated by metavolcanic and metasedimentary rocks ranging in age from 2.9 to 2.7 Ga (Percival, 2012). The Island Gold Deposit occurs on the north central flank of the Michipicoten greenstone belt which has produced over 3 million ounces of gold since the first `gold rush’ in Wawa in the late 1890’s. The greenstone belt has dimensions of 145 km by 45 km in dimensions and lies 150 km south east of the prolific Hemlo Mining Camp. It has been subdivided into three cycles of bimodal volcanism aged 2.9, 2.75 and 2.7 Ga respectively, referred to as the Hawk, Wawa and Catfish assemblages (Turek, 1994, Sage 1994). Volcanism in the lower and upper cycles are interpreted by Sage (1994) to have been Hawaiian and Plinian types respectively. Each of the three cycles is capped by pyrite bearing iron formation with an unconformity separating the volcanic rocks from the overlying iron formation in cycles 2 and 3. The Dore metasediments, composed of wackes and conglomerates, unconformably overlie the 2.7 Ga cycle and are the youngest rocks in the belt. The metamorphic grade varies from upper greenschist to lower amphibolite at the margins of the belt in proximity to the external granitiod rocks. The belt is interpreted to consist originally of a monoclonal sequence that was subsequently underwent complex thrusting, folding and shearing (Sage 1994). Arias and Helmstaedt (1990) interpret the present configuration of the belt as a regional nappe fold refolded about an F2 axis and then imbricately thrusted into numerous slices. 149 �Figure 1: Michipicoten Greenstone Belt (MGB) Figure 2: MGB Geology and Structure 150 �The main structural control on most Au occurrences in the Goudreau-Lochalsh area in which the Island Gold deposit is situated is the Goudreau Lake Deformation Zone (GLDZ). This comprises a 30 kilometer long by 4.5 kilometer wide deformational structure described by Heather and Arias (1992) as a dextral oblique slip structure occurring near the interface between the Wawa and Catfish volcanic cycles. The deformation zone has a gentle sigmoidal shape and is not a focused single structure but a composite of smaller scale structures with development of deformation being more controlled by hosting geological units such as formational contacts where rock competency contrasts exist and less by external regional stress. (Dube B., and Gosselin, P., 2007). The Cradle Lake Deformation is situated south of the GLDZ on the south arm of the Goudreau Lake Anticline near the interface of cycle 2 and 3 volcanic assemblages. The deposits of the Goudreau Lochalsh camp have a strong spatial correlation with a variety of felsic intrusive rocks including trondhjemite stocks and sills, felsic porphyries and dyke complexes which is common to many Superior Province gold deposits (Hodgson and MacGeehan, 1982). Exceptions occur including the Kremzar Deposit (47,000 oz. 1988-90) which is hosted by a regional gabbroic sill. The Island Gold deposit is situated within the cycle 2 volcanic assemblage however there is close spatial association with both mafic, felsic dykes and massive felsic volcanic rocks within the mineralized gold bearing corridor. Competency contrasts and rheologic variations between the more and less competent units provide planes of weakness along which shearing can develop, structural traps can open and gold bearing veins can subsequently be formed. The iron rich chemistry of proximal mafic dykes are also considered to assist in gold formation by providing a chemical trap in the Island Gold system. Other factors include the recognition of an antiform on the north limb of the Goudreau Lake Anticline in which the Island Gold mine sequence appears to be situated in adding to structural complexity of parasitic folding and flexural slippage and associated shearing and thrusting along fold limbs. Figure 3: Goudreau Lake Deformation Zone (GLDZ) 151 �Figure 4: Island Gold Stop Locations STOP 1: The Island Gold Mine Coreshack Location: UTM NAD83 16U 690538mE, 5353660mN Description: Since the beginning of production in 2007 at the Island Gold Mine Richmont Mines has produced 2.1 MT of ore at 6.46 g/t totaling 433,000 ounces of gold The Island Gold Mine is an Archean Lode gold deposit and consists of seven main narrow sub-parallel east-northeast striking, steeply south dipping, stacked quartz vein systems carrying gold mineralization within envelopes of intense sericitization, silicification and carbonatization with 2% to 5% pyrite which can be up to 8 meters wide. (Adam et al., 2015). The mineralized zones are sub parallel and can anastomize thereby merging or splitting along strike. Main veins styles include laminated veins with a defined ribbon appearance (Fig.7), sheeted veinlet zones composed of millimeter to centimeter scale quartz carbonate stringers, massive smokey greyish veins (Fig.5) and several types of milky white extensional veins which can appear a blow outs, gashes or horizontal ladder veins. Quartz veins at depth can be over four meters wide (Fig.7) The mineralized corridor increases from 50 meters wide above the 400L to over 100 meters at depth. Gold mineralization discovered at depth now constitute the principal source of ore and represent the down dip continuation of the mineralized zones which were mined before 2016. As of December 31, 2016, total Proven and Probable Reserves at the Island Gold Mine were 2.5 MT at a grade of 9.17 g/T for 752,200 gold ounces. The estimated Measured and Indicated Resources totaled 0.47 MT at a grade of 5.94 g/T for 91,450 ounces. Estimated Inferred Resources totaled 3.0 MT at a grade of 10.18 g/T for 995,700 ounces. 152 �Diamond drill core will be displayed from intercepts across approximately 2 kilometers of the mine horizon which will include intercepts of mineralization, quartz veining and alteration together with accompanying longsection and cross sections for location and geology. Core representatives of footwall and hangingwall rock will be presented as well as representative rocks types within the sequence. Figure 5: C zone quartz vein Figure 6: GD-14-01C, 19.9 g/t over 3.9 meters true width, -1250 meters 153 �Figure 7: Isometric Diagram Island Gold Mine Looking North East Figure 8: 800 Level quartz veins on 4 meter wide face (42.25g/t – 325.49 g/t uncut) 154 �STOP 2: Heritage Outcrop Location: NAD83 16U 690478mE, 5353601mN Description: The Heritage outcrop is situated in the younger northerly Catfish Volcanic Assemblage (2.9 Ga) which is primarily mafic in character facing north, dipping steeply north and striking 060°-070°. This outcrop is the discovery outcrop of the Kremzar Mine which was in production by Canamax between 1988-1990 and produced a total of 47,000 ounces at a grade of ?. The Kremzar deposit is hosted by a regional gabbroic sill which extends across the property to the Maskinonge Lake fault to the east. Gold was first discovered on the property in 1926 by Patrice Kremzar in a piece of float on the old Pic road. The property contains over 17 gold showings of varying economic potential. The Kremzar orebody occurs within sheared and altered gabbroic intrusive rocks. Kwok (1986) describes the alteration as an outer chlorite-biotite-carbonate zone and an inner zone dominated by sericitebiotite-potassium feldspar-Fe-carbonate. The mineralized sheared areas of rock weather a rusty brown. The main Kremzar mineralization is comprised of two sheared hosted quartz vein systems. The main mineralized zone called the R zone is situated at the west of the exposure and occurs within a 120-140° striking shear zone (dextral movement), dips steeply southwest and plunges 70° north east. A secondary B zone is situated to the north west of the exposure. Outside of the main mineralized shear zone containing the R and B zones the gabbroic rocks have undergone a complex system of brittle fracturing with fracture controlled Fe carbonate alteration and quartz filled fractures dominate the veining with the primary orientation at 30° (sinistral movement) and a secondary orientation of 120-140°, both orientations may contain anomalous gold values. Sharp boundaries exist between the altered sheared rock corridor and the surrounding unaltered gabbro which bound them. Figure 9: Heritage Outcrop R zone 155 �Figure 10: Heritage Outcrop Map A – intensely sheared and altered gabbro B – Undeformed gabbro C – clot size coarse grained amphibole and plagioclase crystals D – Diabase 160° azimuth E- Outer alteration zone, cl-bt-cb F – Inner alteration zone se-bt-Kspar-fe cb G – sharp boundary between alteration and relatively unaltered rocks H – 120°-140° striking shear zone I – R zone vein J – R zone K – B zone L – R and B zone appear to converge M – brittle fracturing N – myriad of narrow quartz filled fracture fillings O – strong fe carbonate P – brittle fractures strongly disrupted in ductile shear 156 �STOP 3: Morrison No.1 Iron Formation Location: NAD83 16U UTM 690775mE, 5352598mN Description: This stop is part of the Goudreau Iron Range, a Michipicoten type iron formation, which occurs at the interface between Cycle 2 (Wawa Assemblage, 2.75 Ga) and Cycle 3 (Catfish Assemblage, 2.9 Ga) volcanics. The iron formation is in sharp conformable contact with the southerly crystal tuffs of the Wawa Volcanic assemblage which it over lies. The Goudreau iron formation is a thickly bedded exhalative chemical sediment which can contain a minimum of five facies (Sage 1994) including a lower siderite or calcite unit, pyrite unit, chert-magnetite unit, chert-wacke unit and an uppermost argillitepyrite-graphite. At the eastern outcrop of the Morrison No.1 a stratigraphic sequence of siderite bearing ironstones (10m) overlain by massive pyrite (3m) in turn overlain by chert-pyrite-magnetite beds (20m) at the top (north) can been seen. Oxygen isotope data indicates iron formation deposition occurred under 200° C (LeSeeleur, 1980) suggesting a distal low temperature regime of formation and that depositing fluids were too low to transport base metals (Edmond et al, 1974). Randomly oriented chloritoid porphyroblasts within the iron formation are prominent and range up to 1 cm in diameter occurring pyroclastic fragments as well as the footwall felsic volcanic rocks to the south. Chlortiod alteration is a common alteration mineral in the Wawa iron ranges. Chloritoid bearing pyroclastic fragments within the overlying the Morrison iron formation are present and common in outcrop. Lockwood (1983,1986)) determined chloritoid-bearing rocks where a stratiform alteration pattern resulted from unfocused fluid discharge and overall low water/rock ratios under low pressure in subaqueous to subaerile conditions. Sage suggests that regionally the great strike extent of chloritoid bearing units indicates a hydrothermal system controlled by rock permeability along a broad aquifer. Shears zones are present throughout the outcrop as narrow discrete zones displaying dextral movement. S and Z symmetry folds are also present in the Morrison No.1 exposure. Z folds may be interpreted as dextral shear folds and the S folds possibly suggesting parasitic folding related to a possible antiformal structure in the area. Minor folds in the iron formation are present and plunge shallowly to the west. Other exposure features of note include boudinaged mafic dykes which `float’ within less competent siderite ironstones but which appear to be one original mafic dyke unit which has been tightly Z folded. The Morrison No.1 has as well been highly Z folded in the immediate area as depicted by geophysical interpretation. Regionally along strike however the iron formation can be highly transposed or distended. The gold mineralization is fracture controlled and quartz filled fractures can be seen developing at the west end of the out crop area. A historic non-compliant 43-101 resource for the Morrison #1 is about 540,000 tons @ .05 g/t gold (T. Foster, 1982). Figure 11: Chloritoid in Mafic Dykes Figure 12: Sulphide Facies Iron Formation 157 �Figure 13: Morrison No.1 West Zone A – footwall volcanic rocks B – abundant chloritoid C – boudinaged mafic dyke in siderite ironstone D – py bearing ironstone + pervasive Fe-carbonate E – strongly siderized tuffaceous rocks F – discrete shears and dextral displacement G – Z symmetry folds H – S symmetry folds I – crenulation cleavage suggesting oblique north side up dextral displacement J – siderite overlain by pyrite overlain by chert-pyrite-magnetite at the top K – narrow quartz veinlets crosscut iron formation 158 �Stop 4: South East Mine Structure Stripped Area Location: NAD83 16U 692318mE, 5352439mN Description: This stripped area is situated at the south eastern end of Goudreau Lake. It represents the only exposed segment of the Island Gold mine structure to date as the producing mine is presently situated beneath Goudreau Lake. The stripped exposure consists of a sequence of fragmental volcanic rock units succeeded to the south by a highly sheared, friable zone of sericitized felsic rocks in turn succeeded by massive felsic volcanic rock units. Local mafic dykes can occur throughout the sequence. Quartz veining is present locally and can be up to a meters in width. The shear has been correlated with the eastern strike extension of the Island Gold mine structure. At this location the structure is dipping steeply north and is striking approximately 70°. Felsic volcanic rocks in the southern part of the exposure area are a generally massive, fine grained, white weathered type and can contain clear colorless rounded quartz eyes up to 5%. The felsic volcanic rocks within the mine sequence to the west have a rock chemistry that straddles the border of the Rhyodacite-Dacite and Andesite field in the Winchester and Floyd (1977) diagram (T.Ciufo, 2017). The felsic rocks place in the dacite field of a TAS diagram (T.Ciufo, 2017). The fragmental volcanic rock unit in the north part of the stripped exposure is generally oligomictic, matrix supported and not well sorted. Variable amounts of whitish lensoid shaped lapilli fragments range in size from 4mm to several centimeters with compositions from felsic to intermediate with rare mafic clasts on occasion. Locally larger subangular to angular fragments up to 20 cm are present. The chloritoid alteration can commonly be seen within the fragments and matrix. Chloritoid bearing fragments can locally been seen suspended in chloritic units of either sediment or dyke rocks. Rock competency contrast is suggested to have created planes of weakness and therefore promoted shear propagation along these boundaries as one part of the mechanism in the gold forming process at Island Gold. In the exposure this seems to be demonstrated as a highly sheared felsic rock zone occurs between a southerly massive felsic volcanic unit to the south and a `less competent’ volcanic breccia unit to the north. Figure 14: Island Gold Mine Structure SE of Goudreau Lake 159 �Stop 5: Webb Lake Sill (optional) Location: NAD83 16U 689841mE, 5351605mN Description: The Webb Lake Stock is a trondhjemite intrusion situated 250 meters north of the Island Gold Mine structure. The stock has been mapped by Sage (1994) as being two rock masses composed of a southerly and northerly lobe which are divided by an intervening unit of a fine grained mafic intrusive or possibly metavolcanic. The main body of the stock extends for a strike length of over 2 km but continues into the Richmont Property to the east with a more sill like geometry traceable for 5 kilometers. The Webb Lake Stock dips 60-70° north and strikes approximately 70°. The intrusive is generally fine grained, hypidiomorphic granular with 65% plagioclase and 35% quartz with minor carbonate, chlorite and tourmaline. Sericite is common and potassium feldspar is rare. The outcrop exposure exhibits Z folds which display shallow 240° oriented fold hinges and north dipping axial planes. This structure is interpreted (K. Jellicoe, 2017) to have formed during D2 deformation during a NW-SE south-up dextral transpression. A late extensional tourmaline vein is present on the outcrop which cross cuts a previously developed D1 foliation as floods the planes of foliation. Figure 15: Z fold on south flank of Webb Lake Sill – Vent Raise Outcrop 160 �STOP 6: Webb Lake Trondhjemite Stock (Optional) Location: NAD83 U16 689416mE, 5351240mN Description: This stop is on the Webb Lake stock and hosting the gold mineralization at the Magino deposit. Gold was first found by Mr. D. J. McCarthy and Mr. W. J. Webb in 1917. Inferred reserves are estimated at 105.4Mt at 0.89g/t Au. The long axis of the north dipping stock is parallel to the regional supracrustal stratigraphy and has a strike length of approximately 2 kilometers and widths up to 300 m. The stock was emplaced in mafic volcanic rocks from the lower part of the Catfish assemblage (Cycle 3). It is called a granodiorite by Argonaut (43-101, 2016) or a trondjhemite by Sage (1994). The gold is hosted in numerous quartz veins and in quartz-flooded bands, parallel to the Goudreau Lake Deformation Zone (Argonaut 43-101, 2016). Silicification, sericitisation and an increase in pyrite content are typical around the mineralized veins and quartz flooded bands. This outcrop exhibits a mostly massive intrusive with numerous thin more ductile bands, and a few thin mineralized quartz veins that can be followed on a few meters. Figure 16: Webb Lake Stock Stop Locations 161 �Figure 17: Webb Lake Trondhjemite, small scale auriferous veins Figure 18: Webb Lake Trondhjemite C/S fabric Stop 7: Iron Formation (Optional) Location: NAD83 U16 688700mE, 5350900mN Description: Upper level of the Michipicoten Iron Formation containing banded layers of chert, magnetite, pyrite with intercalated argillaceous cycle three mafic volcanic rocks. The Webb Lake stock is located less than 30 meters north. 162 �References Adam, D., Bastien, J., Belisle, M., and Poirer, S. 2015 Technical Report and Preliminary Economic Assessment for the Island Gold Lower Zones (according to National Instrument 43-101 and Form 43101F1). Richmont Mines Incorporated, Rouyn-Noranda. Arias, Z., and Helmstaedt, H., 1990. Structural evolution of the Michipicoten (Wawa) greenstone belt, Superior Province, in Geoscience Research Grant Program, summary of research 1989-1990, Edited by V.G.Milne, Ontario Geological Survey, Miscellaneous Paper 150, p. 107-114. Dube, B., and Gosselin, P., 2007. Greenstone-Hosted quartz veins deposits, in Goodfellow, W.D., ed., Mineral Deposits of Canada: A synthesis of Major Deposit Types, District Metallogency, the Evolution of Geological Provinces, and Exploration Methods: Gelogical Association of Canada, Mineral Deposits Division, Special Publication No.5, p.49-73. Heather, K.B., and Arias, Z. 1992. Geological and Structural Setting of Gold Mineralization in the Goudreau-Lochalsh Area, District of Algoma. In Summary of Field Work and Other Activities 1987, by the Ontario Geological Survey. Editied by R.B. Barlow, M.E. Cherry, A.C. Colvine, B.O. Dressler and O.L. White. Ontario Geological Survey, Miscellaneous Paper 137, pp.155-162. Hodgson, C.J. and MacGeehan, P.J., 1982, A review of the geological characterisitics of ‘gold only’ deposits in the Superior Province of the Canadian Shield, in Hodder, R.W. and Petruk, W., eds., Geology of the Canadian Gold Deposits: Canadian Mining and Metallurgy, Special Volume 24, p.211-228. Kwok, Kai-Ming, 1986. Gold Mineralization in Relation to Potassic Alteration, Kremzar property, Finan Township, District of Algoma, Ontario, Canada, unpublished BSc thesis, Laurentian University, Sudbury Ontario, 107p. Lockwood, M.B., 1986. The Petrogenetic and Economic Significance of Chloritoid in the Wawa Greenstone Belt; unpublished Master of Science Thesis, Carelton University, Ottawa, 220p. Percival, J.A., Skulski, T., Sanborn-Barrie, M., Stott, G.M., Leclair, A.D., Corkery, M.T., and Boily, M. 2012. Geology and tectonic evolution of the Superior Province, Canada. Chapter 6 In Tectonic Styles in Canada: The LITHOPROBE Perspective. Editied by J.A. Percival, F.A. Cook and R.M. Clowes. Geological Survey of Canada, Special Paper 49, pp. 321-378. Sage, R.P., 1994. Geology of the Michipicoten Greenstone Belt. Ministry of Northern Development and Mines, Ontario Geological Survey, Open File Report 5888. Turek, A., Smith. P.E., and Van Schmus, W.R. 1984. U-Pb zircon ages and the evolution of the Michipicoten plutonic-volcanic terrane of the Superior Province, Ontario. Canadian Journal of Earth Sciences, 29: 1154-1165. 163 �FIELD TRIP 5 Geology of the Renabie Mine area, Michipicoten greenstone belt Lise Robichaud (Ontario Geological Survey) Jordan McDivitt (Laurentian University) Introduction This field trip focuses on the northeast portion of the Michipicoten greenstone belt and includes the past-producing Renabie Mine, situated within the Missanabie–Renabie gold district (Figure 1). The bedrock geology of the field stops is based on recent mapping of the Renabie Mine area completed by the Ontario Geological Survey (Robichaud, McDivitt and Trevisan 2017; Robichaud et al. 2016; Robichaud et al. 2015; Robichaud, McDivitt and Trevisan 2015; Robichaud and McDivitt 2014). Additionally, detailed outcrop maps for Stops 3, 4, 5, and 15 are modified from a recently completed MSc thesis: Gold mineralization in the Missanabie–Renabie district of the Wawa Subprovince (Missanabie, Ontario, Canada) (McDivitt 2016a). An accompanying Miscellaneous Release of Data (McDivitt 2016b) also provides additional information. Readers should consult publications referenced in this guidebook for additional information and references. The order and number of stops that will be visited during this field trip will be determined by access, time, and weather concerns. Regional Geology The Michipicoten greenstone belt consists of successions of Archean metavolcanic and metasedimentary rocks intruded by Archean granitic rocks (Turek, Smith and Van Schmus 1982) and younger Proterozoic mafic dikes. The supracrustal rocks of the belt have been previously subdivided into 3 distinct volcanic cycles: 2900 Ma, 2750 Ma and 2700 Ma (Sage and Heather 1991; Heather and Arias 1992; Turek, Smith and Van Schmus 1982, 1984; Turek, Van Schmus and Sage 1988). New geochronological data from recent mapping suggest that there is a previously undocumented volcanic cycle ranging in age from 2731 to 2723 Ma (Kamo 2014, 2015, 2016), which indicates that the volcanic history is more continuous than was previously believed. The Michipicoten greenstone belt has been previously interpreted as a continuation of the Abitibi greenstone belt west of the Kapuskasing Structural Zone (Ayer et al. 2010). In comparison, the Abitibi greenstone belt consists of stratigraphically continuous sequences of Archean metavolcanic rocks and metasedimentary rocks ranging from earlier than 2750 Ma to 2695 Ma followed, in Ontario, by 2 sedimentary basins: Porcupine-type at 2690 to 2682 Ma and Timiskaming-type at 2676 to 2670 Ma (Ayer et al. 2002; Ayer et al. 1999a; Ayer et al. 1999b; Ayer, Ketchum and Trowell 2002; Ayer et al. 2005). The second (2750 Ma) and third (2700 Ma) volcanic cycles of the Michipicoten greenstone belt, as well as the 2731 to 2723 Ma volcanism, may represent time-equivalent volcanism similar to the Abitibi greenstone belt; however, no volcanism of 2900 Ma (equivalent to the first cycle) has been described in the Abitibi greenstone belt. Geological Setting of the Renabie Mine Area This field trip targets the northeast portion of the Michipicoten greenstone belt including the pastproducing Renabie Mine. This area is composed of metamorphosed mafic to felsic volcanic and volcaniclastic rocks, interspersed with metasedimentary rocks. Large-bodied, felsic to intermediate granitoids intrude the Archean supracrustal rocks. These include the Missinaibi Lake batholith, which 164 �Figure 1. Simplified geological map of the northeast part of the Michipicoten greenstone belt based on recent mapping by the Ontario Geological Survey (Robichaud, McDivitt and Trevisan 2017; Robichaud et al. 2016; Robichaud et al. 2015). Rennie, Leeson, Stover and Brackin townships are included in the map area. Younging direction symbol indicates stratigraphic facing. Field trip stop locations are indicated on the map. The Renabie road and the Renabie Mine are also denoted for reference purposes. All UTM co-ordinates are provided in Zone 17 using NAD83. 165 �dominates the eastern portion of the area, the Wabatongushi Lake granitoid complex, which occupies the northern portion of Rennie Township, and the Ash Lake pluton, which occupies the southwest portion of Stover Township. Archean felsic stocks, sills, and dikes, as well as Proterozoic gabbroic dikes, are also intrusive to the Archean supracrustal rocks (see Figure 1). The only past-producing gold zones in the Missanabie–Renabie gold district are located in the Renabie Mine area, within tonalite of the Missinaibi Lake batholith near its contact with metavolcanic rocks of the Michipicoten greenstone belt (see Figure 1). The gold zones were mined at the Renabie underground mine and at the C-zone, Nudulama, and Braminco open pits. Mining occurred over a 50 year period, from 1941–1991, and produced roughly 1.1 million ounces of gold at a grade of approximately 6.6 g/t (Turek et al. 1996; Callan and Spooner 1998). The mineralized zones from the Renabie Mine to the Nudulama East outcrop define an east-southeast trend, referred to as the “Renabie trend”, which strikes close to perpendicular to the regional foliation. Assay-defined ore shoots along the Renabie trend have eastern to east-southeastern trends and a moderate westerly plunge (McDivitt 2016a). The Braminco zone, which is located 1.1 km south of the Renabie trend, differs in orientation: the zone is parallel to the regional foliation, which strikes north-northwest and dips steeply to the west. Ore shoots within the Braminco zone plunge moderately to the west, similarly to those along the Renabie trend (Callan and Spooner 1998). Metavolcanic Rocks Mafic metavolcanic rocks are the dominant lithology in the central portion of the area and are interlayered with felsic to intermediate metavolcanic rocks and metasedimentary rocks (see Figure 1). Massive flows are predominant, but pillowed flows are observed locally. The pillows are well to poorly developed and are up to 1 m in diameter. The mafic metavolcanic flows are dark grey to black on fresh surfaces and are typically fine-grained and well foliated. The mafic metavolcanic rocks are typically characterized by greenschist-facies metamorphic mineralogy, although, near the margin of the felsic to intermediate intrusive rocks, the metamorphic mineralogy is representative of amphibolite facies. Where the Archean supracrustal rocks are wedged between the Missinaibi Lake batholith and the Wabatongushi Lake granitoid complex, the rocks are locally migmatized. Felsic to intermediate metavolcanic rocks occur largely in the southern half of Rennie Township and extend south into Stover and Brackin townships (see Figure 1). They are generally fine- to medium-grained volcaniclastic rocks dominated by tuffs and crystal tuffs. Tuff-breccias occur locally in the northern part of Stover Township. The fresh surface of all these felsic to intermediate rocks is light grey and weathers to a lighter grey to beige colour. Bedding is rarely observed and ranges in thickness from finely laminated (a few millimetres) to thickly bedded (up to 1 m). Turek et al. (1996) reported a U/Pb age of 2740±8 Ma for felsic metavolcanic rocks in the southeastern corner of Rennie Township; a U/Pb age of 2730.9±1.2 Ma for the same rocks was reported by Kamo (2016) and Robichaud, McDivitt and Trevisan (2017). A U/Pb age of 2728.6±1.1 Ma was reported for the felsic metavolcanic rocks in Brackin Township (Kamo 2015; Robichaud et al. 2015). A younger U/Pb age of 2704.6±2.1 Ma for the more northern package of felsic metavolcanic rocks, in the north-central portion of Rennie Township was reported by Kamo (2016) and Robichaud, McDivitt and Trevisan (2017). Metasedimentary Rocks Clastic metasedimentary rocks are generally restricted to the southwestern part of Rennie Township, extending to the central part of Stover Township although minor occurrences can also be observed in central Rennie Township. These rocks consist of buff grey, quartz-rich, thinly bedded siltstone, sandstone, and conglomerate. Siltstone is the dominant sedimentary unit, but sporadic beds of sandstone and conglomerate also occur. The conglomerate is matrix-supported, but contains large cobbles of predominantly tonalitic composition with a few mudstone and/or siltstone cobbles. Bedding thickness 166 �varies between sediment types: the siltstone tends to be thinly to thickly laminated, the sandstone beds range from a few decimetres to a metre in thickness, and the conglomerate beds are typically several metres thick. Facing reversals are implied by cross-bedding and graded bedding observed locally in some of the sedimentary rocks, indicating that the area is folded. A U/Pb age of 2695±3 Ma for conglomerates in the northwestern corner of Stover Township was reported by Davis (2016). Iron formation is defined primarily by geophysical evidence (Ontario Geological Survey 1999, 2002a, 2002b, 2003, 2011). Iron formation extends predominantly from the southwestern part of Rennie Township to the central part of Stover Township. Minor iron formation also occurs in east-central Rennie Township and in the southern portion of Brackin Township. The iron formation typically consists of magnetite-rich layers interlayered with siltstone or sandstone, forming beds that range in thickness from thin laminations to thin beds (≤10 cm). Disseminated sulphides were also noted within some of the clastic layers. Archean Intrusions The most extensive Archean intrusion in the map area is the Missinaibi Lake batholith (see Figure 1). It occupies most of the eastern half of the map area and is primarily composed of medium-grained tonalite to granodiorite with minor occurrences of granite. The main mafic mineral is typically biotite, but the presence of hornblende is noted in varying amounts locally. The batholith is generally moderately to well foliated, although some areas are characterized by weak fabric development. Kamo (2015) reports a U/Pb age of 2720.8±1.4 Ma for a tonalitic phase of the Missinaibi Lake batholith in Brackin Township. Turek et al. (1996) reported a U/Pb age of 2741±21 Ma for a tonalitic phase of the Missinaibi Lake batholith in Leeson Township. Younger phases of the Missinaibi Lake batholith occur as plutons, dikes, and stock-like features. The largest occurrence of these younger phases is at the boundary between Leeson and Brackin townships near Crooked Lake (see Figure 1). Younger phases of tonalitic composition tend to be finer grained than the main phase of the batholith and, in some instances, display the regional foliation, although other examples are noted to be massive. A sample of one of these younger tonalitic phases returned an age of 2688±2 Ma (Kamo 2015; Robichaud et al. 2015). The northern portion of Rennie Township is occupied by the felsic intrusive rocks of the Wabatongushi Lake granitoid complex (see Figure 1). These rocks vary in composition from tonalite to granodiorite and typically contain less biotite than rocks from the Missinaibi Lake batholith. There is a 500 m thick boundary between the granitoid complex and the mafic metavolcanic rocks to the south. This boundary is characterized by lit-par-lit injection of the mafic metavolcanic by felsic intrusive rocks. A U/Pb age of 2700±1 Ma for the Wabatongushi Lake granitoid complex was reported by Kamo (2016) and Robichaud, McDivitt and Trevisan (2017). Turek et al. (1996) reported a U/Pb age of 2688 Ma ±14 Ma for the same intrusion. The Ash Lake pluton (see Figure 1) occupies the southwestern portion of Stover Township and is composed primarily of medium-grained tonalite to granodiorite with minor occurrences of granite. The main mafic minerals are typically biotite ± hornblende. The rocks are generally massive to moderately foliated, although the margins of the pluton are characterized by stronger foliation development. Turek et al. (1996) reported a U/Pb age of 2679±5 Ma for the Ash Lake pluton in West Township, situated to the west of Stover Township. Frarey and Krogh (1986) reported a U/Pb age of 2684.5±2.7 Ma for the Ash Lake pluton in Copenace Township, located southwest of Stover Township. The Rennie Lake stock, which occurs in the southwest portion of Rennie Township (see Figure 1), is composed of massive granitic to granodioritic rocks. The main mafic mineral is amphibole with minor biotite. The rocks are light pink on fresh surfaces and weather to creamy pink. A U/Pb age of 2678±4 Ma for the Rennie Lake stock was reported by Davis (2016) and Robichaud, McDivitt and Trevisan (2017). Turek et al. (1996) reported a U/Pb age of ca. 2668 Ma (from a single zircon) for the Rennie Lake stock. 167 �The Ruby Lake stock is interpreted to be temporally related to the Rennie Lake stock. The Ruby Lake stock, which occurs in the southeast portion of Stover Township (see Figure 1), is composed of massive, nonfoliated granitic to granodioritic rocks. The core of the Ruby Lake stock commonly displays potassium feldspar megacrysts, but the texture has also been noted near the intrusion’s contacts. The main mafic mineral is amphibole with minor biotite. The stock is light pink on fresh surfaces and weathers from creamy to dark pink. Turek et al. (1996) reported a U/Pb age of 2661±11 Ma for the Ruby Lake stock. Several smaller mafic and ultramafic intrusions occur in the supracrustal rocks in the area (see Figure 1). The mafic rocks are typically gabbroic in composition and are medium- to coarse-grained, foliated, and typically amphibolitized. The ultramafic rocks differ in composition (from peridotite to pyroxenite and hornblendite) and are variably serpentinized. They are often massive, fine- to coarse-grained, and sometimes display large euhedral crystals of pyroxene or hornblende. Proterozoic Intrusions Proterozoic mafic dikes are late features in the area. Most dikes trend north-northwest; however, a few trend northeast. The Proterozoic dikes generally have sharp, linear aeromagnetic signatures and the dikes shown in Figure 1 are delineated mainly by using airborne magnetic data (Ontario Geological Survey 1999, 2002a, 2002b, 2003, 2011). They are predominantly fine- to medium-grained gabbros, which commonly display diabase texture, and occasionally contain plagioclase phenocrysts or glomerocrysts. The dikes sometimes contain trace amounts of pyrite. Based on their orientation, the north-northwest-trending dikes are interpreted to be part of the Matachewan dike swarm (2454 Ma: Osmani 1991; 2473 Ma: Heaman 1997) and the northeast-trending dikes have been interpreted to be part of the Biscotasing dike swarm (2150 Ma: Osmani 1991). Structural Geology Primary layering (S0) is observed in the form of sedimentary bedding and pillowed flows. Asymmetrical layering, such as graded beds and cross-beds, are sometimes observed within the area. In the eastern portion of the map area, along the boundary with the Missinaibi Lake batholith, facing is to the west. In the northern mafic metavolcanic rocks, just south of the Wabatongushi Lake granitoid complex, facing is to the southwest. The clastic metasedimentary rocks in the southwest corner of Rennie Township face northeastward. The earliest foliation (S1) is only locally defined and is consistently overprinted by the dominant regional foliation (S2). Where it is observable, the S1 foliation is axial planar to F1 folds, which are defined by mineralized laminated quartz veins. Both the S1 foliation and F1 folds are varied in orientation resulting from effects of later overprinting deformation. The S1 foliation and F1 folds are overprinted by tight, generally upright, F2 folds. The regional S2 foliation is axial planar to the F2 folds and is northwest to north-northwest striking with a steep westerly dip in the Renabie Mine area. At the Pileggi No. 1 outcrop, in the northwest corner of Stover Township, the S2 foliation is subvertical and strikes west-northwest (Figure 1). The S2 foliation is well-developed in the Missinaibi Lake batholith, where it is best identified by differentiated domains of biotite, and in supracrustal rocks of the Michipicoten greenstone belt, where it parallels primary layering (S0). The S2 foliation also parallels the intrusive contacts of the supracrustal rocks with the Missinaibi Lake batholith and Ash Lake pluton. A regional L2 mineral stretching lineation is developed in association with the S2 foliation. In the Renabie Mine area the L2 lineation parallels the orientation of ore shoots and has a moderate to steep westerly plunge. 168 �Late, east- to east-southeast-striking, steeply south-dipping shear zones overprint the S2 foliation along the Renabie trend. These shear zones are host to the gold-bearing laminated quartz veins that were the main focus of gold production in the Renabie Mine area. The shear zones formed during a D3 deformation event and are characterized by a strong internal S3 foliation developed within the quartz-sericite-pyrite alteration halos that surround the laminated veins. Shear sense indicators record sinistral-reverse, oblique slip movement along the shear zones during the D3 event. The mineralized D3 shear zones are exposed along the Renabie trend from west to east at the C-Zone, Nudulama, and Nudulama East outcrops (McDivitt 2016a) (see Figure 1). Within the shear zones along the Renabie trend, S3 defines east-northeast-trending, steeply plunging Zshaped drag folds. These folds record a late dextral transcurrent reactivation of the shear zones during a D4 deformation event. Similar Z-shaped F4 folds overprint the S2 foliation at the Pileggi No. 1 outcrop (see Figure 1), suggesting that a late dextral reactivation of mineralized zones with an easterly trend is a recurrent feature in the district (McDivitt 2016a). The study area was also affected by late brittle faulting (see Figure 1). The presence of these late faults is primarily supported by offsets revealed through mapping, topography, and by geophysical interpretation. There are several major faults interpreted in the area, some of which show displacements of up to 1 km. A major sinistral fault trends north-northwest from the southeast corner of Stover Township to the southcentral edge of Rennie Township, transecting the Ruby Lake stock and showing a displacement of approximately 1 km. Another major fault runs underneath Crooked Lake and shows dextral displacement of just under 1 km; however, it does not transect the Ruby Lake stock. There are several faults extending southwestward from the northwest corner of Leeson Township towards the Rennie Lake stock. Some show dextral displacement of 1 km, whereas others do not display obvious offsets. 169 �Field Trip Stops Notes: 1) Some of these outcrops are adjacent to open pits; please refrain from crossing fenced areas to access outcrop. 2) This group is unlikely to visit all of the stops described below during this trip. A number of stop descriptions are included in this guide to provide context and for future visits to the region. Permission must be granted from mining companies for future access. 3) The order and number of stops that will be visited during this field trip will be determined by access, time, and weather concerns. These factors may preclude visiting some stops, and alternative locations may be substituted. 4) All locations are given in NAD 83, Zone 17N UTM coordinates. Stop 1: Braminco open pit Location: UTM Zone 17, 288468E 5360854N The Braminco open pit (see Figure 1) is hosted in the Missinaibi Lake batholith near the contact with the mafic metavolcanic rocks of the Michipicoten greenstone belt. It consists of a shear zone that is 10-15 m wide, striking 135-145° and dipping moderately to the southwest (Sage and Heather 1991). Mineralization at Braminco is similar in character to mineralization along the Renabie trend and consists of shear zone-hosted laminated quartz veins associated with quartz-sericite-pyrite alteration zones and late, hematite-bearing alteration assemblages. However, the Braminco shear zone differs in orientation from the Renabie trend shear zones, which are more easterly in trend. The Braminco shear zone is interpreted to have formed as a sinistral D2 shear zone that underwent a dextral reactivation during D3 (McDivitt 2016a). It was mined primarily for its silica, but gold was also recovered (Wilson, A., OGS, personal communication, 2012). Stop 2: Baltic D outcrop Location: UTM Zone 17, 289990E 5360020N The Baltic D outcrop (see Figure 1) contains gold-bearing laminated quartz veins (Photo 1A) surrounded by sericitized, sulphide-bearing wallrock (Photo 1B) that formed after biotite tonalite of the Missinaibi Lake batholith. The laminated quartz veins are shallow in their dip and define isoclinal, near-recumbent F1 folds. A continuous S1 cleavage is axial planar to the folds and is well-developed within the sericitized wallrock. The F1-folded veins and S1 cleavage are refolded by upright, open, northwest-striking F2 folds (McDivitt 2016a). The regional S2 foliation is axial planar to the F2 folds. The S2 foliation strikes northwest and dips steeply west; it is well-developed in the biotite tonalite surrounding the sericitized wallrock marginal to the veins. The regional L2 stretching lineation occurs in association with S2 foliation and the F2 folds. The lineation plunges moderately northwest, subparallel to the hinges of the F2 folds (McDivitt 2016a). 170 �A B Photo 1. A) Overview photograph of the Baltic D outcrop displaying shallowly-dipping laminated quartz veins. Compass for scale is 6.9 cm wide, with sighting arm pointing north. B) Laminated quartz vein (LQV) adjacent to quartz-sericite-pyrite (QSP) schist, within which S1 is developed. Pencil for scale is 0.8 cm wide and points north. Stop 3: Nudulama East outcrop Location: UTM Zone 17, 289028E 5361881N The Nudulama East outcrop (Figure 2) is the easternmost outcrop defining the Renabie trend (see Figure 1). The outcrop is situated within the Missinaibi Lake batholith, approximately 2 km to the east of the contact with the Michipicoten greenstone belt. Mineralization is characterized by laminated quartz veins (Photo 2A) hosted within steeply dipping, east-southeast-trending D3 shear zones. The shear zones are defined by strongly foliated (S3) quartz-sericite-pyrite schist. Both the laminated veins and schist are mineralized. The veins typically contain 2-5% pyrite with minor chalcopyrite, molybdenite, galena, sphalerite, and rare visible gold. The surrounding schist contains similar amounts of pyrite, but other sulphide minerals are rare to absent. In addition to the laminated veins, late quartz ± chlorite ± epidote ± potassium-feldspar veins associated with hematite-bearing alteration assemblages occur at Nudulama East (Photo 2B, C). Where these veins are abundant, biotite within the host tonalite is pervasively chloritized, and the rock displays a distinct reddish hue because of the presence of hematite. A large quartz-chlorite breccia body is associated with the hematite-bearing alteration zone and cuts across the southern mineralized shear zone (Photo 2D, E). The breccia contains clasts of laminated quartz veins and sheared, hematite-altered tonalite enclosed within a matrix of vuggy quartz locally infilled by chlorite (Photo 2F) (McDivitt 2016a). At Nudulama East, the S2 foliation is oriented roughly perpendicular to the trend of the mineralized D3 shear zones. The S2 foliation and F2 folds defined by laminated quartz veins and felsic dikes are overprinted by F3 folds and transposed into parallelism with the D3 shear zones. The F3 folds have an axial planar S3 foliation, which becomes the main foliation within the shear zones. Mutual overprinting relationships observed between hematite-associated veins and late D3 structures suggest that the hematiteassociated veins are syn-D3 in timing (McDivitt 2016a). Neither the hematite-associated veins nor the quartz-chlorite breccia contain significant gold mineralization. 171 �Figure 2. Detailed geological map of Nudulama East, modified from McDivitt (2016). Lower-hemisphere, equalarea stereographic plots showing the L2 stretching lineation and poles to the S2 and S3 foliations. n = number of measurements. 172 �A B C D E F Photo 2. Field photographs of the Nudulama East outcrop. A) Laminated quartz vein. Pencil for scale is 0.8 cm wide and points north. B) L2-lineated, hematite-bearing altered tonalite. Pencil for scale is 0.8 cm wide and points north. C) Planar hematite-association alteration vein containing quartz-epidote infill mineralogy. Pencil for scale is 0.8 cm wide and points north. D) Hydrothermal breccia with fragments of foliated, hematite-altered tonalite surrounded by a quartz-chlorite matrix. E) A well-defined contact between a laminated quartz vein (LQV) vein within a quartz-sericite-pyrite shear zone, and a quartz-breccia vein containing clasts of hematite-bearing altered tonalite. Compass for scale is 7.1 cm wide, with sighting arm pointing north. F) Irregular-shaped hematiteassociated alteration veins with quartz-chlorite infill mineralogy. Well-developed hematite-bearing alteration zones surrounds the veins. Pencil for scale is 0.8 cm wide and points north. 173 �Stop 4: Nudulama – surrounding outcrop Location: UTM Zone 17, 288623E 5361923N The Nudulama outcrop (Figure 3) is located approximately 400 m west of the Nudulama East outcrop along the Renabie trend (see Figure 1). The outcrop is situated within the Missinaibi Lake batholith, approximately 1.5 km to the east of the contact with the Michipicoten greenstone belt. Mineralization at Nudulama is characterized by laminated quartz veins within an approximately 10 m wide D3 shear zone that consist of strongly foliated quartz-sericite-pyrite schist (Photo 3A, B) (McDivitt 2016a). The laminated veins comprise very fine-grained saccharoidal quartz layers separated by thin sericitic laminations. The veins typically contain up to 5% pyrite with minor chalcopyrite, molybdenite, galena, sphalerite, and rare visible gold; the schist contains similar quantities of pyrite, but a relative absence of other sulphide minerals is noted. Both the laminated veins and schist consistently return significant gold values (i.e., >1 ppm), with higher grade samples containing in excess of 10 ppm gold. Late quartz ± chlorite ± epidote ± potassium-feldspar veins associated with hematite-bearing alteration assemblages occur at Nudulama, but they are less abundant than at Nudulama East. Steeply dipping, Z-shaped drag folds that trend east-northeast are defined by the S3 foliation within the mineralized shear zone. These folds formed during a late dextral transcurrent reactivation of the shear zone during a D4 deformation event (McDivitt 2016a). Figure 3. Detailed geological map of Nudulama, modified from McDivitt (2016). Lower-hemisphere, equal-area stereographic plots of the L2 stretching lineation and poles to S2 and S3 foliations. n = number of measurements. 174 �A B Photo 3. Field photographs of the Nudulama outcrop. A) Mineralized zone shows a laminated quartz vein surrounded by quartz-sericite-pyrite schist. Compass for scale is 6.9 cm wide, with sighting arm pointing north. B) East-facing, vertical wall at the Nudulama open pit (~10 m wide), displaying the same mineralized zone as in A. Photo taken looking north. Stop 5: C-Zone outcrop Location: UTM Zone 17, 287890E 5362050N The C-Zone outcrop (Figure 4) is located approximately 750 m west of the Nudulama outcrop along the Renabie trend (see Figure 1). The outcrop occurs within the Missinaibi Lake batholith, approximately 750 m to the east of the contact with the Michipicoten greenstone belt. Mineralization at the C-Zone outcrop is similar in character to mineralization at Nudulama and Nudulama East outcrops: it consists of laminated quartz veins (Photo 4A, B) hosted within an east-trending, steeply dipping D3 shear zone of variable width (5-10 m). The shear zone consists of strongly foliated (S3) quartz-sericite-pyrite schist. Both the laminated veins and the schist are mineralized. The schist contains 2-5% pyrite; the veins contain similar amounts of pyrite but can also have small amounts of chalcopyrite, molybdenite, galena, sphalerite, and rare visible gold (McDivitt 2016a). A Photo 4. Field photographs of the C-Zone outcrop. A) The main laminated quartz vein. The hammer head points to the east (the hammer is 70 cm long). B) A close up photograph of a laminated quartz vein. Compass for scale is 6.9 cm wide, with sighting arm pointing north. B 175 �Figure 4. Detailed geological map of the C-Zone, modified from McDivitt (2016). Lower-hemisphere, equal-area stereographic plots of the L2 stretching lineation and poles to S2 and S3 foliations. n = number of measurements. Stop 6: No. 2 Shaft and contact outcrop Location: UTM Zone 17, 287246E 5361990N The site of the former Renabie Mine is the westernmost outcrop defining the Renabie trend. At the now defunct site of the No. 2 shaft (see Figure 1), the Shaft Fault offsets the contact between the Archean supracrustal rocks of the Michipicoten greensone belt and the Missinaibi Lake batholith. The fault is not observed at surface and was primarily defined from underground workings. This fault, and others parallel to it, are defined as late, north-trending brittle faults that cut across the Renabie trend. These faults generally appear to have a sinistral offset. The Shaft Fault has been interpreted as an oblique-slip fault having a sinistral strike-slip component as well as a west-side-up (reverse), dip-slip component (Callan and Spooner 1998). On the hill to the south of the No. 2 Shaft, the contact between the mafic metavolcanic rocks of the Michipicoten greenstone belt and tonalite of the Missinaibi Lake batholith can be observed. The contact trends roughly 140°, dips 70° to the southwest and parallels the regional S2 foliation. 176 �Stop 7: Ultramafic metavolcanic rocks Location: UTM Zone 17, 284487E 5366237N Ultramafic metavolcanic rocks (Photo 5) (see Figure 1) are seldom observed in the area. This outcrop displays talc and chlorite alteration and also contains trace amounts of pyrite. The foliation parallels the contact of the supracrustal rocks with the Wabatongushi Lake granitoid complex, which is situated roughly 1.4 km north of the outcrop. Photo 5. Highly altered ultramafic metavolcanic rocks. Stop 8: VMS-style alteration Location: UTM Zone 17, 285343E 5364967N This is a steep outcrop located on the west side of the road (see Figure 1). The main lithology consists of felsic tuffs (Photo 6A) with areas of crystal tuff. At the top of the exposure a thick (~10-20 cm) band of coarse-grained amphibole and garnet can be observed (Photo 6B). Disseminated pyrite can also be found within the alteration area. B A Photo 6. A) Felsic metavolcanic rocks. B) Amphibole and garnet alteration with disseminated pyrite. 177 �Although geochemical analysis did not return particularly good assay values for this outcrop, it does bear some similarities with the Conboy Lake occurrence (MDI42B05NW00021; Ontario Geological Survey 2017). The Conboy Lake occurrence (282745E 5366813N) is located in north-central Rennie Township, approximately 3 km to the northwest of this location. It is an historic zinc, silver and copper occurrence with minor gold mineralization, and is classified in the OGS Mineral Deposit Inventory (MDI) as a developed prospect with reserves. It has a long history of sporadic mineral exploration between 1939 and 2010, as summarized in Ontario Geological Survey (2017). The Conboy Lake occurrence is found within younger felsic metavolcanic rocks (2704.6±2.1 Ma; Kamo 2016; Robichaud, McDivitt and Trevisan 2017) that display pervasive sericite alteration. The zinc mineralization occurs as thick layers of massive sphalerite with disseminated pyrite and chalcopyrite (Robichaud, McDivitt and Trevisan 2015; Riley 1971, p.45-49; Ontario Geological Survey 2017). Stop 9: Felsic crystal tuff Location: UTM Zone 17, 285440E 5362893N The felsic crystal tuffs (see Figure 1) at this outcrop are typical of the area. They are generally fine- to medium-grained with coarser crystals of either quartz or plagioclase. The fresh surface is light grey and weathers to a lighter grey to beige colour. Bedding is rarely observed in the felsic metavolcanic rocks and this outcrop is no different. This outcrop was sampled for geochronology in 2015, with a U/Pb age of 2730.9±1.2 Ma reported by Kamo (2016). Stop 10: Felsic tuff-breccia Location: UTM Zone 17, 283838E 5361108N Felsic tuff-breccias (Photo 7) (see Figure 1) are seldom observed in the area and are very similar to the conglomerates, but differ in their monolithic clast composition. The clasts are composed of felsic metavolcanic material and are angular. In some areas, the clast edges seem to be broken in situ, making it a mosaic breccia. Photo 7. Felsic tuff breccia. Compass for scale is 7.1 cm wide, with sighting arm pointing north. 178 �Stop 11: Pillowed mafic flows Location: UTM Zone 17, 279434E 5361230N This outcrop (see Figure 1) is located approximately 50 m to the north of the Renabie road; it may need to be cleared of the brush to see the features. The outcrop consists of mafic lava flows. Pillows can be observed, ranging in size from 30 to 60 cm. Stop 12: Rennie Lake turnoff conglomerate Location: UTM Zone 17, 279434E 5361230N This polymictic conglomerate is situated on the Renabie road (Photo 8). The conglomerate is matrix supported, but contains large cobbles of predominantly tonalitic porphyry with a few mudstone and siltstone cobbles. The conglomerate is highly foliated and the clasts show an elongation fabric that parallels bedding. Cross-bedding is observed in the more sandy layers and indicates younging to the northeast. A U/Pb age of 2695±3 Ma for the sandy portion of this outcrop was reported by Davis (2016). Photo 8. Highly foliated conglomerate. Compass for scale is 7.1 cm wide, with sighting arm pointing north. Stop 13.1: Baltimore Lake metasedimentary rocks and diabase dike Location: UTM Zone 17, 279076E 5361116N The Baltimore Lake metasedimentary rocks consist of buff-grey, quartz-rich, thinly bedded siltstone intercalated with lesser sandier layers. Small staurolite grains can be observed in some areas. Graded bedding (Photo 9A) can be observed on the outcrop and indicates facing is to the northeast. The diabase dike is located at the southern end of the outcrop, near the water. The dike is interpreted to be of the Matachewan swarm because of its north-trending orientation. The contact between the metasedimentary rocks and the dike can be observed near the water’s edge. The dike contains numerous plagioclase phenocrysts, and in one area, several large glomerocrysts (Photo 9B). On the southern shore of Baltimore Lake, the metasedimentary rocks are more deformed. They are folded, making primary features difficult to discern (Photo 9C, D). The metasedimentary rocks are predominantly sandstones that are finely interbedded with silty layers, but some of the sandstone beds are up to 1 m in thickness. Conglomerates can also be observed on the southern shore of the lake. 179 �A B C D Photo 9. Field photographs of the Baltimore Lake area. A) Graded bedding indicates facing is to the top of the photo. Top of photo is to the northeast. B) Matachewan dike with plagioclase glomerocrysts. C) Interbedded sandy and silty layers showing parasitic axial planar folding. D) Large synclinal fold. Compass for scale is 7.1 cm wide, with sighting arm pointing north. Stop 13.2: Baltimore Lake metasedimentary rocks Location: UTM Zone 17, 278836E 5361004N Approximately 250 m west of the last outcrop sits another large exposure of the same metasedimentary rocks. A very well-polished surface occurs in the middle of the Renabie road and offers the best exposure; however, there are many areas of interest across the outcrop. The metasedimentary rocks consist of buffgrey, quartz-rich, thinly bedded siltstone and sandstone. Siltstone is the dominant sedimentary unit, but sporadic beds of sandstone also occur. The siltstone tends to be thinly to thickly laminated, and the sandstone ranges from a few decimetres to 1 m in thickness. Stop 14.1: Iron formation Location: UTM Zone 17, 277276E 5361723N From the Renabie road, go north past the quarry for approximately 1.3 km until a small east-trending trail is reached. Follow this trail for 1 km on foot (or by ATV); the outcrop will be on the south side of the trail. 180 �The iron formation consists of magnetite-rich layers interbedded with sandstone (Photo 10A, B). Folds can be observed within the banded iron formation, and are especially visible in the central cavity where the third dimension can be observed. Disseminated pyrite and chalcopyrite are also noted. B A Photo 10. A) Banded iron formation displaying magnetite-rich layers interbedded with sandstone. B) Folded banded iron formation. Stop 14.2: Iron formation Location: UTM Zone 17, 276889E 5359035N This iron formation is more readily accessible than that at the previous stop, but is not as well-exposed. This outcrop is located approximately 100 m down the road leading to the Pileggi No. 1 outcrop (Figure 1). There is also another small exposure 200 m further down the road. The outcrop is dominated by the presence of gabbro; the iron formation occupies a small portion (< 1 m2) of this outcrop. The iron formation consists of magnetite-rich layers interbedded with sandstone. Stop 15: Pileggi No. 1 outcrop Location: UTM Zone 17, 277677E 5358487N The Pileggi No. 1 outcrop (Figure 5) occurs approximately 120 m north of the main road (Figure 1). There is a small access trail that can be used to walk up to the outcrop. The outcrop consists of 2 separate exposures (A and B) that connect over a small ridge (see Figure 5 inset). The outcrop consists of metavolcanic rocks of the Michipicoten greenstone belt that have been intruded by feldspar-phyric and amphibole-phyric dikes. Matachewan dikes intrude the metavolcanic rocks. Gold mineralization occurs within deformed laminated quartz veins that returned values up to 22.8 ppm. The laminated quartz veins are overprinted by isoclinal F1 folds associated with an S1 axial planar cleavage. The F1 folds and S1 cleavage formed during an early deformation event, which predated the development of the regional S2 foliation (Photo 11A). The S2 foliation (Photo 11B), which strikes 181 �182 Figure 5. Detailed geological maps of Pileggi No. 1 north (A) and south (B) exposures, modified from McDivitt (2016). Inset map shows the relative location of the two exposures. �approximately 100° and dips steeply, is axial planar to an upright, shallowly plunging F2 fold that occupies roughly half the width of the outcrop. The F2 fold overprints the F1 folded laminated veins and the S1 cleavage. Parasitic folds in the hinge and limbs of the larger F2 fold vary in plunge from subhorizontal to moderately plunging. The S2 foliation is manifested as a disjunctive or crenulation cleavage in the hinges of F2 folds, and as a slaty transposition cleavage along the limbs of the folds. Zshaped drag folds and northwest-trending, subvertical quartz tension gashes overprint the S2 foliation and the laminated veins (Photo 11C). The Z-shaped drag folds trend east-northeast and are steeply plunging. The quartz tension gashes are either straight or Z-shaped sigmoidal. The drag folds and the quartz tensions gashes are attributed to a late dextral reactivation of the mineralized zone during the D4 deformation event (McDivitt 2016a). The variable shape of the tension gashes indicates that they were emplaced throughout the D4 event, with the Z-shaped tension gashes being older than the straight ones. Unlike the laminated veins, the late tension gashes are not gold bearing. B A C Photo 11. Field photographs of the Pileggi No. 1 outcrop. A) Laminated quartz vein overprinted by isoclinal F1 fold and tight F2 folds. The S1 and S2 cleavages are axial planar to F1 and F2 folds, respectively. Facing is to the west. B) F2 folds defined by laminated quartz veins. Note the opposing plunge directions of some of the F2 fold hinges. Pencil for scale is 0.8 cm wide and points north. C) A north-west trending quartz-tension gash. Marginal to the vein, S2 defines a Z-fold, and the vein itself displays Z-asymmetry; both features are supportive of vein emplacement during dextral shearing. Pencil for scale is 0.8 cm wide and points north. Acknowledgements Particular thanks are extended to Ann Wilson and Joseph Walker, for agreeing to lead the field trip in the absence of the authors. The authors appreciates all the support from the staff of the Sault Ste. Marie and Timmins Resident Geologist offices, in particular Anthony Pace and Ann Wilson. Patrick Gervais is thanked for editing figures for this guidebook. Sonia Préfontaine and Marg Rutka are thanked for reviewing the text. 183 �REFERENCES Ayer, J., Amelin, Y., Corfu, F., Kamo, S., Ketchum, J., Kwok, K. and Trowell, N. 2002. Evolution of the southern Abitibi greenstone belt based on U-Pb geochronology: Autochthonous volcanic construction followed by plutonism, regional deformation and sedimentation; Precambrian Research, v.115, p.63-95. Ayer, J.A., Goutier, J., Thurston, P.C., Dubé, B. and Kamber, B.S. 2010. Tectonic and metallogenic evolution of the Abitibi and Wawa subprovinces; in Summary of Field Work and Other Activities 2010, Ontario Geological Survey, Open File Report 6260, p.3-1 to 3-6. Ayer, J.A., Ketchum, J.W.F. and Trowell, N.F. 2002. New geochronological and neodymium isotopic results from the Abitibi greenstone belt, with emphasis on the timing and the tectonic implications of Neoarchean sedimentation and volcanism; in Summary of Field Work and Other Activities 2002, Ontario Geological Survey, Open File Report 6100, p.5-1 to 5-16. Ayer, J.A., Thurston, P.C., Bateman, R., Dubé, B., Gibson, H.L., Hamilton, M.A., Hathway, B., Hocker, S.M., Houlé, M.G., Hudak, G.J., Ispolatov, V., Lafrance, B., Lesher, C.M., MacDonald, P.J., Péloquin, A.S., Piercey, S.J., Reed, L.E. and Thompson, P.H. 2005. Overview of results from the Greenstone Architecture Project: Discover Abitibi Initiative; Ontario Geological Survey, Open File Report 6154, 125p. Ayer, J.A., Trowell, N.F., Amelin, Y. and Corfu, F. 1999a. Geological compilation of the Abitibi greenstone belt in Ontario: Toward a revised stratigraphy based on compilation and new geochronology results; in Summary of Field Work and Other Activities 1998, Ontario Geological Survey, Miscellaneous Paper 169, p.14-24. Ayer, J.A., Trowell, N.F., Madon, Z., Kamo, S., Kwok, Y.Y. and Amelin, Y. 1999b. Compilation of the Abitibi greenstone belt in the Timmins–Kirkland Lake area: Revisions to stratigraphy and new geochronological results; in Summary of Field Work and Other Activities 1999, Ontario Geological Survey, Open File Report 6000, p.4-1 to 4-14. Callan, N.J. and Spooner, E.T.C. 1998. Repetitive hydraulic fracturing and shear zone inflation in an Archean granitoidhosted, ribbon banded, Au-quartz vein system, Renabie area, Ontario, Canada; Ore Geology Reviews, v.12, p.237-266. Davis, D.W. 2016. Geochronology of rocks from northwest Ontario 2015-16. Part 2: LA-ICP-MS Geochronology, internal report for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 100p. Frarey, M.J. and Krogh, T.E. 1986. U-Pb zircon ages of late internal plutons of the Abitibi and eastern Wawa subprovinces, Ontario and Quebec; in Current Research, Part A, Geological Survey of Canada, Paper 86-1A, p.43-48. Heaman, L.M. 1997. Global mafic magmatism at 2.45 Ga: Remnants of an ancient large igneous province?; Geology, v.25, p.299-302. Heather, K.B. and Arias, Z. 1992. Geological and structural setting of gold mineralization in the Goudreau–Lochalsh area, Wawa gold camp; Ontario Geological Survey, Open File Report 5832, 159p. Kamo, S.L. 2014. Report on U-Pb CA-ID-TIMS geochronology on volcanic and plutonic rocks, Superior and Grenville provinces, Ontario; internal report for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 27p. ——— 2015. Report on U-Pb ID-TIMS geochronology on volcanic and plutonic rocks, Superior and Grenville provinces, Ontario; internal report for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. ——— 2016. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: Bedrock mapping projects, Ontario, internal report for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. 184 �McDivitt, J.A. 2016a. Gold mineralization in the Missanabie–Renabie district of the Wawa Subprovince (Missanabie, Ontario, Canada); unpublished MSc thesis, Laurentian University, Sudbury, Ontario, 175p. ——— 2016b. Gold mineralization in the Missanabie–Renabie district of the Wawa Subprovince: Geochemical data and photographs; Ontario Geological Survey, Miscellaneous Release—Data 339. Ontario Geological Survey 1999. Single master gravity and aeromagnetic data for Ontario, Geosoft® format; Ontario Geological Survey, Geophysical Data Set 1036. ——— 2002a. Ontario airborne geophysical surveys, magnetic data, grid data, Geosoft® format, Kapuskasing– Chapleau area; Ontario Geological Survey, Geophysical Data Set 1040b—Revised. ——— 2002b. Ontario airborne geophysical surveys, magnetic data, profile data, Geosoft® format, Kapuskasing– Chapleau area; Ontario Geological Survey, Geophysical Data Set 1040d—Revised. ——— 2003. Ontario airborne geophysical surveys, magnetic and electromagnetic data, grid and profile data, Geosoft® format, Wawa area; Ontario Geological Survey, Geophysical Data Set 1009b. ——— 2011. Ontario airborne geophysical surveys, magnetic and electromagnetic data, grid and profile data (ASCII and Geosoft® formats) and vector data, Magpie River–Missinaibi Lake area—Purchased data; Ontario Geological Survey, Geophysical Data Set 1237. ——— 2017. Mineral Deposit Inventory; Ontario Geological Survey, Mineral Deposit Inventory (March 2017 update), online database. Osmani, I.A. 1991. Proterozoic mafic dike swarms in the Superior Province of Ontario; Chapter 17 in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.661-681. Riley, R.A. 1971. Glasgow, Meath, and Rennie townships, Algoma and Sudbury districts, Ontario; Ontario Geological Survey, Map 2210, scale 1:31 680. Robichaud, L. and McDivitt, J. 2014. Geology and mineral potential of Brackin Township, Michipicoten greenstone belt; in Summary of Field Work and Other Activities 2014, Ontario Geological Survey, Open File Report 6300, p.51 to 5-11. Robichaud, L., McDivitt, J.A., Krystopowicz, N.J. and Szumylo, N. 2015. Precambrian geology of Brackin Township, Michipicoten greenstone belt; Ontario Geological Survey, Preliminary Map P.3797, scale 1:20 000. Robichaud, L., McDivitt, J.A. and Trevisan, B.E. 2015. Geology and mineral potential of Rennie and Leeson townships, Michipicoten greenstone belt; in Summary of Field Work and Other Activities 2015, Ontario Geological Survey, Open File Report 6313, p.5-1 to 5-11. ——— 2017. Precambrian geology of Rennie and Leeson townships, Michipicoten greenstone belt; Ontario Geological Survey, Preliminary Map P.3803, scale 1:20 000. Robichaud, L., Walker, J., West, S.M. and Nywening, A. 2016. Geology and mineral potential of Stover Township, Michipicoten greenstone belt; in Summary of Field Work and Other Activities, 2016, Ontario Geological Survey, Open File Report 6323, p.5-1 to 5-10. Sage, R.P. and Heather, K.B. 1991. The structure, stratigraphy and mineral deposits of the Wawa area; Geological Association of Canada–Mineralogical Association of Canada–Society of Economic Geologists, Joint Annual Meeting, Toronto 1991, Field Trip A6, 118p. Turek, A., Heather, K.B., Sage, R.P. and Van Schmus, W.R. 1996. U–Pb zircon ages for the Missanabie–Renabie area and their relation to the rest of the Michipicoten greenstone belt, Superior Province, Ontario, Canada; Precambrian Research, v.76, p.191-211. 185 �Turek, A., Smith, P.E. and Van Schmus, W.R. 1982. Rb–Sr and U–Pb ages of volcanism and granite emplacement in the Michipicoten belt, Wawa, Ontario; Canadian Journal of Earth Sciences, v.19, p.1608-1626. ——— 1984. U/Pb zircon ages and the evolution of the Michipicoten plutonic–volcanic terrane of the Superior Province, Ontario; Canadian Journal of Earth Sciences, v.21, p.457-464. Turek, A., Van Schmus, W.R. and Sage, R.P. 1988. Extended volcanism in the Michipicoten greenstone belt, Wawa, Ontario; abstract in Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, St. John’s 1988, Program with Abstracts, v.13, p.A127. 186 �Field Trip 6 Kapuskasing Structural Zone and the Borden Lake Gold Deposit Pierre Bousquet M.Sc., P.Geo. Resident Geologist Program, Ontario Geological Survey, Timmins, Ontario and Jason Rickard M.Sc., P.Geo Senior Geologist for Goldcorp Inc.; Timmins, Ontario Introduction The Kapuskasing Structural Zone is a window into the continental crust, which was investigated thoroughly during the past 30 years. Due to its nature, it was suspected to be barren of any mineralization typical of greenstone belts. Recent findings raised more questions about it, and gave geoscience a black eye, especially after the discovery of the Borden Lake gold deposit. This field trip will be an overview of the Kapuskasing Structural Zone (KSZ). The KSZ was the subject of an ILSG field trip back in 1987 (Percival 1987). This field trip will take place within reasonable distance of Highway 101. As a precautionary note, be aware of the traffic when exiting the vehicles and while crossing or walking along the highway or any roads. Field trip participants should also use caution on outcrops and ditches since they might prove slippery and steep. Exploration and mining history Early explorers were mostly looking at the greenstone belts on either side of the Kapuskasing Structural Zone. However, within the zone, exploration was mostly aimed at the carbonatite intrusions. In the late 1950’s, the Nemegosenda Carbonatite Complex was explored by the Dominion Gulf Company. The company performed an aeromagnetic survey which showed a circular anomaly on the east side of Nemegosenda Lake. This anomaly was, in fact, the Nemegosenda Alkaline Complex. Through clever use of a ground magnetic survey in an area where outcrops were scarce, the Dominion Gulf Company singled out highs which required follow-up drilling. From 1954 to 1959, the company drilled over a hundred holes and even opened an adit of 190 metres into the main mineralized zone (D-Zone) which contained niobium (Archibald, 2008). The company also completed diamond drilling in the South-East Area, located 2 km southeast from the adit, which shows endowment in rare earth elements. The property was taken over by Musto Explorations Ltd. in 1987, who pursued the exploration on the South-East Area by retrieving core drilled in the fifties, split and re-assayed the core. The results were similar to those obtained in the fifties. In the early nineties, Placer Explorations did a re-plotting of the D zone and South East Area. These plots were used to calculate a block model of the D zone. In 2008, Sarissa Resources Inc. acquired the Nemegosenda project, and re-evaluated the historic data. Sarissa Resources followed up with a series of diamond drill holes to confirm its location in the D zone. In the fall of 2016, the project was sold to Indo Global Exchange(s) Pte, Ltd. 187 �The Lackner Lake carbonatite alkalic complex is another area which has undergone exploration in the past sixty years. Two patented claims, in McNaught Township, uncovered magnetite mineralization discovered over a century ago. In 1949, prospectors discovered strong radioactivity on these claims. Nemegos Uranium Corporation was formed to explore these and nearby claims. The exploration concentrated on developing apatite, magnetite, and uranium ore deposits. In 1951, an aeromagnetic survey was flown by Dominion Gulf Company, who followed up with staking anomalies located on the east flank of the structure. At the same time, niobium was discovered in samples assayed by Nemegos Uranium Corporation, but the results were not followed up. In 1954, Multi-Minerals Limited demonstrated the extensive niobium-bearing mineralization on the claims formerly owned by Nemegos Uranium Corporation. The company trenched and drilled in the southwest corner of the complex, which uncovered three zones of niobium mineralization. In 1970, Multi-Minerals Limited optioned one zone to Fetio Industrial Developments Limited who proceeded with an evaluation of possible production of titanium-iron and phosphate concentrates. The company shipped 1500 tons of concentrate to the United State for metallurgical testing. The property was then optioned to Mertec Resources Development in 1974, which was terminated in 1975. During 1978 to 1979, Multi-Minerals Limited drilled 6 holes totalling 2176 feet to test uranium potential of various mineralized zones outlined by previous work. 6378366 Canada Inc. and 6070205 Canada Inc. went back to compile data on the Lackner Lake carbonatite in 2008. A series of exploration projects were undertaken, including prospecting, sampling, ground radiometric survey and assaying. It was followed by an aerial magnetic and radiometric survey by Rare Earth Metals Inc. in 2009. The most recent work was completed in 2014 by Gold Crossing Limited and International Explorers & Prospectors Inc. who completed radiometric surveys, geological mapping, sampling, and assaying of rock and past cores. The Shawmere Anorthosite complex was explored back in the 1990’s. Purechem Ltd. did geological mapping, geochemistry, petrography and bulk sampling from 1993 to 1999 for aluminium silica. The property was then purchased by Southern Africa Corporation in 2001, which bulk sampled and performed more tests. In 2002-2007, Avalon Ventures Ltd. acquired the property and completed bulk sampling and sent more than 1000 tons of material for furnace trials and processing of anorthosite. The company, now under the name of Avalon Advanced Metals Inc., still holds the property. The Borden Lake gold deposit was initially explored in the early 1990’s by M. Tremblay. The prospector, with the help of Jacques Robert, completed a series of exploration projects including prospecting, sampling, electromagnetic survey and stripping in an area which featured Timiskaming age conglomerates. Getting positive results, the claims were acquired by Probe Mines Limited in 2010. From 2010 to 2013, Probe Mines Limited drilled over 250 holes on the property. Probe Mines Limited released an indicated resource estimate showing 9,262,000 tonnes at a grade of 5.39 g/t gold, and inferred resources estimate of 3,034,000 tons at a grade of 4.37 g/t gold for an underground operation (cut-off of 2.5 g/t gold), and an indicated 70,301,000 tons at a grade of 1.03 g/t gold and an inferred 247,000 tonnes at a grade of 0.8 g/t gold for an open pit operation (cut-off of 0.5 g/t gold; Dzick 2014). Goldcorp Inc. acquired the property by absorbing Probe Mines Limited in the summer of 2013. Goldcorp Inc. is currently drilling and planning for the possible opening of a mine in the near future. The latest resource estimate show 3.02 Mt at a grade of 5.77 g/t gold of indicated resources, 2.03 Mt at a grade of 5.49 g/t gold of inferred resources (Goldcorp Inc. 2016). 188 �Geologic Setting The Kapuskasing Structural Zone (KSZ) is located in the Superior Province, cutting it in two. It trends in a NNE direction, starting south of James Bay and dying off gradually south of Chapleau. While the rocks on each side are of greenschist to amphibolite facies, mostly composed of volcano-sedimentary belts with intrusives, the rocks of the KSZ are striking in contrast since they are of amphibolite to granulite facies. The KSZ is also characterized by a strong positive gravity and aeromagnetic anomalies that are trending NNE. The west contact of the KSZ is transitional while the east contact is outlined by the Ivanhoe Lake cataclastic zone. The structures within the zone are at an almost right angle with those from the surrounding belts, as pictured in aeromagnetic surveys (Figure 1). The rock compositions within the structure consist of alternating bands of mafic gneiss, paragneiss, tonalitic gneiss and dioritic rocks which are elongated to the northeast (Percival 1981). Two discrete bodies of anorthosite form the Shawmere anorthositic complex (Thurston, Siragusa and Sage, 1977). All the rocks in the KSZ were deformed and metamorphosed to high-grade, and the contact relations among units are still blurry in understanding (Percival 1981). Paragneiss is characterized by garnet, biotite, plagioclase and quartz. It is thought to be of sedimentary origin (Percival 1981). These paragneissic rocks contain up to 20% concordant tonalitic layers, and locally contain hypersthene. Paragneisses were observed to be richer in biotite southeast of the Shenango complex. A small unit of paragneiss west of Tom Smith Lake contains up to 10% muscovite as well as carbonate layers, while a small unit southeast of Carty Lake contains up to 5% graphite (Percival 1981). South of Nemegosenda Lake, along Highway 101, layers of varying composition on the 10-50 cm scale in fine-grained, biotite-poor, biotite-plagioclase-quartz gneiss hints that these rocks have an arkose protolith (Percival 1981). A recurrent component of many paragneiss outcrops are enclaves of mafic gneiss. These enclaves are comparable in composition and texture to the bigger mafic gneiss belts. The variation in composition of the paragneiss suggests an original sedimentary facies change (Percival 1981). Mafic gneiss is distinguished by garnet-clinopyroxene-hornblende-plagioclase-quartz mineral assemblages and generally incorporates concordant tonalite layers (Percival 1981). Centimeter thick layers are made by different proportions of minerals. Discrete 2-4 km wide mafic gneiss belts are found in the Borden-Hellyer Lakes area, in a body adjacent south of the main anorthosite body. Dioritic rocks occur dominantly within paragneiss belts. These homogeneous meta-igneous bodies consist of medium-to-coarse-grained hornblende-biotite-plagioclase rocks with up to 10% quartz as well as some clinopyroxene, orthopyroxene and rare garnet (Percival 1981). The texture of the dioritic rocks varies from foliated to flaser or augen, to gneissic. The largest body occurs in the Chapleau-Nemegosenda Rivers area, which also shows gabbro, hornblendite and pyroxenite layers from 10 cm to 2 m in width. Veinlets of quartz-monzonite constitute 15% of outcrop in the belt northeast of the Lackner complex, which possesses clinopyroxene-bearing leucosome (Percival 1981). 189 �Figure 1: Aeromagnetic map of the Kapuskasing Structural Zone (modified from Gupta 1991). 190 �Tonalitic rocks outcrop as discrete belts of gneissic and xenolithic bodies south of the Shawmere anorthosite complex (Percival 1981). Southeast and south of Carty Lake, there is a body of coarse garnethornblende-biotite-plagioclase-quartz gneissic tonalite hosting mafic and ultramafic granulite enclaves. Inclusions ranging from fist-sized to 1 m, constitute 15% of the rock and are of layered mafic gneiss, amphibolite, and garnet-orthopyroxene-hornblende-biotite rock. Schlieren and units of paragneiss up to 1 km wide are common. The proportion of garnet diminishes and the composition becomes granodioritic to the southwest along the belt. Enclaves encountered are mostly amphibolites with rare bright green cummingtonite-hornblende-biotite rocks (Percival 1981). Three eastward-thinning belts, between Nemegosenda and Lackner Lakes, are comprised of medium-to-coarse grained, thinly-banded gneissic tonalite (Figure 2) containing hornblende, biotite, plagioclase and quartz in variable proportions (Percival 1981). Figure 2: Thinly-banded gneissic tonalite The Shawmere anorthosite complex consists of a large, oval mass of anorthosite to anorthositic gabbro lying adjacent to and ESE of the KSZ underlying an area of 800 km2, and a satellite Y-shape mass underlying an area of 90 km2 (Thurston, Siragusa and Sage, 1977). A strongly foliated quartz diorite and monzonite underlying an area of 260 km2 sits in between the two masses of anorthosite. The complex has an overall dimension of 84 km by 24 km, its main axis pointing in an ENE direction. The complex is divided into zones of contrasting structural and compositional characteristics: a megacrystic gabbroic anorthosite zone, an anorthosite zone, a banded zone and a border zone (Riccio 1981a , Riccio 1981b). The central portion of the complex is occupied by the megacrystic gabbroic anorthosite zone, which extends to the northeastern margin of the intrusion (Riccio 1981a, Riccio 1981b). It is dominated by 191 �gabbroic anorthosite with subordinate anorthosite, anorthositic gabbro and gabbro, and minor melanogabbro (Riccio 1981a, Riccio 1981b). Rocks within this zone have large (1 to 50 cm) phenocrysts (Percival 1981). Grains of pyroxene with amphibole or garnet coronas are common within this zone (Percival 1981). Olivine appears rarely, at the center of orthopyroxene grains (Percival 1981). The northwestern part of the intrusion shows the anorthosite zone, which consists of granular aggregates of plagioclase, with minor amphibole and rare garnet (Riccio 1981a, Riccio 1981b). The southeastern margin of the complex is where the banded zone appears. It is characterised by bands of 1-30 cm thick of pure anorthosite, clinopyroxene-hornblende-plagioclase gabbro, sphene-garnet-amphibole-plagioclasequartz rock, garnet-rich garnet-hornblende-quartz rock and minor ultramafic rocks (Riccio 1981 a, Riccio 1981b). The 50-100 m thick border zone separates anorthositic rocks from country rocks in the CartyHarold Lake area and along the northwestern margin of the complex (Riccio 1981a, Riccio 1981b). This zone consists of foliated to gneissic garnetiferous amphibolite with 5-20% concordant to discordant tonalite bands (Percival 1981). Mafic dikes have been observed in the KSZ, in two different swarms (Percival 1981). The Kapuskasing swarm is characterized by ENE-striking, SE-dipping dikes of 1-10 m wide, consisting of sparsely plagioclase-porphyritic, medium-to-fine-grained, ophitic, green-grey gabbro (Percival 1981). The second swarm is of NE-trending dikes with pitted weathering rusty surface which may be part of the Abitibi swarm (Percival 1981). In the Kapuskasing Structural Zone, four alkalic rock-carbonatite complexes are identified. Three are located within 20 km of Borden Lake: Borden, Nemegosenda and Lackner Alkalic complexes. The fourth complex is the Shenango (Figure 2). All four are easily noticeable on aeromagnetic surveys, appearing like round highs on the maps. The Shenango Alkalic complex contains different rock suites than the Borden, Nemegosenda and Lackner complexes. The Shenango contains diorites, monzodiorites, quartz monzonites and granites (Sage 1987c), which differs from the other carbonatites found in the other intrusives of similar age; it is a silica-saturated series of rocks (Sage 1987c). Figure 3: Alkalic complexes of the Kapuskasing Structural Zone 192 �The Nemegosenda Lake Alkalic Rock Complex is made of fine-to-medium-grained syenitic to malignitic rocks which are limited along the southern border of the complex by an arcuate mass of coarse-grained nepheline syenite (Sage 1987b). The complex is contained within a fenitized envelope produced by metasomatism by alkali-iron-rich aqueous fluids from the crystallizing magma (Parsons 1961). Ijolites occur in the northwest corner of the complex, and a mass of gabbro rests along the northwest margin and an isolated band along the east flank. Parsons (1961) noted fresh unmetamorphosed nature of the gabbro units but classified them as country rocks. Sage (1987b) believed that they are an early phase of the alkalic magmatism, perhaps analogous to the gabbroic margins observed at the Port Coldwell and Killala Lake complexes located north of the northeast corner of Lake Superior. The Borden Township Alkalic Rock Complex intruded Archean biotite-quartz-plagioclase gneisses which show fenitization in proximity to the carbonatite intrusion (Sage 1987a). According to Sage (1987a), sovite, silicocarbonatite and fenitized wall rock alternate throughout the length of the examined drill core. It suggests that the complex likely represents a succession of sovite and silicocarbonatite cone sheets emplaced into brecciated biotite-quartz-plagioclase gneisses (Sage 1987a). The Lackner Lake Alkalic Rock Complex consists of core and peripheral nepheline syenites, separated by a medial, arcuate, partial ring of alkali mafic rocks like ijolite and malignite (Sage 1988). The leucocratic coarse to very coarse-grained, nepheline syenites of the core and periphery cannot be differentiated by mineralogy, textural or crosscutting field relations (Sage 1988). In proximity to the arcuate band of mafic rocks, the syenites are strongly to weakly trachytoid texture, are locally finer-grained, and contain abundant subangular to subrounded mafic xenoliths. The xenoliths, commonly biotite-rich, are derived from the more mafic phases of the complex which are cut by the nepheline syenites (Sage 1988). Dikes of magnetite-apatite cut both the mafic and syenitic phases of the complex, which is also cut by lamprophyre dikes (Parsons 1961). Cataclastic rocks are present in the transition between the Western Abitibi subprovince and the Kapuskasing Structural Zone, especially in the southern Ivanhoe Lake-Ivanhoe River area. The area is characterized by mylonite and cataclasite (Percival 1981). Pseudotachylite and ultra-cataclasite veinlets are developed within the high-grade rocks of the KSZ, including anorthosite, northwest of Ivanhoe Lake (Riccio 1981a, b). The Ivanhoe Lake cataclastic zone, the zone affected by cataclasis, is 1-2 km in width and may have more than one discrete fault zone (Percival 1981). Structure The rocks of the Kapuskasing Structural Zone are characterized by gneissosity, defined by small to major differences in mineralogical composition of individual bands ranging from 1 to 90 cm in thickness (Thurston et al. 1977). The orientation of lithological contacts and gneissosity make up the prominent east-northeast structural grain of the Kapuskasing structural zone. Gneissosity in all rock types is folded or warped about gently-plunging (0-25°) northeast-trending axes (Percival 1987). The folds vary from northwest-facing monoclinal flexures to isoclinal with consistent "Z" sense asymmetry when viewed toward the east. Axial surfaces are rarely accompanied by a foliation defined by flattened quartz grains. The trend of fold axes and lineations is northeast-southwest throughout this part of the Kapuskasing zone, but plunge direction varies on a regional scale from dominantly southeasterly in the south to northeasterly in the north. Between these areas, lineations are within 10° of horizontal and abrupt changes in plunge 193 �direction occur on the 100 m scale. Both regional and local plunge reversals can be related to a gently southeast-plunging warp axis (Percival 1987). Plausible Explanations To this day, the Kapuskasing Structural Zone has seen various explanations for its presence. The origin of the structure has been interpreted as “thinning of the granitic upper crustal layer” (Garland 1950), Mid- to Late Proterozoic rifting (MacLaren et al. 1968; Bennett et al. 1967; Burke and Dewey 1973; Thurston et al. 1977) a suture between the western and eastern Superior Province (Wilson 1968), dextral transcurrent faulting (Goodings and Brookfield 1992), and an east-verging thrust exposing an oblique crustal cross section (Percival and Card 1983; Percival 1986). Percival and West (1994) criticize these models since no simple model is capable of answering all the present features. The suture model does not account for the correlativity in age and geology between the Abitibi and Wawa subprovinces. A horst model describes the geometry of the blocks within the KSZ, but the third dimension shows a better consistency with a compressional origin than extension. Also, the KSZ existed before the emplacement of the carbonatites and alkali rock complexes around 1100 Ma. The dextral transcurrent model is based on northeasterly aeromagnetic and map trends in the KSZ. Other explanations for these trends are possible, and the scale of any transcurrent displacement seemed to be minor (< 20km). The east-verging Early Proterozoic thrust fault seems to explain many features of the KSZ. However, the Matachewan dike swarm displacement noted along the southern end of the KSZ does not seem possible given its semicontinuous nature, which shows 55-70 km west-over-east apparent movement (Percival and West 1994). Percival and West (1994) use a model which follows this explanation: “Supracrustal rocks of the Kapuskasing Zone (2750-2700 Ma) were buried by younger supracrustal rocks, tectonic shortening, and intrusion of mid-crustal tonalities (2700-2660 Ma). Metamorphism began during this period and continued in response to magmatic heat and crustal collapse (2660-2625 Ma). The Borden Lake conglomerate was deposited in a pull-apart basin and transported into the hot deep crust as a sliver along a downward vector on a dominantly transcurrent fault. Kapuskasing rocks remained at depth, where they cooled slowly and underwent intermittent minor deformation (2625-2585 Ma). The Superior Province was eroded by 10 km on average, elevating Kapuskasing levels from approximately 30 to 20 km. Incipient breakup of the Superior craton (2500-2450 Ma) was recorded as new mineral growth at deep structural levels and by Matachewan dyke injection. A few additional kilometres of erosion preceded intrusion of Kapuskasing dykes (2040 Ma) into the intact crustal section. At approximately 1.9 Ga, when Kapuskasing levels had cooled to <300oC, stresses caused by plate collisions at the Superior margin were transmitted into the interior in the form of dextral transpression along northeast-trending and northwest-over-southeast thrusting, elevating Kapuskasing-level rocks […]. Formation of a crustal root accommodated shortening in the lower crust […]. Northwest-dipping normal faults and a conjugate set of strike-slip faults broke the Kapuskasing structure into separate blocks with variable geometry. During the relaxation phase, isostatic rebound reduced topography on the root and probably produced a few more kilometres of uplift along steep structures not coincident with the Ivanhoe Lake Fault.” The recent discovery of the Borden Lake deposit shows that perhaps the story is wrong; maybe we are looking at mineralization put into place prior to the uplift. 194 �Figure 4: Field trip stops Stop 1: Shawmere Anorthosite Zone: 17, Easting: 366922 Northing: 5332558 Access: Take the Warren-Carty main haul road. Travel down the road for about 10 km. The site is accessible from the road through a wet overgrown section which leads to the bulk sample pit created by Avalon Ventures Ltd. back in 2007. The rocks in the bulk sample pit (Figure 5) are a very coarse-grained anorthosite with 1-4 mm diameter grains, with sporadically some 3 cm diameter porphyroblasts although most are in the 1 cm diameter range. The porphyroblasts are xenomorphic, poikiloblastic and disseminated throughout the rocks. The grains in the matrix are equant and may be polygonized. 195 �Figure 3: Shawmere Anorthosite pit Stop 2: Sandy Outcrop Zone: 17, Easting: 360841 Northing: 5318167 The outcrop shows granitic gneiss, with melanosome and leucosome in thin beds of 1 to 3 cm in thickness. The melanosome consists of amphiboles, biotite and calcic feldspars, while the leucosome consists of quartz, potassic and sodic feldspars. The folded and wavy beds underline migmatization of the rock. A diabase dike crosses the outcrop with seeming difficulty. Stop 3: Borden Turnaround Outcrop Zone: 17, Easting: 340126 Northing: 5309885 The outcrop (Figure 6) shows a seemingly partially undifferentiated granitic migmatite. Dark paleosomes are seen on the outcrop, composed of biotite and amphiboles (restites). Garnets are rather invisible, some trace sulphides are present. Behind the outcrop, on boulders broken off during the road construction, it is possible to see an “augen” (Figure 7) made out of a piece of possibly chert, which may suggest that the protolith may be a metasediment. 196 �Figure 4: Turnaround Outcrop Figure 5: "Augen" made of possibly chert, Turnaround Outcrop. Notice the zoning in the eye, which is perpendicular to the gneissossity of the rock. 197 �Stop 4: Borden Lake Conglomerate Zone: 17, Easting: 330441 Northing: 5305051 The Borden Lake Conglomerate is a deformed metaconglomerate (Figure 8) which was dated at 2664 ± 12 Ma (Percival et al. 1981), using tonalitic cobbles. The age date of the zircons is of a later deformationmetamorphic event. Some texts refer to it as a “pebble” conglomerate, but most clasts in the outcrop are more in the “cobble” range of sizes. The rock is vastly clast supported with a 10-15% matrix of quartzbiotite-hornblende-garnet. The cobbles are made of felsic metavolcanics, metasediments, granodiorite, tonalite, plagioclase-porphyritic meta-andesite and amphibolite, with some hornblendite and quartz veins. The foliation is gently north dipping. They are elongated, forming rods lineated at 075oN, and a plunge of 026o (Figures 8 and 9). The metaconglomerate is accompanied spatially by an amphibolite and paragneiss to the south of the highway. Somehow, the amphibolite is the host of the gold mineralization of the Borden Lake deposit. To the south of the outcrop, towards the highway, lies a felsic porphyry unit (Figure 10). The unit is grey in appearance with white “eyes” of quartz porphyroblasts. The porphyroblasts are up to 1 cm in diameter. Reddish garnets can be seen on the outcrop. Figure 6: Borden Lake conglomerate, with elongated cobbles 198 �Figure 7: Borden Lake conglomerate, down plunge view Figure 8: Quartz porphyry, with quartz eyes up to 1 cm in diameter. 199 �Stop 5: Chapleau Truck Stop Outcrop Zone: 17, Easting: 321942 Northing: 5294079 The roadcut (Figure 11) shows sub-vertical orientation of mafic xenoliths in a medium grained tonalite, showing a “salt & pepper” texture. A pegmatitic aplite vein (~10 cm in thickness) crosscuts the outcrop. The abundance of mafic xenoliths seems to increase towards the south of the outcrop along the highway. Figure 9: Truckstop outcrop with subvertical xenoliths and cut by a dark pink aplite vein Stop 6: Borden Lake Deposit The Borden Lake gold deposit lies at the intersection of the Wawa sub-province, the Kapuskasing Structural Zone (KSZ) and the Abitibi sub-province, specifically within the southernmost limits of the KSZ (Dzick 2014). It is located on a peninsula of Borden Lake about 20km from the town of Chapleau. The deposit dips at 40-45° to the north-northeast and strikes to the west-northwest. This trend is bound to the north by an extensive package of metaconglomerates and intersected by a significant east-west striking fault. To the south it is bound by a similar package of metaconglomerates, which then grades into various garnet-boitite gneisses, amphibolites and granulite-grade gneisses. The gold mineralization within the deposit is typically characterized by a higher-grade core surrounded by a lower-grade envelope, within a package of volcano-metasedimentary rocks of variable composition. The main sulphides are pyrite and pyrrhotite, with the latter typically dominating within the gold zone. The west-northwest portion of the deposit is generally lower grade with some higher grade pods. The east-southeast portion of the deposit hosts the High-Grade Zone (Dzick, 2014). 200 �Figure 10 201 �An interpreted geological drill section (Figure 12) intercepts the High-Grade Zone (HGZ) at the south eastern portion of the deposit. The general hanging wall sequence of rocks is an intermixed package of felsic intrusive (quartz-metadiorite to diorite) sills/dikes with felsic metavolcanic-metasedimentary gneisses. This upper sequence of felsic rocks is cross-cut or intermixed with a coarse grained porphyroblastic metagabbro (locally termed amphibole felsic gneiss) and amphibolite units. Sulphide mineralization in the hanging wall is dominantly pyrite in amounts ranging from trace to 1%. The appearance of pyrite and pyrrhotite mineralized amphibolites marks the transition from the hanging wall to the beginning of the gold-bearing zone and the start of the low-grade gold envelope. The low-grade envelope (Figure 13) is typically characterized by intermixed felsic gneisses and amphibolites with an overall increase in sulphide mineralization to 1-2% pyrite/pyrrhotite. The increase in sulphide content generally coincides with a higher degree of strain and development of a foliation that becomes better defined and more intense down sequence. The transition from the low-grade envelope to the beginning of the high-grade core is marked by the presence of a unit that is locally termed the granitic gneiss. The highest grade portions of the gold zone correlate with deformed quartz veins and quartz pegmatites hosted within small sections of garnet-biotite gneiss and amphibolite. All main sequence units within the core can be silicified to varying degrees, with the units proximal to the quartz veins and pegmatites generally undergoing the strongest intensity of alteration. Visible gold is most often associated with quartz veining and areas of intense silicification. Figure 11 202 �The Borden Lake deposit is located on a peninsula on Borden Lake. A volcano-sedimentary package of various composition hosts the gold mineralization, which occurs as a wide zone of disseminated and fracture-controlled sulphides. Pyrite and pyrrhotite are the main sulphides present, with pyrite being the most common. Gold occurs in a high-grade core surrounded by a low to moderate grade envelope, accompanied by minor silver. The grade of the core seems to be improving towards the southeast where it becomes the high-grade zone (HGZ) with average grades above 2.5 g/t gold (Dzick, 2014). The northwest section of the deposit is characterized by sporadic silicification without lithological control or quartz veining. The southeast section, on the other hand, has a well-developed hydrothermal system with quartz flooding and potassic alteration which defines the HGZ. Various host rocks contain the mineralization, usually dominated by metasedimentary horizons and subordinate intrusives of felsic to intermediate composition, who display feldspathic, chloritic and biotitic alterations. Outcrops in the northwest rarely show visible gold, which is the opposite of the HGZ in the southeast, especially in the quartz-rich core. That core is present in the low-grade zone, sometimes attaining very high grades like the HGZ (Dzick, 2014). The deposit showed continuity and has been consistently intersected along strike, reaching a length of 3.7 km (Dzick, 2014). The deposit remained open both to the northwest and southeast directions. The dip and plunge of the deposit are to the northeast and shallow southeast respectively. Gold mineralization seemed to be controlled by a ductile shear zone, which is more apparent in the HGZ. The mineralized zone has been confirmed to a vertical depth of approximately 650 m, and is up to 120 m wide (Dzick, 2014). References Archibald, J.C. 2009. Technical Report on the Nemegosenda Property for Sarissa Resources Inc; Biliken Management Services Inc. Filed on SEDAR.com Bennett, G., Brown, D.D., George, P.T., and Leahy, E.J. 1967. Operation Kapuskasing. Ontario Department of Mines, Miscellaneous Paper 10, 98p. Burke, K., and Dewey, J.F. 1973. Plume-Generated Triple Junctions: Key Indicators in Applying Plate Tectonics to Old Rocks. Journal of Geology, Vol. 81, p. 406-433. Dzick, W. 2014. Probe Mines Limited: Mineral Resources Estimate Update, Borden Lake Project. NI 43101 technical report, June 10, 2014. 179 p. Garland, G.D. 1950. Interpretation of Gravimetric and Magnetic Anomalies on Traverses in the Canadian Shield of Northern Ontario. Publications of the Dominion Observatory (Ottawa), Vol. 16, part 1. Goldcorp 2016. Reserves and Resources estimate table. http://s1.q4cdn.com/038672619/files/doc_downloads/2016/oct/Reserves-and-Resources-TableWebsite_FINAL.pdf Goodings, C.R. and Brookfield, M.E. 1992. Proterozoic Transcurrent Movements along the Kapuskasing Lineament (Superior Province, Canada) and their Relationship to Surrounding Structures. Earth-Science Reviews, Vol. 32, p147-185. 203 �Gupta, V.K. 1991. Shaded image of total magnetic field of Ontario, east-central sheet; Ontario Geological Survey, Map 2586, scale 1:1 000 000. Parsons, G.E. 1961. Niobium-Bearing Complexes East of Lake Superior; Ontario Geological Survey, Geological Report 3, p.33-50, Map 2007, Scale 1 inch to ¼ mile. Percival, J.A. 1981. Geological evolution of part of the central Superior Province based on relationships among the Abitibi and Wawa subprovinces and the Kapuskasing structural zone (Ph.D. Thesis). Queen’s University, Kingston 300 p. Percival, J.A. 1987. The Kapuskasing Uplift: Archean Greenstones and Granulites. Institute on Lake Superior Geology Thirty-Third Annual Meeting; Wawa, Ontario. Vol. 33, Part 5, 54p. Percival, J.A., and Card, K.D. 1983. Archean Crust as Revealed in the Kapuskasing Uplift, Superior Province, Canada. Geology, Vol. 11, p. 323-326. Percival, J.A. and West, G.F. 1994. The Kapuskasing Uplift: A Geological and Geophysical Synthesis. Canadian Journal of Earth Sciences, vol. 31, p1256-1286. Riccio, L. 1981a. Precambrian Geology of the Shawmere Anorthositic Complex (North), District of Sudbury; Ontario Geological Survey Preliminary Map P. 2383, Geological Series, Scale 1:15 840. Geology 1979. Riccio, L. 1981b. Precambrian Geology of the Shawmere Anorthositic Complex (South), District of Sudbury; Ontario Geological Survey Preliminary Map P. 2384, Geological Series, Scale 1:15 840. Geology 1979. Sage, R.P. 1988. Geology of Carbonatite – Alkalic Rock Complexes in Ontario: Lackner Lake Alkalic Rock Complex, District of Sudbury; Ontario Geological Survey, Study 32, 141p. Sage, R.P. 1987a. Geology of Carbonatite – Alkalic Rock Complexes in Ontario: Borden Township Carbonatite Complex, District of Sudbury; Ontario Geological Survey, Study 33, 62p. Sage, R.P. 1987b. Geology of Carbonatite – Alkalic Rock Complexes in Ontario: Nemegosenda Lake Alkalic Rock Complex, District of Sudbury; Ontario Geological Survey, Study 34, 132p. Sage, R.P. 1987c. Geology of Carbonatite – Alkalic Rock Complexes in Ontario: Shenango Township Alkalic Rock Complex, Districts of Sudbury and Algoma; Ontario Geological Survey, Study 35, 119p. Thurston, P.C., Siragusa, G.M., and Sage, R.P. 1977. Geology of the Chapleau Area, Districts of Algoma, Sudbury and Cochrane; Ontario Geological Survey, GR157, 293p. Wilson, J.T. 1968. Comparison of the Hudson Bay Arc with some Other Features. In Science, History and Hudson Bay. Edited by C.S. Beals and D.A. Shenstone. Department of Energy, Mines and Resources, Ottawa, p1015-1033. 204 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2017 Description An account of the resource Institute on Lake Superior Geology. Wawa, Ontario. May 8-12, 2017. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2017 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/e5e2a3aeea1565339a75ff20b1c88207.pdf 2f72c773c08031a7b9bd02c23a9e41a6 PDF Text Text Institute on Lake Superior Geology 64th ANNUAL MEETING May 15-18, 2018 Iron Mountain, Michigan Hosted by: LAUREL G. WOODRUFF, WILLIAM F. CANNON AND ESTHER K. STEWART CO-CHAIRS U.S. GEOLOGICAL SURVEY WISCONSIN GEOLOGICAL & NATURAL HISTORY SURVEY Proceedings Volume 64 Part 1 – Program and Abstracts Edited by Esther K. Stewart �64th INSTITUTE ON LAKE SUPERIOR GEOLOGY VOLUME 64 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD TRIP GUIDEBOOK TRIP 1: ARCHEAN AND PALEOPROTEROZOIC GEOLOGY OF THE FELCH DISTRICT, CENTRAL DICKINSON COUNTY, MICHIGAN TRIP 2: GEOLOGY OF THE HEMLOCK FORMATION TRIP 3: GEOLOGY AND IRON ORES OF THE MENOMINEE IRON RANGE, DICKINSON COUNTY, MICHIGAN TRIP 4: GRANITOID ROCKS OF THE PEMBINE-WAUSAU TERRANE IN NORTHEASTERN WISCONSIN Reference to material in Part 1 should follow the example below: Authors, 2018, abstract title, 64th Institute on Lake Superior Geology Proceedings, v. 64, Part 1, Program and Abstracts, p. xx. Proceedings Volume 64, Part 1: Program and Abstracts, and Part 2: Field Trip Guidebook are published by the 64th Institute on Lake Superior Geology and distributed by the Institute Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca Some figures in this volume were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the volume when it is available on-line at: http://www.lakesuperiorgeology.org ISSN 1042-99 i �Part 1: Program and Abstracts Table of Contents Institutes on Lake Superior Geology, 1955-2018 iii v Sam Goldich and the Goldich Medal Goldich Medal Guidelines vii Goldich Medalists and Goldich Medal Committee ix Citation for Goldich Medal Award to Val Chandler x Honoring the Pioneers of Lake Superior Geology xii Memoriam to William D. Addison xiii Eisenbrey Student Travel Awards xiv Joe Mancuso Student Research Awards xv Doug Duskin Student Paper Awards and Award Committee xvi Board of Directors, Local Committee, and Session Chairs xix Field Trip Leaders xx Corporate and Individual Sponsors of Student Travel and Registration xxi Report of the Chair of the 633rd Annual Meeting xxii Technical Program xxvi Poster Presentations xxxiii Abstracts xxxvi ii �Institutes on Lake Superior Geology, 1955-2018 95 o o 85 o Wabigoon subprovince90 o 80 48 o Wawa-Abitibi subprovince 48o Wawa-Abitibi subprovince o 45 45o Minnesota River Valley subprovince MEETING LOCATIONS Phanerozoic Mesoproterozoic Map by Mark Jirsa 95o Paleoproterozoic o 90 85o Archean Superior Province # Date Place Chairs 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 1955 1956 1957 1958 1959 1960 1961 1962 1963 1964 1965 1966 1967 1968 1969 1970 1971 1972 1973 1974 1975 1976 1977 Minneapolis, Minnesota Houghton, Michigan East Lansing, Michigan Duluth, Minnesota Minneapolis, Minnesota Madison, Wisconsin Port Arthur, Ontario Houghton, Michigan Duluth, Minnesota Ishpeming, Michigan St. Paul, Minnesota Sault Ste. Marie, Michigan East Lansing, Michigan Superior, Wisconsin Oshkosh, Wisconsin Thunder Bay, Ontario Duluth, Minnesota Houghton, Michigan Madison, Wisconsin Sault Ste. Marie, Ontario Marquette, Michigan St. Paul, Minnesota Thunder Bay, Ontario C.E. Dutton A.K. Snelgrove B.T. Sandefur R.W. Marsden G.M. Schwartz & C. Craddock E.N. Cameron E.G. Pye A.K. Snelgrove H. Lepp A.T. Broderick P.K. Sims & R.K. Hogberg R.W. White W.J. Hinze A.B. Dickas G.L. LaBerge M.W. Bartley & E. Mercy D.M. Davidson J. Kalliokoski M.E. Ostrom P.E. Giblin J.D. Hughes M. Walton M.M. Kehlenbeck iii �# 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 Date 1978 1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 Place Milwaukee, Wisconsin Duluth, Minnesota Eau Claire, Wisconsin East Lansing, Michigan International Falls, Minnesota Houghton, Michigan Wausau, Wisconsin Kenora, Ontario Wisconsin Rapids, Wisconsin Wawa, Ontario Marquette, Michigan Duluth, Minnesota Thunder Bay, Ontario Eau Claire, Wisconsin Hurley, Wisconsin Eveleth, Minnesota Houghton, Michigan Marathon, Ontario Cable, Wisconsin Sudbury, Ontario Minneapolis, Minnesota Marquette, Michigan Thunder Bay, Ontario Madison, Wisconsin Kenora, Ontario Iron Mountain, Michigan Duluth, Minnesota Nipigon, Ontario Sault Ste. Marie, Ontario Lutsen, Minnesota Marquette, Michigan Ely, Minnesota 56 2010 International Falls, Minnesota 57 58 59 60 61 62 2011 2012 2013 2014 2015 2016 Ashland, Wisconsin Thunder Bay, Ontario Houghton, Michigan Hibbing, Minnesota Dryden, Ontario Duluth, Minnesota 63 2017 Wawa, Ontario 64 2018 Iron Mountain, Michigan iv Chairs G. Mursky D.M. Davidson P.E. Myers W.C. Cambray D.L. Southwick T.J. Bornhorst G.L. LaBerge C.E. Blackburn J.K. Greenberg E.D. Frey & R.P. Sage J. S. Klasner J.C. Green M.M. Kehlenbeck P.E. Myers A.B. Dickas D.L. Southwick T.J. Bornhorst M.C. Smyk L.G. Woodruff R.P. Sage & W. Meyer J.D. Miller & M.A. Jirsa T.J. Bornhorst & R.S. Regis S.A. Kissin & P. Fralick M.G. Mudrey & Jr., B.A. Brown P. Hinz & R.C. Beard L. Woodruff & W.F. Cannon S. Hauck & M. Severson M. Smyk & P. Hollings A. Wilson & R. Sage L. Woodruff & J. Miller T.J. Bornhorst & J. Klasner J. Miller, G. Hudak, & D. Peterson M. Jirsa, P. Hollings, & T. Boerboom, P. Hinz & M.Smyk T. Fitz P. Hollings T.J. Bornhorst & A. Blaske J. Miller & M. Jirsa R. Cundari & P. Hinz J. Miller, C. Schardt, & D. Peterson A. Pace, A. Wilson, & T.J. Bornhorst L. Woodruff, W. Cannon, & E.K. Stewart �Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970’s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 v �INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL vi �Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. vii �Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. viii �Goldich Medalists 1979 Samuel S. Goldich 1998 Zell Peterman 2016 Mark A. Jirsa 1980 not awarded 1999 Tsu-Ming Han 2017 Philip Fralick 1981 Carl E. Dutton, Jr 2000 John C. Green 1982 Ralph W. Marsden 2001 John S. Klasner 1983 Burton Boyum 2002 Ernest K. Lehmann 1984 Richard W. Ojakangas 2003 Klaus J. Schulz 1985 Paul K. Sims 2004 Paul Weiblen 1986 G.B. Morey 2005 Mark Smyk 1987 Henry H. Halls 2006 Michael G. Mudrey 1988 Walter S. White 2007 Joseph Mancuso 1989 Jorma Kalliokoski 2008 Theodore J. Bornhorst 1990 Kenneth C. Card 2009 L. Gordon Medaris, Jr 1991 William Hinze 2010 William D. Addison & 1992 William F. Cannon Gregory R. Brumpton 1993 Donald W. Davis 2011 Dean M. Rossell 1994 Cedric Iverson 2012 James D. Miller 1995 Gene La Berge 2013 Tom Waggoner 1996 David L. Southwick 2014 Laurel Woodruff 1997 Ronald P. Sage 2015 Rodney J. Ikola 2018 GOLDICH MEDAL RECIPIENT Val Chandler Goldich Medal Committee Serving through the meeting year shown in parentheses. Shannon Zurevinski (2015-2018) Lakehead University Klaus Schultz (2016-2019) U. S. Geological Survey Dan England (2017-2020) Eveleth Fee Office ix �Citation for the Goldich Medal Recipient to Val W. Chandler Val W. Chandler, geophysicist extraordinaire, has been my friend and professional colleague for almost forty years. We have worked together on projects too numerous to mention, beginning in 1979 only a few weeks after Val escaped to the blissful cool of Minnesota after a brief stint at Amoco, Inc. in the heat and humidity of Houston, Texas. His personal accomplishments and contributions to diverse joint projects at the Minnesota Geological Survey, and beyond, are remarkable in their scientific breadth. It is a privilege for me to be Val’s citationist for the 2018 Goldich Medal. Val was born and raised in Indianapolis, Indiana. After graduating from high school in 1967, having excelled academically and also athletically in track and field, he entered Indiana University. There he majored in geology and continued his athletic career as a “weight-man” on the IU varsity track team. In 1970 he won the Big Ten Conference championship in discus and placed second in shot-put. Fortunately for us, he declined the overtures of professional football scouts attracted by his impressive size and strength and instead decided to pursue a graduate education in the earth sciences at Indiana, where he obtained an M.S. in geophysics, and at Purdue, where he acquired a Ph.D. in geophysics in 1978 under the tutelage of Prof. Bill Hinze. His graduate work in both universities involved extensive practical applications of magnetic and gravity methods. Val’s professional contributions to understanding the geological framework of Minnesota and the greater Lake Superior region can be subdivided into three main parts. His first challenge was to plan and then supervise the production of a state-wide, high-definition aeromagnetic map of Minnesota. That project, spread over roughly 12 years, involved negotiating contracts with private-sector geophysical mapping firms, performing quality-control tests of the data, writing progress reports to university and governmental administrators, and securing funding for successive segments of the project from the Minnesota legislature. More or less coincident with all of this, Val oversaw a parallel effort to complete a high-quality gravity survey of the state that involved faculty and students from the University of Minnesota and Northern Illinois University and included important contributions of data from private-sector sources. The net result of these efforts was a set of digital potential-field geophysical maps of the state that were widely acknowledged to be among the very best in North America. After the geophysical mapping of Minnesota was essentially finished, about 1992, Val devoted more and more of his time to geological interpretation of the geophysical data. In this work he collaborated in various ways with geologists in the MGS, such as myself, Mark Jirsa, Jim Miller, and Terry Boerboom, and with many geologists in adjacent states and provinces. Furthermore, he contributed to important national and international geophysical projects such as the development of the gravity anomaly map of North America (1988) and the magnetic anomaly map of North America (2002). All along, Val was assiduous in applying the latest technological advancements to the presentation and interpretation of geophysical data. Among the techniques he perfected is the so-called “SMOG” presentation in which gravity and magnetic anomalies are combined. The SMOG acronym means Superimposed Magnetics On Gravity. A SMOG map shows the first vertical derivative of the magnetic signature (typically in grayscale) draped over the second vertical derivative of the gravity signature (typically shown in bright colors). The value of modern computing power in producing x �these maps and other analytical tools cannot be overemphasized, and Val’s efforts in developing and improving computational applications, such as SMOG maps and various digital modeling methods, have proven to be powerful aids to the geologic mapping of Precambrian terranes beneath glacial cover in Minnesota and the rest of the Lake Superior region. As we all know, the Precambrian rocks of the Lake Superior region host a wealth of metallic mineral resources, and the potential for discovering and developing future economically viable Precambrian mineral deposits in covered areas has long been an attractive possibility to exploration companies and politicians. Indeed, that possibility was emphatically presented to Minnesota policy makers in the late 1970s, during a deep recession in the Minnesota iron-mining industry. It gave rise to a push for “minerals diversification” and created a political environment in which the importance of geophysics to the diversification effort could be successfully argued. That set of conditions brought Val to us, and his presence has produced dividends. Today we can make much better geologic maps of Precambrian terranes than we could before digital aeromagnetic and gravity maps became a reality, and consequently can make more credible assessments of mineral potential. In recent years, however, the sense of urgency expressed in the public and political sectors has changed. Clean water, especially clean groundwater, has supplanted metals as the political “ore of choice”. This is reality. Val, in the third chapter of his career, has pivoted from the geophysical interpretation of Precambrian rocks to the pursuit of techniques that aid three-dimensional hydrogeologic mapping of Quaternary glacial deposits. He has applied passive seismic methods to the estimation of sediment thickness above sub-Quaternary bedrock, an important parameter in aquifer delineation and groundwater management. He continues to perfect passive seismic techniques and works in close cooperation with soft-rock stratigraphers and hydrogeologists at MGS and affiliated state agencies. Last but not least, Val is a teacher. He is an adjunct professor of geophysics in the school of earth sciences at the University of Minnesota. He has taught various undergraduate-level geophysics courses over the years and advised or co-advised several graduate students pursuing M.S. or Ph.D. degrees. He has long been an advocate for advancing the understanding and sensible application of science in the public sphere. Val continues to be fascinated by the geology and geological resources of the Lake Superior region, both solid and liquid. His enthusiasm and his professional contributions to our collective understanding of this area unquestionably qualify him to join the ranks of Goldich Medal recipients. It is my distinct pleasure, therefore, to present Val W. Chandler to the Institute as its 2018 recipient of the Samuel S. Goldich Medal for “Outstanding contributions to the geology of the Lake Superior region”. Submitted by David Southwick Director Emeritus Minnesota Geological Survey xi �Honoring the Pioneers of Lake Superior Geology (Adopted by the Board of Directors, 2016) Preamble At the suggestion of Gene LaBerge, the 2016 executive board agreed to implement a program to recognize historic pioneers in the understanding of geology in the Lake Superior region. Beginning with the 2017 annual meeting, nominations will be accepted from the membership for geologists whose work was conducted primarily before inception of the institute in 1955. Biographical sketches of those pioneers will be presented at future annual meetings so that all might appreciate the value of their contributions. Selection of nominees will be decided in part by the organizing committee of each year's annual meeting, in consultation with the Board, to ensure equitable geographic representation in the selection process. Award Guidelines 1) Nominations from the membership will be submitted via the Institute web site and forwarded to the Chair of the next Annual Meeting. The nominations will be no more than half a page in length and will summarise the contribution of the nominee. 2) The Organising Committee will select one or two individuals to be highlighted at the next Annual meeting and submit those names to the Board for approval. 3) The nominator will be requested to prepare a brief presentation to be given during the next annual meeting with a summary to be included in the Proceedings volume. 4) Unsuccessful nominations will be kept by the Secretary for two years and forwarded to the next meeting Chair; these nominations may be resubmitted at a later date. The Board will review this award every five years. Pioneers of Lake Superior Geology 2017 Douglass Houghton (1809-1845) 2018 not presented xii �In Memoriam William D. Addison October 25, 2017- a day of great loss for The Institute on Lake Superior Geology, its members, and countless others whose professional and personal lives were deeply influenced by Bill Addison. Bill’s many and varied accomplishments are impossible to fully enumerate in a short remembrance. Bill, born in Toronto, lived much of his younger years in Thunder Bay before earning his B.Sc. in Forestry and M.Sc. in Fisheries Biology from the University of Toronto, where he met Wendy Livingston, who would become his wife. Bill and Wendy settled in Thunder Bay, where Bill worked as a fisheries biologist before joining Wendy in a teaching career at Westgate High School, where he taught Biology, Chemistry, and Geology for nearly 30 years. The Institute, and the broader geological community, know Bill for the discovery, made with long-time colleague and co-investigator Greg Brumpton, of the layer of meteor impact debris that was spread across the Lake Superior region because of the great Sudbury impact. Bill first presented the documentation of the impact layer near Thunder Bay in 2005, at the 51st Annual Meeting of the Institute. Papers in the journal “Geology” and a Geological Society of America Special Paper soon followed and led others to discover the debris layer at many other localities around Lake Superior. Bill and Greg received world-wide recognition for their discovery, which spurred a flurry of research by an international group of Earth scientists that continues today. Bill’s search for the Sudbury layer was, remarkably, only one of his many interests in the natural world, although one that he and Greg pursued with great patience and diligence for more than a decade before their final success. Bill is one of a select few to receive both the Goldich Medal and Homer Award from the Institute-- the Goldich shared with Greg Brumpton, recognizing their work on the Sudbury impact-- the Homer entirely a recognition of Bill’s own (mis)deeds. Future generations, to their loss, will know Bill as a name and author of groundbreaking geologic research. But the man-- larger-than-life, congenial, gregarious, and generous, that so many of us had the pleasure of knowing, if for only too short a time, should be remembered and celebrated as well. You could not know Bill for long without feeling that you had made a great new friend—and you would be right. He had seemingly unlimited space in his life and heart for friendship and kindness. He loved sharing his many unique experiences through his raconteurial skills, and had a seemingly limitless trove of fascinating tales of his adventures. Bill was a true lover of nature and supporter of its preservation. He and Wendy traveled the back roads and trails of the world celebrating both its natural beauty and cultural history. A fortunate group of friends received his “Epistles” from the road, an authoritative diary of daily discoveries, beautifully illustrated by his exceptional photography. A life well-lived to the fullest, two loving, accomplished daughters, Michelle and Kirsten who blessed him with four grandchildren, lasting scientific contributions, and a host of friends and colleagues who were fortunate to have known him--this is the legacy of William D. Addison xiii �Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xiv �Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive one-half of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. In fall 2018, the ILSG Board of Directors selected four students to be granted research funding of $500.00 each from the Joe Mancuso Student Research Fund. The awardees were: Chanelle Boucher Lakehead University, MsC, Dept. Geology, cbouche2@lakeheadu.ca TOPIC: Komatiitic units within the Lake of the Woods Greenstone Belt Dustin Andrew Liikane University of Toronto, PhD, Dept. Earth Sciences, dustin.liikane@mail.utoronto.ca TOPIC: Controls on the timing and localization of mineralized intrusions within the Midcontinent Rift Jacqueline L. Drazan University of Minnesota-Duluth, MsC, Dept. Earth and Environmental Sciences, draza004@d.umn.edu TOPIC: Can silicon isotopes of quartz be used to determine chert petrogenesis in VMS-hosting systems in the ~2.7 Ga Abitibi Greenstone Belt, Canada? Margaret Upton University of Minnesota-Duluth, MsC, Dept. Earth and Environmental Sciences, upton040@d.umn.edu TOPIC: Alteration mineral zonation and geochemical characteristics of the Back Forty Deposit, MI—A replacement-style zinc- and gold-rich volcanogenic massive sulfide deposit xv �Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and longtime friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. 2018 Student Paper Awards Committee Latisha Brengman – University of Minnesota-Duluth Robert Cundari – Ontario Geological Survey Esther Stewart – Wisconsin Geological & Natural History Survey xvi �Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Esther Stewart (2018-2021) – Wisconsin Geological & Natural History Survey Anthony Pace (2017-2020) – Ontario Geological Survey Christian Schardt (2016-2019) – University of Minnesota Duluth Rob Cundari (2015-2018) – Ontario Geological Survey Pete Hollings - Secretary (2017-2020) – Lakehead University Mark Jirsa – Treasurer (2015-2020) – Minnesota Geological Survey Local Committee Tom Mroz- BSGE, MSPG, CPG Session Chairs Marcia Bjørnerud - Lawrence University Ben Drenth - U.S. Geological Survey John Esch - Michigan Department of Environmental Quality Daniel Holm - Kent State University Suzanne Nicholson - U.S. Geological Survey Dean Peterson - Natural Resources Research Institute Amy Radakovich - Minnesota Geological Survey Shannon Zurevinski - Lakehead University xix �Field Trip Leaders Field trips have been the mainstay of the ILSG since its inception 64 years ago. We want to give a special thanks to the field trip leaders who volunteered their time and talent in carrying that tradition forward. 1) Archean and Paleoproterozoic Geology of the Felch District, Central Dickinson County, Michigan Bill Cannon, Klaus Schulz, Robert Ayuso – U.S. Geological Survey Tom Mroz – BSGE, MSPG, CPG 2) Geology of the Hemlock Formation Tom Waggoner – Consulting Geologist 3) Geology and Iron Ores of the Menominee Iron Range, Dickinson County, Michigan Tom Mroz – BSGE, MSPG, CPG Bill Cannon – U.S. Geological Survey 4) Granitoid rocks of the Pembine-Wausau Terrane in northeastern Wisconsin Klaus Schulz – U.S. Geological Survey Marcia Bjørnerud – Lawrence University xx �Sponsors The following organizations and individuals made general contributions to the 64th Annual Meeting. We thank them for their commitment to the Institute on Lake Superior Geology. All of the funds contributed this year go toward supporting student travel and registration. INDIVIDUAL CONTRIBUTORS TO STUDENT TRAVEL SCHOLARSHIP MARY KAY ARTHUR STEVEN BAUMANN L. GORDON MEDARIS, JR. With an especially generous donation provided by RON SEAVOY xxi �REPORT OF THE CHAIRS OF THE 63rd ANNUAL MEETING INSTITUTE ON LAKE SUPERIOR GEOLOGY WAWA, ONTARIO The 63rd Institute on Lake Superior Geology (ILSG) was held in Wawa, Ontario, Canada on May 8-12, 2017 with headquarters at the Michipicoten Memorial Community Centre. This was only the second time in the 63 year history of ILSG that the annual meeting has been held in Wawa, the first time in 1987. The meeting was held completely during regular work days (M-F) with technical sessions on Wednesday and Thursday breaking with past tradition of technical sessions being on Thursday and Friday with post-meeting field trips on Saturday. The meeting was cochaired by Anthony Pace (Ontario Geological Survey, Ministry of Northern Development and Mines, Sault Ste Marie, Ontario), Ann Wilson (Ontario Geological Survey, Ministry of Northern Development and Mines, Timmins, Ontario) and Ted Bornhorst (A. E. Seaman Mineral Museum, Michigan Technological University, Houghton, Michigan). Margaret Hanson (A. E. Seaman Mineral Museum, Michigan Technological University, Houghton, Michigan) was registrar for the meeting and co-editor and co-compiler of the two parts of the Proceedings Volume. The meeting was attended by a total of 135 registrants, which exceeded our expected 100 registrants. This included 37 students which is near 30 %, similar to but slightly less than student participation in recent past meetings. The technical sessions and field trips are equally important components of the annual ILSG meeting. The two days of technical sessions included a total of 28 oral presentations (15 by students) and 22 poster presentations (11 by students). The oral presentations included a wide variety of geologic topics from across the Lake Superior region. They were organized to provide a mix of professional and student presentations rather than by themes. There were no student oral presentations scheduled on the afternoon of the second day to facilitate the decision making for the Doug Duskin Student Paper Awards. A separate block of time was set aside during the technical sessions for poster presentations rather than the past practice of poster sessions being held during the social and coffee breaks. The best student oral presentation was by Ross Salerno (University of Minnesota, Duluth) who presented on the Vermilion Granitic Complex of northern Minnesota; the best student poster presentation was by Morgan Sanger (University of Wisconsin, Madison) who presented on seismic interpretation of the Midcontinent rift. We are especially grateful to the members of the Student Paper Awards Committee who must attend each and every talk and truly listen to them! Each year overall presentations by students are improving and this makes the task of identifying the best among them more and more difficult. We thank Mark Puumala (Ontario Geological Survey), Amy Radakovich (Minnesota Geological Survey), and Laurel Woodruff (U. S. Geological Survey) for being willing to judge the student papers. Field trips are an essential and important part of the ILSG annual meeting. All of the field trips were filled to capacity with 85 participating in at least one field trip. There were six field trips for the Wawa meeting, three pre-meeting and three post-meeting. The pre-meeting field trip #1 was a two day trip on May 8 and 9, 2017 that was based in Marathon, Ontario, about 190 km driving distance northwest of Wawa: Archean and Proterozoic geology of the Marathon-Hemlo area led by Allan MacTavish (Panoramic PGMs Canada Ltd), Mark Puumala, Mark Symk, and Tom Muir (Ontario Geological Survery), David Good (University of Western Ontario), and John McBride (Stillwater Canada Inc.). The other two pre-meeting field trips, #2 and #3, were one day xxii �on May 9, 2017 based out of Wawa: More unusual diamond-bearing rocks of the Wawa area led by Ann C. Wilson (Ontario Geological Survey) and Geology of the Wawa gold project led by Jean-Francois Montrueuil, Quentin Yarie, and Conrad Dix (Red Pine Exploration Ltd.). Two of the three post-meeting field trips (#4 and #5) were one day on May 12, 2017 based out of Wawa: Geology of the Island Gold Mine led by Doug MacMillan, S. Comtois-Urban, and Harold Tracanelli (Richmont Gold Mines Ltd.) and Geology of the Renabie area led by Lise Robichaud (Ontario Geological Survey) and Jordan McDivitt (Laurentian University). The other postmeeting field trip (#6) was for one day on May 12, 2017 but relocated late afternoon for overnight in Chapleau, Ontario: Kapuskasing structural zone and Borden Lake Gold deposit led by Pierre Bousquet (Ontario Geological Survey) and Jason Rickard (Goldcorp Inc.). The annual ILSG banquet was held at the Michipicoten Memorial Community Centre on Wednesday evening, May 10 and was attended by 81 individuals. The attendees were treated to a home cooked banquet meal followed by awarding of the the 2017 Goldich Medal to Philip Fralick of Lakehead University. Mark Smyk (Ontario Geological Survey) presented a summary of Phil's contributions to the understanding of Lake Superior geology to banquet attendees prior to awarding him the Goldich Medal. Phil has made significant contributions to the ILSG since 1985; he co-chaired the annual meeting in 2000 and has contributed to more the 75 ILSG abstracts and field trip guidebooks. The banquet presentation was by Johanna Rowe (historian and author from Wawa) who enlightened us on people involved in the long mining history of the Michipicoten area. Several in the local community came to the talk including Mickey Clement who was the first person to bring a sample of diamonds to the Wawa field office; the attendees gave him a round of applause. At the 2016 Board of Directors meeting in Duluth the board adopted a new award, Pioneer of Lake Superior Geology, at the suggestion of Gene LaBerge, 1995 Goldich Medalist and Chair of the 1984 ILSG. The co-chairs selected Douglass Houghton (1809-1845) as the first ILSG "Pioneer of Lake Superior Geology." As the first formal presentation, Ted Bornhorst introduced the new award program and encouraged nominations to be sent to the 2017 co-chairs and followed his introduction of the award by a biographical sketch of honoree Douglass Houghton focused on the attributes that led to his success at such a young age. Pioneers of Lake Superior Geology have contributed to the understanding of geology in the Lake Superior region primarily before the inception of the ILSG in 1955. The Eisenbrey Student Travel Awards are supported by the Institute on Lake Superior Geology and by generous donations by corporations, societies, and individuals. A total of $2,500 US was awarded to students with varying amounts based on distance from Wawa, being senior author on a presentation, traveling with another senior author student presenter, and/or traveling with another student. Only students who applied by the deadline were given an award. This year we thank Argonaut Gold, Geological Society of Minnesota, Mary Kay Arthur, Gordon Medaris, and Ron Seavoy for providing funds, in addition to ILSG, to help us support student participation in the annual meeting of ILSG. The following students were provided financial assistance to attend the meeting in Wawa: Stephen Hanson, Ann Hunt, Ross Salerno, Margaret Upton (University of Minnesota, Duluth), Munira Afroz, Kira Arnold, Brittany Ramsay (Lakehead University), Morgan Sanger, Luke Schranz (University of Wisconsin, Madison), Juliana Olsen-Valdez, and David Wilkes (Lawrence University). xxiii �The Institute’s Board of Directors met on Thursday May 5th to discuss the business of the Institute. The meeting was attended by meeting co-chairs Anthony Pace, Ann Wilson, Theodore Bornhorst, Rob Cundari (2018), Jim Miller (2017), Treasurer Mark Jirsa, Secretary Peter Hollings and guests Esther Stewart, Laurel Woodruff, and Bill Cannon. Secretary Hollings took the minutes of the Board meeting that are as follows: 1. Accepted report of the Chairs for the 62nd ILSG, Duluth, Minnesota; as printed in the Proceeding Volume (Miller), and minutes of last Board meeting, May 5, 2016 (Hollings) 2. Received, discussed, and accepted 2015-2016 ILSG Financial Summary (Jirsa). 3. Received, discussed, and accepted 2015-2016 report of the Secretary (Hollings). 4. Approved Anthony Pace as on-going ILSG Board member. 5. Discussed and approved renewal of Mark Jirsa as Institute Treasurer (end of term 2020). This was later approved by a vote of the membership. 6. Approved Iron Mountain as the site for the 64th annual ILSG meeting. The meeting will be hosted by Esther Stewart, Laurel Woodruff and Bill Cannon. 7. There was discussion as to future meeting locations in Wisconsin with suggested possibilities of Wisconsin Dells and Terrace Bay. Discussion of other locations included mention of Sudbury. 8. Discussed and approved replacing Helene Lukey as the “member from industry” on Goldich Committee (end of term 2017) with Dan England 9. Discussed the possibility of co-hosting ILSG 2020 with the NCGSA meeting – item was tabled pending further discussion with NCGSA organisers. 10. It was agreed that Amy Radakovich (MGS) would be the second signatory on the ILSG accounts. 11. It was agreed that the local chairs would have the final decision as to whether or not to allow silent auctions in support of the ILSG or affiliated student groups. 12. The topic of the A. E. Seaman Mineral Museum serving as registrar (by electronic and check payment) instead of the meeting Chairs running registration through an outside paid service such as Eventbrite was introduced for discussion by Ted Bornhorst. It was agreed that Ted would provide an estimate of the costs involved. 13. It was agreed that the hosting of the ILSG volumes would be relocated from the PRC server to Hollings’ account on Lakehead University servers. Hollings to complete the move ASAP. The dedication and perseverance of the local businesses played an important role in the success of the 2017 ILSG meeting. The co-chairs thanked the community of Wawa through a letter to the Mayor of Wawa, Ron Rody: The staff of the Wawa Economic Development Corporation for providing us with a list of motel accommodations, community contacts and providing all participants with a bag Wawa souvenirs. The staff of the Michipicoten Community Centre, who provided us with the venue to host this event and the staff that worked the bar during our evening social and banquet. We express our sincere gratitude towards Judy Moore and her staff, who catered the event. Many who attended complimented on the food and service she provided. A job well done! The staff of the local Subway shop, who provided the lunches for 5 of the 6 geological field trips during the week. Larry Lacroix of Lloyd’s of Wawa who provided the school bus transportation that was needed for the geological field trips throughout the Wawa and Chapleau areas. Matt Larrett from Michipicoten High School, who provided us with the speaker system for the two days of technical sessions. We thanked the local motels and lastly, Johanna Rowe, who was our guest speaker at the banquet. xxiv �We the 2017 co-chairs would like to again thank all those who continue to make ILSG one of the best regional geoscience meetings in North America: participants, presenters, field trip leaders, session chairs, best student paper committee members, Goldich committee members, ILSG Board members and the incoming 2018 chairs. We appreciated all of support and positive comments about the Wawa meeting and look forward to seeing many of you at the 2018 ILSG in Iron Mountain. Respectfully submitted, Theodore J. Bornhorst, Anthony Pace, and Ann C. Wilson Co-chairs, 63rd Institute on Lake Superior Geology xxv �TECHNICAL PROGRAM TUESDAY MAY 15, 2018 Field trips 1 and 2 begin and end at the Pine Mountain Lodge, Iron Mountain, Michigan 8:00 am - 5:00 pm PRE-MEETING FIELD TRIPS 1) Archean and Paleoproterozoic Geology of the Felch District, Central Dickinson County, Michigan Bill Cannon, Klaus Schulz, Robert Ayuso - U.S. Geological Survey Tom Mroz – BSGE, MSPG, CPG 2) Geology of the Hemlock Formation Tom Waggoner – Consulting Geologist 4:00 pm - 10:00 pm Registration (Pine Mountain Lodge) 7:00 pm - 10:00 pm Welcoming Reception (Pine Mountain Lodge) Poster Session (Pine Mountain Lodge) WEDNESDAY MAY 16, 2018 7:30 am – 11:30 am Registration (Pine Mountain Lodge) 8:00 OPENING REMARKS (Pine Mountain Lodge) Laurel Woodruff, Bill Cannon, Esther K. Stewart, Co-Chairs, 2018 ILSG TECHNICAL SESSION I Session Chairs: Shannon Zurevinski – Lakehead University Ben Drenth – U.S. Geological Survey 8:10 Christian Schardt and Mady David High-technology metal behavior in ore-forming environments and its implication for the Vermilion District, northern Minnesota 8:30 Andrea Reed Pilot study results for potential lithium mineralization on state-managed mineral rights in Minnesota xxvi �8:50 *Matthew W. Matko and Christian Schardt Microanalysis of rock and mineral textures and its relationship to mineralization and ore comminution 9:10 Jeffrey L. Mauk Geochemical signatures of hydrothermal alteration in clastic sedimentary rocks: theory, recognition, and application 9:30 COFFEE BREAK 9:50 Robert Cundari, Mark Smyk, Dorothy Campbell, and Mark Puumala Possible emplacement controls on diamond-bearing rocks north of Lake Superior 10:10 *Joseph Rasmussen, Esther Kingsbury Stewart, John Skalbeck, and Madeline Gotkowitz Modeling the Precambrian topography of Columbia County, Wisconsin using twodimensional models of gravity and aeromagnetic data and well construction reports 10:30 David Southwick, Val Chandler, and Mark Jirsa Geophysical, structural, and tectonic interpretation of the Yellow Medicine and Appleton shear zones, SW Minnesota and SE South Dakota: A work in progress 10:50 Val Chandler, David Southwick, and Mark Jirsa Recent gravity and magnetic investigations of the Minnesota River Valley Subprovince: New insights into ancient problems 11:10 Benjamin J. Drenth, Laurel G. Woodruff, Klaus J. Schulz, William F. Cannon, and Robert A. Ayuso On the source(s) of the Felch-Arnold gravity anomaly, Upper Peninsula, Michigan 11:30 End of Technical Session I 11:30 LUNCH BREAK ILSG BOARD OF DIRECTORS MEETING TECHNICAL SESSION II Session Chairs: Dean Peterson – Natural Resources Research Institute Suzanne Nicholson – U.S. Geological Survey 1:00 *Kira Arnold, Pete Hollings, Seamus Magnus, Shannon Zurevinski, and Robert Creaser Geology and geochemistry of the Terrace Bay Batholith, N. Ontario xxvii �1:20 *Simon Dolega and Philip Fralick Geochemistry of shallow and deep water Archean meta-iron formations and their post depositional alteration in Western Superior Province, Canada 1:40 *Victoria Stinson and Y. Pan Neoarchean to Paleoproterozoic reconstructions using metamorphic core complexes as evidence of continental transform plate motion and their implications in Archean tectonics 2:00 Wouter Bleeker Archean BIF clasts vs. Paleoproterozoic jasper clasts? The proof is in the pudding (stone) 2:20 COFFEE BREAK 2:40 Paul Eger, Courtnay Bot, Dave Meineke, and Dave Adams What to do after the bull has left the china shop- Picking up the community relation pieces 3:00 *Vittoria Smith and Shannon Zurevinski Petrology and 11B composition of tourmaline within the 2685 Ma Ghost Lake Batholith and Mavis Lake Pegmatites 3:20 Klaus J. Schulz, William F. Cannon, and Laurel G. Woodruff Geochemistry of mafic rocks in Dickinson County, Michigan: Evidence for ~2.1 Ga Rifting 3:40 Thomas W. Buchholz, Alexander U. Falster, and Wm. B. Simmons Possible alumotantite from the Nine Mile pluton, Wausau Complex, Marathon County, WI. 4:00 POSTER VIEWING- AUTHORS WILL BE PRESENT AT THEIR POSTERS 5:00 END OF TECHNICAL SESSION II 6:00 RECEPTION AND CASH BAR (Pine Mountain Lodge) 7:00 ANNUAL BANQUET (Pine Mountain Lodge) • Announcement of 65th Annual Meeting Location • • 2018 Goldich Award Presentation to Val Chandler Banquet Presentation - Nancy Langston (Michigan Technological University) Presentation title: Sustaining Lake Superior xxviii �THURSDAY MAY 17, 2018 8:00 OPENING REMARKS, UPDATES (Pine Mountain Lodge) Laurel Woodruff, Bill Cannon, Esther K. Stewart, Co-Chairs, 2018 ILSG TECHNICAL SESSION III Session Chairs: Amy Radakovich – Minnesota Geological Survey Daniel Holm – Kent State University 8:10 Daniel Holm, Terrence J. Boerboom, and Scott Scheiner Reinterpretation of the ages of deposition and folding of Animikie Basin metasedimentary units in east-central Minnesota 8:30 Joshua J. Schwartz, Esther Kingsbury Stewart, and L. Gordon Medaris Jr.+ Detrital zircons in the Waterloo Quartzite, Wisconsin: Implications for the ages of deposition and folding of supermature quartzites in the Southern Lake Superior Region 8:50 Brad Gottschalk, Caroline Rose, and M. Carol Mccartney Geologic history meets the web – online data of the Lake Superior Division of USGS 9:10 William J. Hinze Mapping the Midcontinent Rift System 9:30 COFFEE BREAK 9:50 Jennifer Smith, Wouter Bleeker, Dean Rossell, and Justin Laberge Compositional and geochemical characteristics of the Crystal Lake intrusion, Ontario 10:10 Sean O’Brien, Pete Hollings+, and Jim Miller Geology of the Crystal Lake Gabbro and the Mount Mollie Dyke, Midcontinent Rift, Northwest Ontario 10:30 *Dustin A. Liikane, Wouter Bleeker, Mike Hamilton, Sandra Kamo, Jennifer Smith, Peter Hollings, Robert Cundari, and Michael Easton Controls on the localization and timing of mineralized intrusions within the ca. 1.1 Ga Midcontinent Rift system 10:50 David Good Petrogenesis of mafic magmatism in the Coldwell Complex Part 1. Geochemical model to explain origin of metabasalt by partial melting in the SCLM xxix �11:10 Evgeniy Kulakov, Theodore J. Bornhorst+, Chad Deering, and James B. Moore The youngest magmatic activity of the Midcontinent Rift at Bear Lake, Keweenaw Peninsula, Michigan 11:30 End of Technical Session III 11:30 LUNCH BREAK TECHNICAL SESSION IV Session Chairs: Marcia Bjørnerud – Lawrence University John Esch – Michigan Department of Environmental Quality 1:00 Kelli McCormick, Kevin Chamberlain, and Colin Paterson An 1149 Ma U-Pb baddeleyite crystallization age and geochemistry of gabbroic intrusions at the southwestern margin of the Superior Craton, southeastern South Dakota 1:20 Jim DeGraff and B.T. Carter Thrust Kinematics of the Keweenaw Fault North of Portage Lake, Michigan 1:40 John A. Yellich Michigan Geological Survey six years after assignment to Western Michigan University, where are we today? 2:00 John M. Esch LiDAR Revolutionizing Geological Mapping 2:20 COFFEE BREAK 2:40 Dean M. Peterson Assembling Minnesota: Integration of 140 years of government, academic, and industry geologic studies into a seamless statewide GIS database 3:00 Mark A. Jirsa and others On-going geologic mapping in Minnesota’s Arrowhead Region by the Minnesota Geological Survey 3:20 John M. Esch, Alan Kehew, Sebastian Huot, and John Yellich Surficial geology of the Iron Mountain 7.5 Minute Quadrangle, Dickinson County, Michigan, Florence & Marinette Counties, Wisconsin xxx �3:40 Phil Larson, George Hudak, Al Mactavish, Peter Hinz, Amy Radakovich, Juk Bhattacharyya, Paula Engelhardt, Steve Engelhardt, Brigitte Gelnias, David Good, Emily Gorner, Sheree Hinz, Peter Jongewaard, Deb Kroch, Matt Svensson, and Andrew Tims Land of fire and ice: Summary of the 2017 ILSG field trip to Iceland 4:20 BEST STUDENT PAPER AWARDS STUDENT TRAVEL AWARDS 4:40 END OF TECHNICAL SESSIONS * denotes a student eligible for Best Student Paper Award. To be eligible students must have graduated no more than one month before the ILSG meeting, be first author, and present the paper at the meeting. + denotes author that will present abstract, if different than the first author. xxxi �FRIDAY MAY 18, 2018 8:00am – 5:00pm POST-MEETING FIELD TRIPS Field trips 2 and 3 begin and end at the Pine Mountain Lodge, Iron Mountain, Michigan 3) Geology and Iron Ores of the Menominee Iron Range, Dickinson County, Michigan Tom Mroz – BSGE, MSPG, CPG Bill Cannon – U.S. Geological Survey 4) Granitoid rocks of the Pembine-Wausau Terrane in northeastern Wisconsin Klaus Schulz – U.S. Geological Survey Marcia Bjørnerud – Lawrence University xxxii �POSTER PRESENTATIONS ANDERSON, Eric, SCHULZ, Klaus, DRENTH, Benjamin, CANNON, William, and QUIGLEY, Thomas New gravity and high-resolution aeromagnetic data provide insights into Precambrian geology in the eastern Pembine-Wausau terrane *ASHAUER, Zachary, CURRIER, Ryan, NORFLEET, Mark Textural analyses of rapakivi mantles: Evidence for semi-selective replacement in Proterozoic rapakivi granites AYUSO, R.A., SCHULZ, K.J., CANNON, W.F., WOODRUFF, L.G., VAZQUEZ, J.A., FOLEY, N.K., and JACKSON, J. New U-Pb zircon ages for rocks from the Granite-Gneiss Terrane in Northern Michigan: Evidence for events at ~3750, 2750, and 1850 Ma *BLOTZ, Kaelyn E., LODGE, Robert W.D. Ore petrography of the Flambeau volcanogenic massive sulfide deposit, northwestern Wisconsin: Implications for hydrothermal fluid composition BOERBOOM, Terrence J. Fault-controlled dike emplacement in the Grand Marais, Minnesota area *DRAZAN, Jacqueline, BRENGMAN, Latisha, FEDO, Christopher Preliminary petrographic and geochemical investigation of silicified volcanic rocks and silica-rich exhalative rocks from the ~2.7 Ga Abitibi Greenstone Belt, Canada EASTON, Robert M. GEON 12 to 11 history of the Lake Superior Region and speculation about the relationships between the Midcontinent Rift and the Grenville Orogen ESCH, John M, KEHEW, Alan, HUOT, Sebastien, YELLICH, John Surficial geology of the Iron Mountain 7.5 Minute Quadrangle, Dickinson County, Michigan, Florence & Marinette Counties, Wisconsin *FITZPATRICK, William, HOOPER, Robert, and LODGE, Robert, Gélinas, Brigitte Mineral chemistries of the Tower Mountain Intrusive Complex Au-deposit, Ontario GRAUCH, V.J.S., BEDROSIAN, Paul A., STEWART, Esther Kingsbury, and HELLER, Samuel Inferences on the subsurface distribution of Oronto and Bayfield Groups north and west of the Douglas Fault, Northwestern Wisconsin xxxiii �GREEN, Carlin J., SEAL, Robert, R., II, CANNON, William F., PIATAK, Nadine, and MCALEER, Ryan J. Origin, distribution, morphology, and chemistry of amphiboles in the Ironwood IronFormation, Gogebic Iron Range, Wisconsin, U.S.A. *HAFFTEN, Doug and RADWANY, Molly Geothermobarometry of a Precambrian amphibolite from Cornell WI *HANNACK, Gina, and RADWANY, Molly Hornblende-plagioclase thermometry of the Eau Claire River Complex, western Wisconsin *HONE, Samuel V. and ZIEG, Michael J. Olivine crystal size distribution in the Black Sturgeon Sill, Nipigon, Ontario *JACOBSON, Regan E., LODGE, Robert W.D Reconstructing Paleoproterozoic volcanism in northwestern Wisconsin: Geochemistry of the Flambeau Cu-Zn-Au Mine JIRSA, Mark A., STARNS, Edward C., and SCHMITZ, Mark D. Geology and geochronology of the 2006 Cavity Lake forest fire area, Boundary Waters Canoe Area Wilderness, NE Minnesota KINGSBURY STEWART, Esther, STEWART, Eric D., and ROUSHAR, Kathy New bedrock geologic mapping of Dodge County, Wisconsin provides evidence for Paleozoic reactivation of Precambrian structures *LIIKANE, Dustin A., BLEEKER, Wouter, HAMILTON, Mike, KAMO, Sandra, SMITH, Jennifer, HOLLINGS, Peter, CUNDARI, Robert, and EASTON, Michael Controls on the localization and timing of mineralized intrusions within the ca. 1.1 Ga Midcontinent Rift system MATTOX, Stephen, BOLHUIS, Chris, and SOBOLAK, Christina Using credit-by-exam to connect advanced high school geology courses to university geology departments: Current status of a state-wide program in Michigan *OLSEN-VALDEZ, Juliana and BJØRNERUD, Marcia The Brussels Hill Structure, Door County, Wisconsin: Impact crater, diatreme or other? *OLSON, Maile J., LODGE, Robert W. D. Komatiite-hosted nickel-copper mineralization potential in the eastern Shebandowan Greenstone Belt, Ontario, Canada * ROSE, Katharine; ESSIG, Espree, and THAKURTA, Joyashish Variation trends in sulfur isotope ratios at the Eagle and East Eagle intrusions and the surrounding country and basement rocks of the Baraga Basin, Upper Peninsula, Michigan xxxiv �*RUPP, Kevin, THAKURTA, Joyashish, and MAHIN, Robert Preliminary investigation of the East Eagle Intrusion Gabbro in Marquette County, Michigan. * TYRRELL, C.W., HUBBELL, G.E., and DEGRAFF, J.M. Keweenaw Fault geometry and kinematics along Bête Grise Bay, Michigan *UPTON, Margaret, SCHARDT, Christian, HUDAK, George, QUIGLEY, Eric Alteration mineral zonation and geochemical characteristics of the Back Forty Deposit, MI; a replacement-style zinc- and gold-rich volcanogenic massive sulfide deposit * VALL, Kathryn G., STEINMAN, Byron A., POMPEANI, David P., SCHREINER, Kathryn M., DEPASQUAL, Seth Reconstruction of paleoenvironmental conditions and temporal patterns of ancient mining on Isle Royale using biogeochemical analyses of lake sediment YELLICH, John A. Michigan Geological Survey Six years after assignment to Western Michigan University, Where are we today? ZIEG, Michael J. and HONE, Samuel V. The Origin of Layering in the Olivine Zone, Black Sturgeon Sill, Nipigon, Ontario * denotes a student eligible for Best Student Paper Award. To be eligible students must have graduated no more than one month before the ILSG meeting, be first author, and present the paper at the meeting. xxxv �ABSTRACTS xxxvi �New gravity and high-resolution aeromagnetic data provide insights into Precambrian geology in the eastern Pembine-Wausau terrane ANDERSON, Eric1, SCHULZ, Klaus2, DRENTH, Benjamin1, CANNON, William2, and QUIGLEY, Thomas3 1 US Geological Survey, MS 964, PO Box 25046, Denver, CO 80225 USA 2 US Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA 20192 USA 3 Great Lakes Exploration Inc., Menominee, MI 49858 USA The Pembine-Wausau terrane represents a major Paleoproterozoic belt of metavolcanic and intrusive rocks that formed in an island-arc setting at the southern limit of the Archean Superior Craton (Schulz and Cannon, 2007). The island-arc complex was accreted to the continental margin along the Niagara fault zone during the Penokean orogeny and subsequently intruded by syn- to post-tectonic granitoids. The terrane is known to host a number of significant volcanogenic massive sulfide deposits including the Back Forty deposit. Limited outcrop makes bedrock mapping difficult. In 2016, the USGS contracted a high-resolution aeromagnetic survey over parts of the Pembine-Wausau terrane. The data were collected along north-south flight lines spaced 150 m at a nominal height of 80 m. These data, along with existing and in-fill gravity stations and physical property measurements, are helping improve Precambrian bedrock maps (Figure 1; Sims, 1990; Sims and Schulz, 1993). The complete (terrain-corrected) Bouguer gravity anomaly map shows strong gradients, indicating significant lateral variations in bedrock density. A west-northwest trending linear high with ~5 mGal amplitude occurs along splays of the Niagara fault and correlates with mapped mafic-ultramafic rocks having measured density of ~2.90 g/cm3. South of the Niagara fault zone, a strong gravity gradient trends east-west along the southern mapped extent of the McAllister Formation that consists of basaltic and andesitic rocks with a density of ~2.91 g/cm3. South of the gradient is a linear low that expands to the south and east beneath Phanerozoic cover. A broad high with amplitude of ~15 mGal occurs over the southern extent of the Athelstane and Amberg granites. However, the anomaly differs from geologic map patterns and aeromagnetic magnetic anomalies and so the source is not well understood. A standard reduction-to-pole (RTP) transformation was applied to the aeromagnetic data to better align anomalies with causative sources. The RTP map shows broad positive anomalies with amplitude around 2000 nT north of the Niagara fault. Between fault splays is a series of west-northwest trending linear highs with amplitude around 1000 nT. The linear features are 1.5 to 3 km-long and 500 m-wide; some of these correlate with mapped mafic-ultramafic rocks. The Pembine ophiolite rocks are well imaged by ~2000 nT anomaly high; several additional high amplitude anomalies occur within the Quinnesec Formation. A west-northwest trending aeromagnetic gradient is observed at the southern extent of the McAllister Formation that broadly parallels the strong gravity gradient. To the south are linear north-south to northnortheast trending magnetic highs that extend for more than 10 km. These linear features have 1 �amplitudes between 5 and 60 nT. Near Amberg, they are associated with diabase dikes (magnetic susceptibilities of about 15 x 10-3SI) that cut the late tectonic Athelstane Quartz Monzonite. At the southern end of the mapped Athelstane and Amberg granites is an oval magnetic high that trends northeastward contradictory to geologic map patterns. The amplitude (~325 nT) is significantly less than the anomalies observed over the mafic-ultramafic rocks to the north, suggesting a different source rock composition. The tilt and first vertical derivative maps were derived from the RTP data to accentuate near surface and subtle magnetic features. Both maps show linear trends that change orientations proximal to gravity gradients. The prominent north-south to northeast trending magnetic lineaments do not appear to extend much beyond the gravity gradient into the McAllister Formation. In addition, these lineaments appear to have several subparallel northeast trending discontinuities. Magnetic lineaments in the Pemene Formation parallel mapped trends in the volcanic rocks. Near the Niagara fault the extent of the west-northwest trending magnetic lineaments is much better resolved in the derivative maps than in the RTP maps. Figure 1: Geologic map (A) and reduced-to-pole (RTP) anomaly map (B) of the study area. References Schulz, K.J. and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region, Precambrian Research, 157: 4-25. Sims, P.K., 1990. Geologic map of Precambrian rocks of Iron Mountain and Escanaba 1° x 2° quadrangles, northeastern Wisconsin and northwestern Michigan, U.S. Geological Survey Miscellaneous Investigations Series, Map I–2056, scale 1:250,000. Sims, P.K., and Schulz, K.J., 1993. Geologic map of Precambrian rocks of parts of Iron Mountain and Escanaba 30’ x 60’ quadrangles, northeastern Wisconsin and adjacent Michigan, U.S. Geological Survey Miscellaneous Investigations Series, Map I–2356, scale 1:100,000. 2 �Geology and Geochemistry of the Terrace Bay Batholith, N. Ontario ARNOLD, Kira1, HOLLINGS, Pete1, MAGNUS, Seamus2, ZUREVINSKI, Shannon1, CREASER, Robert3 1 Department of Geology, Lakehead University, Thunder Bay, Ontario P7B5E1 Ontario Geological Survey, Ministry of Northern Development and Mines, Earth Resources and Geoscience Mapping Section, 933 Ramsey Lake Road, Sudbury, ON, P3E 6B5, Canada 3 Department of Earth & Atmospheric Sciences, University of Alberta, 126 ESB Edmonton, Alberta, T6G2R3, Canada 2 The Terrace Bay Batholith is a 25 km long oval shaped granitoid intrusion located in the western portion of the Schreiber-Hemlo greenstone belt, part of the larger Wawa-Abitibi terrane. The pluton, emplaced at 2689+/-1.1 Ma (Kamo, 2016) intrudes circa 2720 Ma metavolcanic rocks, and a nearby pluton of equivalent age intrudes circa 2698-2693 Ma clastic metasedimentary rocks (Kamo, 2016; Davis and Sutcliffe, 2017). Younger plutonism in the region occurred between 2673 and 2667 Ma (Kamo, 2016, Kamo and Hamilton, 2017). This study describes and classifies the Terrace Bay batholith in order to investigate its petrogenesis and related gold and base metal mineralization. The core of the Terrace Bay Batholith is a massive, homogeneous equigranular and locally quartz porphyritic granodiorite. The granodiorite typically consists of medium- to coarsegrained quartz and feldspar phenocrysts with a groundmass of fine-grained feldspars, quartz, amphibole, biotite and disseminated magnetite and sulphide minerals. Multiple outcrops in the center of the batholith host very coarse-grained phenocrysts of feldspar, ranging in size from 1 to 3 cm. An outcrop of diorite was found in the center of the pluton, composed of medium-grained amphibole and plagioclase, with very few quartz crystals. Some areas of the diorite outcrop are monzodioritic, with over 5% potassium feldspar. Thick overburden which covers the contacts with the granodiorite making their relationship uncertain, but the diorite likely represents either a more mafic phase of the granitic magma or possibly an autolith. Geochemically the granodiorite is classified as I-type granite. It has a calc-alkaline signature, characteristic of rocks formed above a subduction zone. On a primitive mantlenormalized trace element profile the samples are LREE enriched with unfractionated HREE and prominent negative Nb-Ti anomalies. The diorite shows a similar trend to the granodiorite suggesting that it was formed from a similar if not the same source. Two distinct alteration styles have been observed in the pluton; a common pervasive potassium-hematite alteration and a less common chlorite and epidote alteration. The chlorite– epidote variety is an intense alteration but is restricted to veins and dykes. The potassiumhematite alteration has been observed across the batholith. In most cases, the groundmass is obliterated and composed of fine- to very fine-grained hematite and potassium feldspars. Phenocrysts of quartz are typically unaltered but relict feldspars have been sericite altered. The chlorite-epidote alteration is generally composed of fine-grained chlorite and epidote in the groundmass with quartz phenocrysts and relict sericite-altered feldspars. The pluton is crosscut by quartz carbonate veins, which locally contains black tourmaline along the vein contacts. Mineralization in the quartz veins includes pyrite, chalcopyrite, galena, molybdenite and arsenopyrite. In contrast, the pluton typically hosts only pyrite and molybdenite. Generally the molybdenite present in the granite is disseminated, but has been 3 �found to occur in coarse pods up to 3 cm wide. Occurrences of molybdenum mineralization are spatially correlative with the gold mineralized occurrences, which are most commonly located in quartz-veined and altered zones near the contacts of the pluton. A molybdenite sample yielded a mineralization age of 2671 +/- 12 Ma. Figure 1. Simplified bedrock geology map of the Terrace Bay batholith and surrounding greenstone belt in Priske, Strey and Syine townships. Modified from Arnold et al. (2017). REFERENCES Arnold, K.A., Hollings, P. and Magnus, S.J. 2017. Geology and mineral potential of the Terrace Bay pluton, western Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2017, Ontario Geological Survey, Open File Report 6333, p.12-1 to 12 Davis, D.W. and Sutcliffe, C.N. 2017. U-Pb geochronology by LA-ICPMS in samples from northern Ontario; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 131p. Kamo, S.L. 2016. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 1: 2015-2016; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. Kamo, S.L. and Hamilton, M.A. 2017. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 2: 2016-2017; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 72p. Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L. and Sage, R.P. 1991. Wawa Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p. 485-539. 4 �Textural Analyses of Rapakivi Mantles: Evidence for Semi-Selective Replacement in Proterozoic Rapakivi Granites ASHAUER, Zachary1, CURRIER, Ryan1, NORFLEET, Mark1 1 Natural and Applied Sciences, University of Wisconsin Green Bay, 2420 Nicolet Dr. Green Bay, Wisconsin 54311 Rapakivi granite complexes are commonly associated with caldera forming eruptions (Karell et al., 2014). Characteristic of these granites is the rapakivi texture, the mantling of plagioclase on K-feldspar meagcrysts. First described in scientific literature by Sederholm in 1891, consensus on a model for rapakivi texture formation remains elusive. Textural analyses that help constrain the mechanism of formation are presented here. The textural analysis consists of a systematic survey of mantle thicknesses in relation to the mantled feldspar radius, and was conducted on two classical rapakivi systems: the Wolf River Batholith (1.48-1.46 Ga; Dewane and Van Schmus, 2007) Wisconsin and the Wiborg Batholith (1.65-1.62 Ga; Rämö, 1991) southeastern Finland. Mantle analysis samples were obtained from dimension stone slabs of the Waupaca Wiborgite in the Wisconsin Capitol Rotunda and of Ylämaa Wiborgite (Baltic Brown) slabs from five Green Bay area businesses. The Wolf River and Wiborg batholiths are A-type intrusive bodies underlying areas >9,000 km2, containing wiborgite variety of rapakivi granite, Waupaca Wiborgite and Ylämaa Wiborgite, respectively. Wiborgite is a porphyritic granite containing ovoidal megacrysts of Kfeldspar ranging upwards of 6 cm in diameter, with over 50% of megacrysts mantled by 1-4 mm of plagioclase. The Waupaca Wiborgite contains a greater population of euhedral K-feldspars than the Ylämaa Wiborgite where nearly all crystals are ovoidal in shape. Mantled feldspar cores and mantles were outlined by hand from high-resolution pictures of slabs. Mantle and core area dimensions were calculated using image analysis software Image J (Schneider et al., 2012) and converted to respective radii for comparison. Results illustrate a trend of thickest mantles developing on the middle size class of crystals, which is consistent across all samples and the two separate systems (Figure 1). Data density plots stretch out along the x-axis; implying larger radius crystals generally have smaller thickness mantles. To properly interpret results, a model was produced replicating variable mantle and crystal radii size scenarios observed from a 2D slice of mantled spheres. The model evaluates three scenarios of mantle thickness in relation to increasing mantled feldspar radius: (1) mantle thickness is variable with crystal radius, (2) mantle thickness is a consistent proportion of crystal radius, and (3) mantle thickness decreases with increasing crystal radius. Scenarios 1 and 2 overestimate mantle thicknesses and display data distributions inconsistent with mantle analysis results (Figure 2). Scenario 3 closely resembles mantle analysis results showing thickest mantles occur on the middle size class of crystals and concentrate closest to the x-axis. Mantle analysis coupled with theoretical modeling suggests mantle thickness has dependence on mantled feldspar size. This is interpreted as forming within a contact melt zone, driven by underplating of hot magma, which resulted in vigorous stirring once a tipping point is reached through buoyant instability. This model thus suggests dissolution-controlled replacement mantle growth in an up-temperature regime, consistent with caldera volcanism. 5 �Figure 1. Mantle thickness (radius of mantled feldspar – radius of core feldspar) plotted against radius of mantled feldspar. A-E independent slabs of Ylämaa Wiborgite, F compiled slabs of Waupaca Wiborgite. Notice middle size class of crystals generally have thickest mantles. Figure 2. Mantle thickness (radius of mantled feldspar – radius of core feldspar) plotted against radius of mantled feldspar. (A) Random thickness mantle scenario, (B) consistent proportion of mantle thickness to mantled feldspar size, and (C) mantle thickness decreases with increasing mantled feldspar size. Notice y-axis scale is double that of mantle analysis results. References: Dewane, T.J., Van Schmus, W.R., 2007. U-Pb geochronology of the Wolf River batholith, north-central Wisconsin: Evidence for successive magmatism between 1484 Ma and 1468 Ma. Precambrian Research 157, 215-234. Karell, F., Ehlers, C., Airo, M., 2014. Emplacement and magmatic fabrics of rapakivi granite intrusions within Wiborg and Aland rapakivi granite batholiths in Finland. Tectonophysics 614, 31-43. Rämö, O.T., 1991. Petrogenesis of the Proterozioic rapakivi granites and related basic rocks of southeastern Fennoscandia: Nd and Pb isotopic and general geochemical constraints. Geological Survey of Finland, Bulletin 355, 161p. Schneider, C.A., Rasband, W.S., Eliceiri, K.W., 2012. NIH Image to ImageJ: 25 years of image analysis. Nature methods 9 (7): 671-675. Sederholm, J.J., 1891. Ueber die finnländischen Rapakiwigesteine. Tschermak's Miner. Petrograh. Mitth. 12, 1-31. 6 �New U-Pb Zircon Ages for Rocks from the Granite-Gneiss Terrane in Northern Michigan: Evidence for Events at ~3750, 2750, and 1850 Ma AYUSO, R.A.1, SCHULZ, K.J.1, CANNON, W.F.1, WOODRUFF, L.G.2, VAZQUEZ, J.A.3, FOLEY, N.K. 1, and JACKSON, J. 1 1 U.S. Geological Survey, Reston, VA 20192, 2 U.S. Geological Survey, Mounds View, MN 55112, 3 U.S. Geological Survey, Menlo Park, CA 94025 Early Archean rocks are part of the granite-gneiss terrane along the southern margin of the Superior craton [1]. Recently, we reported preliminary zircon age data from the Carney Lake gneiss in the granitegneiss terrane of northern Michigan that indicated an Eoarchean component ca. 3750 Ma [2]. Here we report additional sensitive high-resolution ion microprobe (SHRIMP) U-Pb zircon ages for the Carney Lake gneiss that further document the Eoarchean component. We also report new U-Pb zircon ages for the Hardwood gneiss complex, the Peavy Pond complex, and a porphyritic red granite from northern Michigan. Two samples were collected from the Carney Lake gneiss. Zircons were obtained from a granitic K-feldspar-bearing gneiss that is locally pegmatitic; the zircons range from anhedral to subhedral, contain complex irregular growth zoning, and display multiple growth rims. Zircons also were obtained from a banded and folded gray to red granitic gneiss; the zircons are slightly rounded to subhedral. One sample each was collected from the Hardwood gneiss complex, Peavy Pond complex, and porphyritic red granite. The Hardwood gneiss sample is a fine-grained, layered garnet-pyroxene-quartz-magnetite gneiss (granulite grade) that contains brown and mostly anhedral zircons. The Peavy Pond sample is a mediumgrained granite with honey-colored euhedral to subhedral zircons characterized by doubly terminated prisms. The porphyritic red granite is foliated, contains K-feldspar augen, and has clear to pale brown subhedral zircons that commonly display igneous oscillatory bands. SHRIMP U-Pb data were obtained on handpicked zircons. All peaks, including U, Th, Pb, REE, Hf, Ti, and Y, were measured sequentially. Raw data were reduced using the Squid 2 and Isoplot programs [3, 4]. Figure 1: A. Concordia diagram for 129 spot analyses from zircons in the Carney Lake gneiss. B. Concordia diagram for 56 spot analyses from zircons in the Hardwood gneiss. On a Concordia diagram, U-Pb data for Carney Lake show clusters with points ranging from concordant to discordant (Fig. 1A). The predominant data cluster of nearly concordant points has an intercept ca. 2750 Ma; a smaller concentration of nearly concordant analyses occurs at ca. 3750 Ma. One possible data 7 �alignment spans from an upper intercept age of ca. 3750 Ma to the lower intercept age of ca. 2750 Ma; a second possible alignment spans a range from 2750 Ma toward an imprecisely defined intercept around 1000 Ma (Fig. 1A). The ca. 3750 Ma age on zircon cores from the Carney Lake gneiss is evidence of an Eoarchean component in the granite-gneiss terrane (Fig. 1A). The gneiss was affected by igneous and thermal events at ca. 2750 (and younger), which resulted in new zircon crystallization, recrystallization, and formation of overgrowths. U-Pb zircon dates for the Hardwood gneiss yielded evidence of a Neoarchean component (concordant spot analyses) ca. 2750-2500 Ma as well as younger dates ca. 1900 Ma (Fig. 1B). U-Pb data for the Peavy Pond complex range from concordant to discordant and plot along a trend intercepting Concordia at ca. 1850 Ma (a small data cluster plots at ca. 2600 Ma) (Fig. 2A). The majority of spots for the red granite is concordant or plots adjacent to Concordia at ca. 2099 Ma (age of crystallization) (Fig. 2B). Figure 2: A. Concordia diagram for 32 spot analyses of zircons from the Peavy Pond complex. B. Concordia diagram for 18 spot analyses of zircons from the porphyritic Red Granite. The zircons show typical REE chondrite-normalized patterns (LREE-depleted, HREE-enriched), negative Eu anomalies, and positive Ce anomalies. The Carney Lake gneiss zircons have the most diverse REE patterns and widely variable Eu and Ce anomalies. The Hardwood gneiss also has diverse REE patterns. Trace element ratio plots (e.g., U/Yb vs. Hf) [5] suggest a continental magmatic origin for zircons from the Carney Lake gneiss, Hardwood gneiss, and Peavy Pond complex. A continental arc (or enriched mantle?) is implicated for zircons from the red granite. The ca. 3750 Ma age of the Carney Lake gneiss documents the presence of an Eoarchean component in northern Michigan. The ca. 2750 Ma of the Hardwood gneiss indicates the contribution of a Neoarchean component in the region. Igneous intrusive events occurred ca. 2750, 2099, and 1850 Ma. There is no evidence for an older Archean component in the porphyritic red granite. References [1] Peterman, Z.E., Zartman, R.E., and Sims, P.K., 1980: Geol. Soc. America Sp. Paper 182, p. 125–134. [2] Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., and Jackson, J., 2017, Institute on Lake Superior Geology, Proceedings 63rd Annual Meeting, part 1, p. 9-10. [3] Ludwig, K.R., 2009: SQUID 2, Berkeley Geochronology Center Special Publication no. 5, 110 p. [4] Ludwig, K.R., 2012: Isoplot 3.75, Berkeley Geochronology Center Special Publication no. 5, 75 p. [5] Grimes, C.B., Wooden, J.W., Cheadle, M.J., and John, B.E., 2015, Contrib. Min. Pet., 170: 46. 8 �Archean BIF clasts vs. Paleoproterozoic jasper clasts? The proof is in the pudding (stone) BLEEKER, Wouter Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario, K1A 0E8 Email: wouter.bleeker@canada.ca The ca. 2.50-2.25 Ga Huronian Supergroup (Bennett et al., 1991; Young et al., 2001), a largely intra-cratonic rift and regional cover sequence overlying the southern Superior craton (Bleeker and Ernst, 2006), contains one of the more important records of Paleoproterozoic Earth evolution. Among other things, it contains a superb volcanic rift sequence including subaerial and subaqueous basalt flows, and the Copper Cliff Rhyolite and its subvolcanic pluton, the Creighton Granite, now precisely dated at 2459±7 Ma (Bleeker et al., 2015). This lower rift sequence is terminated by conglomerate/sandstone and greywacke turbidites of the Matinenda and McKim formations, the former hosting important detrital pyrite and uraninite concentrations that have been mined for uranium in the Elliott Lake area (Roscoe, 1969). Overlying this lower rift sequence are several cycles of conglomerate, sandstone and silt- to mudstone, and back to sandstones, three of which start out with glacial diamictites. Minor erosional disconformities or unconformities are present at the base of these cycles. Diamictites of the second cycle are overlain by carbonates of the Espanola Formation, which may represent cap carbonates and, thus, may constitute a globally important marker horizon. The third and aerially most extensive of these cycles starts with the Gowganda Formation, which consists of glacial diamictites at its base and hosts the first red-bed sandstones of the Huronian toward its top. The Gowganda Formation is overlain by white to red sandstones and conglomerates of the Lorrain Formation, which hosts a conspicuous member of red jasper clast conglomerate, locally known as “puddingstone” (Fig. 1). The red jasper clasts, typically 0.5-5.0 cm in size, in a white to off-white quartzdominated matrix, make for an attractive rock type that Figure 1: Top, red jasper clast in is sought after as a decorative stone. For the last century, these jasper clasts have “puddingstone”, very fine-grained and with delicate lamination and textures, and no been interpreted as being derived from Archean metamorphic minerals (no magnetite). Bottom, basement to the north, particularly the ca. 2720-2725 typical recrystallized Abitibi BIF, at about the Ma, black to sometimes red, banded iron formations same scale, with recrystallized texture, a fabric, and metamorphic magnetite. (BIFs) of the Abitibi greenstone belt. Good exposures 9 �of such BIFs occur in the Timmins area, around Temagami, and in Wawa. They have been mined for iron ore at a number of localities including Temagami (Sherman Mine), south of Kirkland Lake (Adams Mine), and Wawa (Helen Mine). All of these BIFs have seen pervasive deformation and low grade metamorphism, and contain abundant magnetite. Several observations lead me to question this interpretation: the jasper clasts of the Huronian puddingstone are often angular, i.e. more or less proximal; in typical puddingstone they suddenly become a dominant clast type, again suggesting a proximal source; there are few if any real BIF clasts; the jasper clasts are extremely fine-grained and delicately textured (Fig. 1) and do not contain magnetite, unlike BIF samples from the Abitibi which are noticeably more recrystallized (an order of magnitude coarser in grain size) and invariably contain metamorphic magnetite (Fig. 1). I conclude that the conspicuous jasper clasts of Lorrain puddingstone are not of Archean derivation, but rather represent penecontemporaneous reworking of otherwise poorly preserved Huronian jasper deposits, possibly associated with a minor volcanic or hydrothermal centre that has not yet been identified. Given that the occurrence of puddingstone is strongly concentrated in, if not unique to, the area around Bruce Mines, the source jasper beds were likely local deposits restricted to that part of the Huronian basin, possibly the fine-grained siliceous siltstone and associated layers that have been referred to in some of the early papers on the Huronian as “Bruce Mines Jasper” (Collins, 1925). The jasper clasts constitute a range from dark red to pure white chert, all of which show delicate textures and layering. Among the white chert-like clasts some resemble unrecrystallized agate, also suggesting deep weathering and reworking of Paleoproterozoic volcanic units containing agate nodules. Rare accompanying clasts of quartz porphyry may allow dating of this part of the Huronian succession. If indeed Lorrain-age jasper, these clasts could host important information about the ambient environment at ca. 2.35 Ga. References Bennett, G., Dressler, B.O., Robertson, J.A., 1991. The Huronian Supergroup and associated intrusive rocks. In: Geology of Ontario, Part 1, P.C. Thurston, H.R. Williams, R.H. Sutcliffe and G.M. Stott (eds.), Ontario Geological Survey, p. 549–591. Bleeker, W., and Ernst, R.E., 2006. Short-lived mantle generated magmatic events and their dyke swarms: The key unlocking Earth's palaeogeographic record back to 2.6 Ga. In: Dyke Swarms—Time Markers of Crustal Evolution, E. Hanski, S. Mertanen, T. Rämö, and J. Vuollo (eds), A.A. Balkema, Rotterdam, The Netherlands, p. 3-26. Bleeker, W., Kamo, S.L., Ames, D.E., and Davis, D., 2015. New field observations and U-Pb ages in the Sudbury area: toward a detailed cross-section through the deformed Sudbury Structure. In: Geological Survey of Canada, Open File 7856, p. 151–166. Collins, W.H., 1925. North shore of Lake Huron. Geological Survey of Canada, Memoir 153, 160 p. Roscoe, S.M., 1969. Huronian rocks and uraniferous conglomerates in the Canadian Shield. Geological Survey of Canada, Paper 68-40, 205 p. Young, G.M., Long, D.G., Fedo, C.M., and Nesbitt, H.W., 2001. Paleoproterozoic Huronian basin: product of a Wilson cycle punctuated by glaciations and a meteorite impact. Sedimentary Geology, vol. 141, p. 233-254. 10 �Ore Petrography of the Flambeau volcanogenic massive sulfide deposit, northwestern Wisconsin: Implications for hydrothermal fluid composition BLOTZ, Kaelyn E., LODGE, Robert W.D. Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 The Paleoproterozoic Flambeau Cu-Zn-Au volcanogenic massive sulfide (VMS) deposit is located near the town of Ladysmith, Wisconsin within the Pembine-Wausau Terrane of the Penokean Orogeny (Schulz and Cannon, 2007) and is one of at least 13 other VMS deposits that have been identified in the state (DeMatties, 1994). The Flambeau is the only deposit to have been mined in Wisconsin after it was discovered by Kennecott Minerals Company in 1968 and was exploited due to the unusual grade of the orebody in the supergene enriched cap (May and Dinkowitz 1996). Mining began in 1993 and lasted until 1997 when extraction of the supergene enriched cap was completed. The high-grade copper ore body produced nearly 1.8 million tons of ore with an average of 10% copper and 0.18 ounces of gold per ton before the open pit was completely refilled and the site was reclaimed (Jones and Jones, 1999). The hypogene geology of the Flambeau deposit is characterized by massive to semi-massive Cu-Zn-Pb sulfides hosted in altered intermediate-felsic rocks that were metamorphosed into chlorite-andalusite-biotite schists. Before metamorphism occurred, these rocks were formed in a submarine hydrothermal system and their compositions can provide insight into the mechanisms of gold enrichment at the Flambeau mine. This study focuses on the identification of trace minerals and mineralogical variations within the ore zone at the Flambeau deposit. Samples were collected from drill core stored at the Wisconsin Geologic and Natural History Survey core repository and were then processed into polished thin sections. There are two main types of primary ore: a massive pyrite-chalcopyrite dominated assemblage and a weakly banded sphalerite-pyrite-galena dominated assemblage (May and Dinkowitz, 1996). Using scanning electron microscopy-energy dispersive spectroscopy, trace ore minerals identified in the ore zone include tellurides (hessite, altaite, tsumoite, bismuth), electrum, arsenopyrite, acanthite, bismuthinite, cassiterite, monazite, and an unnamed tungsten mineral. The presence of these minerals is important in determining the physical and chemical characteristics of the hydrothermal fluids since these trace minerals form under specific hydrothermal conditions. The relative abundance of the trace minerals, coupled with the anomalous Cu-enrichment in the Flambeau felsic-intermediate dominated strata, may indicate that this is not a traditional VMS deposit. Preliminary data suggests that there may have been magmatic fluids present in the seawater-dominated hydrothermal system. This interpretation is supported by geochemical characteristics of the alteration assemblages (Blotz et al. 2018). Mass balance calculations suggest a sericite-silica dominated assemblage consistent with argillic alteration. Based on these observations, the Flambeau deposit is possibly an example of a hybrid VMS-epithermal system. 11 �Figure 1: A) Silver telluride throughout pyrite grains and grain boundaries. B) Bismuth telluride within pyrite grain. C) Gold electrum within chalcopyrite. D) Acanthite within sphalerite. Blotz, K.E., Fredrickson, E.T., Lodge, R.W.D., 2018, Characteristics of ore and alteration mineral assemblages at the Flambeau volcanogenic massive sulfide deposit, northwestern Wisconsin. Geological Society of America-North Central Annual Meeting. DeMatties, T.A., 1994, Early Proterozoic Volcanogenic Massive Sulfide Deposits in Wisconsin: An Overview: Economic Geology, v. 89, p. 1122-1151. Jones, C.L., and Jones, J.K., 1999, The Flambeau Mine, Ladysmith, Wisconsin: The Mineralogical Record, v. 30, p. 107-131. May, E.R., and Dinkowitz, S.R., 1996, An Overview of the Flambeau Supergene Enriched Massive Sulfide Deposit: Geology and Mineralogy, Rusk County, Wisconsin, in LaBerge, G.L., ed., Volcanogenic Massive Sulfide Deposit of Northern Wisconsin: A Commemorative Volume: Institute on Lake Superior Geology Proceedings, v. 2, part 2, p. 67-93. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4-25. 12 �Fault-controlled dike emplacement in the Grand Marais, Minnesota area BOERBOOM, Terrence J., Minnesota Geological Survey Over the past few years bedrock field mapping projects have been undertaken in the area around Grand Marais, northeastern Minnesota; these were funded in part by the USGS Statemap mapping program. The most recent round of field work, completed in 2017, was in the Mark Lake 7.5’quadrangle (Fig.1), near the southern margin of the ca. 1,099 Ma Eagle Mountain granophyre (EMG). This work has delineated a network of diabase dikes and sills, some of which were apparently intruded along faults, as evidenced by disruption of a distinctive quartz- and feldspar-phyric rhyolite (QFPR) unit, which in one block has been folded into a southwest-plunging syncline (Figs. 1 and 2). The correlation of the QFPR between the blocks is strengthened by an underlying thin unit of a distinct sparsely but coarsely porphyritic andesite flow along most of its length (Fig. 2). The extent of the QFPR is much greater than previously recognized; it may correlate with the Devil’s Kettle Rhyolite to the east, but correlation of this as well as other interlayered mafic to intermediate volcanic rocks is difficult owing to gaps in outcrop, structural complications, and intervening intrusions. The mafic volcanic rocks within several km of the EMG are typically quite metamorphosed/altered, particularly the more primitive ophitic basalt varieties, and the amygdules contain epidote, local fine fibrous amphiboles, and rarely garnet along with the usual mixtures of chlorite, calcite, and quartz. The diabase dike/sill complex is continuous with the Lake Clara diabase, as previously mapped by the author to varying degrees to the west and southwest of the Mark Lake quadrangle. The Lake Clara complex generally forms an arc from northeast to east-west that mimics the synformal shape of the Sawbill Lake intrusion (Brooker and Miller, 2012), which is part of the Brule-Hovland complex (Fig. 1), and also the general curvature of the volcanic pile. However, several northwest-trending diabase offshoots imply that emplacement was locally controlled by preexisting faults, as the QFPR is clearly offset across these northwest dikes. The margins of the dikes, particularly those of northwest orientation, are flanked by thin zones of intermediate intrusive rocks that commonly show quench textures, and the central part of one of the thickest northwest dikes contains a zoned pod that ranges from ferromonzodiorite at the edge to granophyre in the center (Fig. 2). All of the intermediate to felsic phases associated with the diabase, including the granophyre pod, contain small glassy ‘quartz eyes’ interpreted to be xenocrystic grains derived from melted QFPR implying that melting of rhyolite may have also taken place at some depth. Another small, northwest-trending hybrid dike (NE corner of Figure 2) consists of red fine-grained felsite that contains comagmatic cm-to m-sized, scallop-edg0ed intermediate to mostly mafic enclaves, in nearly equal proportions of felsic to mafic material. The red felsite matrix contains abundant quartz and feldspar phenocrysts, and is essentially identical to the nearby QFPR. This hybrid dike is adjacent to another ‘felsite’ dike that is enclave free, and has only small feldspar phenocrysts. Both are oriented to the northwest, nearly perpendicular to the strike of the hosting volcanic rocks, and are believed to be related to the main Lake Clara diabase dike set (Figure 2). These are outside of the main diabase swarm, but are consistent with melting of the rhyolite at depth and commingling with mafic magma prior to upward movement along fault zones, along smaller incipient faults outside of the main swarm. REFERENCE: Brooker, B.P, and Miller, J.D.,Jr, 2012, Bedrock geologic map of the Sawbill Lake intrusion, Cook County, Minnesota, University of Minnesota Duluth Precambrian Research Center; scale 1:24,000. 13 �Figure 1. Regional geologic context of the Lake Clara diabase complex. Unlabeled areas – Keweenawan volcanic rocks, undivided. Outline of Mark Lake quadrangle shown; geology of the south half of this map from varied 1:24,000 scale maps by Boerboom and others; north half from MGS map S-21. Figure 2. Northern third of the Mark Lake quadrangle showing offset of QFPR across NW trending diabase dike offshoots, marginal intermediate hybrid rocks, and central felsic granophyre pod. 14 �Possible Alumotantite from the Nine Mile pluton, Wausau Complex, Marathon County, WI. BUCHHOLZ, Thomas W.1, FALSTER, Alexander U. 2, and SIMMONS2, Wm. B. 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494; 2Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217. The Nine Mile granite and quartz monzonite pluton is the youngest (≈1505 Ma, Dewane & Van Schmus, 2007) and most silicic of the four intrusions comprising the Wausau Syenite Complex. The Red Rock Granite northeast gravel pit is located in the south-western portion of the Nine Mile Pluton. Pegmatites and aplites are uncommon in this portion of the pluton, but in 2015 a small aplitepegmatite was exposed in the western working face of the northern portion of the pit, and samples were recovered from the talus below the exposure. The arch-shaped dike was approximately 20 cm thick, with a thin pegmatitic zone measuring approximately 5-6 cm thick near a margin of the dike. The center of the pegmatite in several samples had a thin (<0.5 cm) discontinuous band of fine-grained albite. Occurring adjacent to and within the albite band were small zircons, small crystals of columbite-group minerals, and very small (300-400µm) brownish grains of an unusual non-fluorescent niobium-bearing alumotantalate mineral. Associated minor minerals include: almandine-spessartine, columbite-(Fe), tapiolite-(Fe), zircon, hafnian zircon, zoned microlite-pyrochlore and U-rich pyrochlore, betafite, xenotime-(Y), ilmenite, monazite(Ce), and thorite. Chemical analysis of this alumotantalate yields a formula of (Al 0.986 Fe2+ 0.021 Mn2+ 0.001 ) Σ1.008 (Ta 0.803 Nb 0.147 Ti 0.062 ) Σ1.012 O 4.000 which is essentially identical to alumotantite (AlTaO 4 ). The stoichiometry of simpsonite, Al 4 Ta 3 O 18 (OH), effectively rules this species out, and there are no other known alumotantalates. X-ray diffractometry is needed to further confirm the presence of alumotantite, but paucity of material precludes this. Pegmatites of the Nine Mile Pluton are anorogenic in origin; and typically such pegmatites lack Tadominant phases. However, as we have previously reported, Nine Mile pluton pegmatites often contain late-stage Ta-enrichment, resulting in the formation of various Ta-dominant phases, including tantalite-(Mn), tapiolite-(Fe), and microlite. The occurrence of ‘alumotantite’ is noteworthy considering the overall metaluminous nature of the NYF pluton. It seems likely this occurrence resulted from a process of very late-stage fractionation similar to the processes that produced the high-Ta species in other pegmatites in the pluton. The lack of dark mica (annite or siderophyllite) in these dike samples suggests availability of Fe was probably somewhat limited, and the small amount of Fe available for interaction with late-stage fluids was likely consumed in the formation of columbite-(Fe), tapiolite-(Fe), almandine-spessartine and ilmenite, while crystallization of almandine-spessartine suggests development of a peraluminous environment. Pyrochlore, microlite, monazite and albite crystallization probably also reduced concentrations of 15 �Ca, Na, U and other elements. Lacking other cations, remaining Al combined with residual Ta and Nb to crystallize small amounts of probable alumotantite. Another possibility for increased Al availability may be a greisenization trend such as has been observed in one location in the pluton where abundant topaz was found. In an attempt to recover additional material, the site was revisited in June 2017. Little dike material was accessible in the pit wall due to slumping, but aplite-pegmatite samples were recovered from the adjacent floor of the pit. Due to the virtual absence of aplites and pegmatites in this area of the pluton, it is very probable that the samples originated from the same dike as those hosting the probable alumotantite. None of the samples exhibited the thin aplite band noted previously. However, examination of heavy mineral separates from the samples revealed a somewhat similar mineral assemblage partially reflecting the above association. The 2017 samples contained small crystals of dark mica (annite or siderophyllite, unlike the 2015 material with the thin albite band), along with Nb-bearing ilmenite, a Nb-bearing TiO 2 phase, pyrochlore, betafite, monazite, Hf-enriched zircon and columbite-(Fe). Notable is the lack of microlite, tapiolite-(Fe), almandine-spessartine, and probable alumotantite, suggesting the extreme fractionation that produced the unusual phases recovered in 2015 was restricted to a small portion of the dike. REFERENCE: Dewane, T. J., Van Schmus, W. R. (2007): U-Pb geochronology of the Wolf River batholith, north-central Wisconsin: Evidence for successive magmatism between 1484 Ma and 1468 Ma. Precambrian Research, V. 157, pp. 215-234. 16 �Recent gravity and magnetic investigations of the Minnesota River Valley Subprovince: New insights into ancient problems CHANDLER, V.W., SOUTHWICK, D. L., and JIRSA, M. A. Minnesota Geological Survey, University of Minnesota, 2609 Territorial Rd, St. Paul, MN 55114 U.S.A. Geophysical studies have been a long-standing companion to geologic investigations of the gneissic Minnesota River Valley (MRV) subprovince, due in part to limited outcrops and drill cores. Over the last decade various geophysical investigations conducted as part of state and academic programs have provided new insights into the crustal structure and evolution of this ancient and somewhat enigmatic component of the Archean Superior Province. Gravity and magnetic methods have continued to dominate geophysical studies of the MRV subprovince. Recent compilations of regional-scale gravity and magnetic grids and maps have assisted in extending MRV geology into eastern South Dakota (McCormick 2010 a, b; Southwick and others, 2018). At higher resolution, derivative-enhanced grids of gravity and magnetic data have been used extensively for bedrock mapping, which is being conducted in support of the County Geologic Atlas (CGA) Program of the Minnesota Geological Survey. These studies have added considerable detail regarding the internal geology of the blocks comprising the MRV subprovince — which from north to south are the Benson, Montevideo, Morton, and Jeffers. The enhanced gravity and magnetic data have been especially useful in 1:100,000-scale mapping of compositional variations and fold patterns within the gneissic blocks. The enhanced gravity and magnetic grids have also been very helpful in detailed mapping of the structural discontinuities that bound the blocks, including the Great Lake Tectonic Zone, the Appleton Shear zone, the Yellow Medicine shear zone, the Brown County lineament, and the Spirit Lake tectonic zone (SLTZ). Model studies of gravity and magnetic data along selected profiles have been useful in compiling geologic cross-sections for CGA mapping. Similar to earlier modeling at lower resolution, the newer models indicate that the three northernmost structural discontinuities of the MRV subprovince can be suitably approximated by slab-like sources that dip moderately to steeply northwards. Modeling of the interior of the MRV blocks is considerably more challenging; complex fold patterns and strong anomaly interference make interpretation difficult, especially with regard to determining the subsurface geometry of individual anomaly sources. In addition, outcrop evidence in the Minnesota River Valley indicates shallow structural dips of lithologic units locally that may obfuscate geophysical modeling. Nonetheless, model studies in these areas can still be useful for estimating the general range and spatial distribution of density and magnetization values for upper crustal rocks, resulting in improved lithologic identification and mapping. Gravity and magnetic modeling reveals significant differences between the SLTZ and the other structural discontinuities of the MRV subprovince. Firstly, the SLTZ, which forms the southern terminus of the MRV subprovince, is interpreted to dip southwards not northwards. Secondly, using values that are consistent with existing rock property data, most anomaly 17 �signatures of the MRV subprovince can be accommodated by sources within the shallow crust (<10 km. depth), but a prominent magnetic minimum that extends along SLTZ may involve much of the crustal section. Physical property data are not available for lower crustal rocks in Minnesota, but studies elsewhere of long-wavelength magnetic anomalies, crustal xenoliths, and crustal thicknesses indicate that the lower crust of cratonic areas typically is strongly magnetic, most likely reflecting enrichment of magnetite in granulite facies rocks (Langel and Hinze, 1998). Assuming reasonable levels of magnetization for the lower MRV crust (~3.5 SI), much of crustal section to the southeast of the SLTZ is interpreted to be non-magnetic. This apparent loss of crustal magnetization might reflect the deep emplacement of non-magnetic rocks, such as Paleoproterozoic metasedimentary rocks along the tectonic zone, or destruction of magnetic oxides via fluids moving along and above the tectonic zone. Evidence for the former possibility is available from recent magnetotelluric studies, where a conductive zone has been imaged along the southeastern edge of the SLTZ at mid- to deep- crustal levels (Bedrosian, 2016; Yang and others, 2015). Bedrosian suggested that the conductive zone might reflect Paleoproterozoic metasediments, which are known to be associated with prominent conductivity anomalies further north, where these rocks lie at or near the surface. Given the success so far for gravity and magnetic studies of the MRV subprovince, it seems likely that these data will continue to be useful for geologic studies for many years to come. REFERENCES Bedrosian, P. A., 2016, Making it and breaking it in the Midwest: Continental assembly and rifting from modeling of EarthScope magnetotelluric data, Precambrian Research, v. 278, p. 337-361. Langel, R. A., and Hinze, W. J., 1998, The magnetic field of the earth’s Lithosphere, Cambridge University Press, p. 263-268. McCormick, K.A., 2010a, Precambrian basement terrane of South Dakota: South Dakota Geological Survey Program Bulletin 41, 37p. McCormick, K.A., 2010b, Plate 1: Terrane map of the Precambrian basement of South Dakota: South Dakota Geological Survey Program Bulletin 41, External pdf file, compilation scale 1:1,000,000. Southwick, D. L., Chandler, V. W., and Jirsa, M. A., 2018, Geophysical, structural, and tectonic interpretation of the Yellow Medicine and Appleton shear zones, SW Minnesota and SE South Dakota: A work in progress, Institute on Lake Superior Geology 64th Annual Meeting, Part 1, Program and Abstracts, this volume. Yang, B., Egbert, G. D., Kelbert, A., and Naser, M.M., 2015, Three-dimensional electrical resistivity of the north-central USA from EarthScope long period magnetotelluric data, Earth and Planetary Science Letters, v. 422, p. 87-93. 18 �Possible Emplacement Controls on Diamond-Bearing Rocks North of Lake Superior CUNDARI, Robert1, SMYK, Mark1, CAMPBELL, Dorothy1 and PUUMALA, Mark1 1 Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, 435 James St. S., Suite B002, Thunder Bay, ON, P7E 6S7 Canada The recent discoveries of a number of diamond-bearing, ultramafic rocks, including kimberlite, in Archean country rocks of the Superior Province, north of Lake Superior, has provided insights into lithotectonic controls on their emplacement and suggest potential for further discoveries. The Permian to Triassic Pagwachuan kimberlites, 100 km north of Marathon, were discovered by De Beers Canada Inc. in 2015-16 within Neoarchean metasedimentary rocks of the Quetico Subprovince. Five separate kimberlites range in size from 0.5 to 2.5 ha and are reportedly multi-phase, complex pipes (Delgaty et al. 2017). The Paleoproterozoic Rabbit Foot kimberlite, 50 km east-southeast of Marathon, has been explored in recent years by Rio Tinto Canada Diamonds Exploration Inc. It intrudes Neoarchean granitoids of the Pukaskwa batholithic complex and deformed metasedimentary rocks that are probably related to rocks of the Schreiber-Hemlo greenstone belt (Brett and Russell 2016). A number of diamondiferous ultramafic rocks of unknown age have been discovered within the TransSuperior Tectonic Zone (TSTZ) west of Marathon. In 2007, the Madonna alnöite ultramafic lamprophyre dyke was discovered 35 km northwest of Marathon by Rudy Wahl, who also discovered the nearby Prairie Lake paralamproite. Although hosted by Neoarchean rocks, these ultramafic intrusive rocks are spatially associated with Midcontinent Rift (MCR)-related and TSTZ-hosted intrusions, such as the Coldwell and Killala Lake alkalic complexes, the Prairie Lake carbonatite. The Chipman Lake fenites and carbonatites also occur within the northern extension of the TSTZ (Sage 1991), north of the Pagwachuan kimberlites. The Ripple Lake diatreme and associated lamprophyre dykes occur immediately west of the Coldwell complex and are likely associated with the TSTZ. They have also been the focus of diamond exploration. The TSTZ is a north-northeast-trending fault zone that extends for at least 600 km and is locally referred to as the Thiel fault (Sage 1991). The major MCR-related alkalic intrusions emplaced along the TSTZ cluster around 1.0 Ga, although Sage (1983) identified lamprophyre dikes on the Slate Islands emplaced at approximately 300 Ma. The Gravel River fault, traced for over 200 km, is a northeast- to eastnortheast-striking regional fault system that displays an oblique sinistral sense of transcurrent motion (Williams 1989). Several structures related to the TSTZ, the Gravel River fault and a number of northwest-trending faults intersect in the vicinity of the Pagwachuan kimberlite pipes. The discovery of five new kimberlite pipes in the Pagwachuan Lake area highlights the potential for further discovery in the region. In the Geraldton area, a number of discrete magnetic anomalies resemble anomalies related to known kimberlite pipes. The confluence of major, intersecting structures (e.g. TSTZ and Gravel River fault) are proven to be effective pathways for deep-seated magmas, tapping melts well within the diamond stability field. These fault systems are shown to have been activated for extended periods of time (i.e. MCR-related alkalic intrusive rocks ca. 1.1 Ga and the Pagwachuan kimberlite swarm ca. 220 to 252.9 Ma). The occurrence of the Paleoproterozoic Rabbit Foot kimberlite (ca. 1945 Ma; Brett and Russell 2016) suggests that large, crustal-scale faulting and magmatism is long-lived in this part of 19 �the Superior Province, although obvious lithotectonic controls are not as yet identified. Recent discoveries of diamondiferous rocks north of Lake Superior demonstrate its potential and suggest that further discoveries will be made. Figure 1. Geological map showing the location of the Pagwachuan kimberlite cluster and other kimberlitic and ultramafic rocks mentioned in the abstract. Approximate traces of the Gravel River fault after Williams (1989). The abbreviation “TSTZ” indicates the approximate location of the Trans-Superior Tectonic Zone. All UTM co-ordinates provided in NAD83, Zone 16. Bedrock geology from Ontario Geological Survey (2011). References Brett, C.R. and Russell, S. 2016. Indicator mineral and soil geochemical sampling of quaternary cover and microdiamond, indicator mineral, and geochronology of ultramafic intrusive rocks, Oskabukuta property, Ontario, Thunder Bay Mining District; Thunder Bay South District, Assessment Files, AFRO report number 2.56539, 104p. Delgaty, J., Fulop, A., Seller, M., Hartley, M., Zayonce, L., Januszczak, N. and Kurszlaukis, S. 2017. Ontario’s newest kimberlite cluster – the Pagwachuan cluster; poster abstract in 11th International Kimberlite Conference, Gaborone, Botswana, September 18–22, 2017, Extended Abstract No.11IKC-4517, 4p. Ontario Geological Survey 2011. 1:250 000 scale bedrock of Ontario; Ontario Geological Survey, Miscellaneous Release—Data 126–Revision 1. Sage, R.P. 1983. Geology of the Slate Islands; Ontario Geological Survey, Open File Report 5435, 333p. ——— 1991. Alkalic rock, carbonatite and kimberlite complexes of Ontario, Superior Province; Chapter 18 in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.683709. Williams, H.R. 1989. Geological studies in the Wabigoon, Quetico and Abitibi–Wawa subprovinces, Superior Province of Ontario, with emphasis on the structural development of the Beardmore–Geraldton belt; Ontario Geological Survey, Open File Report 5724, 189p. 20 �Thrust Kinematics of the Keweenaw Fault North of Portage Lake, Michigan DeGraff, J.M.1 and Carter, B.T. 2 1 Michigan Technological University, Houghton, MI 49931 2 Consultant, Houston, TX 77027 The Keweenaw Fault (KF) is the most significant fault of the Midcontinent Rift System (MRS) based on its length of 350 km (1), postulated net slip of 9 km (2), and thrusting of copper-bearing Portage Lake Volcanics (PLV, 1.1 Ga) over younger Jacobsville Sandstone (JS) (Fig. 1). This large fault and others figure prominently in ideas about MRS development (3-4) and copper deposits mined until recently along the Keweenaw Peninsula (5). Ideas about the KF (6-7, USGS 1950s maps) largely date to before modern concepts about thrust faults and before advances in cross-section modeling. As a result, many aspects of the fault’s geometry, kinematics, and timing remain unclear or are simply outdated. The prevailing view (3-5) is that the KF began as a steep, rift-bounding, normal fault during crustal extension and later was inverted during compression. This scenario would produce a fault dipping > 45° NW, however published maps and cross-sections show the KF dipping ≤ 45⁰ NW for much of its length and often < 30⁰ (6, USGS 1950s maps). PLV layers locally exhibit slip along their boundaries (6, 8) and near the fault they generally parallel its surface. These observations suggest a thrust fault system detached along PLV layers (Fig. 1b), which is inconsistent with direct inheritance from a rift-bounding normal fault. North of Portage Lake (Fig. 2) good fault exposures occur along crosscutting valleys (5-7), and mining drill holes and workings provide good local control on PLV stratigraphy and structure. One transect northeast of Houghton (Fig. 2, Loc. 1) crosses the Keweenaw and Hancock faults, which bound an anomalous area of gently dipping to horizontal PLV layers (2). USGS geologists interpreted the KF as dipping 22⁰ NW at the surface and possibly connecting to a steeper Hancock Fault. Our data compilation and kinematic modeling show that this geometry can be replicated by thrust motion of a detached master fault, the Hancock Fault being an imbricate thrust with increased dip in its hanging wall. A second transect west of Lake Gratiot (Fig. 2, Loc. 2) crosses the KF and another one to the northwest, possibly analogous to the Hancock Fault but less well defined. USGS geologists described horizontal to shallow dipping PLV layers between these faults based on surface and drill hole data (2). Furthermore, USGS maps show that the KF trace has a prominent reentrant of JS into the area of overthrust PLV layers, which implies a nearly horizontal fault surface based on our 3-point calculations. Forward kinematic modeling of these relationships suggests that the KF propagated upward from a deep detachment and reached a shallow detachment near the top of the JS. The northwest fault may represent an out-of-sequence cutoff of the leading edge of the thrust sheet near the top of a major ramp. We suggest that the KF began as a thrust fault during a post-rift compressional event, its initiation point possibly controlled by deeper, precursor, normal faults. This ongoing research raises many questions answerable with further work. Objectives are to determine layer and fault geometry at the onset of faulting, to infer deformation history of layers displaced by fault motion, and to define subsurface relationships between the Keweenaw and nearby faults. The ultimate goal is to define tectonic conditions leading to origin and evolution of the KF and other major faults in the region. 21 �Figure 1: (a) Major rock units and faults in the Lake Superior area; KF = Keweenaw Fault, DF = Douglas Fault, IRF = Isle Royale Fault (1). Inset map shows extent of Midcontinent Rift System (MRS) from Lake Superior southwest to Kansas (K) and southeast to Detroit (D). Black rectangle is focus area of Figure 2. (b) Cross-section along A-A’ in map showing PLV (red-orange) offset about 9 km by the Keweenaw Fault, and JS (tan) locally deformed in the footwall (2). Figure 2: Focus area north of Portage Lake (adapted from 2). Major faults shown as dark red traces. 1) Dover Creek transect with smaller Hancock Fault northwest of the Keweenaw Fault. 2) Bruneau Creek transect with unnamed fault northwest of the Keweenaw Fault. References 3. 4. 5. 6. 7. 8. 1. Miller, Jr., J.D., 2007, The Midcontinent Rift in the Lake Superior region: a 1.1 Ga Large Igneous Province: IAVCEI Large Igneous Provinces Commission, p. 1-18. 2. Cannon, W.F. and Nicholson, S.W., 2001, Geologic Map of the Keweenaw Peninsula and Adjacent Area, Michigan: United States Geological Survey, Map I-2696, Scale = 1:100,000. Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332. Stein, C.A., Kley, J., Stein, S., Hindle, D., and Keller, G. R., 2015, North America’s Midcontinent Rift: When rift met LIP: Geosphere, v. 11, no. 5, p. 1607-1616. Bornhorst, T.J. and Barron, R.J., 2011, Copper deposits of the Western Upper Peninsula of Michigan: Geological Society of America, Field Guide 24, p. 83-99. Butler, B.S. and Burbank, W.S., 1929, The Copper Deposits of Michigan: USGS Prof. Paper 144, 238 p. Irving, E.D. and Chamberlin, T.C., 1885, Observations on the Junction between the Eastern Sandstone and the Keweenaw Series on Keweenaw Point, Lake Superior: Bull. U.S. Geol. Survey No. 23, U.S. Government Printing Office, Washington, D.C., 58 p. Hubbard, L.L., 1898, Keweenaw Point with particular reference to the felsites and their associated rocks: Geol. Survey Michigan, v. 6, part 2, 155 p. 22 �Geochemistry of Shallow and Deep Water Archean Meta-Iron Formations and their Post Depositional Alteration in Western Superior Province, Canada DOLEGA, Simon1 and FRALICK, Philip1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, Ontario P7B 5E1 Canada (sdolega@lakeheadu.ca) One purpose of studying banded meta-iron formations is to determine the chemical composition of seawater in the Archean ocean and the oxygen content of the Archean oceanicatmospheric system. Geologists use the geochemistry of meta-iron formations to make interpretations on the chemical conditions in the Archean. However, most scientists neglect the possibility of post-depositional alteration affecting the element geochemistry preserved in the meta-iron formations. This thesis explores the role of post-depositional mechanisms on the geochemistry of four banded meta-iron formations. The four different locations hosting Archean meta-iron formations chosen for this study include: meta-iron formations from the Beardmore/Geraldton greenstone belt of the Eastern Wabigoon Domain, Lake St. Joseph greenstone belt of the Uchi Domain, North Caribou greenstone belt of the North Caribou Terrane and Shebandowan greenstone belt of the Wawa Subprovince. The meta-iron formations from the Beardmore/Geraldton and Lake St. Joseph greenstone belts are interpreted to be deposited in a shallow water setting, while meta-iron formations from the North Caribou and Shebandowan greenstone belts are interpreted to be deposited in deeper water environments. This thesis also investigated ocean stratification by comparing the geochemistry of shallow and deep meta-iron formations. The main source iron and silica to the oceans was hydrothermal venting fluids. Iron and silica precipitated out of seawater as iron oxyhydroxides and amorphous silica, which deposited through cyclical processes. Elements dissolved in the Archean ocean were adsorbed onto iron oxyhydroxides and silica during deposition. Crystallization of quartz, magnetite and hematite occurred during diagenesis and magnetite continued to grow during progressive metamorphism. The lack of cerium anomalies, significant Y/Ho ratio values greater than average shales and the non-significant amount of authigenic chromium preserved in the meta-iron formations suggests that the oceans were anoxic. Therefore, in the Archean there was no significant oxygen stratification between the shallow and deeper water environments. Significantly most of the elements were derived from multiple sources, including the siliciclastic phase, seawater or hydrothermal venting fluids, at various proportions. Al 2 O 3 , TiO 2 , Th, V, Nb, U, REEs and Y were determined to be immobile during post-depositional alteration. The rest of the elements may have been isochemical during post-depositional alteration or may have been mobilized during post depositional alteration. 23 �Mobility during diagenesis is clearly exhibited by sodium and potassium in the meta-iron formation samples from the Beardmore/Geraldton, Lake St. Joseph and North Caribou greenstone belts. Sodium was relatively immobile, and potassium was mobilized in the magnetite- and magnetite/grunerite-dominated meta-iron formations during diagenesis. Potassium was relatively immobile, and sodium was mobilized in the hematite-, jasper- and chert-dominated meta-iron formations during diagenesis. If most of the elements remained relatively immobile during post-depositional alteration, then the ocean compositions in the Archean were heterogeneous. Shallow waters were more enriched in K 2 O, Rb and LREEs, while the deeper waters were more enriched in Cs, Na 2 O, CaO, MnO and HREEs. However, if the assumption that these elements were immobile is false, then the meta-iron formation does not preserve the ocean chemistry of the ancient ocean. 24 �Preliminary petrographic and geochemical investigation of silicified volcanic rocks and silica-rich exhalative rocks from the ~2.7 Ga Abitibi Greenstone Belt, Canada DRAZAN, Jacqueline1, BRENGMAN, Latisha1, FEDO, Christopher2 1 Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 1114 Kirby Dr., Duluth, MN, 55812; 2Department of Earth and Planetary Sciences, University of Tennessee, 1621 Cumberland Avenue, 602 Strong Hall, Knoxville, TN 37996 USA. The ~2.7 Ga Abitibi greenstone belt (AGB) is a well-preserved volcanic arc terrane dominated by mafic and felsic volcanic lithologies interstratified with siliceous chemical sedimentary rocks (namely chert and iron formations) and less prevalent clastic rocks (Mueller et al., 2009; Thurston et al., 2008; Gibson et al., 1983). Regionally, the terrane is characterized by sub-greenschist facies metamorphism and locally high SiO2 concentrations due to silicification (post-depositional addition of silica phases; Brengman and Fedo, 2018; Gibson et al., 1983). Typically, silica-rich fluids permeate porous ash, tuffaceous material, or individual volcanic flow units during hydrothermalism to produce silicified rocks of varying composition and SiO2 content. Under conditions of minor replacement, primary textures preserve (ie- phenocrysts, glass shards, amygdules, pumice fragments, volcaniclastic features), allowing field identification of silicified rocks. However, in volcanogenic massive sulfide producing systems, hydrothermal replacement leads to mineralogical and geochemical changes of protolith volcanic rocks, often obscuring primary volcanic textures in the process. This leaves behind a rock with excess SiO2 lacking features indicative of the rocks’ initial genesis. Within these geologic settings, silicified rocks can be difficult to distinguish from other siliceous chemical sedimentary rocks, which precipitate directly from seawater and/or mixed hydrothermal fluids forming discrete units (e.g. Brengman and Fedo, 2018; Thurston et al., 2008). Results from preliminary studies show that the silicon isotope composition of quartz differs between the two siliceous rocks (Brengman et al., 2016). Here we present initial results from a preliminary geochemical investigation of wellpreserved silicified volcanic rocks and an associated exhalite from the Amulet Rhyolite locality near Rouyn-Noranda, QC. We aim to provide a geochemical framework for interpreting preliminary silicon isotope isotopic data of the same rocks used as a tool to differentiate siliceous volcanic rocks from siliceous chemical sedimentary rocks. Composition and texture of the Amulet Rhyolite samples were determined using scanning electron microscopy, transmitted and reflected light microscopy, inductively coupled plasma optical emission spectrometry, and inductively coupled plasma mass spectrometry. The primary volcanic mineral assemblage consists of feldspar, chlorite, and amphiboles, with prevalent zones of epidote and quartz alteration. Based on geochemistry (SiO2 = 55.08–73.1 wt.%; Al2O3 = 12.11–15.36 wt.%; CaO = 0.79–7.67 wt.%; Na2O = 0.18–5.11 wt.%; K2O = 0.2– 5.36 wt.%; Fe2O3(t) = 4.13–18.53 wt.%; MgO = 0.8–5.47 wt.%), mineralogy (amphibole and feldspar micro-phenocrysts, glassy groundmass replaced by chlorite), and texture (quartz-filled amygdules, aphanitic matrix), samples classify as basalts and andesites, with an overabundance of quartz (Figure 1a, b). Mineralogically, samples show alteration features similar to other localities within the AGB: abundant mega- and micro-quartz alteration and patchy epidote alteration with minor dispersed carbonates (Figure 1a,b; Brengman and Fedo, 2018). Overlying one of the amygdaloidal pillowed basalt units is the marker “A” exhalite unit, thought to represent exhalative precipitation (Gibson et al., 1983; Figure 1c,d). The exhalite unit is 25 �characterized by fine banding (Figure 1d) and is principally composed of microcrystalline quartz with minor aluminous mineral phases (Figure 1d). Geochemically this unit is distinct from local volcanic rocks, with higher SiO2 (78.03 wt.%) and K2O content (5.36 wt.%), lower Al2O3 (10.64 wt.%), CaO (0.37 wt.%), and Na2O content (2.33 wt.%) content, and significantly lower Fe2O3(t) and MgO content (1.69 and 0.04 wt.% respectively). Due to the level of preservation, the exhalite unit and underlying silicified volcanic rocks can be differentiated based on petrography and geochemistry making them a good test locality for studying the silicon isotope variability between the two siliceous rock types. These initial geochemical results provide the framework for future silicon isotope analyses on the same sample suite. Figure 1. Representative samples photographs (field and photomicrographs) from samples of the Amulet Rhyolite. (a) Cross-polarized light photomicrograph of amygdaloidal basalt (quartz-filled with chalcopyrite centers). (b) Cross-polarized light photomicrograph of quartz altered andesitic volcanic rock with megaquartz alteration patch. (c) Field photograph of exhalite contact with underlying pillowed basalt unit. (d) Cross-polarized light image of finely banded exhalite unit. REFERENCES Brengman, L.A. Fedo, C.M., Whitehouse, M.J., 2016. Micro-scale silicon isotope heterogeneity observed in >3.7 Ga Isua Greenstone Belt, SW Greenland. Terra Nova: 28, p. 70-75. Brengman, L.A., Fedo, CM., 2018. Development of a mixed seawater-hydrothermal fluid geochemical signature during alteration of volcanic rocks in the Archean (~2.7 Ga) Abitibi Greenstone Belt, Canada. Geochimica et Cosmochimica Acta: 227, p. 227-245. Gibson, H.L., Watkinson, D.H., Comba, C.D.A, 1983. Silicification: Hydrothermal Alteration in an Archean Geothermal System within the Amulet Rhyolite Formation, Noranda, Quebec. Economic Geology: 78, p. 954971. Mueller, W.U., Stix, J., Corcoran, P. L., Daigneault, R., 2009. Subaqueous calderas in the Archean Abitibi greenstone belt: An overview and new ideas. Ore Geology Reviews: 35, p. 4-46. Thurston, P.C., Ayer, J.A., Goutier, J., Hamilton, M.A., 2008. Depositional Gaps in Abitibi Greenstone Belt Stratigraphy: A Key to Exploration for Syngenetic Mineralization. Economic Geology: 103, p. 1097-1134. 26 �On the source(s) of the Felch-Arnold gravity anomaly, Upper Peninsula, Michigan DRENTH, Benjamin J.1, WOODRUFF, Laurel G.2, SCHULZ, Klaus J.3, CANNON, William F.3, and AYUSO, Robert A.3 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver, CO, 80225 2 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN, 55112 3 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA, 20192 Away from the Midcontinent Rift, gravity highs in the Upper Peninsula of Michigan are normally attributed to rocks of the Paleoproterozoic Marquette Range Supergroup. In particular, dense iron formations and mafic volcanic rocks of the Menominee Group produce gravity highs where they reach significant thicknesses, and are juxtaposed against lower-density Archean rocks and sedimentary rocks of the Marquette Range Supergroup (e.g., Klasner at al., 1985). A 13 mGal, E-W trending gravity high lies over Archean and Paleoproterozoic rocks in the eastern part of the Felch trough area near the town of Felch, and extends ~25 km eastward over Paleozoic sedimentary rocks to the vicinity of the town of Arnold (Fig. 1). Bacon (1956) suggested the source is dense units of the Paleoproterozoic Marquette Range Supergroup. However, subsequent mapping (James et al., 1961) in the Felch trough area showed that Archean, not Paleoproterozoic, rocks dominate the area of Precambrian exposures that coincide spatially with the gravity high (Fig. 1). New ground gravity data and density measurements, as well as inspection of spatial relations between gravity anomalies and geologic mapping, show that the most likely candidates for the source of the Felch-Arnold gravity anomaly are the Six-Mile Lake Amphibolite and the Hardwood Gneiss (both long assumed to be Archean). Archean granites and gneisses form most of the surrounding rocks and have mean density of 2700 kg/m3. The Six-Mile Lake Amphibolite (mean density 3020 kg/m3) and the Hardwood Gneiss (mean density 2880 kg/m3) present the best spatial correspondence with the anomaly, and each could have a plausibly large subsurface volume to account for the eastward extension of the anomaly over Paleozoic sedimentary rocks. Other units in the area lack either the density, volume, or spatial distribution required to be candidates for the source. Geochemical similarities between the Hardwood Gneiss and Six Mile Lake Amphibolite suggest that those two units may be related (Schulz et al., this volume). Further, a new radiometric date on the Hardwood Gneiss of ~2.75 Ga (Ayuso et al., this volume) confirms the long-assumed Archean age for that unit. References Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., and Jackson, J., this volume, New U-Pb zircon ages for rocks from the granite-gneiss terrane in northern Michigan: evidence for events at ~3750, 2750, and 1850 Ma: Institute on Lake Superior Geology, Part 1: Program and Abstracts, v. 64. Bacon, L.O., 1956, Relationship of gravity to geological structure in Michigan's Upper Peninsula, in Snelgrove, A.K., ed., Geological Exploration: Institute on Lake Superior Geology 2nd Annual Meeting, Houghton, Michigan, p. 54-58. 27 �Bayley, R.W., Dutton, C.E., and Lamey, C.A., 1966, Geology of the Menominee iron-bearing district, Dickinson County, Michigan, and Florence and Marinette Counties, Wisconsin: U.S. Geological Survey Professional Paper 513, 96 p. Cannon, W.F., and Ottke, D., 1999, Preliminary digital geologic map of the Penokean (early Proterozoic) continental margin in northern Michigan and Wisconsin: U.S. Geological Survey Open-File Report 99-547: http://pubs.usgs.gov/of/1999/of99-547/. Cannon, W.F., Schulz, K.J., Ayuso, R.A., and Mroz, T., this meeting, Archean and Paleoproterozoic geology of the Felch District, central Dickinson County, Michigan: Institute on Lake Superior Geology 64th Annual Meeting Field Guide. James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., 1961, Geology of central Dickinson County, Michigan: U.S. Geological Survey Professional Paper 310, 176 p. Klasner, J.S., King, E.R., and Jones, W.J., 1985, Geologic interpretation of gravity and magnetic data for northern Michigan and Wisconsin, in Hinze, W.J., ed., The Utility of Regional Gravity and Magnetic Anomaly Maps, Society of Exploration Geophysicists, p. 267-286. Schulz, K.J., Cannon, W.F., and Woodruff, L.G., this volume, Geochemistry of mafic rocks in Dickinson County, Michigan: Michigan: Institute on Lake Superior Geology, Part 1: Program and Abstracts, v. 64. Figure 1: Left: Bedrock geology, after James et al. (1961), Bayley et al. (1966), Cannon and Ottke (1999), Ayuso et al. (this volume), and Cannon et al. (this meeting). Right: Complete Bouguer gravity anomalies. Inset (lower right) shows location of study area. 28 �GEON 12 to 11 history of the Lake Superior Region and speculation about the relationships between the Midcontinent Rift and the Grenville Orogen EASTON, Robert Michael1 1 Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, Ontario P3E 6B5 mike.easton@ontario.ca This presentation was inspired by some of the questions posed in the recent Geological Association of Canada Howard Street Robinson Lecture Tour presentation by Dr. P. Hollings on the “Metallogeny and Magmatism of the 1.1 Ga Midcontinent Rift”. It will focus on 3 specific topics related to the evolution of the Midcontinent Rift and whether or not the rift is the result of a typical mantle plume event. These are: 1) regional Geon 12 events that may have had an effect on localizing the rift, 2) the apparent time gap between major tectonic events in the Grenville Orogen in central North America during Geon 11 and the onset of magmatic activity in the Midcontinent rift, and 3) an alternative tectonic setting for generating a Large Igneous Province (LIP) without an active “hotspot” or mantle plume. Regional Geon 12 Events There were 2 attempted rifting events in North America that occurred to the northwest and the east of the Midcontinent Rift during Geon 12. Both produced large, radiating dike swarms, and near their centres, a variety of layered mafic intrusions. These are the circa 1267 Ma Mackenzie dike swarm (LeCheminant and Heaman 1989) and the circa 1238 Ma Sudbury dike swarm (Krogh et al. 1987). The centre of the latter swarm is now buried beneath the Grenville Orogen in western Quebec, approximately 1,800km east of the centre of Lake Superior. The Midcontinent Rift developed in the area between these 2 earlier rifting attempts. Is its location between these 2 earlier “plumes” significant? The Grenville Orogen (Ontario region) during Geon 11 The Grenville Orogen in Ontario is divided into 3 major segments (Carr et al. 2000): 1) the Laurentian margin (which consists of para-autochthonous and para-allochthonous rocks that developed on the margin of Laurentia during the Archean to Mesoproterozoic), 2) the Composite Arc Belt (CAB), which consists of a variety of arc fragments that developed somewhere outboard of North America between 1300 to 1220 Ma and which were stitched together between 1220 and 1190 Ma), and 3) the Frontenac-Adirondack belt (FAB), a continental platform to continental arc that was the site of voluminous anorthosite-mangeritecharnockite-granite (AMCG) magmatism between 1190 and 1140 Ma, and which was stitched to the southern margin of the Composite Arc Belt at circa 1160 Ma (Carr et al. 2000). It is unclear if the Laurentian Margin and the co-joined CAB and FAB were proximal to one another prior to circa 1080 Ma (see discussion in Carr et al. 2000). It is possible also that they were distal to Laurentia at the time the Midcontinent Rift was active. The Ottawan (or “Grenville”) orogeny is the onset of Grenville-wide metamorphism and deformation across all 3 segments of the Grenville in Ontario and is the result of Himalayan-style continent-continent collisional event. Grenville-wide metamorphism and deformation occurs across all 3 segments of the Grenville in Ontario during the Ottawan orogeny. In Ontario, there are remarkably few U-Pb ages (Ontario Geochronology Inventory 2018), and certainly no major magmatic, metamorphic or pervasive deformational events, between the end of AMCG magmatism at circa 1140 Ma and the start of the Ottawan orogeny, a period of approximately 55 million years (1140-1085 Ma). This Ontario Grenville time gap almost exactly coincides with the magmatic time span encompassed by the Midcontinent Rift (Abitibi and other dikes at circa 1140 Ma to late felsic volcanism at circa 1085 Ma on Michipicoten Island). Interestingly, this time gap does not exist in the Grenville in West Texas, where the period 11301110 Ma was characterized by high-grade metamorphism, deformation, and thrusting (Moser et al. 2008), however this was occurring approximately 2,000 km south-southeast of the Midcontinent Rift. 29 �Possible Tectonic Setting An important aspect of the geology of the Lake Superior region is that from circa 1900 to 1350 Ma the southern margin of Laurentia was the loci for repeated arc development, similar to the west coast of North America during the Mesozoic to Cenozoic. The culmination of this arc activity was the formation of a large Andean-style arc, represented by the magmatic rocks of the Eastern Granite-Rhyolite Province and its equivalents (circa 1480-1350 Ma) in the Laurentian Margin of the Grenville Orogen in Ontario. This almost 500 million year period of subduction beneath Laurentia would have greatly modified the lithosphere beneath Laurentia, and left remnants of partly subducted slabs in the mantle beneath the margin of Laurentia. Consequently, the recent history of western North America may provide some insights into how the Midcontinent Rift have formed. Fouch (2012) and Zhou et al. (2018) provide an alternative to the “hotspot” model for the development of the Columbia River basalts, the Snake River Plain, and Yellowstone; one that does not require a “hotspot” or classic mantle plume. Their model involves asthenosphere upwelling following slab-breakoff after subduction of the Farallon and Juan du Fuca plates beneath western North America. There are some similarities between this scenario and the development of the Midcontinent Rift. First, in both areas, there was a long period of arc magmatism prior to the onset of LIP magmatism. Second, magmatism in both areas occurred over a long time interval (more than 30 million years), was geographically widespread, and toward the waning stages became localized with a greater abundance of felsic magmas. Third, the model does not result in radiating dike swarms characteristic of many mantle plumes. Finally, in the case of Yellowstone, this upwelling may eventually lead to a pseudo-plume — something which would explain the latter stages of the Midcontinent Rift and the plume-like geochemistry of the magmas. What is not known is if a transform fault setting is needed after the end of subduction for the down-going slab to break-off and result in asthenosphere upwelling. This was the case for western North America, but presently cannot be confirmed to have occurred in the Ontario Grenville. Nonetheless, a transform would be one means whereby events in the Grenville realm could be isolated from events in Laurentia for a protracted period of time. Even if the tectonic setting envisioned by Fouch (2012) and Zhou et al. (2018) is a viable model for explaining the development of the Midcontinent Rift, there is a key difference. Unlike western North America, in the Mesoproterozoic a continent-continent collision occurred not long after the LIP event was initiated, effectively shutting the system down, likely prematurely. However, this extensive pre-heating of the lower crust in the Lake Superior region may have well been responsible for the ductile and longlived metamorphism present in the Ontario Grenville, as suggested by Carr et al. (2000). References Carr, S.D., Easton, R.M., Jamieson, R.A., and Culshaw, N.G. 2000. Geologic transect across the Grenville Orogen of Ontario and New York; Canadian Journal of Earth Sciences, v.37, p.193-216. Fouch, M.J. 2012. The Yellowstone Hotspot: Plume or Not? Geology, v.40, p.479-480. Krogh, T.E., Corfu, F., Davis, D.W., Dunning, G.R., Heaman, L.M., Kamo, S.L., Machado, N., Greenough, J.D. and Nakamura, E. 1987. Precise U-Pb isotope ages of diabase dikes and mafic to ultramafic rocks using trace amounts of baddeleyite and zircon; in Mafic dike swarms, Geological Association of Canada, Special Paper 34, p.147-152 LeCheminant, A.N. and Heaman, L.M. 1989. Mackenzie igneous events, Canada: Middle Proterozoic hotspot magmatism associated with ocean opening; Earth and Planetary Science Letters, v.96, p.38-48. Moser, S., Helper, M. and Levine, J. 2008. The Texas Grenville Orogen, Llano Uplift, Texas; Fieldtrip guide to the Precambrian Geology of the Llano Uplift, central Texas, Geological Society of America, Annual Meeting 2008, Houston, Texas, 54p. Zhou, Q., Liu, L. and Hu, J. 2018. Western US volcanism due to intruding oceanic mantle driven by ancient Farallon slabs; Nature Geoscience, v.11, p.70-76. 30 �What to do after the bull has left the china shop- Picking up the community relation pieces EGER, Paul1, BOT, Courtnay1, MEINEKE, Dave1 and ADAMS, Dave2 1 Global Minerals Engineering LaPointe Iron Company 2 Today successful mining projects must meet the “triple bottom line”; economic, environmental and social. In 2010 Gogebic Taconite (GTac) acquired a lease option to develop an open pit taconite mine in northwestern Wisconsin. Although the company had mining experience, it was limited to coal near existing mines with minimal opposition. GTac’s proposed development was viewed as either an economic boon or an environmental disaster. GTac’s decision to focus on lobbying at the state level mobilized opposition from the local, environmental and tribal communities. After spending millions of dollars in additional resource evaluation and environmental studies, GTac decided to abandon the project in 2015. When the option to lease was terminated in 2015, LaPointe decided it was important to rebuild local partnerships and begin to develop sound data so that future decisions could be based on science and not fear. These efforts include support for scientific studies, baseline regional water quality monitoring and periodic meetings with community leaders. 31 �Surficial Geology of the Iron Mountain 7.5 Minute Quadrangle, Dickinson County, Michigan, Florence & Marinette Counties, Wisconsin ESCH, John M1, KEHEW, Alan2, HUOT, Sebastien3, YELLICH, John4 1 Michigan Dept. of Environmental Quality, Office of Oil, Gas, and Minerals, P.O. 30256, Lansing, MI 48909, 2 Department of Geological and Environmental Sciences, Western Michigan University, Kalamazoo, MI 49008, 3 Illinois State Geological Survey, Prairie Research Institute, University of Illinois at Urbana-Champaign, Champaign, IL 61820, 4 Michigan Geological Survey, Western Michigan University, Kalamazoo, MI 49008 The Iron Mountain 7.5 minute quadrangle lies within complex glacial deposits of the Green Bay Lobe of the Laurentide Ice Sheet. In 2017 the Michigan Geological Survey mapped the quad as part of a USGS STATEMAP project. Surficial mapping was greatly aided by the availability of LiDAR elevation data. This mapping has provided new, detailed information on the surficial landforms and deposits as well as relationships between the glacial deposits and the underlying bedrock. The complexity of the glacial deposits is in part due to high relief on the bedrock surface and complexity of underlying bedrock formations and structure. Three icemargins were mapped across the quad. A deep bedrock trough mapped as part of an earlier environmental investigation was further defined as well as the bedrock topography and drift thickness mapped across the quad. In addition, the mapping identified ice-walled lake plains, eskers, drumlins and terraces that were not previously mapped. A new OSL age date was obtained for an outwash deposit. The sediments include diamicton (till), sand and gravel, boulders and interbedded silt and clay. The glacial deposits are late Wisconsinan (about 14,500 cal yr BP to 12,500 cal yr BP) in age. Within the Iron Mountain Quad, the west-northwesterly ice flow direction is reflected in the numerous drumlins on the uplands and streamlined bedrock hills. Rock drumlins occur within the quad and others are exposed in the bottom of some sand and gravel pits. The elevation across the map ranges from 919 feet above mean sea level (AMSL) along the Menominee River to 1572 feet AMSL at the top of Millie Hill. Three distinct bedrock-controlled uplands occur north of the Menominee River: Pine Mountain, Millie Hill and Trader Hill. These uplands are mostly cored by diamicton and strewn with boulders. These boulders extend to depth into the subsurface based on the local water wells logs and the three borings drilled for this project. The diamicton is mostly reddish brown to brown. Although the larger foundation for these uplands is bedrock controlled, drift up to 140 feet thick occurs in places. Three ice margins are interpreted within the map. From west to east they are the Winegar-Sagola-Early Athelstane Moraine (about 14,500 cal yr BP), the Middle Athelstane ice margin and the Marenisco-Late Athelstane ice margin (about 13,000 cal yr BP). Under the City of Kingsford lies an extensive pitted outwash plain with elevations ranging from 1140 to 1120 feet AMSL formed west and south of the Middle Athelstane ice margin. An OSL age 12,600 ± 1,000 cal yr was obtained from a sample taken from the edge of this outwash plain. This outwash plain overlies a lacustrine sequence of mostly silts, clays, sands 32 �and some gravels that is as much as 300 feet thick. This lacustrine sequence overlies a deep bedrock trough under the City of Kingsford. A later, short term fluvial event likely occurred over this outwash surface, as evidenced by the sharp, steep, wave cut or fluvially cut scarp at 1140 feet AMSL along the northern side of this surface. As the Green Bay Lobe ice retreated to the east, there must have been successive lowering of the outwash outlets to the south and southeast because of the step-like lowering of the outwash terraces to east along the Menominee River. The lowest of these terraces to the east is 180 feet lower than the large outwash plain to the west. A passive seismic instrument using the Horizontal-Vertical Spectral Ratio (HVSR) method was used to gather additional bedrock control for data on bedrock topography and drift thickness. This technique uses the horizontal-to-vertical spectral ratio method to record ambient seismic noise with 3-component geophones. HVSR calibration readings were gathered at 15 wells and borings of known bedrock depth. These data were used to develop a local HVSR bedrock depth calibration curve. Exploration readings were taken at 44 locations within the map. Very good bedrock depth estimates were made in the outwash areas of the map. In the upland morainal area, however, the method yielded depth estimates that were much too shallow relative to the local bedrock elevation of the area. This disconnect is likely due to buried, overconsolidated dense glacial till which was encountered at depth in the three borings drilled for this project. The HVSR bedrock depth estimates at these three borings match well with the depths to the top of the dense till. A significant gamma-ray log kick was also seen in the borings at or near the top of this dense till. Although the glacial deposits in the Iron Mountain Quadrangle average 40 feet thick, numerous bedrock outcrops exist. The drift is maximally 363 feet over the deep bedrock trough in Kingsford. In many places, the land surface topography is controlled not by the glacial deposits, but by the underlying bedrock and bedrock structure. One important exception is a pronounced buried deep bedrock trough that underlies the large pitted outwash plain in Kingsford. Another buried bedrock trough underlies the lowland along the Menominee River in the southeastern part of the map. A poorly defined bedrock low connects the two troughs north of the Menominee River. There is high relief on the bedrock surface ranging from 730 feet to 1530 feet AMSL across the map. Bedrock outcrops and mounds appear throughout the area, even where nearby borings show over 100 feet to bedrock. The bedrock geology exposed at the surface and underlying the glacial deposits in the Iron Mountain Quadrangle is very complex and has had a significant and controlling effect on the overlying glacial deposits. Underlying the quad are Precambrian complexly faulted and folded Precambrian metasedimentary rocks and metavolcanics and granitic intrusions as well as much later Cambrian Sandstones. References Esch, J.M., and Kehew, A.E., 2017, Surficial Geology of the Iron Mountain 7.5 Minute Quadrangle, Dickinson County, Michigan, Florence & Marinette Counties, Wisconsin, Michigan Geological Survey, Surficial Geologic Map Series SGM-17-04, scale 1:24000. 33 �LiDAR Revolutionizing Geological Mapping ESCH, John M Michigan Department of Environmental Quality, Oil, Gas & Minerals Division, Constitution Hall 2nd Floor South, 525 West Allegan Street, Lansing, Michigan, 48933 LiDAR (Light Detection And Ranging) has fundamentally changed how we view and interpret the landscape and has revolutionized geological mapping. Often subtle features can be seen in the LiDAR topography data that are not visible on aerial photography, topographic maps, or digital elevation models (DEMs). LiDAR is an optical remote sensing technology that emits intense, focused beams of Light at the ground and measures the time it takes for the reflections to be detected by a sensor. This results in a densely spaced (QL2): 2 points/meter network of highly accurate georeferenced elevation points called a point cloud. These elevation points are classified as to what the LiDAR pulse was reflected off (ground, vegetation, water, buildings or other objects). The ground elevation points are used to produce highly accurate Digital Elevation Models that can be used to generate three-dimensional representations of the Earth’s surface and its features. Elevation accuracies are on the order of 10 centimeters. Airborne LiDAR is the most common, but there is also terrestrial LiDAR and bathymetric LiDAR. Terrestrial LiDAR can be used for mapping high cliff and quarry faces to create a virtual outcrop. The most useful airborne LiDAR product is the bare earth digital elevation model (DEM). Other common deliverable LiDAR products are a classified point clouds and intensity Images. A LiDAR attribute that may be of value for geologists is the intensity of the returned pulse, which is the strength of the return or how strongly the laser pulse was reflected back to the sensor. This is usually presented a greyscale .tif image and may be useful for mapping soft ground (wetlands) vs hard ground (potentially bedrock outcrops). Common LiDAR derivative products include DEM hillshade, digital surface models, shaded relief, contours and automated building extraction. The higher resolution topography advantage of 0.6 meter LiDAR DEMs over existing 30 and 10 meter DEMs is obvious. This very dense data coverage allows for seeing subtle geologic features, and cultural features like curbs, plow furrows, and two-tracks. It can also can be used to see what is under tree canopy. Bedrock outcrops often appear distinct from the surrounding topography in the LiDAR data. This is valuable for mapping in remote areas with little known exposure. Hydrologic features like streams, valleys and subtle erosional features are more accurately and easily seen using LiDAR. Many more karst feature like sinkholes, disappearing streams, and solution enhanced joint areas have been identified using LiDAR. Subtle glacial features like ice-walled lake plains are almost never seen on aerial photos or topographic maps 34 �(except for large ones) and were rarely identified in Michigan prior to LiDAR. With LiDAR they are relatively easy to see. Other subtle glacial features seen using LiDAR, but sometimes too small to be seen on topographic maps include small eskers, drumlins, flutes, subtle terraces, fans, deltas, small sand dunes, paleo-shorelines, ice margins, bars, pendants, and erosional scarps. These previously undetected landforms may fundamentally change how one interprets the area geology. Pits and other excavations as well as scarps are easily seen using LiDAR and are helpful for places to investigate. LiDAR allows geologists to be more efficient in the field by allowing them to see the landscape how it really is before going out in the field. It also helps in fine tuning and focusing field work in specific area, features and landforms. It also allows them to map subtle features that may not be accessible due to land ownership permissions. This presentation will show the widely varying applications of LiDAR for geological mapping across Michigan. A B Figure 1 (A) Northwest Albion, Michigan, 7.5 Minute Topographic Map, USGS 1980. 10 Foot Contour Interval. (B) Northwest Albion Area, Calhoun County LIDAR Shaded Relief Map, clearly indicating esker trending SW-NE across map. From Carswell (2014) and Esch (2013). References Carswell, W.J., Jr., 2014, The 3D Elevation Program—Summary for Michigan (ver. 1.2, June 29, 2015): U.S. Geological Survey Fact Sheet 2014–3107, 2 p., Esch, J. M., 2013, Surficial geology of the Northwest Albion 7.5 minute quadrangle, Calhoun County Michigan, Michigan Geological Survey, Surficial Geologic Map Series SGM-1302, scale 1:24,000. 35 �Mineral Chemistries of the Tower Mountain Intrusive Complex Au-Deposit, Ontario FITZPATRICK, William1, HOOPER, Robert1, and LODGE, Robert1, Gélinas, Brigitte2 1 Dept. of Geology, University of Wisconsin Eau Claire, 105 Garfield Av., Eau Claire, WI 54701 Dept. of Geology, Lakehead University, 955 Oliver Rd, Thunder Bay, ON, P7B5E1 2 Hydrothermal fluid systems driven by magmatic activity are some of the most important ore forming processes for many metallic minerals. This project involves characterization of hydrothermal alteration associated with the Archean Tower Mountain Intrusive Complex and its related Au deposit within the Shebandowan Greenstone Belt. The Tower Mountain Intrusive complex (TMIC) consists of a ~1.5km2 monzonite stock cored by a later diorite intrusion. Cross cutting both units are small dikes of monzonite porphyry and meter-scale areas of hydrothermal brecciation. The TMIC intrudes Timiskaming-type assemblages of calc-alkaline volcanic rocks and conglomeritic alluvial-fluvial sedimentary. The main monzonite and syenite stocks have been U-Pb dated to 2690 ± 3 Ma while calc-alkaline volcanics elsewhere in the Shebandowan greenstone belt similar to TMIC host rocks were dated to 2690 Ma (Corfu and Scott, 1998). The contemporary U-Pb dates as well as the lack of a surrounding contact metamorphic zone implies that the Tower Mountain Intrusive complex was syn-volcanic and emplaced at relatively shallow depth (Carter, 1990). The hydrothermal alteration at TMIC requires at least a two stage process. The first alteration results in initial oxidation followed by a mineralizing event with lower fO2. Mineral chemistry data has been collected from numerous primary and alteration minerals using a SEM/EDS detector. The rocks have been heavily altered and few primary minerals remain except for scattered patches of Mg rich-hornblende, Mg-rich augites, magnetite-ilmenite intergrowths and minor monazite, zircon. All primary feldspars have been altered to end-member alkali feldspars (albite and k-spar). The first hydrothermal alteration event resulted in oxidation of the magmatic Fe-Ti oxides with production of hematite and secondary rutile. Other minerals associated with this alteration event include: titanite, epidote, fluoro-apatite (e.g. Ca 4.5 Mn,Fe .5 (PO 4 ) 3 F), and Mg-rich phengite, Mgchlorite which petrographically are seen as sericitization of magmatic alkali feldspars and mafic minerals. Some of the epidotes from this alteration episode have LREE enriched epidote rims and some of the phengites have fluorine in the hydroxyl-site (K 1.7 Na .3 )(Al 3 Mg .8 Fe .2 )(AlSi7 )(OH 3.9 ,F .1 ). This first alteration episode resulted in pervasive pink hematite staining and green (epidote) alteration seen in outcrop (Gélinas et al., 2016). Also related to this alteration episode are scattered, small (~3µm2) barite and celestine grains. The rocks were subsequently altered by a sulfidizing fluid which results in hematite replacement by pyrite, chalcopyrite, and pyrrhotite. Sulfidizing fluid alteration also results in observed Fe-rich chlorite, ferroan-dolomite and the gold mineralization. This model is consistent with Au being transported as an Au-bisulfide complex and precipitated along with the sulfide minerals. Relating the fluid history described above to lithologies seen in the field, the first oxidizing phase of alteration is likely to have initiated with intrusion of the monzonite stocks. Fluids related to alkaline magmatism have long been known to have an association with oxidizing, hematitic alteration (e.g. Greenberg, 1986) and also carbonatizing, halide and LREE-rich metasomatism (e.g Wooley, 2003). Evidence observed in thin section indicates that introduction of reducing, sulfidizing fluids and Au-mineralization is related to later intrusion of the monzonite porphyry 36 �dikes. No magnetite is observed in the monzonite porphyry whereas all other samples contained magnetite in various states of decay along with pyrite. This indicates that sulfidizing fluids were more active in the monzonite porphyry, possibly because of closer spatial association to the intruding magmas. ru A B ru+ttn Ksp Ksp mt ap ttn Mg-chl ab ab hb Mg, Fe chl C D py phen mt ab ttn ab cpy py po ab Ksp Fe-chl Figure 1: A: Magmatic hornblende and magnetite in matrix of Mg-chlorite, albite and k-spar. From monzonite. B: Rutile, titanite clusters with scattered apatite in mg-chlorite, phengite, albite, kspar matrix. From monzonite. C: Highly corroded magnetite separated by Fe-chlorite from pyrite in equilibrium. From hydrothermal breccia. D: Zoned pyrite grain containing rim of pyrrhotite and inclusions of chalcopyrite. From monzonite porphyry. ru: rutile, mt: magnetite, Fe/Mg chl: Iron or Magnesium rich chlorite, py: pyrite, hb: hornblende, Ksp: k-spar, ttn: titanite, cpy: chalcopyrite, po: pyrrhotite Carter, M.W., 1990, Geology of Forbes and Conmee townships, Ontario Geological Survey, Open File Report 5726. Gélinas, B.R., Lodge, R.W.D., Gibson, H.L., 2016, Characterization of the Mineralization and Alteration at Tower Mountain, Conmee Township, Shebandowan Greenstone Belt, Ontario, Ontario Geological Survey, Miscellaneous Release - Data 330. Greenburg, J., 1986, Magmatism and the Baraboo Interval: Breccia, Metasomatism and Intrusion: Geoscience Wisconsin, v. 10, p. 96-112. Woolley, A., 2003, Igneous Silicate Rocks Associtated with Carbonatites: Their Diversity, Relative Abundances and Implications for Carbonatite Genesis: Periodico Mineralogia, v. 72, p. 9-17. 37 �Petrogenesis of mafic magmatism in the Coldwell Complex Part 1. Geochemical model to explain origin of metabasalt by partial melting in the SCLM GOOD, Dave Department of Earth Sciences, Western University, London, ON N6A 5B7 Canada Data for basalt and intrusive mafic rocks from the early stage of evolution of the Midcontinent Rift show well-defined trends for trace element abundances that are consistent with variable degrees of partial melting in a mantle plume source and subsequent fractional crystallization. For example, in a plot of La vs. Zr, early MCR basalt data plot in a field that spans compositions from E-MORB to OIB. The La/Zr values increase from about 0.07 to 0.18 with increasing La, as expected given the relative incompatibility of these elements. Deviations from the E-MORB to OIB-like compositions are explained by interaction of the magma with either SCLM or continental crust during ascent, or by varying depths of partial melting. A comparison of the early stage MCR rocks to the Coldwell mafic rocks shows significant differences between the two groups. First, the ranges of trace element concentrations for Coldwell rocks is one to two orders of magnitude greater than for the entire suite of MCR rocks. Second, Coldwell rocks show significant LREE enrichment relative to HFSE (example, very high La/Zr). Remarkably, Th/Nb for the Coldwell rocks are near mantle values (0.12) signifying enriched LREE is not a result of crustal contamination. It is difficult to explain these differences between Coldwell and MCR rocks by processes such as partial melting or crystal fractionation and some other explanation is required for the decoupling of these highly incompatible trace elements. Trace element abundances for Coldwell metabasalt units 1 to 3b are some of the lowest values in the MCR. Spider diagram patterns for the metabasalt and gabbro units are characterized by strong negative Nb and Zr anomalies that resemble patterns observed for mantle xenoliths from the SCLM at numerous locations, including Bir Ali, Yemen. 38 �A geochemical model that supports generation of Coldwell metabasalt by partial melting of a hypothetical SCLM source was proposed by Good and Lightfoot. The source composition is based on mantle xenolith data from Bir Ali, Yemen and was synthesized by mixing lherzolite with a very small fraction (1-2%) of secondary clinopyroxene or amphibole. This model can explain many features of the Coldwell metabasalt units, however, at this stage, trace element abundances are too low to predict the nature of metasomatism (carbonatite or alkaline) in the source. References Cundari, Robert, 2012. MSc thesis, Lakehead University, 142 pages Davis, Sarah, 2016. BSc thesis, Lakehead University, 63pages Good D.J. and Lightfoot P.C., Submitted to CJES, Feb 2018 Good, D.J., Cabri, L.J. and Ames, D.E., 2017, Ore Geology Reviews, v. 90, p.748-771 Good, D.J., Epstein, R., McLean, K., Linnen, R.L. & Samson, I.M., 2015, Econ Geol v.110, p.983-1008 Keays, R.R. and Lightfoot P.C., 2015, Econ Geol v. 110, p. 1235-1267. Sgualdo, P., Aviado,K., Beccaluva, L., Bianchini, G., Blichert-Toft , J., Bryce, J.G., Graham, D.W., Natali, C. and Siena, F. 2015. Tectonophysics, v. 650, p. 3-17. Sage, R.P. 1994, OGS, Open File report 5888, 592p. Sun, S.S., and McDonough, W.F. 1989. Geological Society, London, Special Publications v. 42, p. 313-345. 39 �Geologic history meets the web – online data of the Lake Superior Division of USGS GOTTSCHALK, Brad, ROSE, Caroline, and MCCARTNEY, M. Carol Wisconsin Geological and Natural History Survey, University of Wisconsin – Extension, caroline.rose@wgnhs.uwex.edu Working out of their headquarters in Madison, Wisconsin from 1882 to 1922, USGS Lake Superior Division geologists laid the groundwork for all subsequent investigations of this region’s Precambrian rocks. Nine monographs, four bulletins, and a professional paper describe the findings of those early geologists. The physical samples and paper records they collected and used to produce those publications comprise the Lake Superior Legacy Collection held by the Wisconsin Geological and Natural History Survey (WGNHS, the Survey). The Survey began digitizing the Lake Superior collection in 2011, when the UW Digital Collection scanned field notes written by Charles Van Hise. The collection contains: more than 30,000 hand samples; over 13,000 thin sections (photographed in plane- and cross-polarized light); 467 field notebooks (321 of which have been scanned); 67 maps; and, 35 catalogs of specimens, lithological descriptions, and more. The web application links all of these components together, presents this image-rich dataset in a visual way, and also provides some historical context. http://data.wgnhs.uwex.edu/lake-superior-legacy/index.html Figure 1: Through the interactive map, samples can be found by location, rock type, sample notes, sample number, field notebook number, and other attributes. 40 �In the Lake Superior Legacy Collection application, researchers and historians can: search for hand sample locations on an interactive map; browse a gallery of thin sections and a list of field notebooks; view and zoom into high-resolution thin section photographs; examine hand-drawn topographic and geologic maps; and review the hand-written field notebooks and lithological descriptions of the pioneering geologists who collected these physical samples. More importantly, they can relate a hand sample to its thin sections and to its location. They will also find the notebook and page number where any specific sample is described – with a link to the online scanned version of those notes. Figure 2: Browse thin section photographs in the gallery; use the viewer to zoom and to fade between plane- and cross-polarized light; follow the link to view details for the related hand section. This long-term data preservation project, completed with the help of the USGS National Geological and Geophysical Data Preservation Program, is allowing today’s geologists to gain access to the work of the early giants, Roland Irving and Charles Van Hise. Additionally, we have been able to present this collection of beautiful thin sections and hand-written notes in a visually appealing application that shares the beauty and history of Lake Superior geology online. 41 �Inferences on the Subsurface Distribution of Oronto and Bayfield Groups North and West of the Douglas Fault, Northwestern Wisconsin GRAUCH, V.J.S.1, BEDROSIAN, Paul A.1, STEWART, Esther Kingsbury3, and HELLER, Samuel2, 1 U.S. Geological Survey, MS 964, Federal Center, Denver, CO, 80225 U.S. Geological Survey, MS 939, Federal Center, Denver, CO, 80225 3 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Rd., Madison, WI 53705 2 The period of extensive sedimentation that largely post-dated magmatic activity related to the 1.1 Ga Midcontinent rift is recorded by the Oronto and overlying Bayfield Groups in northwestern Wisconsin and correlative units in neighboring Minnesota (Fig. 1). The reversesense Douglas fault juxtaposes the two Groups: Oronto Group rocks form a syncline within the St. Croix Horst to the south and east, and the gently inclined Bayfield Group sandstones are to the north and west (Fig. 1). Previous subsurface interpretations north and west of the Douglas Fault have come from geophysical studies (e.g., Ocala and Meyer, 1973; Allen et al., 1997), which relied on recognizing characteristic densities and seismic velocities derived from modeling or sample measurements (Halls, 1969) to identify formation divisions within the Groups. Formations within the Bayfield and Oronto Groups have characteristic velocities and densities that both increase with older age. We have reassessed the previous geophysical work in northwestern Wisconsin using recently obtained digital access to proprietary seismic-reflection data, preliminary results from new airborne electromagnetic (AEM) survey lines, and a reexamination of characteristic seismic velocities correlated to geologic units from legacy refraction data, borehole logs and drill core in the Ashland Syncline area (Fig. 1). Our findings are summarized as follows (refer to Fig. 1). Data processing tests of LS-10 seismic line (westernmost Lake Superior) to find the velocity function that images the sharpest reflections suggest that the top ~2 km is composed of low-velocity (3.2 to 3.7 km/sec) rocks overlying igneous basement. This finding is in general agreement with previous refraction studies (e.g., Ocala and Meyer, 1973), who attribute the low velocities to the Bayfield Group (typical range 2.74-3.5 km/sec from Mooney et al., 1970). However, our geologic correlations from the well data indicate that the upper part of the Freda Sandstone (upper Oronto Group) has velocities from 3.2 to 3.9 km/sec, which overlap with the range expected for Bayfield Group. The low velocities here compared to Freda Sandstone elsewhere are likely caused by the greater volume of siltstones and shales (Halls, 1969). Comparison of lithology and unit picks in drill core to resistivity sections from the AEM lines show consistent correlation of low resistivities (10-50 ohm-m) with siltstones and shales of the upper Freda Sandstone and of moderate to high resistivities (>200 ohm-m) with sandstones of the Bayfield Group. Using these results, we interpret the AEM resistivity sections to show a lakeward thinning wedge of Bayfield sandstones that appear to terminate near the northern shore of the Bayfield Peninsula. The wedge overlies low resistivities typical of the shales and siltstones of the Freda Sandstone. The sequence of velocities found from a refraction site near onshore seismic-reflection line SEI-1 has a velocity-depth pattern generally compatible with overall variations in the velocities observed for Oronto Group units within the Terra-Patrick well (Dickas and Mudrey, 1999). We thus can roughly correlate the refraction velocity profile to reflection packages in 42 �SEI-1 and extrapolate the interpretation to LS-10. Combined with low-resistivities observed below 500 m depth where an AEM line crosses SEI-1, we infer that about 500 m of Bayfield Group overlies about 3.5 km of Oronto Group along SEI-1. Thus, the ~2-km thick, low-velocity, low-resistivity sedimentary section under LS-10 indicates that lower Oronto (Copper Harbor and probably Nonesuch) units are missing in the westernmost part of Lake Superior and Freda strata directly overlie igneous basement. This conclusion is supported by Allen et al. (1997), who used a different line of reasoning from seismic reflection line LS-8. The results imply that lower Oronto strata were either not deposited or were eroded from the area under the westernmost lake up until upper Freda time, whereas the areas to the southeast were accumulating sediments throughout the entirety of Oronto and Bayfield times. References Allen, D. A., Hinze, W. J., Dickas, A. B., and Mudrey, M. G., Jr., 1997, Integrated geophysical modeling of the North American Midcontinent Rift System: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota: Geological Society of America Special Paper 312, p. 47-72. Dickas, A.B., and Mudrey, M.G., Jr., 1999, Terra-Patrick #7-22 deep hydrocabron test, Bayfield County, Wisconsin: Wisconsin Geological and Natural History Survey Miscellaneous Paper 97-1, 117 pp. Halls, H.C., 1969, Compressional wave velocities of Keweenawan rock specimens from the Lake Superior region: Canadian Journal of Earth Sciences, v. 6, p. 555-568. McGinnis, L.D., and Mudrey, M.G., Jr., 2003, Seismic reflection profiling and tectonic evolution of the Midcontinent rift in Lake Superior: Wisconsin Geological and Natural History Survey MP 91-2. Mooney, H.M., Farnham, P.R., Johnson, S.H., Volz, G., and Craddock, C., 1970, Seismic studies over the Midcontientn gravity high in Minnesota and northwestern Wisconsin: Minnesota Geological Survey Report of Investigations 11, 191 pp. Ocala, L.C., and Meyer, R.P., 1973, Central North American Rift System, 1. Structure of the axial zone from seismic and gravimetric data: Journal of Geophysical Research, V. 78, no. 23, p. 5173-5194. Stewart, E.K., and Mauk, J.L., 2017, Sedimentology, sequence-stratigraphy, and geochemical variations in the Mesoproterozoic Nonesuch Formation, northern Wisconsin, USA: Precambrian Research, v. 294, p. 111-132. Figure 1: Regional geology and index map locating geophysical and drillhole information. Modified from Stewart and Mauk (2017). USGS – U.S. Geological Survey. WGNHS – Wisconsin Geological and Natural History Survey 43 �Origin, distribution, morphology, and chemistry of amphiboles in the Ironwood IronFormation, Gogebic Iron Range, Wisconsin, U.S.A. GREEN, Carlin J., SEAL, Robert, R., II, CANNON, William F., PIATAK, Nadine, and MCALEER, Ryan J. U.S. Geological Survey, MS 954, Reston, VA 20192 The Ironwood Iron-Formation, located in the Gogebic Iron Range in Wisconsin, is one of the largest undeveloped taconite resources in the United States. Interest in the development of this resource is complicated by potential environmental and health effects related to the presence of amphibole minerals in the Ironwood Iron-Formation, a consequence of Mesoproterozoic contact metamorphism. The purpose of this study is to provide mineralogical information about these amphiboles to aid regulatory, medical, and mining entities in their evaluation of this potential resource. Optical microscopy, X-ray diffraction, scanning electron microscopy, and electron microprobe analysis techniques were utilized to study the origin, distribution, morphology, and chemistry of amphiboles in the Ironwood IronFormation. The development of amphiboles from Fe-carbonates and Fe-phyllosilicates at temperatures of approximately 300 -340º C has long been recognized as a result of regionally extensive contact metamorphism of the Ironwood Iron-Formation by the Mellen Intrusive Complex, however amphiboles related to the emplacement of diabase or gabbro dikes and sills in low-grade iron-formation were also recognized in this study area. Amphiboles in the Ironwood Iron-Formation most commonly developed in massive and prismatic habits, and locally assumed a fibrous habit. Fibrous amphiboles were locally recognized in the two potential ore zones of the Ironwood Iron-Formation, but were not observed in the portion considered to be waste rock. Massive and prismatic amphiboles show a wide range of Mg# values (0.06 to 0.87), whereas Mg# values of fibrous amphiboles are restricted from 0.14 to 0.35. Factors that influenced the compositional variability of amphiboles in the Ironwood Iron-Formation may have included temperature of formation, the presence of coexisting minerals, morphology, bulk chemistry of the iron-formation, and variations in prograde and retrograde metamorphism. 44 �Geothermobarometry of a Precambrian amphibolite from Cornell WI HAFFTEN, Doug and RADWANY, Molly Department of Plant and Earth Sciences, University of Wisconsin – River Falls, 410 S 3rd Street, River Falls, WI Garnet amphibolite at Cornell, Wisconsin is part of the Chippewa amphibolite complex, a group of amphibolite-facies rocks with diverse protoliths, outcropping in the valley of the Chippewa River and its tributaries in western Wisconsin (Laberge and Myers, 1984). Metamorphism of the region occurred in Precambrian time, due to the Penokean orogeny (Schulz and Cannon, 2007). The Cornell amphibolite contains an ideal mineral assemblage for estimating both temperature and pressure of metamorphism, making it useful for understanding the effect of the Penokean Orogeny on the Marshfield Terrane. We obtained whole-rock geochemical data for the sample using a Bruker S8 Tiger X-ray fluorescence spectrometer at University of Wisconsin – Eau Claire. Ratios of immobile trace elements (Zr/Ti=0.02 and Nb/Y=0.14) indicate a mafic-intermediate protolith for the amphibolite. Petrographic analysis of the Cornell amphibolite reveals an assemblage of 20% hornblende, 45% plagioclase, 20% quartz, and 15% garnet with accessory apatite and zircon. We used a JEOL 8900 electron microprobe at the University of Minnesota to obtain mineral compositions. Which revealed ferro-tschermakitic hornblende, almandine-rich garnet and plagioclase An 40-65 . To determine the temperature of metamorphism, we used the Holland and Blundy (1994) hornblende-plagioclase thermometer. The results of this method suggest temperatures between 606-646°C. To determine the pressure of metamorphism, we applied the Kohn and Spear (1990) garnet-hornblende-plagioclase-quartz geobarometer. The range of pressures determined using this method is 5.74-6.64 Kbar. These results reveal more specific information about the Chippewa amphibolite complex and the dynamic Precambrian past of Wisconsin. REFERENCES Holland, T and Blundy, J, 1994, Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry, Contributions to Mineralogy and Petrology, Vol. 116, p. 433447 Kohn, M.J. and Spear, F.S., 1990, Two new geobarometers for garnet amphibolites, with applications to southeastern Vermont, American Mineralogist, Vol. 75, p. 89-96 Laberge, G.L. and Myers, P.E., 1984, Two early Proterozoic successions in central Wisconsin and their tectonic significance, Geological Society of America Bulletin, Vol. 95, p. 246 Schulz, K.J. and Cannon, W.F., 2007, The Penokean orogeny in teh Lake Superior region, Precambrian Research, Vol. 157, p. 4-25 45 �Figure 1: Backscatter electron image showing the main four phases in the Cornell amphibolite. Mineral compositions were used to determine pressure and temperature of metamorphism. Grt: garnet, Pl: plagioclase, Ts: tschermakite, Qz: quartz. 46 �Hornblende-Plagioclase thermometry of the Eau Claire River Complex, western Wisconsin HANNACK, Gina, and RADWANY, Molly Department of Plant and Earth Sciences, University of Wisconsin River Falls, 410 S. 3rd Street River Falls, WI The Eau Claire River Complex is a metamorphosed and deformed layered mafic intrusion that outcrops in Big Falls County Park near Eau Claire, Wisconsin. The unit is part of the larger Chippewa amphibolite complex within the Precambrian Marshfield Terrane (Cummings, 1984). The unit is characterized by a compositional layering that alternates between mafic and feldspathic compositions. We distinguish two types of rock - mafic amphibolites, containing ~80% hornblende, and feldspar-rich amphibolites, containing ~15% hornblende. The vertical, north-south striking foliation is defined by compositional layers, likely inherited from primary, igneous layering. Lineations, defined by hornblende, are at a high angle to compositional layering (approximately east to west) and represent crystallization of hornblende during metamorphism of the complex. Cummings (1984) established field evidence of a granulite facies metamorphic event represented by garnet porphyroblasts. Overprinting occurred during a second event at amphibolite facies and resulted in pseudomorphism of garnet by hornblende. The second event was pervasive and accompanied by dynamic recrystallization of plagioclase. Finally, there was a third, greenschist facies metamorphic event, resulting in crystallization of epidote, zoisite, biotite, and chlorite. Mafic amphibolites consist of approximately 80% hornblende, 15% plagioclase and 5% accessory minerals ilmenite and apatite. The feldspar-rich amphibolites are comprised of 80 to 85% plagioclase, 10 to 15% hornblende and 5% ilmenite and apatite. Greenschist facies retrograde assemblage consists of minor epidote, zoisite, chlorite, and biotite. Pyrite also occurs in several samples, especially in association with epidote veins. We determined mineral compositions for one mafic amphibolite using a JEOL 8900 electron microprobe at University of Minnesota. Plagioclase compositions range from An 42 to An 85 . The amphibole present is Figure 1: Microprobe image containing hornblendemagnesio-hornblende with Mg/(Mg+Fe) plagioclase thermometer pairs between 0.55 to 0.59. Microprobe data 47 �was utilized in the application of the Holland & Blundy (1994) edenite-richterite thermometer (Figure 1). We chose this thermometer based on the minor amount of quartz found in petrographic analysis. Results of this thermometer show that the Eau Claire Complex underwent upper amphibolite facies metamorphism at temperatures ranging from 719 °C to 769 °C (assuming P of 10 kb; Figure 2). Our results provide insight into the mineralogic and structural effects of the Penokean Orogeny on the crystalline rocks of the Marshfield terrane and a quantitative estimate of thermal conditions during this collisional event. Figure 2: Temperatures of metamorphism of the Eau Claire River Complex, as determined by application of the Holland and Blundy (1994) edenite-richterite thermometer. References Cummings, ML, 1984, The Eau Claire River Complex: a metamorphosed Precambrian mafic intrusion in western Wisconsin, Geological Society of America Bulletin, Vol. 95, p. 75-86 Holland, T and Blundy, J, 1994, Non-ideal interactions in calcic amphiboles and their bearing on amphibole-plagioclase thermometry, Contributions to Mineralogy and Petrology, Vol. 116, p. 433-447 48 �Mapping the Midcontinent Rift System Hinze, William J. Department of Earth, Atmospheric, and Planetary Sciences, Purdue University, West Lafayette, IN 47906 Mapping the ca. 1100 Ma old Midcontinent Rift System (MRS) which is arguably the most significant non-orogenic structure of the North American midcontinent has been an important, but challenging objective for the past half century because it is hidden for most of its ~2500 km length by relatively flat-lying Phanerozoic sedimentary rocks. Even where it is crops out in the Lake Superior region, it is largely covered by sedimentary rocks of late-stage Keweenawan rift basins and Pleistocene glaciation sediments. Mapping of the buried rift system is based on interpretation of gravity and magnetic anomaly data that are not definitive, poorly distributed deep seismic profiling, and limited basement drill holes. As a result, uncertainty exists in the geographic location, extent, and configuration of the buried MRS. Additional ambiguity is caused by confusion in defining the required characteristics of continental rifts. A review of the available data on the MRS and Mesoproterozoic rocks of the North American midcontinent provides insight into the likely geographic location, configuration, and extent of the rift system and identifies portions of the rift’s configuration that are most problematic. Generally, the MRS is shown extending in a southerly-open arc along several rift units from the western end of Lake Superior into Kansas and a shorter, eastern, branch of less intense rifting that continues south from the eastern end of the Lake into southeastern Michigan. Studies of basement rocks from Kansas southwestward indicate that magmatic activity similar in age to the MRS occurred in this region as it did broadly over the present-day midcontinent. However, no rift basins have been found or are indicated by geophysical and deep drilling data to the south of Kansas or elsewhere where magmatic activity occurred outside of the trend of the MRS. Perhaps this magmatic activity is evidence of incipient rifts, that is regions where intrusions have accompanied lithospheric extension which failed to reach an intensity where surface faulting and rift basins developed. The termination of the other branch of the rift in Michigan is complicated by its intersection with geologic structures resulting from the Grenville orogeny whose latest activity is slightly younger than the MRS. Gravity and magnetic anomalies suggest that the MRS terminates at the Grenville Front in southeastern Michigan, but originally it may have extended to approximately the border with Ontario. Late stage Grenvillian overthrusts may overlie the extreme terminus of the MRS in Michigan. North/south striking gravity positive anomalies extending through Ohio along the eastern margin of the Grenville Front perhaps as far as Alabama have fostered the hypothesis that the rift system extends south from southeastern Michigan. The positive gravity anomalies have been purported to be the locus of rift basins of volcanic rocks. However, because rift basins do not occur at the basement surface coincident with the gravity highs, they have been interpreted as relic rifts that exist beneath Grenvillian overthrusts that post-date the MRS. Unfortunately, seismic reflection profiling does not confirm the presence of these rifts beneath the overthrusts observed 49 �in the seismic reflection data. Alternatively, integrated interpretation of the Grenville Front Tectonic Zone gravity and magnetic anomalies, seismic reflection profiling, and basement drillhole samples indicate that the Grenville Front positive gravity anomalies are caused by upthrusted high-grade metamorphic rocks from upper and mid-level crustal rocks of the Grenville orogen. According to the latter interpretation the lithic-arenites, the Middle Run formation, that occur unconformably below the Paleozoic sedimentary formations west of the Grenville Front in Ohio and adjacent states were deposited in a foreland basin primarily from erosion of the Grenville highlands to the east rather than in a marginal late-stage rift basin adjacent to a rift trough. This conclusion is supported by reported dating of detrital zircons from the Middle Run formation. Especially problematic in mapping the MRS is the location of the eastern margin of the terminus of the rift in the Northern Peninsula of Michigan and the connection of the Lake Superior Rift with the Cross-Michigan Rift segment to the south. A “third branch” of the MRS remains elusive, but the most likely candidate is the Nipigon Embayment that did not develop into a rift basin such as found beneath Lake Superior and the eastern and western branches of the MRS. 50 �Reinterpretation of the ages of deposition and folding of Animikie Basin metasedimentary units in east-central Minnesota HOLM, Daniel1, BOERBOOM, Terrence J.2, and SCHEINER, Scott1 1 Dept of Geology, Kent State University, Kent OH 44224 dholm@kent.edu 2 Minnesota Geological Survey, boerbo001@umn.edu Animikie Basin (AB) sedimentary rocks in Minnesota have historically been interpreted solely as Penokean (1875-1835 Ma) foredeep deposits. Yet detrital zircons as young as ca. 1770 Ma (Heaman and Easton, 2005) in the northern reaches of the AB (Rove Formation in Ontario) indicate that deposition in the upper part of the basin may have occurred during Yavapai orogenesis (1800-1700 Ma). Much of the AB sequence is only very weakly metamorphosed and mildly deformed. However, along its southern margin in east-central Minnesota, deformed AB sedimentary rocks show an increase in metamorphism and strain southward toward a mid-crustal plutonic-gneiss dome terrane largely exhumed during Yavapai orogenesis. Holst (1984) recognized two distinct structural zones; a northern once-deformed region characterized by upright folds and a single well-developed cleavage, and a southern twice deformed terrane characterized by refolded recumbent fold nappes and two cleavages. Holst interpreted the map trace separating these two deformation zones to be a Penokean thrust fault (Fig. 1A) and assumed all of the sedimentation and deformation to be Penokean. Given the wealth of geologic and geochronologic data which now document Yavapai-age magmatism, metamorphism, sedimentation and deformation overprinting the Penokean orogeny, we reinterpret Holst’s Line as a possible Yavapai age angular unconformity that separates the once/twice-deformed units (Fig. 1B), implying that only the southern terrane sediments and the early recumbent nappes are Penokean. In this model, rapid exhumation of the entire Penokean/Yavapai internal zone resulted in rapid erosion rates and renewed Yavapai orogenic sedimentation into the Animikie Basin followed by folding of both sedimentary sequences. Bedrock mapping, geophysical data, and geochemical/isotopic analyses of the metasedimentary rocks along the southern margin of the AB in Carlton County Minnesota, briefly described below, are at least consistent with this new hypothesis. A B Fig. 1. Schematic synopsis of tectono-sedimentary interpretations along the southern margin of the Animikie Basin, east-central Minnesota. A: Late Penokean thrust fault cuts Penokean Animikie foredeep deposits (after Holst, 1984). B: Yavapai age unconformity separates Penokean (south) from Yavapai orogenic deposits (north). Remapping and relogging of cores and cuttings (Boerboom, 2009) and aeromagnetic data reveals lithologic differences south and north of Holst’s Line. The southern units are characterized by a moderate gravity high and a belt of discontinuous aeromagnetic anomalies interpreted as ‘sulfidic horizons’ with large cubic pyrite porphyroblasts. The sulfidic horizons may be similar to those at the base of the Baraga basin in Michigan (the Bijiki Iron Formation) and possibly to portions of the iron rich layers near in the Cuyuna South Range. The sulfidic horizons are absent north of Holst’s Line. 51 �Geochemical analyses of samples collected across the contact reveal a concentrated grouping of trace element data from the southern samples and a larger spread from the northern samples, suggesting a more variable source for the northern sedimentary sequence. More importantly, Nd isotopic data across the contact reveal a more juvenile source for rocks south of (i.e., below) the contact (ƐNd (T)1.77Ga = 9.2 and 1.2) and an older more enriched Archean source directly north of (i.e., above) the contact (ƐNd (T)1.85Ga = -0.4 and -3.9). We interpret the southerly sequence to be Penokean and derived from the newly accreted juvenile arc terrain. However we propose that the northerly sequence is Yavapai in age and derived from an older more enriched Archean or mixed source such as is presently exposed in the plutonic-gneiss dome corridor. We propose a simplified model for tectono-sedimentary formation of the AB (Fig. 2). At the base of the basin the Penokean unconformity is shown as a nonconformity above Archean basement. To the south, the basal Penokean unconformity becomes an angular unconformity separating pre-Penokean sedimentary and volcanic rocks to the south from Penokean foredeep rocks to the north. Only the southern portion of the AB closest to the orogenic zone experienced Penokean deformation. During the Yavapai orogeny, the southern margin of the basin experienced uplift and erosion, followed by Yavapai orogenic deposition resulting in the formation of a disconformity in the north and a Yavapai angular unconformity in the south. Rapid Yavapai age exhumation of the plutonic-gneiss dome terrane led to copious amounts of sediment being shed into the basin. Near the end of the Yavapai orogeny, deformation resulted in folding of the Yavapai foredeep deposits and refolding of Penokean and pre-Penokean rocks in the south. Deformation waned to the north away from the orogenic zone. If correct, our reinterpretation has important ramifications for interpreting the inventory of structures in the upper Great Lakes region. For instance, the sedimentary rocks and the late open upright second-generation folds exposed at Thomson Dam, may both be manifestations of Yavapai orogenesis and not classic features of Penokean orogenesis. 1770 Ma zircons -εNd +εNd Fig. 2. Simplified schematic synopsis of tectono-sedimentary formation of the Animikie Basin in east-central Minnesota. Boerboom, T.J., 2009, Plate 2 – Bedrock Geology, pl. 2 of Setterholm, D.R., Project manager, Geologic Atlas of Carlton County, Minnesota, Minnesota Geological Survey County Geologic Atlas Series C-19, pt. A, 7 pls, scale 1:100,000. Heaman, L.M., and Easton, R.M., 2005, Proterozoic history of the Lake Nipigon area, Ontario, constraints from UPb zircon and baddeleyite dating (Abs.): Institute on Lake Superior Geology 51, Part 1 – Program and Abstracts, p. 24-25. Holst, T.B., 1984, Evidence for nappe development during the Early Proterozoic Penokean Orogeny, Minnesota: Geology, v. 12, p. 135-138. 52 �Olivine Crystal Size Distribution in the Black Sturgeon Sill, Nipigon, Ontario HONE, Samuel V. and ZIEG, Michael J. Department of Geography, Geology, and the Environment, Slippery Rock University, Slippery Rock, PA 16057 Recent years have seen widespread acceptance of the idea that igneous intrusions are often emplaced in multiple phases or pulses (Miller et al., 2011). To test whether textural data could be used to distinguish between these pulses, we examined olivine from the Black Sturgeon sill, a 256 m thick diabase intrusion located southwest of Lake Nipigon in Ontario. We collected samples from a continuous drill core through the sill and subdivided it into several zones using modal mineralogy and textural data. The most primitive and olivine-rich zone is from 120-200 m above base. In this study, we investigated the use of crystal size distributions (CSDs) to discriminate between distinct populations of olivine in this olivine-rich zone. CSDs are a well-established method of quantifying the textures of igneous rocks (Zieg and Marsh, 2002). We performed this analysis by collecting images from 42 samples throughout the olivine zone and stitching them together into large mosaics. We manually traced at least 200 olivine crystals from each sample, then entered the list of crystal sizes into the software package CSDCorrections 1.5 (Higgins, 2000) to calculate the CSDs. The slope and intercept of best-fit lines to the CSDs are used to quantify the texture (Zieg and Marsh, 2002). Two typical CSDs with best-fit lines are shown in Figure 1. Cluster analysis of the slope-intercept data reveals four well-defined textural groups, which could correspond to distinct populations of olivine (Fig. 2). These groups are typically found in separate parts of the olivine zone (Fig. 3), with sharp boundaries between the different groups. As an example, the textures of samples 264 and 265.5 are clearly qualitatively and quantitatively distinct (Fig. 1), even though they are found within 1.5 m of each other. Using the slope and intercept of CSDs, along with cluster analysis, we can identify separate populations of olivine in the olivine zone of the Black Sturgeon sill. While we have not found any consistent variation within groups, the breaks between the populations are sharp and coincide roughly with compositional changes (Zieg and Hone, this issue). We interpret these breaks as evidence of the episodic emplacement of the Black Sturgeon sill. CSDs have proved an effective complement to other methods in identifying discrete magma pulses in this sill. The episodic nature of magma emplacement in the Black Sturgeon sill also raises the possibility that other layered mafic intrusions formed by a similar process. References Higgins, MD, 2000. Measurement of crystal size distributions. American Mineralogist, 85: 1105-1116. Miller, CF, Furbish, DJ, Walker, BA, Claiborne, LL, Koteas, GC, Bleick, HA, and Miller, JS, 2011. Growth of plutons by incremental emplacement of sheets in crystal-rich host: Evidence from Miocene intrusions of the Colorado River region, Nevada, USA. Tectonophysics, 500: 65-77. Zieg, MJ, and Marsh, BD, 2002. Crystal size distribution and scaling laws in the quantification of igneous textures. Journal of Petrology, 43, 1: 85-101. 53 �1 mm 1 mm Figure 1. Representative textures. a) Photomicrograph of sample 264(5% olivine). b) CSDs of samples 264 and 265.5. c) Photomicrograph of sample 265.5 (26.5% olivine). The finer-grained texture (c) has a higher intercept and steeper slope. Figure 2. Identification of four textural groups based on the CSD slope and intercept. Figure 3. CSD slope (a) and intercept (b) profiles through the olivine zone. 54 �Reconstructing Paleoproterozoic volcanism in northwestern Wisconsin: Geochemistry of the Flambeau Cu-Zn-Au Mine JACOBSON, Regan E., LODGE, Robert W.D. Department of Geology, University of Wisconsin – Eau Claire: Eau Claire, WI 54702-4004 The Flambeau mine is located 1 mile southwest of the town of Ladysmith, located in Rusk County, WI. The mine is a part of the Wisconsin Magmatic Terrane and is a group of volcanic and plutonic rocks that formed during volcanism associated with the accretion of the Pembine-Wausau terrane onto the southern end of the Superior craton during the Paleoproterozoic Penokean orogeny (May and Dinkowitz, 1996). The Flambeau is one of a number of volcanogenic massive sulfide (VMS) deposits, but it is the only one that has been mined. The mined portion of Flambeau deposit is part of an enriched zone where relatively little is known about the original hypogene geology. The ore-hosting rocks consists of metamorphosed variably-altered volcanic rocks and cherty iron-formations that are now sericite to quartz-sericite schists, and biotite-andalusite-chlorite schists (DeMatties, 1994). The Flambeau was mined from 1993-1997 and produced 181,000 tons of copper, 334,000 ounces of gold, and 3.3 million ounces of silver contributing over a billion dollars in state revenue. Mining of the enriched orebody was completed in 1997 (Jones and Jones, 1999). Since then, the site has been reclaimed and revegetated. Therefore, the only remaining rock for the Flambeau site is preserved in drill core stored at the Wisconsin Geological & Natural History Survey core repository. VMS ore deposits are formed in submarine environments where high temperature hydrothermal fluids react with cold sea water to cause the precipitation of sulfide minerals. Usually, the characteristics of the volcanic system influence the composition of ore and alteration mineral assemblages. However, ongoing studies of Flambeau ores and hydrothermal alteration do not align with previous interpretations for the Flambeau volcanic system (Blotz et al. 2018). Historic research focuses largely on the mined portions of the deposit as the remainder of the strata is covered by thick Quaternary glacial deposits. This study utilizes major and trace element geochemistry of least-altered host rocks to assess the magmatic and tectonic affinity of the rocks hosting the Flambeau deposit. This study is the first trace rare earth data set for the Flambeau deposit. Assessing the magmatic affinity of these rocks will provide insight to the petrogenesis of arc magmatism/collision and the magmatic/tectonic controls on VMS mineralization during ocean-continent collision. Previous stratigraphic interpretations of the hangingwalll rocks indicate three units consisting of quartz-augen schist, metadacite, and chlorite schist (May and Dinkowitz, 1996). Geochemical data produced in this study indicates that the Flambeau deposit is primarily hosted by intermediate volcanics locally interleaved with quartz-phyric rhyolitic rocks (Fig. 1). Preliminary data reveal a bimodal distribution and all show arc-like characteristics on primitive-mantle normalized plots with light REE enrichment and negative Nb and Ti anomalies. Felsic rocks show FI to FII type (Lesher et al. 1986) trace element characteristics inicative of formation at moderate crystal depths. Ongoing trace element analysis will more clearly document the magmatic affinity of this volcanic suite. 55 �Figure 1 This shows Zr/Ti and Nb/Y ratios of units 3a, 5, and 2a. These ratios are representative of the host rocks in the Flambeau VMS deposit. Plot from Pearce (1996). References Dematties, T.A., 1994, Early Proterozoic Volcanogenic Massive Sulfide Deposits in Wisconsin; an overview: Economic Geology, v. 89, p. 1122–1151, doi: 10.2113/gsecongeo.89.5.1122. Jones, E.L., and Jones, J.K., 1999, The Flambeau Mine, Ladysmith, Wisconsin: The Mineralogical Record, v. 30, p. 107-131 May, E.R., and Dinkowitz, S.R., 1996, An Overview of the Flmabeau Supergene Enriched Massive Sulfide Deposit: Geology and Mineralogy, Rusk County, Wisconsin, in LaBerge, G.L., ed., Volcanogenic Massive Sulfide Deposit of Northern Wisconsin: A Commemorative Volume: Institute on Lake Superior Geology Proceedings, v. 2, part 2, p. 67-96 Lesher, C.M., Goodwin, A.M., Campbell, I.H., Gorton, M.P., 1986. Trace-element geochemistry of ore-associated and barren, felsic metavolcanic rocks in the Superior Province, Canada. Canadian Journal of Earth Sciences 23, 222-237. Blotz, KE, Fredrickson, ET, Lodge, RWD. (2018) Characteristics of ore and alteration mineral assemblages at the Flambeau volcanogenic massive sulfide deposit, northwestern Wisconsin. Geological Society of America, Abstracts with Programs. Pearce, J.A., 1996. A user's guide to basalt discrimination diagrams. In: Wyman, D. A. (Eds.), Trace element geochemistry of volcanic rocks; applications for massive sulphide exploration. Geological Association of Canada, short Course Notes, p. 79-113. 56 �On-going Geologic Mapping in Minnesota’s Arrowhead Region by the Minnesota Geological Survey JIRSA,* Mark A., Minnesota Geological Survey (MGS) (jirsa001@umn.edu) *The large number of collaborators engaged in various components of this endeavor precludes complete acknowledgement here. Consult MGS Open-File Report OFR2016-04 for authorship details. This presentation describes geologic mapping by the MGS in northeastern Minnesota, with an emphasis on the bedrock geology. We are in year 3 of a 6-year process to create County Geologic Atlases for St. Louis and Lake Counties. The “Arrowhead Project” area includes parts of the Boundary Waters Canoe Area Wilderness, Voyageurs National Park, Superior National Forest, and several State forests. It also encloses the easternmost part of the Mesabi Iron Range, the “Cu-Ni District,” and the Duluth metropolitan area (the 4th largest in the state). County atlases contain maps and other imagery depicting bedrock geology, geophysical data, bedrock geochronology, bedrock topography, depth to bedrock, surficial sediments, and subsurface sediment layers; together with the extensive digital data sets used to construct them. The atlases are designed to provide regional 4-D geologic framework in digital and print formats to support ongoing and future studies related to land, water, and mineral resources. Figure 1. Generalized map of northeastern Minnesota showing the location and geologic setting of county-scale mapping (modified from MGS State Map Series S-21). Because these are two of the largest counties in the state, we’ve divided them for mapping purposes into subareas referred to here as the Central, Southern, and Northern Arrowhead. Work in each subarea involves 1 or more seasons of field mapping by 5 geologists, rotary-sonic drilling, trenching, and acquisition of drill hole, petrographic, geochronologic, and geophysical data. Work in the Central Arrowhead subarea is complete (e.g., Jirsa and others, 2017), and components of the other subareas are in various stages of completion. Preliminary products for all subareas are published as they become available in MGS Open-File Report OFR2016-04. This open-file report will remain the primary repository for on-going mapping. Once preliminary work in all subareas is published, the data will be recombined into county geologic atlases. Creation of these products is a team effort, involving 14 staff members from MGS, and 57 �several from partner agencies. Staff of the Minnesota Department of Natural Resources will use these maps and data sets to conduct regional groundwater studies, including assessments of flow systems, aquifers, groundwater chemistry, and an assessment of sensitivity to pollution, to produce Part B of the County Geologic Atlases. Support is provided by the U.S. Geological Survey National Cooperative Geologic Mapping Program, the Environment and Natural Resources Trust Fund (as recommended by Legislative-Citizens Commission on Minnesota Resources—LCCMR), and the Boards of Commissioners of St. Louis and Lake Counties. The region’s bedrock includes portions of three subprovinces of the Archean Superior Province, Paleoproterozoic strata including the Biwabik Iron Formation, and Mesoproterozoic volcanic and intrusive rocks of the North Shore Volcanic Group and Duluth Complex (Fig. 1). The latter hosts polymetallic mineral deposits under consideration for new mining. The bedrock in much of the region has been mapped to varied levels of detail in the past, driven in part by minerals exploration. Despite this, our current effort identified many areas that escaped previous mapping or were mapped in minimal detail, and it attempts to fill those voids and integrate the disparate sources of information. Mapping is conducted primarily at 1:24,000-scale, with printable products that are generalized from the companion digital data sets to scales of 1:100,000 to 1:200,000. One of the more geologically interesting aspects of recent work is the recognition that chemical weathering of bedrock prior to glaciation played a fundamental role in shaping the region’s topography, bathymetry, hydrogeology, and ecology. In this region where bedrock is at and close to the land surface, recently acquired empirical evidence indicates that differential erosion of saprolitic bedrock reflects both the compositional and structural attributes of the rock. Essentially, the bedrock surface in much of the area reflects the somewhat transitional boundary between fresh and weathered rock. In areas where outcrop, drill hole data, and access are limited, lidar imagery can be employed to infer both compositional and structural trends in bedrock. In addition, the presence of varied thicknesses and compositions of saprolite likely contributes to hydrogeologic characteristics, though further study is needed. The most recent bedrock mapping (Northern Arrowhead subarea) includes a component of geochronologic analyses. Although the results are not yet published, all the samples submitted are inferred to be Neoarchean—an era of rocks in Minnesota historically lacking extensive highresolution geochronologic data. Three main temporal objectives are attempted with these samples: 1) establish ages of successor basin deposits in the Wawa subprovince of the Superior Province; 2) establish ages of intrusions emplaced into several geologic settings; and 3) establish ages of major neosomatic components of migmatitic rocks that comprise the Quetico subprovince. If successful, these data will refine the temporal framework for deposition, magmatism, deformation, and metamorphism that will contribute to understanding the region’s tectonic evolution. As with all products derived from the Arrowhead project, the geochronologic results will be published in the open-file report mentioned above. REFERENCE Jirsa, Mark A., Boerboom, Terrence J., Radakovich, Amy L., Chandler, Val W., Peterson, Dean M., Schmitz, Mark D., Dengler, Elizabeth L., Wagner, Kaleb G., Lively, Richard S., and Setterholm, Dale R., 2017, Geologic mapping in the Central Arrowhead Area, northeastern Minnesota: 63rd Institute on Lake Superior Geology Proceedings, v. 63, Part 1, Program and Abstracts, p. 46-47. 58 �Geology and geochronology of the 2006 Cavity Lake forest fire area, Boundary Waters Canoe Area Wilderness, NE Minnesota JIRSA1, Mark A., STARNS2, Edward C., and SCHMITZ3, Mark D. 1 Minnesota Geological Survey, 2609 W. Territorial Road, St. Paul, MN 55114-1009 ConocoPhillips Alaska, Inc. 700 G St., Anchorage, AK 99501 3 Department of Geosciences, Boise State University, 1910 University Drive, Boise, ID 83725-1535 2 The bedrock geology in this part of the Boundary Waters Canoe Area Wilderness (BWCAW) is incredibly diverse and remarkably well exposed. Parts of the area were mapped to varied levels of detail in the 1930’s, and more recently in the 1970’s and 1980’s, with efforts focused primarily along waterways. A devastating wind storm in 1999 flattened trees in much of the region, and a delayed result in 2006 was the Cavity Lake forest fire. The fire exposed bedrock and allowed comparatively unencumbered access to interior parts of the map area, creating a unique and time-sensitive opportunity for mapping. Field work and compilation of prior mapping was conducted in 2007-2008 with funding from the USGS National Cooperative Geologic Mapping Program, and a preliminary map was produced (Jirsa and Starns, 2008). That map has been revised to incorporate subsequent geochronologic analyses and field work by the lead author and students of University of Minnesota, Duluth, Precambrian Research Center field camps (2007-2015). The map at scale 1:24,000, and companion data are published as Minnesota Geological Survey Miscellaneous Map M-193 (Jirsa and others, 2017). One previously unpublished geochronologic date by coauthor Schmitz is presented on the map and discussed here. Figure 1. Generalized bedrock geology of northeastern Minnesota showing the location of Cavity Lake map area (bold black outline). Inset map shows location of Neoarchean rocks within the Wawa subprovince of the Superior Province. M-193 portrays bedrock that represents crustal evolution spanning the Neoarchean to the Mesoproterozoic (Fig. 1), with an emphasis on structural and stratigraphic relationships in the Neoarchean portion. Neoarchean greenstone-granite terrane of the Wawa subprovince of Superior Province is represented by a succession of mostly mafic to ultramafic metavolcanic and hypabyssal intrusive rocks (ca 2700 Ma); unconformably overlain by hornblende-phyric, calc-alkalic volcanic and volcaniclastic rocks (ca 2690 Ma; newly published age), and intruded by the Saganaga Tonalite (also ca 2690 Ma). This succession was uplifted, chemically weathered (Driese and others, 2011), and subaerially 59 �eroded to provide detritus to one or more successor basins. Based on correlation with Neoarchean terrane along strike in Canada (Corfu and Stott, 1998), the latter sequence of clastic strata is thought to have been deposited at about 2684-2682 Ma—after emplacement of the Saganaga Tonalite (2690 Ma), and before the primary regional deformation and metamorphic event at about 2680 Ma (Boerboom and Zartman, 1993). All of these rocks were cut by mafic dikes inferred from field relationships to be both Paleoproterozoic and Mesoproterozoic. The Neoarchean rocks and some dikes are unconformably overlain by Paleoproterozoic metasedimentary strata of the Animikie Group (ca 1880-1830 Ma), which includes the Gunflint Iron Formation. The uppermost layers of iron-formation are intensely deformed and overlain east of this map area by thin lenses of ejecta from a meteorite impact that occurred near Sudbury, Ontario, at ca 1850 Ma (Jirsa and others, 2011). Mesoproterozoic rifting is manifest in hypabyssal dikes and sills known collectively as the Logan intrusions (ca 1115 Ma), and several intrusive phases of the Duluth Complex (ca 1100 Ma) emplaced into both Neoarchean and Proterozoic rocks. One of the most notable geologic features in this region is the local preservation of 4 major unconformities; two within Neoarchean rocks, and two in and at the base of Paleoproterozoic strata. The Neoarchean rocks in the central Boundary Waters Canoe Area Wilderness are inferred here to comprise a Timiskaming-type extensional basin and its apparent wall- and floor-rocks. The geologic units in the region were parceled by Gruner (1941) into eight structural segments separated by anastomosing shear and fault zones, and this map exposes parts of the eastern four of those segments. Although rock types are comparatively pristine within each segment, correlation of units from one fault-bounded block to another is challenging. Each block was uplifted, down-dropped and tilted differently, which results in different stratigraphic levels of exposure and repetition of strata locally. This map attempts to “unstrain” the rocks within each segment to reveal stratigraphic variations that may reflect fluctuations in basin geometry and progressive erosional dissection of basin wall rocks. This contributes to a regional tectonic model that involves basin development during late stages of terrane accretion. REFERENCES Boerboom, T.J., and Zartman, R.E., 1993, Geology, geochemistry, and geochronology of the central Giants Range batholith, northeastern Minnesota: Canadian Journal of Earth Sciences, v. 30, p. 2510-2522 Corfu, F., and Stott, G.M., 1998, Shebandowan greenstone belt, western Superior Province: U-Pb ages, tectonic implications, and correlations: Geological Society of America Bulletin v.110, p.1467-1484. Driese, S.G., Jirsa, M.A., Ren, M., Sheldon, N.D., Brantley, S.L., Parker, D., and Schmitz, M., 2011, Neoarchean paleoweathering of tonalite and metabasalt: Implications for reconstructions of 2.69 Ga early terrestrial ecosystems and paleoatmospheric chemistry: Precambrian Research, v.189, p. 1-17. Gruner, J.W., 1941, Structural geology of the Knife Lake area of northeastern Minnesota: Geological Society of America Bulletin, v.52, p.1577-1642. Jirsa, M.A., Fralick, P.W., Weiblen, P.W., and Anderson, J.L.B., 2011, Sudbury impact layer in the western Lake Superior region, in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 147–169. Jirsa, M.A., and Starns, E.C., 2008, Preliminary bedrock geologic map of the 2006 Cavity Lake fire area, parts of Ester Lake, Gillis Lake, Munker Island, and Ogishkemuncie Lake 7.5-minute quadrangles, northeastern Minnesota: Minnesota Geological Survey Open-File Report OF08-05, scale 1:24,000. Jirsa, M.A., Starns, E.C., and Schmitz, M.D., 2017, Bedrock geology of the 2006 Cavity Lake fire area, Boundary Waters Canoe Area Wilderness, northeastern Minnesota: Minnesota Geological Survey Miscellaneous Map M-193, scale 1:24,000. 60 �The youngest magmatic activity of the Midcontinent Rift at Bear Lake, Keweenaw Peninsula, Michigan KULAKOV, Evgeniy1, BORNHORST, Theodore J.2, DEERING, Chad3, and MOORE, James B.4 1 Centre for Earth Evolution and Dynamics, University of Oslo, Oslo, Norway A. E. Seaman Mineral Museum, Michigan Tech, Houghton, MI 49931 3 Department of Geological and Mining Engineering and Sciences, Michigan Tech, Houghton, MI 49931 4 Moore Rock Farm, Keweenaw Peninsula, MI 2 The Bear Lake igneous body is located near Michigan's McLain State Park in the southwest side of the Keweenaw Peninsula (sections 24 and 25, T56N, R34W) and represents the youngest known magmatic activity of the Midcontinent Rift. Bear Lake was mapped as an intrusive rhyolite (Cornwall and Wright, 1956) cross-cutting the Freda Sandstone of the Oronto Group. It is a dark gray to red colored fine-grained igneous rock with micro-phenocrysts of K-feldspar, biotite, hornblende, quartz, apatite, and iron oxides and is flow-banded in many outcrops. Bear Lake is, however, not a rhyolite, but instead intermediate in composition with an average (N=6) of 59.1 wt. % SiO 2 , 1.2 wt. % Na 2 O and 6.6 wt. % K 2 O. This composition falls in the trachyandesite field of LeMaitre for fresh rocks (2002 Cambridge University volcanic rock classification) and would be further subdivided as a latite. While altered, the alkaline tendency of the rock is confirmed by the high concentration of the immobile elements with 1273 ppm Zr, 270 ppm La, and 496 ppm Ce. The grain size and flow banding texture is consistent with the igneous body having formed either as a shallow intrusive or an extrusive flow. The interpretation that this is an extrusive deposit is supported by the results of drilling within the body by Johnson et al. (1980). They described glacial overburden underlain by 9 m of highly altered fragmental rocks in turn underlain by 8 m of siltstone and coarse-grained arkose over the igneous rock body. The contact between the igneous body and beds of the Freda Sandstone is not exposed. Those beds nearest to the contact are altered with visible calcite. The igneous body is variably altered and contains veinlets of calcite, quartz, and heulandite, which are more prominent near an exploration shaft in the west central side of the body dug about 1917. Strong conductors and anomalous copper content of about 190 ppm (Snider and Parker, 1979) led to geophysical surveys and drilling (Johnson et al.,1980). Minor amounts of native copper were reported in the drill core (Johnson et al., 1980). Importantly, establishing a geologically meaningful absolute age for the Bear Lake latite would constrain the rate of rift-centric sedimentation and the true end of known rift magmatism. Early attempts at establishing an age were discussed by Morey and Van Schmus (1988) who reported a Rb-Sr age of 1062+/-34 Ma compared to 1007+/-25 Ma by Chaudhuri (1975), but concluded the Rb-Sr isotopic system did not represent the emplacement age. Subsequent dating has shown that the 1060 Ma age is comparable to the 1060 to 1050 Ma age of widespread hydrothermal alteration associated with the native copper deposit (Bornhorst et al., 1988). In 1984, Bornhorst and a student, D. Wall, completed field work along Seven Mile Creek which bisects the body near the contact with the dual objective of developing a better understanding of the age relationship with the Freda Sandstone, intrusive versus extrusive origin and selecting a 61 �sample for zircon U-Pb dating. A sample was submitted to the Royal Ontario Museum for zircon separation and U-Pb dating. However, petrographic observations of the zircons revealed that they were likely xenocrysts and, therefore, no follow up analytical work was completed. Recently, the University of Oslo was able to extract three small zircons that yielded a statistically consistent concordia age of 1091.4 +/-1.7 Ma by ID-TIMS. The Bear Lake body intersects the Freda Sandstone and is about 1,500 m above the base of the Nonesuch Shale (Fig. 1) dated at 1081 +/- 9 Ma (Pb-Pb isochron; Ohr, 1993). Stratigraphically about 1,000 m under the Nonesuch is the Lake Shore Traps dated at 1087.2+/-1.6 Ma (U-Pb age on zircons; Davis and Paces, 1990). Fairchild et al. (2017) reported a 1.6 Ma younger age for the Lake Shore Traps and also reported an age of about 1084 Ma for rocks of comparable stratigraphic position from Michipicoten Island. Thus, the 1091 Ma age on the Bear Lake igneous body is not consistent with other published radiometric ages and, therefore, is not geologically meaningful; i.e. the zircon grains are xenocrysts as suspected decades prior. Figure 1: Stratigraphic column. We have not yet exhausted our efforts to obtain a geologically meaningful radiometric age for Bear Lake and continue to search for primary igneous zircons. References Cited Bornhorst, T.J., Grant, N.K., Obradovich, J.D., and Huber, N.K., 1988, Age of native copper mineralization, Keweenaw Peninsula, Michigan: Economic Geology, v. 83, p. 619-625. Chaudhuri, S., 1975, Geochronology of upper and middle Keweenawan rocks of Michigan: of the 21st Annual Institute on Lake Superior Geology, Proceedings, v. 1, p. 32. Cornwall, H. R., and Wright, J. C., 1956, Geologic map of the Hancock quadrangle, Michigan: U. S. Geological Survey Mineral Investigations Field Studies Map MF 46. Davis, D.W., and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system: Earth and Planetary Science Letters, v. 97, p. 54-64. Fairchild, L. M., Swanson-Hysell, N.L., Ramezani, J., Sprain, C.J., and Bowring, S. A., 2017, The end of Midcontinent Rift magmatism and the paleogeography of Laurentia: Lithosphere, v. 9, no. 1., p. 117-133. Johnson, A., Parker, B., Snider, D., Van Alstine, J., 1980, Petrology of the Bear Lake intrusive, Keweenaw Peninsula, Michigan: 26th Institute on Lake Superior Geology Proceedings, v.1, p. 67. Morey, G. B., and Van Schmus, W. R., Correlation of Precambrian rocks of the Lake Superior region, United States: U. S. Geological Survey Professional Paper 1241-F, 32p. Ohr, M., 1993, Geochronology of diagenesis and low-grade metamorphism in pelites: Ph.D. dissertation, The University of Michigan, Ann Arbor, MI.,161 p. Snider, D. W., and Parker,, B. K., 1979, Geochemical and geophysical anomalies associated with the Bear Lake intrusive, sections 24 and 25, T56N, R34W, Houghton County, Michigan: 25th Institute on Lake Superior Geology Proceedings, v. 1, p. 38. 62 �Land of Fire and Ice: Summary of the 2017 ILSG Field Trip to Iceland LARSON1, Phil; HUDAK2, George; MACTAVISH3, Al; HINZ4, Peter; RADAKOVICH5, Amy; BHATTACHARYYA6, Juk; ENGELHARDT7, Paula; ENGELHARDT8, Steve; GELNIAS9, Brigitte; GOOD10, David; GORNER9, Emily; HINZ4, Sheree; JONGEWAARD11, Peter; KROCH, Deb; SVENSSON10, Matt; TIMS12, Andrew 1 Vesterheim Geoscience, PLC, Duluth, MN Natural Resources Research Institute, University of Minnesota - Duluth, MN 3 Panoramic PGMs (Canada) Limited, Thunder Bay, ON 4 Ontario Ministry of Northern Development and Mines, Thunder Bay, ON 5 Minnesota Geological Survey (MGS), University of Minnesota-Twin Cities 6 Department of Geography, Geology and Environmental Science, University of Wisconsin – Whitewater 7 HydroGeo Solutions LLC, Green Bay, WI 8 Green Bay, WI, independent photographer 9 Department of Geology, Lakehead University, Thunder Bay, ON 10 Department of Earth Sciences, Western University, London, ON N6A 5B7 Canada 11 Cliffs Natural Resources - Retired 12 Northern Mineral Exploration Services, Thunder Bay, ON 2 ABSTRACT The Institute on Lake Superior Geology recently conducted a field trip to Iceland between July 26 and August 5, 2017. The 11-day trip was led by Phil Larson, George Hudak, Al MacTavish, and Peter Hinz, and included 16 participants, including 10 professional geologists and 4 graduate students. The trip traversed roughly the south one-half of the island (Fig 1). Stops covered a wide variety of topics – from volcanology, igneous petrology, and magmatic-hydrothermal ore deposits to the glacial geology and geomorphology of Iceland. Iceland is the subaerial expression of the crust that forms the floor of the Atlantic Ocean. Though all of its rocks are younger than about 25 Ma, the geology of Iceland bears many similarities to that of the 1.1 Ga Midcontinent Rift in North America. The Mid-Atlantic Ridge steps ~100 km east across the island and allowed us to see a modern-day expression of rift-related volcanic and hypabyssal intrusive mafic (to rarely felsic) rocks, which continue to form today. The onset of the most recent Ice Age in the Pleistocene created spectacular ice sheets, valley glaciers, and glacial sediment deposits, as well as the unique landforms reflecting the interaction of volcanism and glacial ice, all of which we observed on the trip (Thordarson & Hoskuldsson, 2014). 63 �Figure 1. Shaded relief map (Google Images, 2017) of Iceland, showing glaciers in white. Colored lines show the 11-day trip route, starting and ending in Reykjavik. This presentation summarizes highlights from our trip and draws comparisons of Iceland’s geology to the Midcontinent Rift. Featured highlights of the bedrock geology will include both subaerial and subglacial volcanic deposits such as mafic tuffs, pillowed basalts, cinder cones, columnar basalts, peperites, pillowed dikes, pumice and ash deposits, moberg, tuyas, and spectacular aa and pahoehoe fields. We will discuss the surface expression of the Mid Atlantic Ridge at the Krafla Lava Fields, the Snaefellsnes Peninsula Volcanic Zone, and the historic Láki Flow. The presentation will also showcase vast outwash plains created by glacial jokulhaups, moraine deposits, proglacial lakes, and crevasse fields. REFERENCES Thordarson, T., & Hoskuldsson, A. (2014). Iceland Second Edition (2nd ed., Classic Geology in Europe 3). Edinburgh: Dunedin Academic. 64 �Controls on the localization and timing of mineralized intrusions within the ca. 1.1 Ga Midcontinent Rift system LIIKANE, Dustin A.1, BLEEKER, Wouter2, HAMILTON, Mike3, KAMO, Sandra3, SMITH, Jennifer2, HOLLINGS, Peter4, CUNDARI, Robert5, and EASTON, Michael6 1 Department of Earth Sciences, University of Toronto, Toronto, Ontario; dustin.liikane@mail.utoronto.ca Geological Survey of Canada, Ottawa, Ontario 3 Jack Satterly Geochronology Laboratory, Dept. of Earth Sciences, University of Toronto, 22 Russell St., Toronto, Ontario, Canada, M5S 3B1 4 Department of Geology, Lakehead University, 955 Oliver Rd, Thunder Bay, Ontario, Canada, P7B 5E1 5 Ontario Geological Survey, 435 James Street South, Thunder Bay, Ontario, P7E 6S7 6 Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, Ontario, Canada, P3E 6B5 2 The 1.1 Ga Mid-Continent Rift (MCR) represents one of the largest and best-preserved intra-continental rift systems of Precambrian age (Davis and Green, 1997; Miller and Nicholson, 2013; Swanson-Hysell et al., 2014; Bleeker et al., 2018). It is host to the Duluth Complex (second largest layered intrusion in the world), which contains significant low-grade mineralization of Ni-Cu-Co and platinum group elements (PGEs), and possibly reef-style PGE mineralization. Higher-grade Ni-Cu mineralization has also been identified within the rift, localized in smaller, conduit-type intrusions (e.g., Eagle deposit). Numerous mineralized intrusions are associated with the MCR on both sides of the border (e.g., Coldwell Complex near Marathon, Ontario; Tamarack near Duluth, Minnesota; Eagle near Marquette, Michigan), with many being actively explored by a number of companies. Many intrusions related to the MCR have been dated by U-Pb methods (Figure 1); however, many of the ages were obtained prior to the introduction of routine chemical abrasion techniques on single zircon grains. This improvement often leads to better precision and accuracy, allowing for sub-million-year age resolution. With this technique, we can better constrain (for the first time, in some cases) the age of emplacement of several MCR-related intrusions. This will allow us to understand the dynamics of the MCR’s plumbing system, and how it evolved over time. At the deposit scale, the high-precision ages may allow us to recognize whether these intrusions are long-lived conduits or intrusions emplaced in one single magmatic pulse. Furthermore, high-precision ages, along with lithogeochemistry, will allow us to link these individual intrusions to distinct stages of the flood basalt sequence. It may also reveal temporal constraints on the formation of mineralized intrusions within the MCR. References: Bleeker, W., Liikane, D.A., Smith, J., Hamilton, M., Kamo, S.L., Cundari, R., Easton, M., and Hollings, P., 2018, Controls on the localization and timing of mineralized intrusions in intra-continental rift systems, with a specific focus on the ca. 1.1 Ga Mid-continent Rift system: in Targeted Geoscience Initiative: 2017 report of activities, v. 2, (ed.) N. Rogers; Geological Survey of Canada, Open File 8373, p. 15–27. https://doi.org/10.4095/306594. Davis, D. W., and Green, J. C., 1997, Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution: Canadian Journal of Earth Sciences, v. 34(4), p. 476–488. doi:10.1139/ e17-039. 65 �Miller, J., and Nicholson, S.W., 2013, Geology and mineral deposits of the 1.1 Ga Midcontinent Rift in the Lake Superior region – an overview: In Field guide to the copper-nickel-platinum group element deposits of the Lake Superior Region: Edited by Miller, J. Precambrian Research Center Guidebook, v. 13-01, p. 1–49. Swanson-Hysell, N.L., Burgess, S.D., Maloof, A.C., Bowring, S.A., 2014, Magmatic activity and plate motion during the latent stage of Midcontinent Rift development: Geology, v. 42, p. 475-478. Figure 1: Summary map of the Midcontinent Rift (adapted from Miller and Nicholson, 2013, and references therein), highlighting all of the rift-related intrusions. Undated or poorly dated intrusions, and/or ages that are problematic, are identified by stars with yellow outline. High-precision U-Pb ages on volcanic rocks are shown for reference. Summary of lithostratigraphic columns from across the Midcontinent Rift are integrated into this map nearest to their approximate geographic locations (adapted from Swanson-Hysell et al., 2014, and references therein). For complete references to all the ages, see Bleeker et al. (2018). 66 �Microanalysis of rock and mineral textures and its relationship to mineralization and ore comminution MATKO, Matthew W. , and SCHARDT, Christian Department of Earth and Environmental Sciences, University of Minnesota Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 Research on ore bodies has typically focused on macro-scale processes, such as fluid migration, chemical and metal transfer, as well as physical and chemical changes (Cathles, 1981; Zientek, 2012). The results of this research are then typically been applied to large-scale features based on field observations, laboratory experiments, and theoretical assumptions to create models for ore deposits. However, it is not well understood how micro-scale properties such as: mineral grain size, grain shape, grain orientation, or fracture characteristics may influence various ore deposit formation. Mineralization styles such as: disseminated, next-texture, porphyry, vein-style may be partially controlled, if not dictated, by these features. It is therefore important to gain a better understanding of the role of these features in larger-scale processes. To obtain small-scale rock property information, multiple analytical methods (x-ray computed tomography - XRCT, Electric Pulse Disaggregation - EPD, Mineral Liberation Analysis - MLA) were used to examine selected ore samples (porphyry, Mississippi Valley type, volcanic massive sulfide, liquid magmatic sulfides). Sample cores were scanned using XRCT and spatial reconstructions were produced using 3D image processing software. This technique yields in-situ information such as grain size parameters, porosity distribution, or spatial orientation of mineral aggregates. Hand samples were disaggregated into individual mineral grains using EPD by sending repeated electric pulses through the material, causing mineral separation preferentially along mineral grain boundaries. This technology allows material separation while preserving mineral grain morphology (Cabri et al., 2008) and may provide a true alternative to traditional methods of ore comminution. The resulting aggregate material was analyzed with scanning electron microscopy using MLA software. This software yields mineral liberation data from the EPD technology, in addition to grain shape, size, mineral associations, and mineral abundances. Results show that XRCT data can be utilized to locate small-scale melt migration pathways such as micro-fractures (fig. 1a) in addition to mineral grain morphology and size distribution. Identification of these micro-scale features has the potential to greatly assist in our understanding of how ore textures and mineral deposits form. In addition to visual analysis of insitu material, data can be exported to construct vector graphs, which enable the visualization of ore grain orientations to determine existing ore grains orientation patterns within the rock (fig. 1b). Results also indicate very good mineral separation of silicates from sulfides using EPD compared to traditional mechanical processing, especially at < 250-150 µm. Separation efficiency was confirmed by running MLA on the samples, where it was determined that the average liberation yield for ores present ranged from 70 % to 80 % (fig. 1c). There does appear to be some dependence on ore grain size and deposit type that can affect these values. EPD technology offers unique opportunities for ore processing such as the pre-weakening of material before being subjected to more traditional methods of crushing, or an initial ore-silicate separation phase to remove the bulk of the gangue material. 67 �Figure 1 3D reconstruction of Cu-sulfide ore from a porphyry deposit that highlights a planar feature created by mineralization within a micro-fracture (a). Data for particles within the planar feature graphed to display long axis grain orientation (b). A composite image showing a representative sample of sulfide mineralization from disaggregated Eagle Mine material using MLA software. It highlights the high degree of Cu-sulfide liberation from silicates while also showing minor Cu-sulfides association with Ti-oxides (c). References Cabri, Louis J., Rudashevsky, N. S., Rudashevsky, V. N., & Oberthür, T. (2008). Electric-pulse disaggregation (Epd), hydroseparation (Hs) and their use in combination for mineral processing and advanced characterization of ores. Canadian Mineral Processors, 40th Annual Meeting, Ottawa, Proceedings. v. 211, p. 211-235. Cathles, L.M. (1981) Fluid flow and ore genesis of hydrothermal ore deposits: Economic Geology 75th Anniversary Volume, p. 424 – 457 Zientek, M.L. (2012) Magmatic Ore Deposits in Layered Intrusions - Descriptive Model for Reef-Type PGE and Contact-Type Cu-Ni-PGE Deposits: U.S. Geological Survey Open File Report 2012-1010, p. 48 68 �Using Credit-by-Exam to Connect Advanced High School Geology Courses to University Geology Departments: Current Status of a State-wide Program in Michigan MATTOX, Stephen1, BOLHUIS, Chris 2, and SOBOLAK, Christina 3 1 Department of Geology, Grand Valley State University, Allendale, MI, 2Hudsonville High School, Hudsonville, MI; 3St. Fabian Middle School, Farmington Hills, MI To address the national shortage of geologists and to diversify the geoscience workforce we are constructing a seamless path from rigorous high school geology classes taught by well-trained teachers to geoscience departments at state universities and colleges. This program is modeled on an existing high school – university collaboration that has added a significant number of students to the career pipeline. Although geologists have vigorously lobbied for an AP course and exam in geology, the College Board has resisted because they perceive that demand will be low. The lack of an AP course has made it difficult for those high schools nationwide that offer advanced (i.e., college-level content) geology courses to obtain appropriate recognition for their students’ accomplishments. Mattox designed a test, with input from my GVSU peers (and reviewed by faculty at eight other university geology departments), that includes 70 multiple choice, 10 essay questions, a map test with skills and landforms, and a rock and mineral exam (essentially what we use in our introductory physical geology course). We give the exam over 5-6 hours on two different days. The support of two NSF grants allowed us to build the network of universities and then extend the size of the program. NSF supported ended in 2016. Universities awarding credit: Central Michigan, Eastern Michigan, Grand Valley, Lake Superior State, Michigan Tech, Northern Michigan, University of Michigan-Dearborn, Wayne State University, Western Michigan, and Hope College. Montana State has awarded credit to two students. Mattox has started new discussions with Ferris State University and nearby 2YCs: Muskegon Community College and Grand Rapids Community College. Additional colleges and universities are invited to join each year. Participating colleges sign a MOU to award credit. Since 2001, 1334 students have taken a rigorous high school geology/Earth science course. Of these 777 students have passed, about 58%. With NSF support the program has grown and now about 250 students are tested each year at 9 or 10 high schools. Commonly, 20 students request credit and 5-7 start university as declared geology or Earth science majors, usually at CMU, GVSU, MTU, NMU, or WMU. Insights from this project include: • • • Administrators were surprisingly receptive to and supportive of a rigorous high school geology credit by exam course. Ideally teachers with a B.S. in Geology and additional course work or M.S. teach the course. However, motivated science teachers without an Earth Science B.S. can be successful. The high school course can be a single semester, two trimesters, or a full year. 69 �• • • • Pass rates tend to improve over 3-4 years and then stabilize. Some schools never had a student pass. About two new high schools join the program each year. Currently, high school teachers are working to align the Next Generation Science Standards to the content/skills of the college physical geology class. We have demonstrated the feasibility of establishing a statewide network of universities to award college credit for passing a credit by exam during a high school geology course. Growth of credit-by-exam in Michigan over the last six years. High schools in the program: HUD (Hudsonville); GH (Grand Haven); GPS (Grosse Point South); OK (Okemos); BR (Black River); DA (Dream Academy); DIT (Detroit Institute of Technology); HF (Henry Ford Academy); MA (Multicultural Academy); PHS (Pioneer); HHS (Huron); SHS (Sturgis); RHS (Roscommon HS); WOHS (West Ottawa HS); FHC (Forest Hills Central); and KH (Kenowa Hills). This material is based upon work supported by the National Science Foundation under OEDG Grant No. NSF 08-605 1006652. Any opinions, findings, and conclusions or recommendations expressed in this material are those of the author(s) and do not necessarily reflect the views of the National Science Foundation. 70 �Geochemical signatures of hydrothermal alteration in clastic sedimentary rocks: theory, recognition, and application MAUK, Jeffrey L. 1 USGS, MS-973 Denver Federal Center, P O Box 25046, Denver, CO 80225-0046, USA Much of the terminology to describe hydrothermal alteration came from classic studies of porphyry copper deposits, which popularized terminology such as potassic, argillic, advanced argillic, phyllic, and propyllitic (e.g., Lowell and Guilbert, 1970; Sillitoe, 2010, and references therein). This terminology was developed for plutonic and volcanic igneous rocks that, when fresh, contain unaltered feldspar and mafic minerals. The main driver of hydrothermal alteration, hot water, promotes reactions that convert feldspar and mafic minerals to phyllosilicate minerals. In contrast, clastic sedimentary rocks form from weathered material, and that weathering results in degradation or destruction of feldspar and mafic minerals to form clay minerals such as smectite. During diagenesis, these clay minerals transform to higher rank clay minerals, such as interstratified illite-smectite and illite. Diagenetic reactions can ultimately lead to formation of micas such as muscovite. Therefore, normal weathering and diagenetic reactions destroy many igneous minerals, and form minerals that are similar to those that occur in hydrothermally altered igneous rocks. This can make it difficult to recognize hydrothermal alteration in sedimentary rocks. This abstract describes some key styles of hydrothermal alteration, and evaluates geochemical methods that can help to identify these types of alteration in clastic sedimentary rocks. Chemical sedimentary rocks, such as limestone and dolomite, are not considered. In igneous rocks, propylitic alteration commonly covers extensive areas; it is characterized by chlorite, epidote, and calcite, with minor pyrite. Chlorite, calcite, and pyrite are common diagenetic minerals, so propylitic alteration is an excellent example of a style of alteration that is readily identifiable in igneous rocks, but difficult to impossible to recognize in sedimentary rocks. Furthermore, geochemical studies of altered igneous rocks show that major elements are relatively immobile during propylitic alteration, and the main components that are gained are sulfur and carbonate. Again, because carbonate minerals and pyrite are common in sedimentary rocks, geochemical analyses are unlikely to provide diagnostic evidence of propylitic alteration. In contrast, alkali metasomatism, which includes K and Na alteration, can produce diagnostic minerals and distinct whole rock geochemical compositions. These reactions can produce Kfeldspar albite, illite, or Na-mica, or a combination of these minerals. Mass changes associated with K and Na metasomatism can be evaluated graphically using plots of molar (2Ca + Na + K)/Al versus molar K/Al to evaluate K-metasomatism, and molar (2Ca + Na + K)/Al versus molar Na/Al to evaluate Na-metasomatism. These plots allow identification of important hydrothermal minerals, and reflect alteration processes by showing trends from unaltered toward altered rocks. Sodium metasomatism is a hallmark of many sediment-hosted Cu deposits, but albite is also a common diagenetic mineral, so geochemical testing must include a sufficiently large suite of rocks to test how common and widespread albite is on a regional basis, and whether Na metasomatism stems from diagenesis or mineralization, or both. Phyllic alteration is characterized by quartz, sericite, and pyrite. Where intensely and pervasively developed, this alteration produces white rocks that are rich in pyrite; the pyrite may form up to 10% of the volume of the rock. Phyllic alteration is accompanied by leaching of Mg, Na, and Ca, and enrichment of K and S, so it is readily characterized by major element and S data. 71 �Argillic alteration can be texturally destructive, and produces clay minerals such as montmorillonite, illite, cholorite, and kaolinite. Advanced argillic alteration is very texturally destructive; it results from more aggressive acid leaching of rock that produces quartz or vuggy silica, plus alunite and kaolinite. In both argillic and advanced argillic alteration, Mg, Fe, Ca, Na, and K are leached from the rock. Silica may appear to be enriched, but that is due to loss of other elements rather than actual addition of SiO 2 . Silicification is the addition of silica to the rock, typically as quartz in veins, or in pore-filling cement, or both. In some cases, quartz replaces precursor minerals in the rock. All styles of silicification may be well-developed in host rocks around hydrothermal veins. Recognition and quantification of veins is relatively easy, but identification of silicification by pore filling or mineral replacement is more difficult. This is best exemplified in sandstones and quartzites, where beds that are naturally coarser-grained and more well-sorted would be harder and may have a more vitreous luster, leading them to be classified as silicified. Alas, geochemistry offers little assistance here, because the natural variability of clastic sedimentary rocks ensures a wide range of SiO 2 concentrations, and it is exceptionally difficult to document Si gain except under extreme conditions. Sulfidation is the addition of sulfide minerals—typically pyrite—to rock. This is common around many hydrothermal veins, and can be intense and pervasive around sediment-hosted massive sulfide deposits. Sulfur addition is readily characterized by whole rock geochemistry, provided that, as noted above, the addition of sulfur is sufficient to exceed background concentrations from diagenetic pyrite. Bleaching is used where rock is a lighter color than normal, and has two main styles: (1) a general lightening in color, such as dark green to light green, and (2) changing of a red rock to a white rock. The former is common around some hydrothermal veins. The latter occurs in some sediment-hosted copper deposits, and is spectacularly displayed in redbeds in the four corners region of the U.S., where the bleached zones reflect pathways of oil and gas migration. Bleaching is a nearly isochemical process, although in some cases (2) results in Fe loss. Carbonate alteration is common and widespread around many hydrothermal veins, and around many stratiform and stratabound orebodies. Some deposits show pronounced zonation of carbonate minerals, from Fe-rich near deposits, to Ca-rich in distal areas. The carbonate alteration can occur as veins and veinlets, particularly in the inner alteration zones, but disseminated carbonate is more common. Carbonate alteration can be quantified by geochemical analyses of carbonate C, provided that the addition of carbonate is sufficient to exceed background concentrations from diagenetic carbonate. In summary, whole rock geochemistry can provide a means to evaluate and quantify many, but not all types of hydrothermal alteration. Major element analyses, plus total S, carbonate C, and organic C, are the most important for geochemical evaluation of clastic sedimentary rocks. The most significant sediment-hosted deposits in the Midcontinent Rift are sediment-hosted Cu deposits, and for those, alkali metasomatism is the most promising alteration indicator. References Lowell, J. D., and Guilbert, J. M., 1970, Lateral and vertical alteration-mineralization zoning in porphyry ore deposits: Economic Geology, v. 65, p. 373-408. Sillitoe, R. H., 2010, Porphyry copper systems: Economic Geology, v. 105, p. 3-41. 72 �An 1149 Ma U-Pb baddeleyite crystallization age and geochemistry of gabbroic intrusions at the southwestern margin of the Superior Craton, southeastern South Dakota McCORMICK, Kelli1, CHAMBERLAIN, Kevin2, and PATERSON, Colin3 1 Department of Mining Engineering and Management, South Dakota School of Mines and Technology, Rapid City South Dakota 57701; 2Department of Geology and Geophysics, University of Wyoming, Laramie, WY 82071, Faculty of Geology and Geography, Tomsk State University, Tomsk 634050 Russia; 3Department of Geology and Geological Engineering, South Dakota School of Mines and Technology, Rapid City South Dakota 57701 Gabbroic intrusions have been intersected in drill holes in southeastern South Dakota along the southwestern margin of the Superior Craton (Fig. 1). In order to better constrain the ages of the basement terranes in this region, several samples from one set of intrusions, the Corson diabase, were analyzed for the presence of datable minerals. For this study, samples of one Corson intrusion analyzed by Meyers (2013) was sent for mineral separation. Approximately 40 baddeleyite grains were separated by U. Söderlund (Lund University). Dating of the baddeleyite was by U-Pb isotope dilution thermal ionization mass spectrometry at the University of Wyoming. A U-Pb baddeleyite crystallization age of 1149.4 + 7.3 Ma from this Corson diabase sample (McCormick et al., 2017) is interpreted to represent an early stage of the Midcontinent Rift (MCR). McCormick et al. (2017) suggest that Corson diabase intrusions represent a failed rift arm (Fig. 1). The inferred NE trend of the Corson diabase, considered together with the trends of other possible MCR-related intrusions, are also consistent with the model of a mantle plume origin for the MCR (Hutchinson et al., 1990). Sampling of another Corson diabase core for U-Pb mineral dating is in progress. Several samples from two cores (03-W-01 and 03-W-02) intersecting a large gabbroic intrusion(s) near the town of Wakonda (Fig. 1) were analyzed in this study for the presence of datable minerals. A sample from 03-W-01 sent for mineral separation yielded approximately 50 baddeleyite grains, some more than 100 µm. Dating of these minerals is in progress. In conjunction with the dating, samples were taken from the Wakonda cores for geochemical analysis and compared with existing geochemistry of Corson diabase. The Wakonda and Corson samples are tholeiitic and generally olivine normative. A plot of TiO 2 vs Mg# (Fig. 2) shows the Wakonda gabbros to be somewhat geochemically distinct from the Corson diabase, but similar to MCR intrusions around the Thunder Bay region. REFERENCES: Cundari, R.M., Carl, C.F.J., Hollings, P. and Smyk, M.C., 2013, New and compiled whole-rock geochemical and isotope data of Midcontinent Rift-related rocks, Thunder Bay Area. Ontario Geological Survey Miscellaneous Release—Data 308. McCormick, K. A., Chamberlain, K. R, and Paterson, C. J., 2017, U–Pb baddeleyite crystallization age for a Corson diabase intrusion: possible Midcontinent Rift magmatism in eastern South Dakota. Can. J. Earth Sci., Published at www.nrcresearchpress.com/cjes on 10 October 2017, 7 p. Myers, J., 2013. A petrographic analysis of mafic intrusions of an unknown age from Southeastern South Dakota. Unpublished senior thesis, Department of Geology and Geological Engineering, South Dakota School of Mines and Technology: 17 p. Hutchinson, D.R., White, R.S., and Cannon, W.F., and Schulz, K.J., 1990. Keweenaw hot spot: geophysical evidence for a 1.1 Ga mantle plume beneath the Midcontinent rift system. J. of Geophys. Res., 95, p. 10,869-10,884. 73 �Figure 1: Map encompassing the intrusions discussed in this study. Diamonds are Corson diabase, triangles are other gabbros. MCR = Midcontinent Rift; SBZ = Superior boundary zone; SLTZ = Spirit Lake tectonic zone; SQ = Sioux Quartzite (subsurface extent); SRA = Superior rift arm (proposed). Becker embayment after the NICE working group. Figure 2: TiO 2 vs Mg# plot of southeastern South Dakota gabbroic intrusions and Thunder Bay area MCR intrusions from Cundari et al., 2013. Filled diamonds are Corson diabase, filled squares are Wakonda gabbro core 03-W-01, and large filled circles are Wakonda gabbro core 03-W-02. 74 �Geology of the Crystal Lake Gabbro and the Mount Mollie Dyke, Midcontinent Rift, Northwest Ontario O’BRIEN, Sean1, HOLLINGS, Pete1, and MILLER, Jim2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada; 2Department of Earth and Environmental Sciences, University of Minnesota - Duluth, 1114 Kirby Drive, 223 Heller Hall, Duluth, MN 55812. 1 The Crystal Lake Gabbro (CLG) is a Y-shaped, up to 750 m wide, layered mafic intrusion with a 5 km long northern limb and a 2.75 km long southern limb, with localized Cu-Ni mineralization. The Mount Mollie Dyke (MMD) is an arcuate, 60 to 350 m wide, macrodyke that lies on trend east of the CLG and extends for 35 km toward Lake Superior. Both intrusions are part of the 1.1 Ga Midcontinent Rift (MCR) and were emplaced into the Paleoproterozoic Rove Formation of the Logan Basin, approximately 50 km south of Thunder Bay. Current U-Pb age determinations imply a ~10 m.y. age difference with CLG being formed at 1099.6 ± 1.2 Ma and the MMD being formed at ~1109.3 ± 6.3 Ma (Heaman et al., 2007; Hollings et al., 2010). However, this age difference is at odds with both intrusions being normally polarized (an attribute of MCR rocks younger than 1102 Ma; Davis and Green, 1997) and their being on trend with each other. The CLG profiled in a drill core from its southern limb can be broadly divided into Upper, Main, and Lower Zones with further subdivisions of the Main and Lower Zones based largely on geochemistry. The Lower Zone occurs between two xenoliths of an early MCR (~1115 Ma) plagioclase porphyritic Logan Sill diabase. The Lower Zone consists of subophitic to ophitic troctolite, augite troctolite, and olivine gabbro and can be subdivided into an upper and basal marginal subzone as well as an interior subzone. Both marginal subzones host disseminated sulphides. An overall bottom-up-directed fractional crystallization of the Lower Zone is suggested by the progressive decrease in Fo content of olivine, Mg# of clinopyroxene, and whole-rock MgO upsection. Above the upper Logan Sill xenolith, the Main Zone similarly consists of subophitic to ophitic troctolite, augite troctolite, olivine gabbro, and gabbro. Petrography, lithogeochemistry, and mineral composition was used to subdivide the Main Zone into five subzones: a basal marginal subzone, upper margin subzone, and three interior cycles that display cryptic variations indicative of fractional crystallization and magma recharge events. Like the margins of the Lower Zone, the Upper Zone as well and the basal marginal subzone of the Main Zone contain disseminated sulphides and are characterized by relatively high Fo content olivine and low incompatible trace element concentrations. These mineralized zones are interpreted to have crystallized from the same initial pulse of magma into the CLG, which was sulphur-saturated. Cyclical cryptic variations in the internal subzone of the Main Zone are interpreted to indicate upward directed fractional crystallization, interrupted by emplacement of additional magma pulses into the core of the intrusion. All rocks of the Main Zone are olivine and plagioclase orthocumulates indicating that fractional crystallization was not particularly efficient. Throughout the evolution of the CLG, the differentiation of the magma was limited as it did not result in clinopyroxene and Fe-Ti oxide becoming cumulus phases. This was likely due to magmatic recharge and inefficient fractional crystallization. 75 �Texturally and geochemically, the MMD can be broadly divided into an Upper and Main Zones, with a subdivision of the Main Zone into an upper and lower sequence and a pegmatitic segregation subzone. The Upper Zone consists of ferrodiorite and likely represents the end product of extensive fractionation. The Main Zone is characterized by troctolite, augite troctolite, olivine gabbro, and gabbro with MgO, CaO, Al 2 O 3 , and Ni concentrations decreasing upwards and SiO 2 , TiO 2 , K 2 O, Na 2 O, P 2 O 5 , and incompatible trace element concentrations increasing, consistent with bottom-up fractional crystallization. Strong differentiation of the MMD magma is indicated by the habit change of clinopyroxene from ophitic (intercumulus) to granular (cumulus), which is the basis for the subdivision of the lower and upper sequences. The lower sequence of the Main Zone also hosts a 24 m thick interval containing 1 to 2 m wide gabbroic pegmatite layers. These pegmatites are interpreted to be the result of localized enrichment of magmatic volatiles. The presence of an evolved core in the MMD surface expression, coupled with the mineral composition of olivine, plagioclase, and clinopyroxene, remaining at relatively constant Fo, An, and Mg# values, respectively, below the pegmatitic layers suggests that there was some degree of lateral crystal fractionation as well as bottom up fractionation. The well-defined fractionation sequence as well as an absence of abrupt geochemical changes suggests that the MMD fractionally crystallized from a single pulse. Liberation of external sulphur from the surrounding Rove Formation, is suggested by the greater than mantle S/Se values as well as δ34S values between +4.0 and +21.0‰ of the sulphides within the CLG. The addition of external sulphur evidently resulted in sulphur saturation during initial emplacement of the CLG magmas. Primitive mantle normalized multi-element diagrams and trace element ratios provide supporting evidence for a localized shallow level of crustal contamination, as well as a deeper more widespread contamination component of both the CLG and MMD magmas. The estimated parental magma compositions and average primitive mantle normalized trace element concentrations of the CLG and MMD suggest that they shared similar, if not the same, magma source. The CLG parental magma was slightly more evolved than the MMD suggesting that the magmas were sourced from a fractionating staging chamber. The estimated parental magma compositions of the CLG and MMD closely resemble those of the Layered Series intrusions of the Duluth Complex, supporting previous speculation that the CLG may be a satellite intrusion of the Duluth Complex. Despite current geochronology data to the contrary, the results of this study strongly suggest that the CLG and the MMD are petrogenetically linked, if not parts of the same intrusive system. REFERENCES Davis, D. and Green, J. (1997) Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences 34, 476-488. Heaman, L., Easton, R., Hart, T., Hollings, P., MacDonald, C. and Smyk, M. (2007) Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences 44, 1055-1086. Hollings, P., Smyk, M., Heaman, L.M. and Halls, H. (2010) The geochemistry, geochronology and paleomagnetism of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada. Precambrian Research 183, 553-571. 76 �The Brussels Hill Structure, Door County, Wisconsin: Impact crater, diatreme or other? OLSEN-VALDEZ, Juliana and BJØRNERUD, Marcia Geology Department, Lawrence University 711 E Boldt Way, Appleton, WI 54911 USA Brussels Hill (44.759°N, 87.593°W) is a localized area of intensely fractured and faulted bedrock in a region of otherwise undeformed lower Silurian dolostone. It was first identified as a geologic anomaly by Kluessendorf (2011), who suggested that the site was an impact crater. More recently, Lawrence University students have carried out geological and geophysical surveys of Brussels Hill (e.g., Zawacki & Bjørnerud, 2014), and in 2017 we obtained a ca.100 m core from the center of the structure. While some characteristics of the site are consistent with an eroded impact crater, others are at odds with this hypothesis. The area of disturbed rock coincides with a distinctive, nearly circular, flat-topped topographic high, ca. 2 km in diameter, which stands 40 m above the surrounding landscape and is ringed by rugged, tree-covered slopes. Around the edges of the hill, most prominently on the north side of the structure, the Silurian bedrock dips gently (15-20°) inward. Glacial till lies above the dolostone at the top of the hill. A quarry near the central part of the disturbed area provides excellent three-dimensional exposures of the most intensely deformed bedrock. In the quarry, bedding orientations vary dramatically over distances of meters. Coherent structures are difficult to discern, and the rocks are fragmented at every scale. In places, primary layering can be traced for tens of meters, while elsewhere the rocks are pervasively brecciated to centimeter- or smaller-sized clasts. Some breccias are monomict but most are polymict, containing dolostone and chert clasts with a variety textures and hues. Both types of breccia commonly contain subspherical vugs 1-5 mm in diameter. The breccias lack any sort of internal stratigraphy and are thus unlikely to represent fall-back ejecta, as first suggested by Kluessendorf (2011). No shatter cones have been observed, even though the finely crystalline host dolostone would be favorable for their formation. Silurian dolostone is the only bedrock normally exposed in this area, but meter-scale lensoid and tabular bodies of fine-grained glauconite-bearing sandstone occur at Brussels Hill. These sandstones have a carbonate matrix with spherical, sub-mm-scale voids. Many also have fine laminae that parallel the edges of the bodies. Most of the grains in the sandstones are highly rounded, but about 15% have unusual shapes: shard-like, crescentic, and irregularly concavo-convex. No shock lamellae have been observed in the quartz grains. The mature, rounded grains and presence of glauconite suggest that these sands were derived from Cambrian or possibly Ordovician strata that normally lie 350 to 400 m in the subsurface in Door County, and their presence rules out a karstic collapse origin for the disturbance. We conducted a gravity survey of Brussels Hill using a LaCoste-Romberg gravimeter and Trimble differential GPS system (Edwards, 2016). N-S and E-W transects with readings every 200 m were made across the hill, extending on the north and west into areas where bedrock is undisturbed. Using the GPS ‘base and rover’ system, station locations were sited to 1 cm precision. After free-air, terrain, and tidal corrections of the time-stamped and geolocated data, the resulting Bouguer anomaly map revealed a small but significant positive gravity anomaly of 0.85 mGal near the center of the hilltop. Because the brecciated rocks exposed at the surface have a lower density (ca. 2.5 g/cm3) than the surrounding pristine dolostone (2.8 g/cm3), the observation of a positive gravity indicates that there must be relatively dense rocks in the subsurface below the center of the Brussels Hill structure. 77 �We have recently logged a 103-m drill core from the quarry, obtained in cooperation with the Wisconsin Geological and Natural History Survey. The rocks in the core are brecciated to varying degrees to a depth of about 70 m, and the core intercepted the Upper Ordovician Maquoketa Shale at about 66 m. The brecciated zones are similar to those exposed at the surface, with vug size broadly correlated with the size of clasts. Thin sections of apparently intrusive veins of brecciated material show crude size sorting. Chert clasts in the core have a wide range of colors – not only white and grey, common in the Silurian units -- but also beige, green, dark red and brown, perhaps from Ordovician strata. Some brown cherts are shattered dilatantly in a manner that suggests explosive decompression. Collectively, these observations point to the involvement of a gas phase that forcefully propelled broken rock upward from significant depth and into a complex network of fractures. The biggest surprise from the drilling was that the lowest rocks in the core – about 30 m of the Maquoketa Shale -- are almost undeformed. This was unexpected, given that material from underlying Cambrian/Ordovician units is intermingled with the Silurian dolostones above; the exotic sandstones and cherts must have passed through the Maquoketa level en route to the surface. Our provisional interpretation is that the Maquoketa Shale, which acts as a regional aquitard in the modern groundwater system, behaved in a similar way in the face of the gas pressures during the explosive event at Brussels Hill. In an impact scenario, carbon dioxide released by shock-related devolatilization of carbonate rocks in the deep subsurface may have been trapped by the low-permeability Maquoketa Shale, which then failed locally, providing isolated conduits for deep-seated rocks to be brought to the surface by over-pressured gases. The drilling site was apparently not one of those spots. However, recent experiments on carbonates in shock metamorphism (Bell, 2016) show that the pressures required for devolatilization of calcite and dolomite exceed 20 GPa – higher than the transient pressures needed to form shatter cones and planar deformation features in quartz (ca. 8 and 12 GPa, respectively), which are absent at Brussels Hill. Other inconsistencies with the impact hypothesis are 1) the inward dips of the beds around the disturbed zone and 2) the height of the hill, which at 40 m is about 10 times higher than the expected central uplift for a crater of 2-3 km diameter (Cintala & Grieve, 1992). We therefore speculate that the disturbance was caused by an overpressured gas phase that came from below, perhaps from a kimberlite or similar intrusion. In this case, the gases that formed the vuggy rocks and carried rocks up from lower stratigraphic levels may have left a distinctive geochemical signature. Preliminary XRF and cathodoluminescence analyses do suggest that some of the carbonate material in the vuggiest breccias and intrusive sandstone bodies is chemically distinct, with elevated Sr values and blebby occurrences of calcite (in rocks that are otherwise entirely dolomitic). The Brussels Hill structure lies about 160 km south of the Jurassic Lake Ellen kimberlite in Iron County, MI (Cannon & Mudrey, 1981; Zartman et al., 2013). References cited Bell, M., 2016. Meteoritics and Planetary Science 51, 619-46. Cannon, W.F. & Mudrey, M., 1981. USGS Circular 842. Cintala, M. & Grieve. R., 1992. Geol. Soc. Am. Special Paper 293, 51-60. Edwards, K., 2016. Lawrence University Senior Thesis, unpub. Kluessendorf, J., 2011. Geol. Soc. Am. Abstr. 43.1, 117. Zartman et al., 2013. Journal of Petrology, 54, 575-608. Zawacki, E. & Bjørnerud, M., 2014. Geol. Soc. Am. Abstr. 46.6, 707. 78 �Komatiite-hosted nickel-copper mineralization potential in the eastern Shebandowan Greenstone Belt, Ontario, Canada OLSON, Maile J1., LODGE, Robert W. D1., and HINZ, Sheree2 1 Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI 54702-4004 2 Ontario Geological Survey, Thunder Bay, ON, Canada, P7E 6S8 Komatiite-hosting strata in Archean greenstone belts are important exploration targets because of the potential for hosting nickel-copper (Ni-Cu) magmatic sulfides (e.g. Houlé 2011). The 2.72 Ga Shebandowan greenstone belt, which is part of the Wawa-Abitibi terrane (Stott et al. 2010), has known to host such deposits in the western part of the belt at the Shebandowan Nickel Mine (Morton 1982). Despite abundant komatiite deposits in the Eastern part of the belt, no other magmatic sulfide deposits have been discovered in this region to date. Our research continues to refine geochemical and textural data that provides evidence of assimilation and supports interactions between komatiites and silica- and sulfur-rich sedimentary rocks. This project explores the potential for Ni-Cu deposits in this region. Furthermore, petrographic and geochemical study of these rocks can improve our limited understanding of Archean tectonic processes. Komatiites were formed during the Archean when young Earth had enough heat to produce large volumes of mantle-derived magmas. Because these komatiitic melts have such an extremely high temperature, they thermally mechanically erode the base of the flow deposit, carving out a channel for itself, giving the melt the ability to assimilate the host rock. When ultramafic magmas assimilate sulfur-bearing crustal sedimentary rocks, they can form Ni-CuPGE deposits. Detailed field mapping of bedrock exposures was completed in the Bateman property exploration trenches, dug in 2008 by Linear Metals Corporation and expands on previous research by Hinz and Hollings (2015). These exploration trenches improved and expanded the available outcrop surface, giving the opportunity to observe the stratigraphy of komatiite flows and how they interact with the surrounding strata. Original textures have largely been preserved in this area due to minimal deformation and metamorphism so textural evidence can be used as a good indicator of komatiite-sediment interaction. Preliminary results from textural, petrographic, and geochemical analyses provide indication of komatiite-sediment intermingling and the presence of Ni-Cu-PGE sulfides. Fig. 1.A-B shows komatiite-sediment contact textures in outcrop with the lighter colored chert being brecciated in contact with the ultramafic flow. The bedding in the chert are being truncated and the edges of the breccia fragments are rounded and potentially thermally eroded (Fig. 1.A). Transmitted-light petrography, and whole rock geochemical analyses were used on the collected samples. Fig. 1.B shows the chert and komatiite have a very chaotic contact zone with arms and irregular blobs of each rock type. In some areas, the contact is diffuse. Other textures include variolitic glassy margins, rounded sedimentary inclusions, and disruption of chert laminations, suggesting the two rock types had interactions prior to lithification. Fig. 1.C-D are petrographic photos of thin sections from the samples and have jigsaw brecciation as well as rounded edges of brecciated sediment clasts from komatiite assimilation. Fig. 1.D also has a diffuse contact that distinctly presents evidence for komatiite-sediment mingling and potential partial melting of the sedimentary rock. The irregularity of the contact and brecciation expresses that the fracture mechanism is not tectonic but is due to hot komatiites shattering colder 79 �sediments. Geochemical diagrams show the compositional array of komatiites deflects towards the calc-alkalic part of the diagram which is unusual for these magmatic suites and is likely caused by contamination. Melt modeling diagrams showing komatiite-sediment interactions with potential melt compositions display geochemical ranges that cannot be explained by fractional crystallization but instead seems to have a mixing pattern with the sedimentary rock. Figure 1: A-B) Outcrop photographs illustrating the various contact relationships between light colored metasedimentary and dark colored mafic to ultramafic units exposed in the Bateman Property trenches. C-D) Petrographic photographs illustrating the same contact relationships. References Stott, G.M., Corkery, M.T., Percival, J.A., Simard, M. and Goutier, J. 2010. A revised terrane subdivision of the Superior Province; in Summary of Field Work and Other Activities 2010, Ontario Geological Survey, Open File Report 6260, p.20-1 to 20-10. Hinz, S. and Hollings, P. 2015. Preliminary description of the ultramafic metavolcanic rocks in the eastern part of the Shebandowan greenstone belt, northwestern Ontario; in Summary of Field Work and Other Activities 2015, Open File Report 6313, p. 16-1 to 16-7. Houlé, M.G., Lesher, C.M., 2011. Komatiite-associated Ni-Cu-(PGE) mineralization in the Abitibi Greenstone Belt, Ontario. Reviews in Economic Geology 17, 89-121 Morton, P., 1982. Archean volcanic stratigraphy, and petrology and chemistry of mafic and ultramafic rocks, chromite, and the Shebandowan Ni-Cu Mine, Shebandowan, northwestern Ontario. Carleton University, p. 346. 80 �Assembling Minnesota: Integration of 140 Years of Government, Academic, and Industry Geologic Studies into a Seamless Statewide GIS Database PETERSON, DEAN M.1 1 Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Duluth, Minnesota 55811-1442. dmpeters@d.umn.edu Over the last 140 years, the search for ores and the mining of mineral deposits has played a huge role in revealing the geology of Minnesota. Dozens of companies utilized classic exploration techniques (geological mapping, geophysical surveying, geochemical studies, test pitting, drilling, and shaft sinking) to target, develop, and mine ore deposits. The early successes (1880s) of these endeavors drove home to forward thinking individuals in government and at the University of Minnesota the need to understand and characterize the geology of the state in a broad context. These developments included regulations to manage lands and archive company data, long-lived programs in mineral processing and metallurgy research, and dedicated programs to map the bedrock geology of the state (Figure 1). All told, a vast amount of information exists on the geology of Minnesota in archives of state agencies, at colleges and universities, and within the United States (USGS) and Minnesota geological surveys (MGS). However, much of this information is currently still archived in file cabinets in analog form (paper maps, documents, folios), though vast amounts of these data have been scanned and are accessible online. Although great strides have been made to integrate these historic datasets into ongoing digital geologic products, major gaps exist and the standardization of how to capture and digitally archive the geologic facts these data hold is by no means complete. To encourage the development of, and risk assessment tools for, an environmentally sound mining industry, government agencies need to put forth both attractive and competitive policies as well as robust geological information. This is particularly true for the mineral exploration component of the mining industry, for without exploration activities the eventual development and extraction of minerals will not take place. Therefore, the University of Minnesota’s Natural Resources Research Institute (NRRI) has begun a copyrighted internal initiative to create a seamless digital GIS compilation (Table 1) that preserves, integrates, and interprets all of the known and trusted bedrock geological data for the entire state of Minnesota, i.e., Assembling Minnesota. The completion of such a compilation in a format prepared for integrated modeling via spatial analysis is a formidable task, and in the end can take many years to complete. This geological compilation is designed to preserve the observed facts generated by geologists/geophysicists/ geochemists in the field over the last 140 years in a way that future geologists may use to make new geological interpretations long into the future. Figure 1. Timeline of outcrop mapping and mineral exploration/development in Minnesota. 81 �Table 1. Listing of the current datasets in the Natural Resources Research Institute’s, Assembling Minnesota geological GIS compilation, © 2018 Regents of the University of Minnesota. All rights reserved. 82 �Modeling the Precambrian topography of Columbia County, Wisconsin using twodimensional models of Gravity and Aeromagnetic data and Well Construction Reports RASMUSSEN, Joseph1, KINGSBURY STEWART, Esther2, SKALBECK, John1, and GOTKOWITZ, Madeline2 1 University of Wisconsin-Parkside, 900 Wood Road, Kenosha, WI 53141 Wisconsin Geologic and Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 2 The Cambrian-Ordovician aquifer is the primary source of groundwater for high-capacity wells across much of Wisconsin. This prominent groundwater system is over 2000 feet thick in some areas and is impacted by the underlying crystalline Precambrian basement (Leaf et al., 2014), which includes many irregularities, the most prominent of which is the Baraboo Syncline of Columbia and Sauk Counties. This project is a continuation of previous work done in Dodge and Fond du Lac Counties by the University of Wisconsin-Parkside and the Wisconsin Geological & Natural History Survey (MacAlister et al., 2016). The goal of this project is to produce an updated Precambrian topographic map of Southern Wisconsin by using gravity and aeromagnetic data to interpret Precambrian topography away from outcrop and boreholes. This will improve definition of the lower extent of the aquifer, aiding water supply management efforts. Modeling of gravity and aeromagnetic data from the United States Geological Survey (Snyder and Daniels, 2002) was conducted using GM-SYS 3D modeling software in Geosoft Oasis Montaj. Grids of subsurface layers were created from the data and constrained by well and drilling records as well as outcrop maps that were digitized using ArcMAP (Dalziel and Dott, 1970). The Precambrian basement underlying Columbia County is comprised of ca 1.75 Ga granites and rhyolites that are non-conformably overlain by <1.71 Ga quartzite, slate, and ironformation of the Baraboo interval (Medaris et al., 2003). The Baraboo interval metasedimentary rocks and underlying granite and rhyolite was subsequently folded and faulted (e.g. Dalziel and Dott, 1970). The folded layer of iron-formation provides a telltale signatures that aids construction of geophysical models because it has an average magnetic susceptibility of 53000 µcgs, compared to the average susceptibility of the rest of the bedrock of around 1500 µcgs. We use geologic mapping and cross-sections, drill core, magnetic susceptibility and density measurements, petrography and geochemistry, and well construction reports to refine physical modeling constraints. Preliminary results indicate (1) regional Precambrian geology may be interpreted from geophysical data and (2) Precambrian topography is controlled by Precambrian geology and is therefore somewhat predictable. 83 �Figure 1 – Gravity (left) and Aeromagnetic (right) anomaly maps of Columbia County showing the location of 2D models. Dalziel, I. W. D. and Dott, R. H.,1970. Geology of the Baraboo District, Wisconsin: A description and field guide incorporating structural analysis of the Precambrian rocks and sedimentological studies of the Paleozoic strata. Wisconsin Geological and Natural History Survey Information Circular 14. Leaf, A. T., Gotkowitz, M. B., and Dunning, C. P., 2014. A groundwater flow model for Columbia County, Wisconsin. Wisconsin Section of American Water Resources Association Program and Abstracts, Wisconsin Dells, p. 69. Macalister, E. A., Skalbeck, J. D. and Stewart, E. K., 2016. Estimating the subsurface basement topography of Dodge County, Wisconsin using three dimensional modeling of gravity and aeromagnetic data. AGU abstract 184894. Medaris, L.G., Singer, B.S., Dott, R.H., Naymark, A., Johnson, C.M., Schott, R.C., 2003. Late Paleoproterozoic climate, tectonics, and metamorphism in the southern Lake Superior region and Proto-North America: Evidence from Baraboo interval quartzites. Journal of Geology 111, 243257. Snyder, S. L. and Daniels, D. L., 2002. Wisconsin Aeromagnetic and Gravity Maps and Data: A website for distribution of data. USGS Open File Report 02-493. 84 �Pilot study results for potential lithium mineralization on State-managed mineral rights in Minnesota REED, Andrea Minnesota Department of Natural Resources, 1525 3rd Avenue East, Hibbing, MN 55746 The Minnesota Department of Natural Resources (DNR) manages mineral rights on approximately 12 million acres of land. The royalties and rentals generated from these lands help fund Minnesota’s School and University Trusts, the state General Fund, and local governments. To improve the earnings for these entities, the DNR maintains and collects mineral exploration data. The DNR also seeks opportunities to diversify Minnesota’s nonferrous mineral portfolio. With lithium recognized as an element critical for clean energy development (U.S. Department of Energy 2010) and an increasing demand for it, the DNR decided to conduct a pilot study on the potential for lithium occurrences in Minnesota. Of the different lithium deposit types, granitic pegmatites seem to show the most promise of hosting lithium occurrences in Minnesota. Igneous and metamorphic rocks cover a significant portion of the state and are the host rocks for pegmatites (London 2008). Pegmatites host known lithium occurrences in the Quetico Subprovince (the Georgia Lake pegmatites) and Wabigoon Subprovince (Mavis Lake pegmatite group) in Ontario (Selway et al. 2005). A minor amount of lithium is known to occur in the abandoned Rader Mine near Lake of the Woods, on the Minnesota side of the Wabigoon Subprovince (Zamzow & Morey 1991). The pilot study site is located in the Quetico Subprovince on Public School Trust Land northeast of Orr, MN. Little was initially known about the site other than it contained a large pegmatitic granite outcrop. Bulk samples (roughly 4 kg each) of rock were taken from a single pegmatite dike to determine the type of granite and identify the presence of lithium and other trace elements. Small chip samples of feldspar were taken from multiple places along the length of multiple dikes to assess overall fractionation trends. Similar methods are generally accepted for rare element-bearing granitic pegmatite exploration (e.g., Černý 1991, Selway et al. 2005). In addition, nearby glacial till was sampled to see if Laser-Induced Breakdown Spectroscopy (LIBS) could be used to link sand fraction sediments to the pegmatite outcrop. The results for the pegmatite sampling are presented here. Three pale pink monzogranite-pegmatite dikes were identified in the course of fieldwork. The dikes range from 1 to 15 meters in width in outcrop, strike slightly north of east, and have variable apparent dips to the south. Textures range from aplitic to pegmatitic (up to 10 ⨯ 20 cm crystals), with the most common grain size being medium- to coarse-grained granite. Mineralogy is principally composed of feldspars and quartz with trace amounts of magnetite, biotite, and apatite (in order of decreasing abundance). Using the methods and descriptions of Frost et al. (2001), Whalen et al. (1987), London (2008), Černý (1991), whole rock analysis of the bulk samples indicate a mixed A- and I-type signature (tending more towards A-type) and a weakly peralkaline to metaluminous nature, suggesting the sampled dike should be categorized as an NYF granite. Trace element analysis revealed low lithium and REE content, slightly increasing fractionation to the west in the Rb/Sr, Rb/Ba, and La/Yb ratios, confirmation of the non-peraluminous nature of the dike in the Zr/Hf ratio, and confirmation of the A-type signature in the behavior of the REE pattern. Electron microprobe analysis of k-feldspar in collected perthite chip samples reveals that the ratios of Rb/Ba, Rb/Sr, and K/Rb, as well as the distribution of P, (London 2008, London & 85 �Černý 1990) show a strong increasing fractionation trend of this dike set to the northwest. It also shows a slight increasing fractionation trend to the west, confirming the trend pattern seen in the bulk samples. In general, the fractionation trend of these granitic dikes is oriented approximately perpendicular to their strike. Overall, the results suggest that lithium is unlikely to be a significant component in any rocks related to this specific granitic system, even if the identified fractionation trend were to be followed beyond the bounds of the pilot site. References Černý, P. (1991). Rare-element granitic pegmatites. Part II: regional to global environments and petrogenesis. Geoscience Canada, 18(2), pp. 68-81. Frost, B., Barnes, C., Collins, W., Arculus, R., Ellis, D., and Frost, C. (2001). A geochemical classification for granitic rocks. Journal of Petrology, 42(11), pp. 2033-2048. London, D. (2008). Pegmatites. Special publication 10 of The Canadian Mineralogist. Québec, QC: Mineralogical Association of Canada. 347 p. London, D. and Černý, P. (1990). Phosphorus in alkali feldspars of rare-element granitic pegmatites. The Canadian Mineralogist, 78, pp. 771-786. Selway, J., Breaks, F., and Tindle, G. (2005). A review of rare-element (Li-Cs-Ta) pegmatite exploration techniques for the Superior Province, Canada, and large worldwide tantalum deposits. Exploration and Mining Geology, 14(1-4), pp. 1-30 U.S. Department of Energy (2010). 2010 Critical Materials Strategy Summary. U.S. Department of Energy. 4 p. Retrieved from: https://energy.gov/sites/prod/files/edg/news/documents/Critical_Materials_Summary.pdf Whalen, J., Currie, K., and Chappell, B. (1987). A-type granites: geochemical characteristics, discrimination and petrogenesis. Contributions to Mineralogy and Petrology, 95 pp. 407419. Zamzow, C., and Morey, G. (1991). M-074 Reconnaissance geologic map of the Northwest Angle, Lake of the Woods County, Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://conservancy.umn.edu/handle/11299/60043 86 �Variation trends in sulfur isotope ratios at the Eagle and East Eagle intrusions and the surrounding country and basement rocks of the Baraga Basin, Upper Peninsula, Michigan ROSE, Katharine1; ESSIG, Espree2 and THAKURTA, Joyashish1 1 Department of Geological and Environmental Sciences, Western Michigan University, 1903 W. Michigan Ave. Kalamazoo, MI 49008 2 Eagle Mine, Lundin Mining Corporation, 4547 County Road, Champion, MI 49814 The Eagle Ni-Cu sulfide deposit in Marquette County, Michigan is a magmatic sulfide deposit composed of massive, semi-massive and disseminated sulfide minerals hosted in conduit-shaped peridotitic intrusive rocks in the Baraga Basin (Ding et al., 2010). The intrusion has been dated at 1.1 Ga and has been interpreted to be a part of the magmatism associated with the Mesoproterozoic Midcontinent Rift event. Although the ore-grade Ni-Cu sulfide mineralization is located in the sulfide rich part of the intrusive bodies (Figure 1 and 2), relatively small amounts of sulfide minerals are dispersed throughout the intrusion and in the immediate country rocks of the metamorphosed Paleoproterozoic Michigamme Formation and further in the Archean granite-gneiss which forms the basement rock of the Baraga Basin area of UP Michigan (Ding et al., 2012; Hinks, 2016). δ34S ‰ (V-CDT) values have been determined from sulfide minerals of the Eagle and East Eagle intrusions. The distribution trends in the isotope ratio has been studied with respect to spatial directions as well as rock compositions. Peridotitic Rocks Semimassive Sulfides Michigamme Formation Peridotitic Rocks Michigamme Formation Semi-massive and Massive Sulfides Gabbro Massive Sulfides Archean Basement Granite-Gneiss Figure 1: 3-D model of main Eagle intrusion, drill holes EA0300 and EA03301, looking to the east. (Source: Eagle Mine) Figure 2: 3-D model of Eagle East intrusion, looking to the north, with drill hole 17EA364 and 17EA364A. (Source: Eagle Mine) 87 �Based on previous (Hinks, 2016) and present studies, δ34S values of pyrrhotite, chalcopyrite and pentlandite in the massive, semi-massive and disseminated sulfides vary within a range of 0‰ to 5‰. Disseminated pyrite and pyrrhotite in Michigamme Formation slates display δ34S values from 6‰ to 32‰. Disseminated pyrite grains in the Archean basement rocks also display a wide range of δ34S values from -11‰ to 7‰. The distribution of δ34S data in the country and basement rocks, when observed from a directional standpoint along depths of drill-cores show a variation from 10‰ to 2‰ towards the intrusion from the surrounding rocks. However, variations in δ34S values are more distinct from one rock type to another. New data from this study (Table 1) show that within the spatial domain of the sulfide deposit the δ34S changes are more uniform and within a narrower range between 2‰ and 3‰ for the Eagle intrusion. Drill Hole Sample ID δ34S‰ 17EA364 KGR-03-a 32.6 17EA360D KGR-18-a 2.6 17EA360D KGR-18-b 2.5 17EA360D KGR-26-a 6.1 EAUG 0300 KGR-37-a 2.6 EAUG 0300 KGR-37-b 2.5 EAUG 0300 KGR-38-a 2.9 EAUG 0300 KGR-38-b 2.8 EAUG 0300 KGR-39-a 3.1 EAUG 0300 KGR-39-b 3.0 EAUG 0300 KGR-40-a 3.0 EAUG 0300 KGR-41-a 2.9 MF-Michigamme Formation SMSU-Semi-Massive Sulfide Unit BSMT-Archean Basement Granite Unit MF Gabbro Gabbro BSMT SMSU SMSU SMSU SMSU SMSU SMSU SMSU SMSU Table 1: Sulfur isotope ratios determined from sulfide minerals in the Eagle intrusion, the surrounding country, and basement rocks of the Baraga Basin. REFERENCES Ding, X., C. Li, E. M. Ripley, D. Rossell, and S. Kamo, 2010, The Eagle and East Eagle sulfide orebearing mafic ultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution, Geochem. Geophys. Geosyst., 11, Q03003, doi:10.1029/2009GC002546. Ding, X., E.M. Ripley, S.B. Shirey, C. Li (2012), Os, Nd, O and S isotope constraints on country rock contamination in the conduit-related Eagle Cu-Ni-(PGE) deposit, Midcontinent Rift System, Upper Michigan: Geochim. Cosmochim. Acta, 89, pp. 10-30. Hinks, B., 2016, Geochemical and petrological studies on the origin of nickel-copper sulfide mineralization at the Eagle intrusion in Marquette County, Michigan, MS Thesis, Western Michigan University 88 �Preliminary Investigation of the East Eagle Intrusion Gabbro in Marquette County, Michigan RUPP, Kevin1, THAKURTA, Joyashish1, and MAHIN, Robert2 1 Department of Geosciences, Western Michigan University, 1903 W. Michigan Ave. Kalamazoo, MI 49008 2 Eagle Mine, Lundin Mining Corporation, 4547 County Road, Champion, MI 49814 The Eagle deposit is a high-grade, mafic to ultramafic Ni-Cu-bearing sulfide deposit located in Michigan’s Upper Peninsula in Marquette County. The Eagle and East Eagle intrusions are associated with the ~1.1 Ga Midcontinent Rift System and are also associated with the east-west trending Marquette-Baraga dike swarm. Current proven and probable reserves for Eagle are 4.8 million tonnes with an average grade of 2.8% Ni, 2.4% Cu, 0.1% Co, 0.3 gpt Au, 3.4 gpt Ag, 0.7 gpt Pt, and 0.5 gpt Pd (Clow et al., 2017). Recent drilling programs have intersected a vertical gabbroic rock unit in contact with the high-grade mineralization zone of the East Eagle conduit (figure 1). Initial geochemical analysis indicate that the gabbro is depleted in Cu and PGE and becomes more enriched in MgO and FeO with depth. Intrusions depleted in metals often overlie massive sulfide deposits due to the preferential accumulation of metals within sulfide minerals. This is significant in that the gabbroic unit could indicate another massive sulfide deposit at the base of the intrusion. This study attempts to determine the relationship between the gabbroic unit and the known Eagle intrusions based on petrological and geochemical data. Primary objectives of this project are: (1) age determinations using the U-Pb zircon/baddeleyite method, (2) comparison of the gabbroic samples with Eagle and East Eagle based on petrography, whole and trace element geochemistry, and mineral compositions, and (3) comparison of sulfur isotope values with the known values for the Eagle and East Eagle intrusions. The Eagle intrusions were radiometrically age dated and determined to be 1107.3 ± 3.7 Ma (Ding et al, 2010). If these ages are similar, the prospect for sulfide mineralization in a lower staging chamber will be heightened. Previous geochemical studies on the Eagle intrusion show FeO/MgO ratios and the Al 2 O 3 contents of parental magmas to be within the range of picritic basalts erupted during early-stages of the Midcontinent Rift. Whole and trace element geochemical analysis, along with microprobe analysis, will aid in determining the genetic relationship between Eagle and the gabbroic unit. Textural and isotopic characteristics of disseminated sulfides hosted within the gabbro will also be analyzed using reflected light microscopy and a Delta V Mass Spectrometer. Preliminary samples show high degrees of sericitic, propylitic, and carbonate alterations which decrease with depth away from the East Eagle intrusion. Pervasive alteration in many of the samples makes distinguishing individual mineral phases and textures difficult, but primary relict textures (mainly olivine) are seen throughout the samples. Most samples resemble a medium to fine-grained, olivine magnetite gabbro. Plagioclase (50-60%) occurs as subprismatic to lath-like grains that are moderately to strongly altered (up to 60% alteration minerals) to sericite. Olivine (2-8%) are distinguish by subprismatic to subhedral relict grains that altered (90100% alteration minerals) to serpentine and iron-rich oxides. Subprismatic pyroxenes (10-20%) show varying degrees of chlorite alteration to chlorite. Disseminated sulfide mineralization is observed with major sulfide minerals consisting of pyrite, pyrrhotite, chalcopyrite, and pentlandite. Electron microprobe analysis is needed to determine the specific mineral compositions. 89 �Elev (z) -600 Massive sulfides Peridotite gabbro -800 Archean basement -10000 Figure 1: A North-facing 3D-model of the gabbroic unit adjacent to the massive sulfide deposit of East Eagle. Drill core 17EA360 is shown intersecting the gabbroic unit and continuing down through the Archean Basement (image courtesy of Lundin Mining Corporation. Special thanks to Espree Essig and the exploration team) REFERENCES Clow, G. G., Lecuyer, N. L., Rennie, D. W., Scholey, B. J. Y. (2017) NI 43-101 Technical Report on the Eagle Mine, Michigan, USA. Report for Lundin Mining Corporation, dated April 26, 2017, pp. 1-306. Ding, X., C. Li, E. M. Ripley, D. Rossell, and S. Kamo (2010), The Eagle and East Eagle sulfide orebearing maficultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution, Geochem. Geophys. Geosyst., 11, Q03003, doi:10.1029/2009GC002546. 90 �High-technology metal behavior in ore-forming environments and its implication for the Vermilion District, northern Minnesota. SCHARDT, Christian and DAVID, Mady Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 1114 Kirby Dr. Duluth, MN 55812 High-technology metals (HTM), such as In, Ge, Ga, and Tl, are increasingly important for essential industrial applications as well as renewable energy technology. They typically occur in very low concentrations (~ 1 ppm; Terashima 2001) and may reach concentrations of up to 0.15 % in some deposits (e.g., Li et al., 2015). As they do not form ore minerals, they substitute for other metals (Cu, Zn, Sn) in ore mineral such as sphalerite, chalcopyrite, and stannite (Johan, 1988, Pavlova et al, 2015). As a consequence, these metals are sourced as byproducts from other ore deposits (Ishihara and Endo, 2007; Pavlova et al. 2015). While the formation of these ore deposits is relatively well understood, it remains unclear why these metals are restricted to certain ore deposits. This is due to our poor understanding of their general thermodynamic behavior, sourcing, transport, and enrichment mechanisms in selective ore deposition environments. In fact, there is a surprising lack of data regarding the concentration of these metals in various geological settings and their sourcing in ore-forming systems. To better understand the behavior of these metals and gain insight into their enrichment, crustal abundances and concentrations in other sources (seawater, rivers, hydrothermal fluids) were collected along with available data from deposits with confirmed HTM enrichment (volcanogenic massive sulfides, Mississippi Valley Type, Sedimentary-Exhalative, tin granites). In addition, existing geochemical data from the Vermilion district (Peterson, 2001), assumed to host potential massive sulfide mineralization, have been supplemented by available till analysis provided by Larson (2018), and new analysis of various drill holes located within in the Vermilion district. In, Ge, Ga, and Tl show lowest average values in seawater and river waters (< 1 ppb to 0.5 ppm), well below average crustal abundances (up to 15 ppm; see figure 1.). Hydrothermal fluids show concentrations between 0.1 and 10 ppm, close to average crustal averages. HTM values for continental crust (metamorphic, sedimentary, magmatic) show In having the lowest average values (0.1 - 0.5 ppm), increasing to 0.7 ppm for Tl, followed by Ge (1.4 ppm) and Ga (18 ppm; see figure 1). Limited data for the Vermilion district and the Duluth Complex are comparable to trends of other magmatic and volcanic averages, respectively. Results suggest that these HTM are not typically enriched in surface waters (< 1 ppm). Data for hydrothermal fluids show Tl enrichment (up to 20-fold) while In shows average concentrations. Ge and Ga, however, are significantly depleted (Ge: 5-fold; Ga: 2-fold) compared to average crustal abundances. This poses the question where and how HTM get enriched to values recorded in some sulfide deposits (≥ 1000 ppm) and why this process seems to be restricted to certain geological environments. To study this issue, geochemical data from HTM-bearing ore deposits have been analyzed to determine if differences in the substitution behavior of HMT exist between different ore deposit types. Initial results indicate differences in the substitution behavior of HTM in host rocks (inset figure 1) as well as ore minerals (not shown), which may point to a) variable HTM sourcing, b) mineralization conditions, and/or c) different hydrothermal fluid chemistries as a function of formation environment. Further analysis is underway to determine if this also applies to the Vermilion district and its potential to host significant HTM concentrations. 91 �Figure 1 Minimum, average, and maximum concentrations of In, Tl, Ge, and Ga in crustal rocks and fluids. The Vermilion district and the Duluth Complex show patterns similar to other volcanic and magmatic rock data (not shown). Inset: Cu-Ga-Zn ternary plot for whole-rock data from major HTM-bearing deposits (VMS - volcanogenic massive sulfides). Differences exist between mafic/felsic volcanic, plutonic (tin deposits), and sedimentary settings (siliciclastic). Similar trends are also observed for other element combinations, including non HTM elements. References Ishihara, S., and Endo, Y. (2007) Indium and other trace elements in volcanogenic massive sulfide ores from the Kuroko, Besshi and other types in Japan. Bulletin of the Geological Survey of Japan, v.58, p. 7 - 22 Johan, Z, 1988, Indium and Germanium in the Structure of Sphalerite: an Example of Coupled Substitution with Copper. Mineralogy and Petrology, v. 39, p.211 - 229 Larson, P., 2018, personal. communication Li, Y., Tao, Y., Feilin, Z., Mingyang, L., Feg, X., and Xianze, D., 2015, Distribution and existing state of indium in the Gejiu Tin polymetallic deposit, Yunnan Province, SW China. Chinese Journal of Geochemistry, v. 34, p. 469 - 483 Pavlova, G.G., Palessky, S.V., Borisenko, A.S., Vladimirov, A.G., Seifert, T., and Phane, L.A. (2015) Indium in cassiterite and ores of tin deposits. Ore Geology Reviews, v. 66, p. 99–113 Peterson, D.M., 2001, Development of Archean Lode-Gold and Massive Sulfide Deposit Exploration Models using Geographic Information System Applications: Targeting Mineral Exploration in Northeastern Minnesota from Analysis of Analog Canadian Mining Camps; University of Minnesota Ph.D. thesis, 503 pages, 12 plates, 1 CD-ROM. Terashima, S. (2001) Determination of Indium and Tellurium in Fifty Nine Geological Reference Materials by Solvent Extraction and Graphite Furnace Atomic Absorption Spectrometry. Geostandards Newsletter, v. 25, p. 127 - 132 92 �Geochemistry of mafic rocks in Dickinson County, Michigan: Evidence for ~2.1 Ga Rifting SCHULZ, K.J.1, CANNON, W.F.1, and WOODRUFF, L.G.2, 1 U.S. Geological Survey, Reston, VA 20192, 2 U.S. Geological Survey, Mounds View, MN 55112 Mafic rocks of purported Archean and Paleoproterozoic age are a significant and widespread component of the bedrock geology in Dickinson County, Michigan (James et al., 1961). For this study we have sampled mafic rocks that occur in the Carney Lake Gneiss and other Archean gneisses of the county as well as mafic rocks in the Dickinson Group and the Hardwood Gneiss. Field relations of the mafic rocks in the Archean gneisses are often ambiguous; some are clearly dikes but whether they are Archean or Paleoproterozoic in age is often uncertain. Mafic rocks in the Carney Lake Gneiss and other Archean gneisses in the region range from highly deformed amphibolite inclusions in granitic gneiss to less deformed “salt and pepper” amphibolites to metadiabase dikes with no penetrative fabric and lower metamorphic grade. We have analyzed ten samples of the “salt and pepper” amphibolites and found two basaltic compositional types. Group 1 samples, two of which are from identified dikes, have relatively low MgO (~4 to 6 wt. %), moderately fractionated incompatible trace element patterns, and negative Nb-Ta anomalies on a primitive mantle normalized (PMn) trace element plot (Figure 1A). In contrast, Group 2 samples, for which field relations are ambiguous, have higher MgO (~6 to 10 wt. %), lower trace element contents than the first group, and distinctive flat PMn trace element patterns (Figure 1A). The Dickinson Group is composed of the basal East Branch Arkose overlain by the Solberg Schist and Six Mile Lake Amphibolite (James et al., 1961). Two samples of amphibolite collected along strike in the East Branch Arkose have very similar tholeiitic basalt compositions characterized by moderately enriched light REE and no Nb-Ta anomalies on a PMn trace element plot (Figure 1B). A sample of a metadiabase dike cutting Archean granitic gneiss north of the East Branch Arkose sample location and an amphibolite from a road cut to the south near Felch are similar in composition except for a positive Th anomaly when normalized to primitive mantle, which is likely the result of crustal contamination. Samples of mafic rocks from the Solberg Schist range from basalt to andesite (~45 to 56 wt. % SiO2; ~4 to 12 wt. % MgO), are more enriched in light REE than the amphibolite in the East Branch Arkose, and have negative Nb-Ta anomalies on a PMn trace element plot (Figure 1B). Samples of the Six Mile Lake Amphibolite, in contrast to the amphibolites in the East Branch Arkose and Solberg Schist, have much lower trace element contents and flat PMn trace element patterns much like the Group 2 amphibolites sampled in the Carney Lake Gneiss (Figure 1B). In addition, a large metagabbro body and a dike sampled in the Solberg Schist are similar in composition to the Six Mile Lake Amphibolite. This supports the interpretation that the Six Mile Lake Amphibolite is the upper, youngest part of the Dickinson Group (James et al., 1961). Samples of mafic gneiss in the Hardwood Gneiss complex are generally similar in composition to the Six Mile Lake Amphibolite with similar low trace element contents and 93 �relatively flat PMn trace element patterns (Figure 1C). One mafic gneiss sample is enriched in light REE and has a large negative Nb-Ta anomaly that is likely the result of contamination by felsic crustal rocks. Three Paleoproterozoic dike swarms, the Marathon, Kapuskasing, and Fort Frances, which outcrop around the northern margin of Lake Superior and range in age from 2126 to 2067 Ma, are attributed to a long-lived mantle plume event that accompanied rifting along the southern margin of the Superior craton (Halls et al., 2008). Like the mafic rocks in the Dickinson Group, the older dikes (Marathon and Kapuskasing) show enriched and fractionated incompatible trace element patterns while the youngest (Fort Frances) are relatively depleted and have flat trace element patterns. The overlap in composition of the mafic rocks sampled in Dickinson County with the Paleoproterozoic dikes on the north side of Lake Superior suggests the Dickinson County mafic rocks also may be related to the final rifting of the Superior and Wyoming cratons. This is supported by the presence of 2.1 Ga detrital zircons in the East Branch Arkose (Craddock et al., 2013). Figure 1. Primitive mantle normalized trace element patterns for mafic rocks from Dickinson County. References Craddock, J.P., Rainbird, R.H., Davis, W.J., Davidson, Cam, Vervoort, J.D., Konstantinou, Alexandros, Boerboom, Terry, Vorhies, Sarah, Kerber, Laura, and Lundquist, Becky, 2013, Detrital zircon geochronology and provenance of the Paleoproterozoic Huron (~2.4–2.2 Ga) and Animikie (~2.2–1.8 Ga) basins, southern Superior Province: Journal of Geology, v. 121, p. 623–644. Halls, H.C., Davis, D.W., Stott, G.M., Ernst, R.E., and Hamilton, M.A., 2008, The Paleoproterozoic Marathon large igneous province: New evidence for a 2.1 Ga long-lived mantle plume event along the southern margin of the North American Superior Province: Precambrian Research, v. 162, p. 327–353. James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., 1961, Geology of Central Dickinson County, Michigan: U.S. Geological Survey Professional Paper 310, 176 p. 94 �Detrital Zircons in the Waterloo Quartzite, Wisconsin: Implications for the Ages of Deposition and Folding of Supermature Quartzites in the Southern Lake Superior Region SCHWARTZ, Joshua J.1, STEWART, Esther K.2, and MEDARIS, L. Gordon Jr.3 1 Geological Science, California State University, Northridge, California 91330 Wisconsin Geological and Natural History Survey, Madison, Wisconsin 53705 3 Department of Geoscience, University of Wisconsin–Madison, Madison, Wisconsin 53706 2 Proterozoic supermature quartzites of the Baraboo Interval are a prominent and significant Precambrian feature of the southern Lake Superior region, covering an area of ~175,000 km2. The Waterloo Quartzite in SE Wisconsin has long been correlated with other quartzites of the Baraboo Interval, based on similarities in sedimentary characteristics, geological setting, and chemical composition. Metapelites in the Baraboo and Waterloo sequences are among the most chemically mature sedimentary rocks in the geological record, having Chemical Indices of Alteration of 99.6 and 97.6, respectively. Although similar to the Baraboo Quartzite in many respects, the Waterloo Quartzite differs in having experienced pervasive K–metasomatism (Table 1). In addition, axial-planar cleavage is more strongly developed in quartzite at Waterloo compared to Baraboo, and Waterloo exhibits a higher grade of metamorphism, with Waterloo metapelite containing andalusite (amphibolite facies) and Baraboo metapelite containing pyrophyllite (greenschist facies). Seven samples of Waterloo quartzite and pebbly quartzite in Dodge county were collected for detrital zircon analysis to evaluate sources of sediment and to determine possible relationships to other supermature quartzites in the region. Samples include five quarried blocks in the Michels Materials Waterloo Quarry and individual samples from outcrops at Hubbleton and Mud Lake. Bedding orientations overlain on an aeromagnetic anomaly map of the area suggest that quartzite at the Michels quarry may be in a lower stratigraphic position than those at the Hubbleton and Mud Lake localities. A relative probability plot for detrital zircons (filtered for dates <10% discordant) in the Waterloo Quartzite at the Michels quarry displays a strong geon 16 (Mazatzal) and geon 17 (Yavapai) signal, and diminished geon 18 (Penokean) and geon 25–27 (Algoman) signals (Fig. 1A). Maximum Ages of Deposition calculated from the youngest statistically homogenous population (MSWD ≤ 1.0) are 1759±13 (n=21), 1694±11 (n=36), 1671±14 (n=21), 1669±14 (n=21), and 1643±11 Ma (n=42). In contrast, Waterloo quartzites at Hubbleton and Mud Lake are characterized by an absence of geon 16 zircons, a pronounced Penokean population, and a subdued, but distinct Algoman population (Fig. 1A). Quartzites in the Baraboo Range, Sauk County, consist of a stratigraphically lower fluvial facies and an upper shoreface marine facies, whose detrital zircon populations differ from each other (Fig. 1B). Fluvial quartzites display a predominant geon 17–18 signal, and shoreface marine quartzites contain a pronounced geon 18–19 population and a distinct geon 25-27 population. 95 �The Algoman, Penokean, and post– Penokean (Yavapai) populations of detrital zircons in the Baraboo quartzites are consistent with derivation from the proximal post– Penokean Montello Batholith and more distal, northerly Penokean and Archean basement; postPenokean zircons are more abundant in the stratigraphically lower fluvial facies than in the higher shoreface marine facies, which was deposited after burial of the 1750 Ma Montello Batholith. Detrital zircons in Waterloo quartzites at Hubbleton and Mud Lake were also derived from the proximal Montello Batholith and more distal, northerly Penokean and Algoman terranes. In contrast, Waterloo Quartzite at the Michels quarry is characterized by a relative abundance of geon 16 zircons and a pronounced population peak at 1700 Ma. Clearly, the provenance of this quartzite was different than that of the other analyzed quartzites. Geon 16 juvenile crust is absent in the southern Lake Superior region, but occurs in the subsurface south of Wisconsin as part of the transcontinental 1.68–1.60 Mazatzal belt (Whitmeyer and Karlstrom, 2007). Thus geon 17 and geon 16 zircons in quartzite at the Michels quarry were likely derived by northerly transport from the proximal Yavapai terrane and more distal Mazatzal terrane to the south. The shift in detrital zircon age populations between the Michels quarry and Hubbleton and Mud Lake outcrops reflects a change in transport direction from north to south. Deposition of quartzite at the Michels quarry is bracketed between ca. 1643 Ma, its youngest maximum age of deposition, and 1452 Ma, the 40Ar/39Ar cooling age of muscovite in folded metapelite (Medaris et al., 2003). Because the Michels quarry lies stratigraphically below the Hubbleton and Mud Lake outcrops, the depositional age for quartzites at these localities must also be no older than ca. 1640 Ma. The extreme chemical maturity of Waterloo metapelite requires derivation from a region of subdued relief that experienced intense chemical weathering. Such chemical maturity, combined with geon 16 Maximum Ages of Deposition for quartzite at the Michels quarry, is consistent with deposition of the Waterloo Quartzite after the 1630 Ma Mazatzal Orogeny, with depositional space for the thick (~1000 m) quartzite sequence being provided by post-Mazatzal rifting. If this scenario is correct, it requires that deformation and folding of the Waterloo Quartzite occurred during the geon 14 Wolf River tectonomagmatic event. References Medaris et al., 2003, Journal of Geology, v. 111, p. 243–277. Van Wyck and Norman, 2004, Journal of Geology, v. 112, 305-315. Whitmeyer and Karlstrom, 2007, Geosphere, v. 3, 220-259. 96 �Compositional and geochemical characteristics of the Crystal Lake intrusion, Ontario SMITH, Jennifer1, BLEEKER, Wouter1, ROSSELL, Dean2 and LABERGE, Justin2 1 2 Geological Survey of Canada, 601 Booth Street, Ottawa, Canada; email:jennifer.smith6@canada.ca Rio Tinto Exploration Canada Inc. 1300 Walsh Street, Thunder Bay, Canada The 1.1 Ga failed rift system hosts a range of mafic-ultramafic, carbonatitic and alkaline intrusions (Bleeker et al., 2018), many of which are actively being explored for a range of commodities (e.g., Ni, Cu, PGE, Co, Cr, V, Nb). The discovery of the high grade, massive sulphide, Ni-Cu Eagle deposit in 2002, has resulted in a surge of exploration activity and interest in the Ni-Cu-PGE potential of the MCR. Early rift (1117 to 1106 Ma) conduit-type, ultramafic intrusions (e.g., Tamarack, Eagle), remain the most attractive but challenging exploration targets (Heaman et al., 2007). The 1099±1 Ma Duluth Complex and similar large, sheet-like intrusions (e.g., Sonju Lake, Mellen Complex, Echo Lake, Crystal Lake, Coldwell Complex) still remain prospective, although typically contain lower metal tenors (Ripley, 2014). The 1099.1±1.2 Ma (Heaman et al., 2007) Crystal Lake layered intrusion, located 47 km southwest of Thunder Bay, contains low-grade Ni-Cu-PGE sulphide mineralisation and uneconomical chromite occurrences (Geul, 1970; Smith & Sutcliffe, 1987). Although mineralisation was first discovered in the 1950s and has been extensively explored since, the intrusion remains a prospective exploration target with Rio Tinto undertaking more recent drill programs in 2014-15 (Figure 1). This intrusion outcrops as a prominent Y-shaped body, intruding S-bearing shales, argillites and greywackes of the Paleoproterozoic Rove Formation. Geochemically, the Crystal Lake Figure 1. Crystal Lake intrusion and location of Rio Tinto’s 2014intrusion can be distinguished from the more 2015 boreholes. Adapted from Geul 1970. primitive conduit-type bodies by: olivine composition (Fo 51-79 ), low Ni/Cu and Pt/Pd ratios (<1), higher REE abundances, LREE enrichment and minimal fractionation of HREEs (Gd/Yb <2; Thomas, 2015). Previous work divided the intrusion into four discrete zones (Smith & Sutcliffe, 1987). The Basal Zone contains an aphanitic chill zone, with inclusions of S-bearing Rove sedimentary rocks. The overlying Lower Zone is characterised by medium to pegmatitic, vari-textured gabbro with irregular, coarse segregations of Cr-bearing leucogabbro and anorthosite. The Middle Zone marks the beginning of phase layering and comprises four magmatic cycles. Each cycle corresponds to an influx of magma (Cogulu, 1993a) and consists of a basal Cr-spinel bearing troctolite/olivine gabbro and an upper anorthositic gabbro. The Cr-spinel occurs in discrete layers and is recognised within orthocumulate and adcumulate rocks where it constitutes 8 to 36 modal%, respectively. Compositional differences in Cr-spinel occurrences have been attributed to the effects of in-situ re-equilibration (Cogulu, 1993a). The Upper Zone is marked by the disappearance of Cr-spinel and anorthositic layers. This unit consists of coarse-grained olivine gabbro and medium-grained troctolite. Low-grade, Ni-Cu-PGE sulphide mineralisation is developed throughout the Lower and Middle Zones, with the Upper Zone barren of sulphides. The Lower Zone is characterised by disseminated to massive sulphides which are mainly concentrated towards the basal contact and within late pegmatitic zones. The association of sulphides with pegmatitic 97 �phases is not unique to the Crystal Lake intrusion, also being recognised in the Coldwell Complex and other world-class deposits (e.g., Merensky Reef). Cogulu (1993b) noted that pyrrhotite dominates the basal assemblages. Middle Zone sulphides are disseminated and closely associated with volatiles. Here, assemblages are Cu-rich with lesser proportions of pentlandite and pyrrhotite (Cogulu, 1993b). From Rio Tinto’s 2014-15 dill holes the following observations are made. The Lower and Middle Zones are characterised by low Pt/Pd (<0.3), Ni/Cu (<1) and mantle-like Cu/Pd values (103-104). Whilst the Ni/Cu increases into the Middle Zone, the Cu/Pd ratio decreases along with incompatible element concentrations. The Upper Zone is more homogeneous with higher Pt/Pd (0.5-1), Ni/Cu (often >1), and higher than mantle Cu/Pd ratios (>104). Ni/Cu decreases through the Upper Zone whilst Pt/Pd, Cu/Pd, and incompatible elements increase. The implications and cause of these geochemical trends has yet to be fully constrained. The addition of crustal S is considered critical in the genesis of many of the MCR Ni-Cu sulphide deposits (Ripley, 2014). Preliminary δ34S data indicate a strong crustal component throughout the Crystal Lake intrusion with δ34S ranging from 1.4-16.5‰ (Thomas, 2015). The majority of data resides outside the mantle range of 0±2‰. Thomas (2015) argues that the intrusion was emplaced as a series of S-saturated magma pulses, with δ34S variability attributed to contamination by different S-bearing horizons. S/Se ratios however, are more consistent with in-situ contamination with a footwall influence evident in the Lower Zone. The Middle and Upper Zones exhibit lower than mantle S/Se ratios, showing no evidence of a crustal control, which various processes may have masked (e.g., S-loss, upgrading, increased R-factor). To date, no proposed model accounts for all of these features. The Crystal Lake intrusion remains an interesting deposit. Whilst the mineralisation shows many parallels to those observed at the Duluth and Coldwell Complexes, various questions remain regarding the source characteristics and range of magmatic processes involved in their development. References Bleeker, W., Liikane, D.A., Smith, J., et al. 2018. Activity NC-1.3: Controls on the localisation and timing of mineralised intrusions in intra-continental rift systems, with a specific focus on the ca. 1.1 Ga Midcontinent Rift (MCR) system. Geological Survey of Canada, Open File 8373, 15-27. Cogulu, E.H., 1993a. Factors controlling postcumulus compositional changes of chrome spinels in the Crystal Lake intrusion, Thunder Bay, Ontario. Geological Survey of Canada, Open File 2748. Cogulu, E.H., 1993b. Mineralogy and chemical variations of sulphides from the Crystal Lake Intrusion, Thunder Bay, Ontario. Geological Survey of Canada, Open File 2749. Geul, J.C., 1970. Geology of Devon and Pardee Townships and the Stuart Location. Ontario Department of Mines, Geological Report 87. Heaman, L.M., Easton, R.M., et al., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences, 44(8), 1055-1086. Ripley, E.M., (2014). Ni-Cu-PGE mineralisation in the Partridge River, South Kawishiwi, and Eagle intrusions: A review of contrasting styles of sulphide-rich occurrences in the Midcontinent rift system. Economic Geology, 109(2), 309-324. Smith, A.R., & Sutcliffe, R.H., 1987. Keweenawan intrusive rocks of the Thunder Bay area. Ontario Geological Survey Miscellaneous paper 137. Thériault, R.D., Barnes, S-J., & Severson, M.J., 1997. The influence of country rock assimilation and silicate to sulphide ratios on the genesis of the Dunka Road Cu-Ni-PGE deposit, Duluth Complex. Canadian Journal of Earth Science, 34, 375-389. Thomas, B., 2015. Geochemistry, sulphur isotopes and petrography of the Cu-Ni-PGE mineralised Crystal Lake Intrusion, Thunder Bay, Ontario. M.Sc. Thesis. 98 �Petrology and 11B Composition of Tourmaline within the 2685 Ma Ghost Lake Batholith and Mavis Lake Pegmatites SMITH, Vittoria and ZUREVINSKI, Shannon Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada The Ghost Lake Batholith (GLB) and derived Mavis Lake Pegmatite group are an example of a granite-pegmatite system in which both the parent granite and least- to- most evolved pegmatites are visible and accessible. The GLB has been divided into eight internal units based on mineralogy and texture, while the Mavis Lake Pegmatite group is divided into three broad zones based on the mineralogy of the pegmatite bodies (Breaks and Moore, 1992). Samples collected from the biotite granite phase (GLB-3) of the Ghost Lake Batholith have the mineral assemblage typical of an S-type peraluminous granite. The granite mineralogy is made up of mostly quartz, albite, potassium feldspar and biotite with accessory muscovite, garnet, zircon, and blue apatite. Pegmatitic segregations within the parent granite consist of potassium feldspar, quartz, and biotite, and accessory garnet. Apatite within this unit is typically associated with or included directly within the biotite and grains analyzed via SEM have been found to be LREE-enriched. Pegmatite bodies within the beryl-columbite zone of the Mavis Lake Pegmatite group are hosted within mafic metavolcanic rocks and show considerable variation in mineralogy and texture. Pegmatites range from potassic to albitic, and garnet is occurring as an accessory phase. Tourmaline is common in various pegmatitic units within the beryl-columbite and spodumene-beryl-tantalite zone. Due to its commonality within the Mavis Lake group, tourmaline has been sampled and studied using major and trace element techniques to assess its usefulness as an indicator of fractionation between and within individual pegmatite bodies. Tourmaline core compositions within the pegmatite throughout the Mavis Lake Group range between schorlitic in composition to dravitic (Fig.1). In the case of the Taylor emerald occurrence, the tourmaline species within the pegmatite range from dravitic in the border zone, to schorlitic within the pegmatite body, reflecting a decrease in Fe and an increase in Mg from rim to core. Tourmaline zoning profiles from tourmaline within the Taylor emerald occurrence show significant substitution between Fe and Mg in the Y- site and an inverse relationship between increasing Na contents and decreasing vacancies within the crystals X-site (Fig.2). Similarly, tourmaline from a nearby potassic pegmatite show similar progression in decreasing Na from core to rim with an increasing amount of vacancies. In contrast, the Fe contents in the Y-site steadily increase from core to rim, and Mg shows a closely inverse reaction to the Al contents, suggesting a proton-loss substitution. Boron isotope data collected in situ via SIMS report δ11B values from pegmatites within the contact beryl zone between 8.1‰ to 13.9‰. Variable mineralogy and major- and trace-element mineral chemistry within the Beryl-columbite zone suggest that the degree of host-rock interaction highly influences the tourmaline chemistry. This supports previous work by Breaks and Moore (1992), who had previously suggested that the high Mg contents within the Taylor pegmatites may be related to metasomatic transfer with the metavolcanic host units. While tourmaline is widely considered a petrogenetic indicator for the degree of fractionation within pegmatitic systems, this concept does not seem to apply to a system like Mavis Lake where tourmaline is restricted to the border zones of the pegmatites and is highly influenced by host-rock interaction. 99 �Figure1. Tourmaline speciation diagram for tourmaline within the Mavis Lake Group following the classification scheme of Henry et al., 2005. Fig. 2: Tourmaline zoning profiles for the X and Y sites within a euhedral crystal of tourmaline in potassic pegmatite. REFERENCES: BREAKS, F.W. AND MOORE, J.M., 1992. "The Ghost Lake batholith, Superior Province of northwestern Ontario: a fertile, S-type, peraluminous granite-rare-element pegmatite system." Canadian Mineralogist v. 30, p.835-835. 100 �Geophysical, structural, and tectonic interpretation of the Yellow Medicine and Appleton shear zones, SW Minnesota and SE South Dakota: A work in progress SOUTHWICK, David, CHANDLER, Val, and JIRSA, Mark Minnesota Geological Survey, University of Minnesota, 2609 Territorial Rd, St. Paul, MN 55114 U.S.A. The Yellow Medicine and Appleton shear zones (YMSZ and ASZ) are prominent geophysical features of the Minnesota River Valley (MRV) subprovince of the Superior craton. Maps of the first vertical derivative of the magnetic anomaly and the second vertical derivative of the gravity anomaly show that the two zones converge into a single strand in east-central South Dakota, and that the combined fault strand continues west-southwest as least as far as the east margin of the Paleoproterozoic Trans-Hudson orogen. Faulting in the YMSZ and ASZ is thought to have begun in the Sacred Heart accretionary event (ca. 2600 Ma) in which the MRV subprovince was amalgamated to the south margin of the Superior craton. Fault motion may have peaked during Yavapai tectonism, between ca. 1785 and 1775 Ma, in concert with a major episode of granitic magmatism and orogenic uplift. The Minnesota segment of the YMSZ consists of an axial zone where there are multiple anastomosing sub-zones of concentrated fault damage and km-wide flanking zones of dispersed fault damage. The axial zone is well defined geophysically; the zones of dispersed fault damage are not. Drill cores reveal that the axial zone contains heterolithic crush breccia, fine crush breccia, crush microbreccia, protocataclasite, cataclasite, and protomylonite that were derived from identifiable quartzofeldspathic orthogneiss, foliated garnet-quartz-hornblende paragneiss, amphibolite, and plagiogranite. Graphite-rich fault rocks encountered in boreholes toward the east end of the YMSZ, near the west-northwest- verging tectonic front of the Penokean orogen, may be tectonically dismembered slices of Penokean metasedimentary rocks caught up in Yavapai faulting. Pseudotachylyte is relatively abundant in the axial zone of the YMSZ and in the narrow faults in the flanking zones of dispersed fault damage (Craddock and Magloughlin, 2005). Diabase dikes of the Kenora-Kabetogama/Fort Frances swarm (ca. 2070 Ma) are offset by and/or terminated against the YMSZ and the ASZ, whereas hornblende andesite and ferrodiorite dikes that cut the 1792 Ma and younger intrusions of the composite East-Central Minnesota Batholith (ECMB) transect the YMSZ and ASZ without deviation. A U-Pb zircon age of ca. 1780 Ma inferred for one of the hornblende andesite dikes (Schmitz et al., 2018, in prep.) limits the timeframe of geophysically discernable fault motion to the period between 2070 Ma (pre-Penokean) and 1780 Ma (mid-Yavapai). Geophysical patterns suggest that a considerable component of fault displacement on the YMSZ system was left-lateral strike slip that on a regional scale shifted the Morton block, south of the YMSZ, eastward relative to the Montevideo block on the north. We speculate that this displaced a NNW-trending piece of the southern Trans-Hudson orogen eastward from central South Dakota into far WSW Minnesota, where NNW geophysical trends are evident and as yet unexplained. This regional interpretation is based on potential-field images that were upward-continued to five km in order to even out resolution differences among the various data sets that were compiled in the source magnetic and gravity maps of North America (North American Magnetic Anomaly Group (NAMAG), 2002; Committee for the Gravity Anomaly Map of North America, 1988). Our interpretation in eastern South Dakota is submitted as an alternative to an earlier interpretation presented by McCormick (2010a, b) that was based on the original unleveled magnetic and gravity maps. 101 �Vertical displacement on the YMSZ, up on the north, is indirectly inferred from Yavapai K-Ar ages from rocks in the Montevideo block and the absence of K-Ar ages younger than late Neoarchean in rocks in the Morton belt (Goldich et al., 1961). These observations suggest that K-Ar systematics were reset in Montevideo rocks that were hotter and deeper in the crust prior to Yavapai convergence and associated uplift above the north-dipping YMSZ, whereas the Morton rocks remained higher in the crust and relatively cool. The possibility of mafic underplating having had a role in Montevideo reheating and uplift in mid-geon 17 is suggested by the observed higher-gravity signature of the Montevideo block, particularly toward its east end. REFERENCES Committee for the Gravity Anomaly Map of North America, 1988, Gravity anomaly map of North America: Geological Society of America, 5 sheets, scale 1:5,000,000. Craddock, J.P., and Magloughlin, J.F., 2005, Calcite strains, kinematic indicators, and magnetic flow fabric of a Proterozoic pseudotachylyte swarm, Minnesota River valley, USA: Tectonophysics, v. 402, p. 153-168. Goldich, S.S., Nier, A.O., Baadsgaard, H., Hoffman, J.H., and Krueger, H.W., 1961, The Precambrian geology and geochronology of Minnesota: Minnesota Geological Survey, Minneapolis, Minnesota, Bulletin 41, 193 p. McCormick, K.A., 2010a, Precambrian basement terrane of South Dakota: South Dakota Geological Survey Program Bulletin 41, 37p. McCormick, K.A., 2010b, Plate 1: Terrane map of the Precambrian basement of South Dakota: South Dakota Geological Survey Program Bulletin 41, External pdf file, compilation scale 1:1,000,000. North American Magnetic Anomaly Group (NAMAG), 2002, Magnetic anomaly map of North America: U. S. Geological Survey Open File Report OFR 02-414 (On line only) (http://pubs.usgs.gov. /of/2002/of02-414/) Schmitz, M.D., Southwick, D.L., Bickford, M.E., Mueller, P.A., and Samson, S.D., 2018, in prep., Neoarchean and Paleoproterozoic events in the Minnesota River Valley subprovince, with implications for southern Superior craton evolution and correlation: Submitted to Precambrian Research March 2018. 102 �New bedrock geologic mapping of Dodge County, Wisconsin provides evidence for Paleozoic reactivation of Precambrian structures KINGSBURY STEWART, Esther, STEWART, Eric D., and ROUSHAR, Kathy Wisconsin Geologic and Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 We present a preliminary 1:100,000-scale bedrock geologic map of Dodge County, Wisconsin (fig. 1). Bedrock is mostly buried beneath 6 to 18 meters (20 to 60 feet) of glacial deposits that locally exceed 60 meters (200 feet) within bedrock channels in the eastern part of the county. Due to the significant glacial deposits, mapping is based on integration of geophysical data (gravity and aeromagnetic anomaly data, passive seismic readings), subsurface data (drill core, downhole geophysical logs, geologic logs based on cuttings from municipal wells, and well construction reports of private water wells), and observations collected at sparse outcrops and quarries. To produce the map, we interpolated a bedrock elevation surface from the top of bedrock contact recorded in 3,831 geolocated wells. Depth-structure maps of the base of Paleozoic map units and the top of the Precambrian surface were gridded from map unit contacts picked in 882 wells. The intersection of the depth-structure maps and the bedrock elevation surface defines map unit contacts. The bedrock geology is comprised of a Precambrian bedrock surface characterized by regional-scale folding and topographic relief overlain by upper Cambrian siliciclastics and Ordovician through Silurian dolostone and siliciclastics. The Paleozoic section thickens from west to east towards the Michigan Basin such that western Dodge County is underlain by the Cambrian through Middle Ordovician sandstone and dolostone while eastern Dodge County is underlain by Silurian dolostone of the Niagara Escarpment. Results from the first three years of this four-year effort clarify the stratigraphy and structure of the Precambrian units as well as the influence of Precambrian structure on deposition of the overlying Paleozoic sediments. The Precambrian rocks include folded metasediments of the Baraboo interval (<1.7 Ga) that were intruded by ca. 1.4 Ga granite (Medaris et al., 2011). A bedrock core drilled as part of the mapping effort encountered a likely altered banded ironformation that is known to be present ~40 miles (64km) to the northwest within the Baraboo interval stratigraphy. We tie this core to a characteristic, curvilinear aeromagnetic anomaly and extrapolate to calibrate the regional aeromagnetic data in Dodge County and thus map the distribution of Precambrian units. Map patterns of the Precambrian surface demonstrate that the Baraboo-interval metasediments were folded into east-northeast-trending, doubly-plunging anticlines and synclines with ~30km (18.6 mile) wavelength. Map patterns further demonstrate that the Waterloo quartzite, which outcrops in a broad syncline in southwestern Dodge County, is distinct from, and likely stratigraphically above, the Baraboo quartzite. Precambrian topography was mostly infilled by Cambrian sandstone such that the thickness of the Cambrian Elk Mound Group sandstone can vary by >82 meters (270 feet) over several miles while the thickness of the overlying Cambrian Tunnel City and Trempealeau Groups are relatively consistent. The Paleozoic units were then folded into broad, east-west trending, gentle anticlines and synclines with lengths of 13.6 km (8.5 miles) to 40 km (25 miles), widths of about 8 to 10.5 km (5 to 6.5 miles), and amplitudes of 20 to 100 meters (65 to 328 feet). Data from well cuttings and drill core suggest faulting locally uplifted the Precambrian basement through early Ordovician Prairie du Chien Group. The overlying Middle Ordovician Ancell Group unconformably overlies the Prairie du Chien Group. The overlying Sinnipee Group is gently folded with no clear evidence for fault offset. Sulfide mineralization is present throughout the Paleozoic section in Dodge 103 �County and is preferentially located along faults near fold axes (Brown and Maas, 1992, this study). Fold geometry and preferential sulfide mineralization along fold limbs observed in Dodge County is similar to fold geometry and mineralization reported by Heyl et al. (1959) for the Upper Mississippi Valley Lead-Zinc District, suggesting similar controls on deformation and mineralization for southwestern and southeastern Wisconsin. Figure 1. Generalized 1:1,000,000-scale bedrock geologic map of Dodge County showing data sources for 1:100,000-scale mapping. Inset map locates Dodge County (blue) in Wisconsin. Modified from Mudrey et al. (1982). References Brown, B. A. and R.S. Maas. 1992. A reconnaissance survey of wells in eastern Wisconsin for indications of Mississippi Valley Type Mineralization: Wisconsin Geological and Natural History Survey Open File Report 92-3, 31p. Heyl A. Jr., A.R. Agnew, E.J. Lyons, and C.H. Behre Jr. 1959. The geology of the Upper Mississippi Valley ZincLead District: US Geological Survey Professional Paper 309, 310p. Medaris, L.G., Jr., R.H. Dott, Jr., J.P. Craddock, and S. Marshak. 2011. The Baraboo District- A North American classic in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-Continent of North America: Geological Society of America Field Guide 24, p. 63-82. Mudrey, M.G., Jr., Brown, B.A., and Greenberg, J.K. 1982. Bedrock geologic map of Wisconsin: Wisconsin Geological and Natural History Survey State Map 18, scale: 1:1,000,000. 104 �Neoarchean to Paleoproterozoic reconstructions using metamorphic core complexes as evidence of continental transform plate motion and their implications in Archean tectonics STINSON, V.R. and PAN, Y. Department of Geological Sciences, University of Saskatchewan, 114 Science Place, Saskatoon, SK S7N 5E2 Canada Paleotectonic reconstructions use geological, geophysical, and paleontological information to piece together cratons, continents, and supercontinents. Metamorphic core complexes commonly have transform, transpression, and transtension components which are underrepresented in paleotectonic reconstructions and may also be used to provide evidence of subduction, collision, and exhumation. Due to the nature of the medium to high-grade metamorphic and felsic plutonic igneous lithologies in the footwall they are typically well-preserved and yield robust minerals used in geochronology. In this study we have combined literature review and field mapping, geochronology, and petrology to investigate the potential for reconstructing Archean cratons in transpressive to trantensional tectonic settings in the Neoarchean. This study recommends the use of metamorphic core complexes as evidence for transpressionaal to transtensional plate motions in paleotectonic plate reconstructions including reconstructions for the Proterozoic and Archean eons as data is sparse or poorly preserved. Further multi-disciplinary tectonic studies are necessary to broaden our understanding of Archean tectonics and by using metamorphic core complexes as analogues for transform plate boundaries we may greatly enhance paleotectonic reconstructions. The paleo-northeast-directed oblique collision between the Minnesota River Valley terrane with the southern Superior craton in the Neoarchean created predominantly dextral transpression in the Minnesota River Valley terrane and regional sinistral and dextral transpression to local sinistral transtension throughout the southern Superior craton. The rigid, Paleoarchean to Neoarchean Minnesota River Valley terrane collided into the recently formed, and rheologically weaker, southern Superior craton forming sinistral-oblique regional structures in the western Superior craton to the formation of metamorphic core complexes the eastern Superior craton suggesting transtension increased towards the east. This tectonic evolution from the formation of 2.700 Ga and 2.67 Ga MORB and arc to subduction to collision at 2.65-2.63 Ga to exhumation and lateral escape at 2.63 to 2.60 Ga is present in numerous Archean cratons worldwide with indistinguishable lithologies, structures, igneous and metamorphic petrology, geochemistry, geochronology, and mineral deposits. Evidence of this collision to transtensional tectonics and strike-slip deformation in the eastern Superior craton is preserved in the Archean cratons worldwide. Evidence of collision is preserved in the Paleoarchean to Neoarchean Minnesota River Valley terrane, Wyoming, Gawler, and West Antarctica and transpression to sinistral transtension due to lateral escape is preserved in Neoarchean Baltica (Central Kola Belt), Zimbabwe and Kaapvaal (Limpopo Belt), Yilgarn (Tropicana Gneiss), Eastern Antarctica, North China and Dharwar cratons. The development of craton, continental, or supercontinent breakup may have been triggered due the subduction of a transform plate boundary in the Wawa and Abitibi subprovinces, the size or rheology contrast of the colliding plates, the angle of the collision, and the formation of the metamorphic core complexes, lack of decoupling between the mantle and lithosphere boundary triggering mafic dyke swarms, plutonism, and (super) continental breakup. 105 �Keweenaw Fault Geometry and Kinematics along Bête Grise Bay, Michigan Tyrrell, C.W.1 Hubbell, G.E. 1, and DeGraff, J.M. 1 1 Michigan Technological University, Houghton, MI 49931 The Keweenaw Fault (KF) extends 350 km along the southern margin of the Midcontinent Rift System (MRS) from northwestern Wisconsin to near the tip of the Keweenaw Peninsula in Michigan (1). Reverse movement on the fault has thrust and tilted Portage Lake Volcanics (PLV, 1.1 Ga) over younger Jacobsville Sandstone (JS) (Fig. 1). The northeast portion of the fault near Keweenaw Point has been a matter of some interest since the USGS mapping campaign of the 1950s. Based on geophysical evidence, some have proposed that the fault continues offshore along an arc curving to the right by 90° to a southeasterly direction (1, 3-4). The farthest northeast location where the Keweenaw Fault can be directly examined is along the south side of the Keweenaw Peninsula from Bête Grise Bay eastward (Figs. 1, 2A). Here USGS maps from the 1950s show five shoreline areas where PLV and JS strata are juxtaposed (56). Our detailed mapping under the USGS EdMap program reveals that the previously mapped fault trace, based in part on aeromagnetic data and showing all PLV-JS contacts as faulted, oversimplifies the geologic relationships in this area. The anomalously sinuous, single fault trace mapped in the 1950s consists of at least five fault segments, generally striking ESE and forming a left-stepping pattern along the shoreline (Fig. 2B). At least three PLV-JS contacts previously mapped as faulted instead exhibit an unconformity between basal JS strata and older PLV lava flows. At one location, slightly deformed JS strata unconformably overlie fault breccia and gouge cutting PLV strata, indicating that one period of major slip on this KF segment occurred before local JS deposition. At other locations, JS strata are clearly cut and deformed along faulted contacts with PLV lavas, providing evidence for a second period of slip on the KF system after some or all JS deposition. Along shore near the Bare Hill rhyolite, PLV strata dip moderately to steeply SE to SSE for at least 3 km, a reversal of normal northerly dip that suggests an anticline developed north of this KF segment. Faulted PLV-JS contacts in the area generally dip > 80° N but locally dip steeply south. Geologic relationships across one fault segment suggest a significant component of dextral strike slip, and secondary faults have surface markings that indicate a mix of strike-slip and dip-slip motion. Ongoing work to quantify these relationships is designed to determine the degree of strikeslip to dip-slip partitioning along this portion of the KF system. The regional trace of the KF changes direction by over 70⁰ from NNE near Houghton to ESE at Bête Grise Bay, which mimics the change in strike of PLV layers over the same distance (Fig. 1). Paleomagnetic work (7-8) suggests that this direction change is a primary geometric attribute of the fault and not a result of bending around a vertical axis. Large crustal-scale faults often curve and split into segments near their terminations (9-10). Our mapping results thus imply that the KF system may terminate near the end of the peninsula in a series of fault splays, possibly transferring slip to other faults farther east. We hypothesize that slip on the KF changes from dominantly reverse dip-slip movement along its NE-trending portion near Houghton to dominantly dextral strike-slip near the tip of the Keweenaw Peninsula, and that slip magnitude decreases over this same distance. Acknowledgements: We appreciate the USGS funding this work and the timely field visit and comments last year by Bill Cannon, Klaus Schulz, and Laurel Woodruff, which does not imply their agreement or disagreement with these results. 106 �Figure 1: Keweenaw Peninsula where Portage Lake Volcanics are thrust over Jacobsville Sandstone. Black rectangle along the Keweenaw Fault near the tip of the peninsula marks focus area of Figure 2. (adapted from 2). Figure 2: Focus area along the Keweenaw Fault from Bête Grise Bay eastward (adapted from 5-6). Major faults shown as dark red traces. A) USGS maps from 1950s. B) Status of current fault mapping overlaid on prior maps. References 1. Miller, Jr., J.D., 2007, The Midcontinent Rift in the Lake Superior region: a 1.1 Ga Large Igneous Province: IAVCEI Large Igneous Provinces Commission, p. 1-18. 2. Cannon, W.F. and Nicholson, S.W., 2001, Geologic Map of the Keweenaw Peninsula and Adjacent Area, Michigan: United States Geological Survey, Map I-2696, Scale = 1:100,000. 3. Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332. 4. Hinze, W.J., Allen, D.J., Braile, L.W., and Mariano, J., 1997, The Midcontinent Rift System: a major Proterozoic continental rift: in Ojakangas, R.W., Dickas, A.B., and Green, J.C. (eds.), Middle Proterozoic to Cambrian Rifting, Central North America: Boulder, Colorado, Geological Society of America Special Paper 312, p. 7-35. 5. Cornwall, H. R., 1954, Bedrock Geology of the Lake Medora Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-52, scale 1:24,000. 6. Cornwall, H.R., 1955, Bedrock Geology of the Fort Wilkins Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-74, scale 1:24,000. 7. Hnat, J.S., van der Pluijm, B.A., and van der Voo, R., 2006, Primary curvature in the Mid-Continent Rift: Paleomagnetism of the Portage Lake Volcanics (northern Michigan, USA): Tectonophysics, v. 425, p. 71–80. 8. Kulakov, E.V., Smirnov, A.V., and Diehl, J.F., 2013, Paleomagnetism of 1.09 Ga Lake Shore Traps (Keweenaw Peninsula, Michigan): new results and implications: Can. J. Earth Sci., v. 50, no. 11, p. 1085-1096. 9. Boyer, S.E. and Elliott, D., 1982, Thrust systems: AAPG Bulletin, v. 66, p. 1196-1230. 10. Brozovic, N. and Burbank, D.W., 1999, Dynamic fluvial systems and gravel progradation in the Himalayan foreland: Geological Society of America Bulletin, v. 112, no. 3, p. 394-412. 107 �Alteration Mineral Zonation and Geochemical Characteristics of the Back Forty Deposit, MI; a Replacement-style Zinc- and Gold-rich Volcanogenic Massive Sulfide Deposit UPTON, Margaret1, SCHARDT, Christian1, HUDAK, George2, QUIGLEY, Eric3 1 Department of Earth and Environmental Sciences, University of Minnesota - Duluth, 1114 Kirby Drive, 223 Heller Hall, Duluth, MN 55812 2 Natural Resources Research Institute, University of Minnesota - Duluth, 5013 Miller Trunk Hwy, Duluth, MN 55811 3 Project Geologist, Aquila Resources, 414 10th Avenue, Menominee, MI 49858 The Aquila Resources Back Forty zinc- and gold-rich polymetallic volcanogenic massive sulfide (VMS) deposit is located adjacent to the Menominee River near Stephenson in the Upper Peninsula of Michigan. VMS deposits are created in submarine environments when heated seawater circulates through oceanic crust and precipitates base and precious metals at or near the seafloor due to both cooling and neutralization of the ore fluid. In the process, host rock mineralogy and geochemistry are modified by both downwelling and upwelling hydrothermal fluids, which produces distinct alteration mineral assemblages and metasomatic changes within the host rock (Shanks and Thurston, 2012). Alteration mineral assemblages and their spatial distribution can be used to unravel the geochemical evolution of the system, and help locate mineralization. The relationship between host rock and alteration mineralogy is not well understood or documented at the Back Forty Deposit but essential for understanding its genesis. The main objectives of this work are to identify physical, mineralogical, and geochemical characteristics of hydrothermal alteration mineral assemblages associated with the Back Forty Deposit. Alteration mineral and geochemical characterization include drill core logging, lithogeochemical, petrographic, and SEM analysis to better understand detailed mineral assemblages and mineral species’ chemical attributes. Core from nine drill holes (~ 2,950 meters), were logged to identify alteration mineral assemblages, intensity, and their textural characteristics. The deposit, hosted in felsic pyroclastic rocks, shows mostly sericite alteration, which was used to establish an alteration intensity scale of 1-4 (1: weak, 4: intense). Major alteration mineral assemblages observed were sericite ± silica ± chlorite. Sericite alteration is pervasive throughout the deposit (2-3) with silica alteration intensity ranging from 1-2 and a few areas of silica flooding (3-4). Weak chlorite alteration occurred throughout the deposit within the host rhyolite crystal tuff units as spotty chlorite associated with sulfide mineralization (1-2). Whole rock and trace element lithogeochemistry will be evaluated using the alteration box plot (Large et al., 2001) and mass balance analysis, such as the ISOCON method (Grant, 1986; see fig. 1). Results will be essential to assess quantitative chemical changes associated with alteration mineral assemblages and their spatial distribution to identify likely hydrothermal fluid flow pathways and mineralization vectors within the deposit. Using petrographic analysis as well as structural and lithological data, cross sections identifying alteration mineral zonation and its relative extent will be created to determine the relationship between massive sulfide mineralization and alteration mineral assemblage presence and intensity. From these results, it 108 �may be possible to use alteration mineralogy and geochemistry to determine the direction of hydrothermal fluid flow associated with mineral deposition and aid in future exploration efforts to locate additional mineralization on the Back Forty Deposit property, as most VMS deposits occur in clusters (Galley et al., 2007). Concentration of Intense Sericite Altered Sample, 108410 50 Hg Ta 40 Sn Yb V Hf K2O TiO2 SiO2 W 30 Zr Al2O3 Ga P2O5 Nb Ni Mo Ba 20 Y As MgO Pb S Ag Co 10 Fe2O3(T) 0 Na2O Cr 0 10 Sr Au MnO Cu 20 30 CaO 40 CO2 Zn 50 Concentration of Least Altered Sample, LK-348 Figure 1. ISOCON plot of selected elements used to compare elemental gains and losses between least and most altered samples. Isocon line of best fit is defined by immobile elements (Hf, Nb, Ta, Zr). Components above the line are enriched; below are depleted (modified from Ross, 2011). References Aquila Resources, 2017. Back Forty: Zinc- and Gold-rich Deposit. http://www.aquilaresources.com/projects/backforty-project/#! (accessed March 2018). Galley, A., Hannington, M., Jonasson, I., 2007. Volcanogenic Massive Sulphide Deposits. Geological Survey of Canada, Special Publication 5, p. 141-161. Grant, J. A., 1986. The Isocon Diagram: A Simple Solution to Gresens' Equation for Metasomatic Alteration. Economic Geology, v. 81, p. 1976-1982. Large, R. R., Gemmell, B.J., Paulick, H., 2001. The Alteration Box Plot: A Simple Approach to Understanding the Relationship between Alteration Mineralogy and Lithogeochemistry Associated with Volcanic-Hosted Massive Sulfide Deposits. Economic Geology, v. 96, p. 957-971. Shanks, W.C.P., Thurston, R., 2012. Volcanogenic Massive Sulfide Occurrence Models. USGS Scientific Investigations Report 2010–5070–C, 363 p. Ross, C., Hudak, G., Morton, R., Quigley, T., and Mahin, B., 2011, Preliminary stratigraphy and physical volcanology associated with the Paleoproterozoic Back Forty VMS deposit, Menominee County, Michigan [abstract/poster]: Institute on Lake Superior Geology, v. 57, Part 1, p. 70-71. 109 �Reconstruction of paleoenvironmental conditions and temporal patterns of ancient mining on Isle Royale using biogeochemical analyses of lake sediment VALL, Kathryn G.1, STEINMAN, Byron A.1, POMPEANI, David P.2, SCHREINER, Kathryn M.3, DEPASQUAL, Seth4 1 Earth and Environmental Sciences, Large Lakes Observatory, University of Minnesota Duluth Department of Geography, Kansas State University, Manhattan, KS 66506 3 Chemistry, Large Lakes Observatory, University of Minnesota Duluth 1049 University Dr, Duluth MN 55805 4 Cultural Resources, Isle Royale National Park, 800 E Lakeshore Dr, Houghton MI 49931 2 Isle Royale and the Keweenaw Peninsula of Michigan are home to some of the oldest examples of native North American metalworking and land use. The overarching objective of this research is to produce a reconstruction of the timing, spatial patterns, and environmental impacts of mining activities on Isle Royale through sedimentological and biogeochemical analysis of lacustrine sediments. We also seek to produce a parallel record of paleoenvironmental conditions in order to assess the potential impacts of environmental change on ancient mining cultures. In 2016, we collected a 7.5 m long sediment core sequence from Lily Lake on Isle Royale, MI. Lily Lake lies approximately 100 m above the current water level of Lake Superior, and formed approximately ~11,000 years before present following the retreat of the Laurentide ice sheet. Lily Lake has been exposed to very little human land use change relative to other lakes on Isle Royale (e.g. there are no ancient mine pits in the immediate catchment), and thus is well suited for reconstructing past environmental changes. We analyzed weakly sorbed metal concentrations using ICP-MS to test hypotheses on the timing and transport mechanisms of potential metal pollution derived from ancient mining activities. In addition, we conducted EAIRMS analysis (including carbon/nitrogen ratios, and the isotopic composition of organic C and N) on bulk organic sediment to provide a record of natural paleoenvironmental changes. Preliminary results from the metals analysis provide evidence of Middle Archaic mining activity that is temporally consistent with radiocarbon dated artifacts and similar evidence from other lakes located adjacent ancient mine pits on Isle Royale and the Keweenaw Peninsula of Michigan. Additional work is required to assess the relative influence of natural versus anthropogenic processes that may have influenced metal concentrations in Lily Lake sediment and to determine a transport mechanism for the putative mining related pollution. This study will provide a record of spatial/temporal patterns of mining activity and paleoenvironmental change in the Great lakes region that will aid in our understanding of large scale continental climate patterns, environmental responses, and the potential influence of climate/environmental variability on ancient land use and mining practices. 110 �Michigan Geological Survey Six years after assignment to Western Michigan University, Where are we today? John A. Yellich, CPG, Director, Michigan Geological Survey The Michigan Geological Survey functionality was reduced in 1978 to conducting minimal research, scientific publications and data management. For the next 30 years, Michigan went through multiple oil and gas booms and busts, Superfund authorization, Leaking Underground Storage Tanks, Brownfields, some mining development, yet no funding for a functioning geological survey. Where could you go to get up to date geologic research or information? What was and is still being used are special publications from the USGS, associations or academia and a 1982 Surficial Geological mapped based on 1915 field mapping, 1955 updates and a color change with some soils in 1982, this is 1915 surface geology only. The Michigan Geological Survey (MGS) was assigned in 2011, by legislation from the DEQ - Office of Oil Gas and Minerals to the Geological and Environmental Sciences (GES) Department at Western Michigan University (WMU), with no funding. MGS has functioned at the GES Department with WMU funding for two years, grants and a Special Appropriations (SA) from the Michigan Legislature in 2016, and has strived to establish a scientific value of a functioning geological survey by presenting programs and projects associated with the current day natural resources needs of Michigan. MGS surveyed the stakeholders and has been assessing some of the components of the noted societal needs, an integral segment of any survey today. A functioning geological survey is not the same as it was 25 or more years ago. Consequently, the users of geologic based data are not just the geoscientist, but regulators, county planners and development organizations, engineers, environmental scientists, extractive and land development industries, citizen scientists, anyone that has “boots” that touch the ground. MGS has completed a significant portion of the demonstration process and presented results to all the State of Michigan functional departments and has received letters of recommendation to support the continued geological research and mapping efforts conducted to date. MGS was also recognized and a resolution submitted to the Governor by the twelve sovereign Michigan tribes as needing a funded functioning geological survey to map and assess the water resources of Michigan. Michigan and many other states have a new contaminant, PFAS, and MGS has presented a geological approach to assess the aerial magnitude of this impact, geology. All these products and projects are scientific societal needs, however, at this time, the Legislature and Governor’s office has not seen fit to have an annual funding mechanism for the Michigan Geological Survey. The Natural Resources of Michigan are and have been an economic foundation and provided societal benefits to Michigan since 1840’s, over a 175 years. The identification and protection of these resources needs sufficient geologic information to assess, protect and manage all the components associated with any natural resources. 111 �Since 2011, the MGS has published 14 quadrangles (4-UP; 10-LP) with four in process within those 6 years, having one full time staff, some faculty and two contractors. There are critical need areas of Michigan that need to be mapped, but it takes a commitment by the State and requests by society for funding. MGS has provided geologic guidance on water and chloride issues in Ottawa County and MGS projects and research have strongly supported geological science in all aspects of identifying, managing and accessing the water resources of Michigan. MGS was instrumental in support of Statewide airborne LiDAR and encouraged Michigan to develop a program to contact the users and identify the benefits. This airborne effort was then done at a reduced cost and with nearly half of the State flown we now have LiDAR that will provide greater benefits to scientists and the public. MGS initiated and completed research utilizing standard geophysical methods and is utilizing new methods and remote sensing to support the 3D mapping projects and derivative geological products. Tromino Passive Seismic, NASA Gravity Recovery and Climate Experiment (GRACE) and Interferometry to assess, bedrock depth and topography, water storage and surface movements, respectively. For example, a City of Portage bedrock valley mapping for water resources, GRACE projections of increased water storage in Cass, St. Joseph, Kent and Ottawa counties has presented scientific research projects that support Michigan natural resources and yet no full time funding. These successful scientific demonstrations have also supported students in MS theses and PhD dissertations. MGS has strongly supported regionally and in Washington, DC the USGS geologic, geophysical and FEDMAP programs for airborne and ground surveys to assess the buried geology, near surface geology, water resources and shoreline stability issues of the Great Lakes areas, to name a few. The geologic community has a voice that has not been loud enough to be heard in Lansing and also, Washington, DC. Geoscientists must tell everyone that to understand our world today and tomorrow, we need geology. For example here we are in Michigan six years later, not knowing if we have sufficient water resources for some areas, questioning scientific data with “Wikipedia” type information, needing an updated geologic map of the UP bedrock and glacial systems and you as geoscientists not loudly proclaiming that validated geology needs to be done in priority areas of Michigan. You are the experts, what should the Geological Survey be doing? 112 �The Origin of Layering in the Olivine Zone, Black Sturgeon Sill, Nipigon, Ontario ZIEG, Michael J. and HONE, Samuel V. Department of Geography, Geology, and the Environment, Slippery Rock University, Slippery Rock, PA 16057 Layering in mafic intrusions is one of the most interesting, but also most controversial, aspects of igneous petrology. In this study, we explore the role of phenocrysts (entrained crystal cargo) versus in-situ fractionation in controlling layer development. Samples were taken from a continuous drill core through the Black Sturgeon sill (BSS), a 250 m mafic intrusion with a welldeveloped olivine zone from 120-200 m above its base. We analyzed bulk-rock geochemistry, modal mineralogy, and textures at 0.5 m intervals through this range, then used the resulting data set to investigate the petrogenetic processes responsible for the geochemical and petrographic variations in this part of the sill. Principal component analysis was used to characterize and summarize variations in the standardized and log-normalized major and minor oxide abundances. The first three principal components are the most significant, accounting for over 90% of the system variance. Based on these components, the rocks in the olivine zone fall into four distinctive compositional groups (Fig. 1). We interpreted the petrogenetic significance of the principal components by comparing them to trace elements, modes, norms, textures, and fabrics. In this initial comparison, we identified strong correlations between: the first principal component (PC1) and Ni-Sr (Fig. 2a); the second principal component (PC2) and incompatible element abundances, particularly Cu (Fig. 2b); the third principal component (PC3) and Sc (Fig. 2c). Thus, we conclude that PC1 is controlled by the ratio of olivine to plagioclase, PC2 is controlled by the abundance of a fractionated interstitial liquid component, and PC3 is controlled by augite abundance. Three fundamental observations are critical for understanding the significance of our results. (1) The olivine zone consists of four segments (a-d) defined by variations in the first principal component (Fig. 3a), which reflect the relative importance of olivine and plagioclase. (2) The olivine zone has a single coherent Z-shaped profile for PC2, controlled by smooth variations in incompatible element abundances (Fig. 3b). (3) PC2 and PC3 are positively correlated in Segments a, c, and d; they are negatively correlated in Segment b (Fig. 3c). Each of the four segments represents a distinct batch of magma, with its own characteristic phenocryst assemblage. The average ratio of olivine to plagioclase generally increased upwards, suggesting either an upward increase in source primitivity or “subcretion” of increasingly evolved magma pulses. All segments intruded rapidly compared to solidification time; after emplacement was complete, crystal-mush compaction drove evolved interstitial liquids from Segment b up into Segment d. The relationships between PC2 and PC3 suggest that augite crystallized after compaction-driven redistribution of evolved liquids in Segments a, c, and d. In Segment b, however, it was part of the crystal cargo, and Sc was not depleted (as Cu was) by the expulsion of interstitial liquids. In conclusion, the emplacement of multiple pulses of magma, each entraining a unique crystal cargo, controlled the basic layering structure and the major-oxide variability of the BSS olivine zone. Compaction-driven redistribution of interstitial liquids significantly modified trace element abundances, producing cryptic compositional layering incongruent with the modal assemblages. Although our results only address the formation of layering in this specific intrusion, the procedures we have developed can be applied to any system. 113 �Figure 1. Compositional groups. Four compositional groups can be distinguished. Figure 2. Interpretation of PCs. (a) PC1 reflects the olivine:plagioclase ratio. (b) PC2 reflects incompatible abundance. (c) PC3 reflects augite abundance. Figure 3. Stratigraphic profiles. (a) The olivine zone can be divided into four distinct segments. (b) Incompatible element abundances suggest compaction-driven redistribution of interstitial liquids. (c) Augite is correlated with incompatibles in Segments a, c, and d, but not in Segment b. This suggests that augite was a phenocryst phase in Segment b, but not in Segments a, c, or d. 114 � https://digitalcollections.lakeheadu.ca/files/original/b037a57a11eb6107fb3048857e8d3eee.pdf 399d169e2e63eb43833bac42fcefd842 PDF Text Text �Institute on Lake Superior Geology 64th ANNUAL MEETING May 15-18, 2018 Iron Mountain, Michigan SPONSORED BY: U.S. GEOLOGICAL SURVEY AND WISCONSIN GEOLOGICAL AND NATURAL HISTORY SURVEY Meeting Co-Chairs Laurel Woodruff, William Cannon, and Esther Stewart Proceedings Volume 64 Part 2: Field Trip Guidebooks Compiled by William F. Cannon Cover Photos: 1. Sandstone of Munising Formation (Upper Cambrian) lying on the Vulcan Iroin-formation at Groveland Iron Mine. Seen on trip 1. Photo by William Cannon. 2. Pillowed basalt of the Hemlock Formation at Way Dam, Michigan. Seen on Trip 2. Photo by Thomas Waggoner., 3. Dave’s Falls on the Pike River near Amberg, Wisconsin. Bedrock is the Athelstane Quartz Monzonite cut by diabase dikes. Seen on trip 4. Photo by William Cannon. 4. Quinnesec Iron Mine on the Menominee Iron Range, Michigan. Seen on Trip3. Photo by William Cannon. �64TH INSTITUTE ON LAKE SUPERIOR GEOLOGY VOLUME 64 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD TRIP GUIDEBOOK TRIP 1: ARCHEAN AND PALEOPROTEROZOIC GEOLOGY OF THE FELCH DISTRICT, CENTRAL DICKINSON COUNTY, MICHIGAN TRIP 2: GEOLOGY OF THE HEMLOCK FORMATION TRIP 3: GEOLOGY AND IRON ORES OF THE MENOMINEE IRON RANGE, DICKINSON COUNTY, MICHIGAN TRIP 4: GRANITOID ROCKS OF THE PEMBINE-WAUSUA TERRANE IN NORTHEASTERN WISCONSIN Reference to material in Part 1 should follow the example below: Authors, 2018, abstract title, 64th Institute on Lake Superior Geology Proceedings, v. 64, Part 1, Field Trip Guidebook, p. xx. Proceedings Volume 64, Part 1: Program and Abstracts, and Part 2: Field Trip Guidebook are published by the 64th Institute on Lake Superior Geology and distributed by the Institute Secretary: Peter Hollings Department of Geology Lakehead University Thunder Bay, ON P7B 5E1 CANADA peter.hollings@lakeheadu.ca Some figures in this volume were submitted by authors in color, but are printed grayscale to conserve printing costs. Full color imagery will appear in the digital version of the volume when it is available on-line at: http://www.lakesuperiorgeology.org ISSN 1042-9964 ��Part 2: Field Trip Guidebooks Table of Contents Trip 1: Archean and Paleoproterozoic Geology of the Felch District, Central Dickinson County, Michigan 1 Trip 2: Geology of the Hemlock Formation 39 Trip 3: Geology and Iron Ores of the Menominee Iron Range, Dickinson County, Michigan 69 Trip 4: Granitoid Rocks of the Pembine-Wausua Terrane in Northeastern Wisconsin 107 �FIELD TRIP 1 Tuesday May 15, 2018 ARCHEAN AND PALEOPROTEROZOIC GEOLOGY OF THE FELCH DISTRICT, CENTRAL DICKINSON COUNTY, MICHIGAN William F. Cannon, Klaus J. Schulz, Robert A. Ayuso, U.S. Geological Survey Thomas H. Mroz, BSGE, MSPG, CPG INTRODUCTION This trip examines the stratigraphy, structure, and economic geology of Precambrian rocks near the town of Felch in central Dickinson County of northern Michigan. The location of the area is shown on Figure 1 relative to other well-documented structures and iron ranges of the region. Precambrian rocks range in age from Archean to Paleoproterozoic, and outliers of Cambrian sandstones are also widespread. Relationships along the basal Cambrian unconformity are included in the trip. Much of the interest in the area, both geologically and economically, has been focused on the Felch trough (Figure 2) where Paleoproterozoic rocks of the Chocolay and Menominee Groups form a complex syncline that is infolded and infaulted with Archean gneisses. Minor amounts of iron ore were produced in the early history of the area and a major concentrating-grade mine, the Groveland, was active in the 1960’s and 1970’s. The geology of the area was mapped and described in detail by the U.S. Geological Survey in 1950’s and published in 1961 (James et al., 1961). The information in that report is the basis for much of the descriptive material in this guide. A few more recent studies have added significant additional data and geochronological constraints that clarify some of the relationships between various rock units and the events that formed them. An underlying theme of the trip is that a substantial amount of additional research is warranted to fully understand this complex region and that relatively abundant outcrops, along with newly acquired high-resolution geophysical data, make this a very attractive region for a wide variety of research. Stratigraphy. The long-accepted stratigraphic sequence of Archean and Paleoproterozoic rocks that was defined by James et al. (1961) requires some modification based on more recent radiometric age determinations. Some of these are discussed in more detail below in individual stop descriptions. A tentative correlation chart in Figure 3 reflects these suggested changes, but additional work seems warranted to clarify some aspects before proposing formal changes to stratigraphic nomenclature and correlations. 1 �Figure 1. Generalized geologic map of part of the Upper Peninsula of Michigan showing the location of central Dickinson County in relation to major structures and iron ranges of the region. Modified slightly from James et al. (1961, plate 1). The oldest rocks of the region are complex gneisses of the Carney Lake Gneiss and other probably correlative units in structurally separated fault blocks. They have long been considered Archean. Recent geochronology of samples collected south of the fieldtrip area using the USGS/Stanford Sensitive High Resolution Ion Microprobe (SHRIMP) produced U-Pb data on zircons that confirm an Archean age (Ayuso et al., 2017; 2018). Two samples were collected for radiometric dating from the southern half of the Carney Lake complex: 1) sample 1 is from a granitic K-feldspar-bearing gneiss that is locally pegmatitic; 2) sample 2 is from a banded and folded gray to red granitic gneiss. Abundant zircons (70-200) were obtained from sample 1 that range from anhedral to subhedral, contain complex igneous and irregular growth zoning, and multiple growth rims; these zircons have irregular to pyramidal overgrowths. The zircons from sample 2 range from slightly rounded to subhedral and are otherwise mostly similar to zircons from sample 1. A total of 129 cores and rims were analyzed. Individual zircons have older ages near their cores (mostly discordant) and younger ages near their rims (Figure 4A). 2 �3 �Figure 3. Proposed correlation chart for the area of Field trip 1. Modifications from previous correlations are based on recent radiometric determinations that provide direct ages for some units and place constraints on the ages of others. On a concordia diagram, U-Pb data plot as clusters of data points ranging from concordant to discordant and suggest several chords and intercepts that are common to both samples from the Carney Lake Gneiss (Figure 4B). That study identified cores of individual zircons as old as 3.8 Ga. The most common age for individual zircons and for rims on older grains is about 2.75 Ga and records a younger major event in the late Archean. James et al. (1961) defined the Dickinson Group, consisting of the East Branch Arkose, Solberg Schist, and Six Mile Lake Amphibolite, as Archean based on field relationships between the Six Mile Lake Amphibolite, the upper formation of the group, and presumed Archean gneisses to the south. However, U-Pb dating of detrital zircons in the East Branch Arkose has shown that it must be 2.1 Ga or younger (Craddock et al., 2013). We suggest that the Six Mile Lake Amphibolite is also Paleoproterozoic because of chemical similarities between it and other Paleoproterozoic mafic rocks. The Archean age proposed by James et al. (1961) was based on field relationships 4 �between amphibolites and Archean granites and gneisses that require further evaluation. The Solberg Schist is intruded by a gabbro that has the distinctive composition of the Six Mile Lake Amphibolite (see Stop 6 for discussion). James et al. (1961) described the contact between the East Branch Arkose and Solberg Schist as transitional. Further study of the Dickinson Group is clearly required to resolve the age relationships of its three formations. Figure 4. A- BSE (back scatter electron) image of a zircon from the Carney Lake Gneiss showing ages of four analyzed spots. B- Concordia diagram for 129 spot analyses from zircons in the Carney Lake Gneiss. A distinctive granite in the area referred to informally as the “porphyritic red granite” by James et al. (1961), has now been dated using SHRIMP U-Pb on clear and pale brown zircons at 2.1 Ga (Ayuso et al., 2018; and discussion of Stop 11). The porphyritic red granite provides a local source for the relatively common detrital zircons of that age found in the East Branch Arkose of the Dickinson Group (Craddock et al., 2013). These zircons also suggest an erosional interval after 2.1 Ga to unroof the granite prior to deposition of the Dickinson Group. The Dickinson Group is overlain by the Hemlock Volcanics northwest of the field trip area. Those volcanics are dated at 1.875 Ga (Schneider et al., 2002) and establish a minimum age for the Dickinson Group. The Chocolay Group, consisting of the glaciogenic Fern Creek Formation, Sturgeon Quartzite, and Randville Dolomite, is well constrained to have been deposited between 2.3 Ga, the age of detrital zircons within it, and 2.1-2.2 Ga, the age of xenotime cement in the sediments (Vallini et al., 2006). The Menominee Group, including the classic stratigraphy of the Vulcan Iron-formation and underlying Felch Formation, is nowhere in contact with the Dickinson Group although the two are not far separated in the Felch area (see Figure 2). They may be largely time equivalent in spite of rather pronounced lithologic differences. The Michigamme Formation lies at the top of the stratigraphic succession across the region and has been widely interpreted to have an unconformable contact with the underlying units. The Michigamme and underlying Hemlock Volcanics are intruded by the Peavy Pond Complex to the 5 �northwest of the field trip area. A recent SHRIMP U-Pb date on honey-colored euhedral to subhedral zircons for the Peavy Pond Complex yielded an age of 1.85 Ga (Figure 5) (Ayuso et al., this volume). The results place a minimum age on at least the lower part of the Michigamme Formation. This presents something of an enigma in that elsewhere in the Upper Peninsula the Sudbury impact layer, which was deposited at 1.85 Ga, lies at or near the base of the Michigamme Formation (Cannon et al., 2010). It seems, therefore, that rocks assigned to the Michigamme Formation in the Felch area could be substantially older than the Michigamme Formation in much of the larger region of the Upper Peninsula. Figure 5. A- BSE (back scatter electron) image of a zircon from the Peavy Pond Complex showing the SHRIMP U-Pb age of an analyzed spot. Bright spots are residual gold coating from SHRIMP analyses. B- Concordia diagram for 32 spot analyses of zircons from the Peavy Pond Complex. Tectonics. Rocks in the Felch region record multiple orogenic events from 3.8 Ga to younger than 1.83 Ga. Archean rocks are very complexly deformed and metamorphosed by at least two Archean orogenies, one at 3.8 Ga and a later 2.75 Ga event. They also record one or more deformational and metamorphic events in the Paleoproterozoic. Although the work of James et al. (1961) mapped and described much of the Archean of the area, detailed structural studies have not been done, and much remains to be learned about the sequence of rock-forming and deformational events recorded in these complex gneisses. More study has been devoted to the tectonics of Paleoproterozoic rocks and at least an outline of their tectonic history is known, although much additional detail seems likely to be decipherable with futther examination. Only two published structural studies based on substantial new field data have been published since 1961. Ueng and Larue (1988) defined four tectonic terranes and six phases of deformation within Paleoproterozoic strata. Klasner and Sims (1993) proposed a somewhat different sequence and suggested that the set of faults that dominate the map pattern near Felch were backthrusts formed in later phases of the Penokean orogeny at about 1.85 Ga. Both of those studies were conducted prior to significant geochronologic constraints determined over the past two decades. They ascribed all deformational events to the Penokean orogeny, 6 �which by that definition was a prolonged and complex event including both lateral and horizontal shortening phases. More recently, several geochronologic studies show that the tectonic sequence is much more complex and that the region bears the imprint of two other orogenies after the Penokean. A sequence of Archean-cored gneiss domes surrounded by Paleoproterozoic sedimentary and volcanic rocks across central Minnesota, northern Wisconsin, and northern Michigan, once thought to be late Penokean structures, was documented to have formed in mid geon 17 (Schneider et al., 2004). This gneiss dome corridor was interpreted to have been related to subduction during the Yavapai orogeny at about 1.75 Ga. Archean masses near Felch are part of the gneiss dome corridor and, although now dismembered by later faulting, record that Yavapai basement doming event rather than Penokean deformation. A distinctive aspect of the Felch region is a set of easterly and ENE-trending faults that cut all other structures (see Figure 2). They result in a sequence of fault slices that dominate the map pattern of the region. A comparable array of such faults is not known elsewhere in the region. This faulted domain is bounded on the north by the Bush Lake fault, across which there is a sharp discordance in structural trends, nearly 90° in places. The age of the faulting can be constrained by the fact that it offsets the 1.85 Ga Peavy Pond Complex to the west of the field trip area. It also has been shown to offset metamorphic isograds (Attoh and Klasner, 1989) of the Peavy node where peak metamorphism has been dated at 1.83 Ga (Schneider et al., 2004; Holm et al., 2007). Attoh and Klasner (1989) also documented a change in metamorphic pressures across the Bush Lake fault with peak pressures of 4.8 kbar to the south and 3.3 kbar to the north. This suggests that fault motion was south side up with vertical displacement on the order of five kilometers. This set of faults also partly dismembers the gneiss dome structures related to the mid-geon 17 Yavapai orogeny and suggests that they are post-Yavapai structures. An imprint of the geon 16 Mazatzal orogeny in this region was documented by Romano et al. (2000), who showed a heating event of 300-350oC across the Felch region in late geon 16. Whether this heating was accompanied by deformation has not been determined in the Felch region, but to the south, in Wisconsin, strong deformation of Baraboo interval quartzites shows conclusively that strong Mazatzal-aged deformation was widespread. It seems reasonable, with our present level of understanding of the tectonics of the Felch region, that the easterly-trending set of faults is the northernmost manifestation of Mazatzal deformation. Metamorphism. In a classic work on regional metamorphism, James (1955) identified three metamorphic nodes within Paleoproterozoic rocks in the Upper Peninsula of Michigan. One of those, the Peavy node, encompasses the area of this field trip. Most stops lie within the staurolite zone and the effects of recrystallization are obvious in outcrop. Since James’ work, radiometric dating and more modern metamorphic petrology studies have further defined the character of the metamorphism of the region. The age of peak metamorphism of the Felch node has been determined as circa 1.83 Ga from dates derived from metamorphic monazite (Schneider et al., 2004; Holm et al., 2007) measured in Archean gneiss near Foster City. Peak temperature and pressure of metamorphism were about 600-650oC and 5 kbars (Attoh and Klasner, 1988). A younger amphibolite facies metamorphism was documented at 1.78-1.74 Ga related to Yavapai accretion and gneiss doming (Holm et al., 2007). Much older metamorphism, probably culminating at about 2.75 Ga, is evident within Archean rocks, many of which are gneisses and migmatites. The Hardwood Gneiss, seen at Stop 7, 7 �records granulite facies conditions with estimated pressures of 8.2-11.6 kb and temperatures of ~770°C for an initial event, and conditions of 6.0-10.1 kb and temperatures of 610-740oC for a second event inferred to be Paleoproterozoic but “pre-Penokean” (Peterson and Geiger (1990). Economic geology. Iron, mostly related to the Vulcan Iron-formation, has been the principal commodity of interest in the Felch region. Early exploration and attempts to develop mines are summarized by James et al. (1961). The first indication of ore in the Felch trough was a description of a ridge of high grade iron ore in eastern Iron County by Jacob Houghton and reported to William Burt, a government land surveyor, in Marquette County in 1846 and was recorded in Senate Documents for the 31st Congress (Jackson, 1849). At what later became the Groveland mine a veneer of oxidized Vulcan Iron-formation was found on the south side of the ridge that drew the attention of early developers. But once operations started stripping a pit, the veneer was found to be only several feet thick and did not extend to depth as originally thought. The first development was an economic disaster as monies had been invested in town sites and the extension of a railroad to the Escanaba iron furnaces. Being not far from more prospective areas such as the Menominee Range to the south, the Felch area was heavily prospected and numerous attempts at mining occurred between 1880 and 1913. Only four mines of any significance were developed in the Felch area and produced a total of about 625 million tons of ore, mostly of low grade. These direct-shipping ores were soft masses of hematite and goethite that were found directly beneath the unconformity between the Vulcan Iron-formation and overlying Cambrian sandstones. They are widely accepted to be paleosupergene deposits formed by late Precambrian and/or early Cambrian weathering. The Felch region differs from much more productive nearby regions, such as the Menominee Range and Iron River-Crystal Falls district, in being substantially metamorphosed. Rocks, including the Vulcan Iron-formation, are coarse-grained as a result of metamorphic recrystallization and thus are less susceptible to the paleo-weathering than comparable iron-formations in other nearby districts where large paleosupergene ore bodies were formed. With the advent of large-scale concentrating and pelletizing technology in the 1950’s, portions of the Vulcan Iron-formation became targets for concentrating-grade ore production. The Groveland area proved to have sufficient tonnage of iron-formation accessible by open-pit mining to allow development of the Groveland mine, a significant iron-producer in the 1960’s and 1970’s. The potential for mineral production in this area has dropped in recent decades as exploration projects, such as for uranium and diamonds, have failed to find deposits of current economic viability. There are active gravel pit operations, reprocessing of Groveland Mine waste rock dumps for crushed stone, and small quarries in the Randville Dolomite for decorative stone. This limited activity is the current extent of mineral resource exploitation in central Dickinson County 8 �FIELD TRIP STOPS Stop 1. Groveland mine. (45.988°N, 87.981°W) The geology within the long-abandoned open pit of the Groveland mine is not accessible for field trips because of flooding, slumping of pitwalls, and safety concerns of current owners. The mine property is fenced and is not accessible to the public. We have been given access to the property to examine material on large waste-rock piles that show good examples of the various lithologies of the Vulcan Ironformation, and provide views into parts of the flooded pit. Figure 9 shows the present surficial character of the mine area and the location of the waste piles available for observation and sampling. Figure 6. The Groveland Mine has a long history that began in the late 1880’s and continued into the 1980’s. This view, probably from the 1970’s, is looking south with the plant in the upper central part of the air photo. Source M. A. Hanna Company. Archives of the Michigan Department of Environmental Quality. History. Mining of iron ore at Groveland began as an underground operation in 1891 on outcrop of the Vulcan Iron-formation identified in 1846 by assistants to William Burt, the government land surveyor. The operation was abandoned after only a few years of operation due to the lack 9 �of direct shipping ore. The mine was reopened in 1901 and mined for four years by Corrigan, McKinney & Company. In 1907 the mine was again started up by the Groveland Mining Company. and had production through a 294 foot deep, three compartment shaft with levels at 70, 140,` and 210 feet. Iron content remained a problem and production ceased in 1913. It was reported that the last shipment had unacceptable iron content and was dumped into Lake Erie. Several companies gained ownership of the properties and in 1926 test pits and trenches were completed by an independent developer, Mr. R. M. Adams. In 1948 the properties were consolidated through leases by M. A. Hanna Co. and in 1951 the Grovelend became their first taconite project. A pilot project ran for six months to develop grinding and concentration processes of the jasper ores, but it took seven years to develop a viable process to treat the complex ore, which is unique because of its mineralogy and metamorphism. The ore is very coarse grained and was defined by the operators as three types; magnetite, magnetite silicate, and specular hematite (Figure 7). In 1957 construction of the concentrating and pelletizing plant began and in 1959 the plant became operational with an output of 700,000 tons of concentrate annually. The mine became the second concentrating grade (taconite) operation in Michigan, following closely the opening of the Humbolt mine on the Marquette range. In 1963 a traveling grate pellet plant was completed with a capacity of 1.25 million tons per year. The $35 million expansion also included a concentrator upgrade and production was increased to 1.6 million tons annually. In 1968 a fourth line was added and resulted in an annual capacity of 2 million tons. 1977 was a record year with the output of iron concentrate reaching 2.1 million tons. Overall, the expenditures totaled $70 million, and employed 530 resulting in an investment of $132,000 per employee. Annual payroll was $12.5 million and taxes provided $1.332 million of revenue to the state and local governments. Operating services and supplies were $31.5 million for the local economy. In 1980 the mine closed after producing about 36 million tons of pelletized iron concentrate. Portions of the fresh water ponds have been developed into recreational fishing areas for the public. Geology. The most complete geologic description of the Groveland deposit is by Cumberlidge and Stone (1964), two geologists with M.A. Hanna Mining Company, and was based on extended observations as the present pit was developed. They showed that the ore body formed the keel of a complex doubly plunging syncline that was overturned to the south. The thickest extent of the Vulcan Iron-formation was along the south limb as shown in the plan map and cross section, (Figure 8). The north overturned limb was capped by Cambrian sandstone prior to stripping, and initial underground development was probably in the iron-formation subcrop below the sandstone contact. The Vulcan Iron-formation is divided into three informal members at the Groveland mine (Figure 8): 1) the lower Vulcan consisting mostly of hematitic jasper, 2) the middle Vulcan is dominantly even-bedded magnetite-silicate iron-formation with lesser hematitic jasper, and 3) the upper Vulcan Iron-formation is uniformly bedded magnetite-silicate iron-formation. The mine is located in the staurolite zone of regional metamorphism, and iron silicates in the Vulcan Iron-formation are commonly coarse-grained reflecting that intense recrystallization. The most common silicate minerals identified by Cumberlidge and Stone (1964) are Ca-Mg hornblende, tremolite, and actinolite. Pyroxene, biotite, and cummingtonite are less abundant and garnet is rare. Some representative photographs of the geology of the mine area is shown in Figure 7. 10 �Figure 7. Photographs from the Groveland mine. A- Unconformity between flat-lying sandstone of the Munising Formation and weathered Vulcan Iron-formation on the north wall of the Groveland pit. B- Folded jasper–specularite iron-formation, C-Lenticular beds of oolitic jasper with specularite interbeds, D-Silicate-magnetite iron-formation with large sheaves of amphibole. . 11 �Figure 8. Geologic map and cross section (with magnetic and gravity profiles) of the Groveland mine area. The Groveland pit was developed in the thickest part of the Vulcan Iron-formation in the eastern part of the map. Source: M.A. Hanna Mining Company from archives of the Michigan Department of Environmental Quality. Stratigraphic section summarizes descriptions of the Vulcan Iron-formation and related units as reported in Cumberlidge and Stone (1964) provided by Thomas Waggoner (personal communication). Unit names are informal mine terminology. . 12 �Figure 9. False color LiDAR image of the area of the Groveland mine showing the location of Stop 1 and outcrops as mapped by James et al. (1961) prior to mine development. 13 �Figure 10. Location of Stops 2 and 3 and the location of holes drilled for uranium exploration near Stop 2. Image from Google Earth. Stop 2. Archean gneiss at Gene’s Pond. (46.058o N, 87.855 o W) At the boat launch site at the west end of Dixon Road are several outcrops of Archean granitic gneiss and mafic dikes that cut the gneiss. The granitic rocks are mostly plagioclase porphyritic rocks with a moderately to well developed, nearly vertical shear foliation (Figure 11A). The mafic dikes cut the foliation. They appear largely massive and undeformed in outcrop but are thoroughly amphibolitized and the amphiboles show a weak alignment in thin section (Figure 11B). Based on landforms, we interpret that there are outliers of Munising Formation (Upper Cambrian sandstone) both east and west of this locality, and further interpret that the present land surface here is very nearly the exhumed unconformity at the base of the Cambrian. The unusual reddish hue of much of the granite may be a reflection of weathering or alteration along the unconformity (see Figure 11A and 12). 14 �Figure 11. A- Moderately sheared Archean granite. B- Massive mafic dike with blocky fracture. Figure 12. Photomicrographs of sheared granitic rock at Stop 2. Rock is mostly plagioclase with moderately developed cataclastic textures. Nearly all plagioclase grains are stained with submicroscopic hematite (?). A- Plane polarized light, B- Crossed Nichols. A rather extensive exploration effort for unconformity type uranium deposits was undertaken here in the early 1980’s by Minatome Corporation (Hunter, 1986; Lehman, 1987). This included twenty shallow drill holes shown on Figure 10, one of which was only a few tens of meters south of these outcrops. Uranium, occurring as pitchblende, was found as open-space fillings along with calcite, hematite, and minor chlorite. The drilling defined an E-W brittle fault that dips about 60o north and is subparallel to an older mylonitic fault. The mineralized assemblage heals breccias that are most common in the hanging wall of the brittle fault. The drill holes were located to test the depth extent of surface radioactive occurrences but found that no mineralization extended more than 85 m below the surface and that the surface extent of mineralization of individual occurrences was about of equal vertical dimensions. U-Pb dating of the pitchblende (Lehman, 1987) yielded a range of results, all of which were of Paleozoic or younger age. A likely conclusion is that the mineralization formed in Paleozoic or younger times just below the unconformable contact of Cambrian sandstone and Archean gneiss. This exploration and its results have never been described in detail, but drill logs and core for most of the holes are available for study at the Michigan Geological Sample Repository in Harvey, Michigan. 15 �Stop 3. Randville Dolomite at Gene’s Pond. (46.072 o N, 87.866 o W.) A small lakeside outcrop just south of the boat launch at Gene’s Pond public access site displays many of the typical features of the Randville Dolomite in the northern part of the Felch area. In much of the Felch area strong metamorphism has converted the Randville Dolomite into coarsely crystalline white to gray marble. Much of the primary structure is obliterated. However, Stop 3 lies north of the highest grade metamorphic zones and metamorphic recrystallization is only minor with abundant primary features preserved. Here, the Randville underlies a large area between the Bush Lake fault on the north and the Norway Lake fault on the south. It has been described in some detail by Clark (1961, Chapter C of James et al., 1961). Unfortunately, the more illustrative Randville outcrops described in detail by Clark are not easily accessible for field trips. The small outcrop at Stop 3 is shallowly dipping and well bedded dolomite with undulose, generally upward-domed, bedding that is likely stromatolitc structures (Figure 13) Figure 13. Randville Dolomite at Stop 3 showing undulose bedding, probably reflecting weakly developed stromatolitic mounds. The general description of the Randville provided by Clark (1961, p. 107-109) is “The Randville dolomite is apparently divisible into three members: (1) an upper member and (2) a lower member of dolomite with minor interbedded slate, separated by (3) a slate member with minor interbedded dolomite…. The total thickness is more than 800 feet in the Norway Lake area. Neither the top nor the bottom of the formation is exposed, and the character of the rocks that immediately underlie and overlie the Randville dolomite is not known. …… Most of the dolomite 16 �is massive to thin bedded, and stromatolites (algal structures) and intraformational conglomerate are common. The dolomite is light gray to red on fresh surfaces and weathers white to light brown. It has a fine sugary texture. Grains of quartz sand, most of which show undulatory extinction, are abundant in some beds and in some places comprise more than 50 percent of the rock. No oolites were found. The stromatolites, in sections normal to bedding planes, are concentrically banded structures with domal or columnar form, and in sections parallel to bedding planes they are concentrically banded elliptical forms. Most are 1 to 3 inches in diameter and 2 to 6 inches high. The banding of the stromatolites is convex upward, providing a reliable criterion for tops of beds. Where the structures are partly replaced by chert the forms are accentuated on weathered surfaces. Intraformational conglomerates or breccias are present ….. Most of the pebbles are dolomite, but a few pebbles of dolomitic slate occur. The pebbles are 1 to 4 inches in diameter and are well rounded. No strong dimensional orientation is evident. The matrix is dolomite with intermixed coarse quartz sand.” The slates, as described by Clark (1961) are dark gray to gray-green and are composed largely of sericite with some quartz, chlorite, and microcline. Graded beds are common. Stops 4, 5, and 6. The Dickinson Group The three formations originally defined as the Dickinson Group (James et al., 1961), 1- the East Branch Arkose (Stop 4), 2- the Solberg Schist (Stop 5), and 3- the Six Mile Lake Amphibolite (Stop 6) will be examined from north to south, the originally interpreted stratigraphic order. Stratigraphy and age- The Dickinson Group was originally defined by James et al. (1961) to include three formations, from presumed oldest to youngest, the East Branch Arkose, the Solberg Schist, and the Six Mile Lake Amphibolite, which they interpreted to be in conformable contact with each other. The East Branch Arkose is a sequence of arkosic conglomerate and sandstone with interbedded mafic volcanic rocks. The Solberg Schist consists of finer clastic and volcanic rocks with at least one interbedded banded iron-formation, the Skunk Creek Member. The Six Mile Lake Amphibolite was interpreted to be highly metamorphosed mafic volcanic rocks with abundant granitic intrusions. James et al. (1961) provided detailed geologic maps and descriptions of each unit. This discussion is based largely on their work. Each unit has a maximum thickness of 2,000 to 4,000 feet, and James et al. (1961) state that a total thickness of 10,000 to 12,000 feet for the group is indicated. This original work was done before radiometric ages were available so relative ages were based on field relationships. The supposed lower formation of the group, the East Branch Arkose, seems clearly to lie unconformably on Archean gneisses to the north as indicated by numerous gneissic pebbles in conglomerates of the East Branch Arkose. To the south, James et al. (1961) believed that the Six Mile Lake Amphibolite was intruded and altered by a large batholith of Archean age, so ascribed a late Archean age to the group. More recently, the acquisition of radiometric data has modified the permissible age range for the Dickinson Group. All units of the group were recrystallized during regional metamorphism related to the Peavy metamorphic node (James, 1955). That metamorphism has been dated at 17 �approximately 1.83 Ga, the age of metamorphic growth of monazite (Holm et al., 2007), thus providing a minimum age for the group. Detrital zircons reportedly from the East Branch Arkose show a spectrum of ages (Craddock, et al, 2013) that includes a strong, well defined, peak at approximately 2.1 Ma which would establish a maximum age for the East Branch. Unfortunately the coordinates given by Craddock et al. (2013) for the sample correspond to a roadcut of Solberg Schist, not East Branch Arkose, putting some uncertainty on interpretation. But if the Solberg is conformable with the East Branch, as interpreted by James et al. (1961) the results still provide a constraint for the group as a whole even if the sample is from the Solberg. Thus, the available radiometric age constraints place the East Branch Arkose and possibly Solberg Schist sedimentation and volcanism within the range of 2.1 to 1.83 Ga, similar to other Paleoproterozoic strata of the region. We suggest a reinterpretation of the Dickinson Group in which it is entirely Paleoproterozoic. The age of the Six Mile Lake Amphibolite is problematic in this interpretation in that James et al. (1961) described a southward transition of amphibolites into gneisses that are clearly of Archean age. Whether these amphibolites are a part of the Six Mile Lake or some older sequence is not clear at present and requires further evaluation. Lithology- The following lithologic descriptions are summarized entirely from the detailed descriptions provided in James et al. (1961). East Branch Arkose: As described by James et al., (1961. p. 13-14), “The formation consists of thick-bedded arkose with many beds of coarse conglomerate, interbedded with metamorphosed tuffs and basic volcanic flows. The conglomerates, though not the dominant rock type, are the most striking feature of the formation. The beds typically are 10 to 30 feet in thickness. The pebbles in the conglomerate have been drawn out into lenses that, on a horizontal surface across the nearly vertical beds, have a length-to-width ratio of about 3:1. In most parts of the area this shearing is parallel to bedding, that is, eastward but in a few places it is at an angle. Linear structure is not pronounced; most of the flattened pebbles have a length in vertical section about equal to that in horizontal section. In a few places a nearly vertical linear structure marked by grooving of the pebbles is evident. Vitreous quartzite is the dominant rock type among the pebbles, with granite gneiss, slate or schist, and quartz being of lesser abundance. Some of the quarzite pebbles show well-defined bedding. The arkose is pink to gray, massive, and abundantly cross-bedded. In many outcrops and in hand specimens it closely resembles a granite gneiss, but the well-defined crossbedding is complete proof of its sedimentary origin. In the more southerly outcrops of the East Branch arkose, dark-gray fine-grained tuffs are interbedded with the arkose; the rock consists principally of quartz and untwinned feldspar, with scattered grains of epidote, biotite, and carbonate. Rounded grains of opalescent quartz are present in some layers. Metabasalt flows are not uncommon. The rock is black and hornblende-rich. In the outcrops in the NW sec. 17, T. 42 N., R. 28 W., the originally scoriaceous top of one of these flows can be seen. Some of the metamorphosed flows are moderately magnetic and give rise to the aeromagnetic anomalies shown on the general map of the district.” Solberg Schist- James and et al. (1961, p. 17-18) describe the Solberg Schist as follows. “The Solberg schist lies immediately south of the East Branch arkose. A considerable amount of interbedding of arkose and schist is evident in the outcrops, so that location of the contact between the units is somewhat arbitrary. . . . The more northerly exposures of the unit consist 18 �chiefly of dark fine-grained hornblende and biotite schists. Locally, muscovite is an abundant constituent. Some outcrops show a banding, essentially parallel to the foliation, which may represent original layering. In one place, near the south edge of sec. 13, T. 42 N., R. 29 W., the schist is interbedded with coarse clastic material similar to rock of the East Branch arkose. The more southerly exposures of the Solberg schist consist of quartz-mica schist, parts of which might be better termed micaceous quartzite. This rock is exposed in the north part of sec. 24, T. 42 N., R. 29 W., and secs. 21 and 2, T. 42 N., R. 28 W. In general, the rock is massive, gray, and well banded. The banding, which consists of alternation of quartzitic and micaceous layers, almost certainly represents original bedding. Linear structure is strongly developed in all the schists, especially in the hornblendic varieties. It is marked by orientation of hornblende needles and by biotite. The lineation is in the plane of foliation and in general plunges eastward at a low angle.” The Skunk Creek member of the Solberg schist is a bed of iron-formation. This bed gives rise to a very strong magnetic anomaly, by means of which the iron-formation can be traced in a belt across most of the map area. . . .The Skunk Creek member has been penetrated by drill holes in several places. It is chiefly from this drill core that information concerning the lithologic character has been obtained, although none of the holes has cut the entire unit. The distinctive part of the formation is a thin-banded rock consisting of alternating layers of granular quartz (probably originally chert), magnetite, and various mixtures of hornblende, biotite, grunerite, garnet, and epidote. This material grades into biotite-hornblende schist, containing magnetiterich layers, by interbedding at both the upper and lower contacts. The iron-formation is cut by many thin dikes of coarse pegmatite; garnet, and tourmaline are commonly developed near the contacts. The thickness of the Skunk Creek member is somewhat uncertain because of the small amount of data available, but it is about 100 feet.” Six Mile Lake Amphibolite- James et al. (1961, p. 18-19) described the Six Mile Lake Amphibolite as follows. “In outcrops the amphibolite is a dark almost black massive fine- to medium-grained rock in which hornblende is the major constituent. In thin section feldspar (andesine or oligoclase) is abundant; in hand specimens it is less noticeable but gives a faint salt-and-pepper appearance to the rock, especially on surfaces broken across the foliation. Compositional layering is evident in some places, but in general the rock is homogeneous. The more southerly outcrops approach banded gneiss in character. Foliation parallel to the compositional layering is generally present, but may be subordinate to a strong linear structure that characterizes the rock. The lineation, which is marked by orientation of hornblende needles, plunges eastward at a low angle. In almost every outcrop the amphibolite is cut by dikes, pods, or irregular bodies of younger pegmatite.” Structure- The Dickinson Group appears to form a thick south-facing monocline that dips vertically to steeply southward. Virtually all the numerous stratigraphic top determinations from cross bedding in the East Branch are south facing and show no indication of fold repetition (James et al., 1961). Such determinations are lacking in the Solberg Schist, but based on the presumed conformable relationship to the East Branch, at least the lower portions of the Solberg seem likely to be south-facing. A series of magnetic anomalies, most notably that produced the Skunk Creek Member, can be followed for many kilometers as markedly straight traces, 19 �indicating that large scale folds within the Solberg are lacking on the scale of the James et al. study and that it, like the East Branch, is entirely south-facing. On a smaller scale, a strong schistosity is developed in both units and outcrop-scale folds with shallow plunges are fairly common. Whether these features reflect a much larger structure, only partly preserved, or are the largest-scale structures that formed as the units were rotated to their near-vertical orientation is a matter for speculation with our present understanding of the region. Correlation and tectonic setting of deposition- The radiometrically indicated age range of the Dickinson Group suggests that it is a possible time-correlative of the Menominee and/or Baraga Groups exposed nearby and across the southern Lake Superior region. Volcanic rocks interlayered with iron-formations of the Menominee Group were formed at 1.875 Ga (Schneider et al., 2002) and a diabase sill intruded into iron-formation in the Marquette iron range has been dated at about 1.890 Ga (Peitrzak-Renaud and Davis, 2014). The ejecta layer from the Sudbury impact at 1.850 Ga lies near the upper contact of the Menominee Group and the base of the overlying Baraga Group (Cannon et al., 2010). Post tectonic granite plutons that intruded the Baraga Group at 1.833 Ga (Schneider et al., 2001) provide a minimum age. It is possible, therefore, that the Dickinson Group was also deposited during or slightly before the MenomineeBaraga sequence, but its lithology and apparent tectonic setting are unmatched by any other sequence in the region. The combined stratigraphic thickness of 2-3 kilometers of predominantly clastic sediments with interlayers of mafic volcanic rocks argues for deposition in a rapidly subsiding basin in which a generally fining-upward sediment sequence was deposited. The East Branch Arkose, based on lithology and bedding features, seems clearly to be fluvial. Yet, the occurrence of banded cherty iron-formation in the Solberg Schist argues that the basin evolved into marine conditions by the later parts of deposition of that formation. It seems reasonable that the Dickinson Group was deposited in an extensional rift basin in the back-arc basin phase of the Penokean orogeny at about 1875 Ma as defined by Schulz and Cannon (2007). Stop 4. East Branch Arkose. (46.044o N, 87.840o W) Good exposures of the East Branch Arkose can be seen just west of Spring Hill Road. Please respect private property immediately south of the area examined by this trip. The area was mapped in detail (James et al., 1961) and a portion of the map is reproduced in Figure 14. All lithologies are exposed here including coarse, cross-bedded arkose (Figure 15A), quartzite pebble conglomerate (Figure 15b, c, d) and massive basalt. They can be seen in sequence along a short south-to-north transect. All units have nearvertical dips and face south. Slightly to the west, a set of diabase dikes cuts the units at a low angle. Of interest in the conglomerates is the great preponderance of pebbles of white to pink quartzite, a small percentage of which have well preserved bedding. The only known source for such clasts in the region is the Sturgeon Quartzite, which is the basal unit of the Paleoproterozoic sequence in most of the area. Pebbles of granite are common but make up only a small percentage of the pebble-sized clasts. They were used by James et al. (1961) as further evidence of an Archean source, but considering the new age for the porphyritic red granite of 2.1 Ga, the clasts could be from that unit rather than the Archean. Quartzite pebbles typically have flattened shapes that are more likely to be original shapes rather than caused by deformational flattening. A suggestion of imbricate structures can be seen locally. The abundant sand-sized grains of K- 20 �feldspar in the arkosic beds further indicate a granitic (Archean or porphyritic red granite?) source for much of the formation. Figure 14. Map of the outcrop area of East Branch Arkose from James et al., (1961). Darker shades are areas with abundant outcrops; lighter shades are covered. The approximate transect for the field trip is shown near the east edge of the area and is about 150 meters west of Spring Hill Road. 21 �Figure 15. Photographs of the East Branch Arkose. A- Cross-bedded coarse-grained arkose. BConglomerate consisting mostly of quartzite pebbles. Imbricate pebbles indicate current flow from right to left. C- Conglomerate with clast of quartzite with relict bedding. D- Close-up of conglomerate containing both quartzite and granite pebbles. Two samples of the basalt (amphibolite) collected along strike have very similar tholeiitic basalt compositions (~7% MgO, ~50% SiO2) with intermediate TiO2 (~1.4-1.9%) and FeOtotal (~12.5%). The trace elements are characterized by moderately enriched light rare earth elements (REE) and no negative Nb-Ta anomaly when normalized to primitive mantle (Figure 16). A sample of one of the dikes cutting the Archean granitic gneiss at Stop 2 has a trace element content identical to the basalt in the East Branch Arkose except for an enrichment in Th (Figure 16B); it also has higher SiO2 (~54%) than the basalt and may have been contaminated by the granitic gneiss. An amphibolite sampled in a road cut on the west side of Felch also has a composition very similar to the basalt in the arkose (Figure 16). The East Branch mafic rocks are very similar in composition to basalts found in continental rifts. 22 �Figure 16. Chondrite normalized REE plot (A) and primitive mantle normalized trace element plot (B) of basalt in the East Branch Arkose, a metadiabase dike in Archean granitic gneiss at Stop 2, and an amphibolite (sill?) from a road cut on the west side of Felch. Stop 5. Solberg Schist. (46.04oN, 87.82oW) This roadside outcrop on west side of County Road 581 is representative of much of the Solberg Schist. The rock is interlayered amphibolite and biotite-garnet schist with a prominent near-vertical foliation. The unusual ribbed appearance of the outcrop surface may reflect subtle bedding although lithologic changes across the ribs are not obvious (Figure 17B). If they are bedding, the outcrop displays a westward-plunging antiform. Radiometric data on detrital zircons are reported by Craddock et al. (2013) from a location whose coordinates correspond to this outcrop. The spectrum of ages ranges from about 3.8 Ga to about 2.1 Ga. The significant peak of ages at 2.1 Ga places a maximum age on the unit. Unfortunately the sample was described as East Branch Arkose by Craddock et al. (2013), so there is some uncertainty as to where the dated sample was collected and what it represents. This outcrop, and many others, show intense, small-scale deformation structures and commonly nearvertical foliation and bedding. Although these suggest that the unit is complexly deformed and may include significant repetition of stratigraphy, the disposition of the Skunk Creek Member is enigmatic in that it has a nearly straight outcrop trace for more than 20 kilometers (see Figure 2) and shows no indication of fold repetition. 23 �Figure 17. Outcrop photographs of the Solberg Schist. A-Well-bedded schist with interlayers of micaceous schist and more quartzose layers. Thin stringers of granite in upper half. BShallowly-dipping beds (?) cut by vertical foliation. Compositionally, the Solberg mafic schist ranges from basalt (~6-12% MgO, ~45-50% SiO2) to andesite (~4% MgO, ~56% SiO2) with relatively high TiO2 (~1.7-2.4%) and FeOtotal (~13-16%). Samples have steep light REE-enriched chondrite normalized patterns and negative Nb-Ta anomalies when normalized to primitive mantle (Figure 18). They are compositionally distinct from the amphibolites in the East Branch Arkose and the Six Mile Lake Amphibolites. Figure 18. Chondrite normalized REE plot (A) and primitive mantle normalized trace element plot (B) for samples of the Solberg Schist. Field for basalt of the East Branch Arkose and related amphibolites (shown in Figure 16) shown for comparison. 24 �Stop 6. Six Mile Lake Amphibolite. (46.02oN, 87.846oW) A small exposure on the west side of Wickman’s Marsh Road is typical of the Six Mile Lake Amphibolite. Here it is a rather uniform schistose amphibolite with small deformed granitic stringers. It is cut by two undeformed pegmatite dikes (Figure 19) The Six Mile Lake Amphibolite was originally described and named by James et al. (1961), as summarized above, who placed it as the uppermost formation in the Dickinson Group and ascribed an Archean age. As also discussed above, that is now in question because of 2.1 Ga detrital zircons in the lower part of the group. At least a portion of the rocks included in the Six Mile Lake by James et al. (1961) are intruded by granitic rocks of Archean age and appear to grade southward into banded gneiss the makes up much of the Archean in that area. Whether these latter amphibolites are truly a part of the Six Mile Lake or are an older amphibolite unit that lies adjacent to the Six Mile Lake is a subject for further evaluation. Compositionally the Six Mile Lake Amphibolite is a tholeiitic basalt characterized by low TiO2 (<1.5 wt. %) and trace element content. Unlike most of the amphibolites in the region, which are characterized by enriched light REE and negative Nb and Ta anomalies when normalized to primitive mantle, the Six Mile Lake Amphibolite has a flat chondrite normalized REE pattern and no Nb and Ta anomalies (Figure 20). It should be noted that the large metagabbro body in the Solberg Schist has a similar composition to that of the Six Mile Lake Amphibolite (Figure 20). Amphibolites from the Carney Lake Gneiss complex and the Hardwood mafic gneiss also have compositions similar to the Six Mile Lake Amphibolite (Figure 20). Figure 19. Six Mile Lake Amphibolite. A.- Hornblende schist with granitic stringers. B.- Schist cut by pegmatite dikes. 25 �Figure 20. Chondrite normalized REE plot (A) and primitive mantle normalized trace element plot (B) for samples of Six Mile Lake Amphibolite and related rocks. Fields for some other amphibolites from the region shown for comparison Stop 7A. Mafic granulite of the Hardwood Gneiss. (45.969oN, 87.711oW) Roadcuts on both the north and south sides of Highway M-69 are good examples of typical granulite of the Hardwood Gneiss. The rocks consist of various assemblages of hornblende, pyroxene, plagioclase, microcline, and garnet. They are very strongly foliated (Figure 21A) and, in places, show prominent compositional layering (Figure 21B). The Hardwood Gneiss was recognized and named by James et al. (1961) as an unusual unit of very highly metamorphosed rocks that are exposed over an area of only about 5 square kilometers at the eastern edge of their study area (see Figure 2). To the south, the Hardwood is in fault contact with the Paleoproterozoic Michigamme Formation. To the west and north it is in contact with Archean granite and granitic gneisses, but the nature of the contact is not known. The area of exposure is bounded on the east by Cambrian and younger strata. It is quite possible that the Hardwood underlies a substantially larger area beneath that cover. The general structure of the Hardwood in the area of exposure is the keel of a gently east-plunging synform defined by the gneissic foliation, so it is likely that the gneisses continue eastward in the pre-Cambrian basement. James et al. (1961, p.22-23) described the Hardwood as follows “In general, the gneiss is strongly layered, with individual layers ranging from a fraction of an inch to a few feet in thickness. The dominant rock type is a dark-medium-grained gneiss composed of hornblende, plagioclase, and pyroxene. Interlayered with this rock are beds of dark gneiss but of different grain size, beds of dark vitreous-lustered rock with alternating light and dark laminae, garnetquartz-mica schist, and light-colored rock that resembles quartzite. Some of the layers are rich in magnetite. …. The gneiss appears to have been originally a volcanic sequence, at least in part tuffaceous, with some inter- bedded sedimentary rocks, intruded by gabbro sills. The rocks have been dynamothermally metamorphosed under conditions that resulted in the alteration of most of the original pyroxene in the igneous facies to hornblende and garnet, and the development of mica, hornblende, and plagioclase in rocks that appear to have originally been acidic volcanics. The Hardwood gneiss, as seen in the exposures, is folded along axes that plunge eastward at low angles, and the general structure appears to be an eastward plunging syncline.” 26 �Figure 21. Hardwood Gneiss at stop 7A. A- Typical mafic granulite with strong, somewhat anastomosing, foliation. B- Straight-banded granulite gneiss. Compositional layering is expressed as variations in plagioclase:mafic mineral ratio. C and D are photomicrographs of two contrasting textures. C- Granoblastic textured interlayered quartz-microcline-plagioclase rock (center) and hornblende with minor (ortho?)pyroxene. D- Foliated rock with pyroxene (larger light grains) and small garnets (dark layer near top) in quartz-sericite matrix. Additional lithologies that have been included in the Hardwood gneiss are metasediments (Stop 7B), including quartzite and pelitic schist. The schists include garnet-biotite-sillimanite-kyanite mineral assemblages. Chemical analyses of mafic rocks in the Hardwood Gneiss are generally similar to those of the Six Mile Lake Amphibolite with similar low TiO2 and trace element content (Figure 22). One mafic gneiss sample has enriched light REE and a large negative Nb-Ta anomaly is likely the result of contamination by felsic crustal rocks. 27 �Figure 22. Chondrite normalized REE plot (A) and primitive mantle normalized trace element plot (B) for mafic gneiss samples from the Hardwood Gneiss. Field for Six Mile Lake Amphibolite shown for comparison. Metamorphism of the Hardwood Gneiss was studied by Peterson and Geiger (1990) who determined conditions of its metamorphism based on mineral assemblages. They defined two distinct episodes of metamorphism. Geothermobarometry indicates conditions of 8.2-11.6 kbar and ~770°C for the earliest event, and conditions of 6.0-10.1 kbar and 610-740oC for the second. They proposed that the older event was Archean and contemporaneous with a high-grade metamorphic event recorded in the Minnesota River Valley. They interpreted the younger event as probably Paleoproterozoic and pre-Penokean, with metamorphic conditions more intense than those generally ascribed to the Penokean orogeny in Michigan. Although they recognized both events in the typical layered gneisses, only the second event is recorded by the metasedimentary units. These extremely high metamorphic pressures are unique in the southern Lake Superior region and are comparable to or slightly higher than those of the Kapuskasing structure in Ontario. For comparison, the metamorphic pressures interpreted for the Peavy node, the nearest Paleoproterozoic metamorphic node, are less than 5 kbars (Attoh and Klasner, 1989). The metamorphic conditions recorded by the Hardwood Gneiss are equivalent to temperatures and pressures of the lowermost crust. The Hardwood Gneiss has been widely accepted to be of Archean age and to be part of the gneiss complex that forms the basement for the Paleoproterozoic sedimentary and volcanic sequence, although no radiometric ages have been previously determined. Our new SHRIMP UPb zircon data reveal a group of concordant to nearly concordant zircon spot analyses at ca. 2750-2500 Ma (Ayuso et al., 2018) documenting a Neoarchean component (Figure 23). A second group of spot analyses (Figure 23) document a younger period of zircon growth ca. 1900 to 2200. A thermal event of this age has not been recognized previously in the region. Metamorphism of the Peavy node, which encompasses the area of the Hardwood Gneiss, has been dated at ca. 1830 Ma at outcrops within a few kilometers west of the Hardwood (Schneider et al., 2004; Holm et al., 2007) using Pb-Pb ages of monazite. Furthermore, metamorphic pressures of the Peavy node are estimated at ca. 5 kb (Attoh and Klasner, 1988) in contrast to the much higher pressures determined for the Hardwood (Peterson and Geiger (1990). An additional significant feature of the Hardwood Gneiss is that we have detected no Eoarchean components, unlike the nearby Carney Lake Gneiss, which has yielded numerous Eoarchean spot analyses. Thus, the Hardwood Gneiss is presently enigmatic in terms of its parentage, metamorphic history, and kinematics of emplacement relative to surrounding rock units. 28 �Figure 23. A- BSE (back-scatter electron) image of two anhedral zircons from the Hardwood Gneiss showing SHRIMP U-Pb analyzed dates for two spots. B- Concordia diagram for 56 spot analyses for the Hardwood Gneiss. Stop 7B. Pelitic schist of the Hardwood Gneiss. (45.967oN, 87.729oW) A low roadcut on the west side of Swan Peterson Road, just north of its intersection with M-69 is garnet- biotitesillimantie-kyanite schist. These are among the structurally (stratigrahically?) lowest parts of the Hardwood Gneiss. Garnet porphyroblasts are common (Figure 24A). A strongly developed foliation dips about 30 degrees east. White elongate masses (Figure 24B) consist of quartz and both kyanite and sillimanite. They are elongated within the foliation but have slightly flattened shapes in cross section (Figure 24B). They define a prominent lineation that plunges about 35° east. Figure 24. A- Photomicrograph of biotite schist with porphyroblasts of eudedral to subhedral garnet. B- Hand sample of schist. Upper surface is the schistose parting and shows elongate somewhat contorted white masses of quartz and aluminosilicates, both sillimanite (fibrolite) and kyanite. Cut front face shows that these are somewhat flattened, rod-like masses that define a prominent lineation along fold axes. 29 �Stop 8. Cambrian/Paleoproterozoic unconformity. (45.997oN, 87.841oW) Roadcuts on the north side of Highway M-69 west of the village of Felch expose the unconformity between the Munising Formation (Cambrian) and Paleoproterozoic strata of the Vulcan Iron-formation and Randville Dolomite. Along most of the roadcut nearly flat-lying sandstone of the Munising Formation is at the base of the exposures (Figure 25A), but the higher hills to the north are underlain by Paleoproterozoic rocks indicating that substantial topography existed along the unconformity. Near the west end of the outcrop (Figure 25B) the unconformity is a steeply eastward dipping contact across which the flat-lying Munising abuts against a topographic knob composed of clasts of Vulcan Iron-formation, some of which have indications of secondary iron enrichment (Figure 25C). Munising sand fills the voids in the rubble. The material appears to have been a loosely packed pile of iron-formation rubble at the time of Cambrian transgression. Voids were filled with sand that infiltrated the pile. The angular shape of clasts indicates that material was not transported any substantial distance. The material may have been talus accumulated along the base of a south-facing slope. Farther east along the roadcut the lower part of the exposures are again angular rubble of both Vulcan Iron-formation and Randville Dolomite with voids filled with Munising sand. Exposures higher on the hill just to the north are entirely the Paleoproterozoic units indicating that the rubble is a thin carapace of material lying against a steep south-facing slope that existed during Cambrian transgression. In some places the rubble has intervals of well-bedded sandstone within it (Figure 25D); an indication that the rubble was being delivered to the site during sandstone deposition. These relationships suggest that the rubble of Paleoproterozic rocks was a talus deposit along the base of a marine shoreline cliff during deposition of the Munising Formation. The relationships seen here between Precambrian rocks and the Cambrian sandstone emphasize that considerable topography existed on the surface over which the Cambrian seas transgressed across the area. Similar relationships are seen on Field Trip 3 along the Menominee Range. The relationships also suggest that outliers of Cambrian sandstone may be much more widespread under low areas than has been shown on previous maps, and that outcrops of Precambrian rocks on the present land surface were likely only at shallow depth beneath the unconformity before being exposed by younger erosion. Much of the topography presently seen over Precambrian areas is likely to be largely relict topography of the Cambrian landscape. 30 �Figure 25. A- Flat-lying red sandstone of the Munising Formation (foreground) passing laterally into rubble of Vulcan Iron-formation at far left. View looking northwest. B- Steeply dipping unconformity shown by dashed line between Munising Formation (right) and iron-formation rubble (left). C- Clast of specularite-jasper iron-formation within rubble surrounded by Munising sand. D- Rubble of iron-formation and dolomite with thin, nearly horizontal interbeds of Munising Formation. Stop 9. Randville Dolomite and post-Cambrian breccia. Along highway M-69, near the intersection with County Road 11, a roadcut on the north side of the highway exposes coarsely recrystallized carbonate and quartz. We interpret this to be the Randville Dolomite which, in previous mapping (James et al., 1961), was traced to within about 500 meters of this newer roadcut. An unusual feature of this exposure is a zone of breccia about 10 meters wide composed of fragments of both the Randville and of the Munising Formation. The Munising Formation occurs as coherent clasts of red sandstone (Figure 26), not unconsolidated sediments as at Stop 8, indicating that the breccia formed sometime after deposition and lithification of the Munising Formation. The Randville Dolomite occurs as angular clasts as large as about a meter diameter. The cause of this late brecciation is not clear. One possibility is that it is a karst collapse breccia formed by solution of the Randville. The Munising fragments must have been transported downward at least a few meters assuming that the Cambrian unconformity was originally slightly above the present land surface. Another possibility is that the breccia is related 31 �to kimberlite intrusion, although we have found no indication of igneous rocks in the breccia matrix. Post-Ordovician kimberlites are known at numerous localities in the region and most have clasts of Ordovician carbonates that have fallen downward in the pipes. (See discussion in Field Trip 2 of this volume). Other areas of disturbed Cambrian sediments have been described in the area (James et al., 1961). The Munising has been observed with dips as high as 60o locally. Most of these disturbed areas are along Precambrian faults and were interpreted to be caused by Paleozoic or later reactivation of those faults. The faults were inferred and, in one instance documented, to have normal offset. As mapped by James et al. (1961), a fault does pass about 100 meters south of this exposure so there is some possibility that the brecciation seen here is simply related to reactivation of that fault. . Figure 26. Sandstone of the Munising Formation with steep eastward dip in carbonate breccia. Stop 10. Banded gray gneiss. (46.004oN, 87.900oW) The term “banded gray gneiss” was applied informally to a belt of Archean rocks that occur mostly immediately north of the Felch trough. Roadcuts on both sides of Highway M-69 display representative lithologies of this unit. (Note that the present position of M-69 is substantially different from that shown in James et al, (1961) because of relocation related to the Groveland mine development.) This exposure is migmatitic gneiss composed mostly of mafic material with minor granitic stringers (Figure 27A). Most layering is nearly vertical and numerous tight isoclinal folds are present (Figure 27B). An unusual feature is a large inclusion of coarse-grained plagioclasehornblende gneiss shown on the south cut. The inclusion has a strong foliation that lies at a distinctly lower angle than that of the surrounding gneiss. 32 �Figure 27. A- Typical amphibolitic gneiss with numerous granitic stringers. Many stringers are undeformed and are apparently post- or late-tectonic injections or segregations. B- Tightly folded migmatitic gneiss. The general description of the banded gray gneiss in this vicinity, given by James et al. (1961, p. 121-122) is as follows: “Most of the gneiss is light gray, or alternating light and dark gray, and is thinly layered. The rock typically contains thin rather discontinuous biotitic and hornblendic layers, not more than a millimeter thick, alternating with quartz-feldspar layers several millimeters thick. Both on fresh breaks and on weathered surfaces the gneiss is somewhat mottled as a result of segregation of feldspar into patches a few millimeters across and 5 or 6 mm long. Near some dikes of granite pegmatite, the gneiss contains layers a few millimeters wide composed of pinkish feldspar, some of which cut across the gneissic layering. In some places, particularly near the southern margin of the gray gneiss belt, the gneiss contains lenses and pods of amphibolite as much as 100 feet thick and a quarter of a mile or more long. The light-colored layers of the gneiss are composed chiefly of feldspar, quartz, and biotite. The texture is exceedingly irregular and the abundance of the various constituents varies widely from place to place. The larger grains quartz, plagioclase, and microcline vary in size from about 0.5 mm to 1.5 mm. The microcline, which usually occurs as smaller crystals than the other constituents, is unaltered. The plagioclase in two specimens is oligoclase, whereas in a third it is probably albite. Other minerals, present in small quantities, are muscovite, chlorite, epidote, zircon, leucoxene, apatite, and iron oxides. The amphibolite that makes up the dark layers and pods in the banded gray gneiss is virtually identical to the Six-Mile Lake amphibolite previously 33 �described. The rock is medium grained, with strong preferred orientation of the minerals. Hornblende, plagioclase, and quartz are the chief constituents.” An analysis of a sample of the gray gneiss from this road cut indicates it is basaltic with ~50% SiO2, ~5% MgO, and relatively high TiO2 (~1.9%) and FeOtotal (~14.5%). The trace elements are characterized by enriched light REE and negative Nb-Ta anomaly when normalized to primitive mantle (Figure 28). The mafic gray gneiss composition overlaps with that of the Solberg Schist (Figure 28) although the Solberg is likely of Paleoproterozoic age. Figure 28. Chondrite normalized REE plot (A) and primitive mantle normalized trace element plot (B) for mafic gray gneiss. Field for Solberg Schist is shown for comparison. Stop 11. Porphyritic red granite. (46.009oN, 88.061oW) Roadcuts on both the east and west sides of Highway M-95 are examples of typical porphyritic red granite, which was recognized as a map unit by James et al. (1961). The porphyritic red granite is a ferroan potassium-rich granite with A-type within-plate chemical characteristics. As mapped, it occurs in two elliptical bodies as shown on Figures 2 and 29. However, outcrops in the area are very sparse and the margins of the granite are poorly constrained. Likewise, its relationship to surrounding units is not clear. The surrounding area was designated “Dickinson group undivided” based largely on the westward extension of magnetic anomalies, mostly in the Solberg Schist, from areas of better exposure to the east. Because the magnetic anomaly produced by the Skunk Creek Member, a medial bed in the Solberg Schist, passes to the south, it is likely that the porphyritic red granite is surrounded by the lower parts of the Dickinson Group. Whether the granite bodies are intrusive into the Dickinson Group or are domes of pre-Dickinson basement is not clear from available exposures. The foliation is approximately parallel to the inferred margins of the granite bodies and dips steeply outward from their centers indicating that they are domal structures. As described by James et al. (1961), “The rock is generally homogeneous and coarse grained. Inclusions are rare, but dark schlieren and compositional layering locally are present. . . . Lenticular to tabular pink feldspars about half an inch long are abundant and impart a porphyritic appearance to the rock. Some of the feldspars are euhedral, but most are in fact augen, and on horizontal surfaces a foliation produced by oriented feldspars is faint to distinct. In vertical sections the structure is easily seen. In most exposures, steeply plunging lineation, 34 �marked by orientation of both microcline augen and biotite, is well developed, and in places it is the dominant structure. The feldspars form about two-thirds of the rock; quartz and biotite make up the remaining onethird. Quartz itself makes up 10 to 20 percent of the rock. The cores of many of the microcline phenocrysts are white or colorless and are mantled with feldspar that is salmon pink or red. In thin section the microcline crystals have indefinite borders against a finer grained aggregate of microcline, oligoclase that is reddish and kaolinized, and quartz. The relationships suggest granulation of the borders of original tabular microclines followed by recrystallization of the granulated material.” The preferred interpretation of James et al. (1961) was that the granite is pre-Dickinson group and therefore likely to be of Archean age, based on their assignment of an Archean age for the Dickinson Group. A new radiometric age (SHRIMP U-Pb data for zircon) determined for the granite (Ayuso et al., 2018) is ca. 2.099 Ga. (Figure 30). This well constrained age is rather surprising in that no other granites of comparable age are known in the region. Granites of this age may provide a local source for the rather abundant 2.1 Ga detrital zircons reported from the nearby Dickinson Group (Craddock et al., 2013). The date further strongly suggests that the porphyritic red granite is the basement on which the Dickinson Group was deposited and was uplifted in domal structures after Dickinson deposition. Figure 29. A portion of Plate 2 west from James et al. (1961) showing two bodies of porphyritic red granite occurring as cores of domes surrounded by undivided strata of the Dickinson Group. Red dots are aeromagnetic anomalies and red lines are inferred connections of magnetic beds between measurement points, solid where probable, dashed where uncertain. 35 �Figure 30. A- BSE (back scatter electron) image of zircon grain from the porphyritic red granite showing age of spot analyzed by SHRIMP. B- Concordia diagram for 18 spot analyses of zircons from the porphyritic red granite. References Attoh, K. and Klasner, J.S., 1989, Tectonic implications of metamorphism and gravity field in the Penokean orogeny of northern Michigan, Tectonics, v. 8, p. 911-933 Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vasquez, J.A., and Jackson, J., 2017, Evidence for the presence of Eoarchean crust in northern Michigan, Institute on Lake Superior Geology, Proceedings of 63rd annual meeting, part 1: Program and abstracts, p. 910. Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., and Jackson, J., 2018, New U-Pb zircon ages for rocks from the granite-gneiss terrane in northern Michigan: Evidence for events at ~3750, 2750, and 1850 Ma: Institute on Lake Superior Geology, Proceedings of 64th annual meeting, part 1: program and abstracts. Cannon, W.F., Schulz, K.J., Horton, J. W., Jr., and Kring, D.A., 2010, The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan USA, Geological Society of American Bulletin, v. 122, p. 50-75. Clark, L.D., 1961, Chapter C, Precambrian geology of the Norway Lake area: in James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., Geology of central Dickinson County, Michigan, U.S. Geological Survey Professional Paper 310, p. 97-113. Craddock, J.P., Rainbird, R.H., Davis, W.J., Davidson, C., Vervoort, J.D., Konstantinou, A., Boerboom, T., Vorhies, S., Kerber, L., and Lundquist, B., 2013, Detrital zircon geochronology and provenance of the Paleoproterozoic Huron (~2.4-2.2 Ga) and Animikie (~2.2-1.8 Ga basins, southern Superior Province, Journal of Geology, v. 121, p. 623-644. 36 �Cumberlidge, J.T., and Stone, J.G., 1964, The Vulcan Iron-formation at the Groveland Mine, Michigan, Economic Geology, v. 59, p. 1090-1106. Holm, D.K., Schneider, D.A., Rose, S., Mancuso, C., McKenzie, M., Foland, K.A., and Hodges, K.V., 2007, Proterozoic metamorphism and cooling in the southern Lake Superior region, North American and its bearing on crustal evolution, Precambrian Research, v. 157, p. 106126. Hunter, J., 1986, Uranium mineralization at the Felch prospect, Upper Peninsula, Michigan, United States of America (summary): Uranium deposits in magmatic and metamorphic rocks, Proceedings of a technical committee meeting, Salamanca, p.213-215. Jackson, C. T., 1849, Report on the geological and mineralogical survey of the mineral lands of the United States in the State of Michigan, U.S. 31st Cong., 1st sess., S. Doc. 1, p. 371-935. James, H.L., 1955, Zones of regional metamorphism in the Precambrian of northern Michigan, Geological Society of America Bulletin, v. 66, p. 1455–1488. James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., 1961, Geology of Central Dickinson County, Michigan, U.S. Geological Survey Professional Paper 310, 176 p. Klasner, J.S., and Sims, P.K., 1993, Thick-skinned, south-verging backthrusting in the Felch and Calumet troughs area of the Penokean orogeny, northern Michigan, U.S. Geological Survey Professional Paper 1904-L, 28 p. Lehman, G.A., 1987, U-Pb dating of pitchblende from Dickinson County, upper Michigan, suggests reactivation of Precambrian structures during formation of the Michigan basin, Proceedings of the Institute on Lake Superior Geology, part 1, Proceedings and Abstracts, v. 33, p. 37-38 Nicholson, S.W., Shirey, S.B., Schulz, K.J., and Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development, Canadian Journal of Earth Sciences, v. 34, p. 504–520. Peterson, J.W., and Geiger, C.A., 1990, The Harwood gneiss: evidence for high P-T Archean metamorphism in the Southern Province of the Lake Superior region, Journal of Geology, v. 98, p. 273-281. Pietrzak-Renaud, N, and Davis D., 2014, U-Pb geochronology of baddeleyite from the Belleview metadiabase: age and geotectonic implications for the Negaunee Iron-Formation, Michigan: Precambrian Research, v. 250, p. 1-5. Romano, D., Holm, D., and Foland, K., 2000, Determining the extent and nature of Mazatzalrelated overprinting of the Penokean orogenic belt in the southern Lake Superior region, north-central USA, Precambrian Research, v. 104, p. 25–46, doi: 10.1016/S03019268(00)00085-1. 37 �Schneider, D., Bickford, M., Cannon, W., Schulz, K., and Hamilton, M., 2002, Age of volcanic rocks and syndepositional iron-formations, Marquette Range Supergroup: implications for the tectonic setting of Paleoproterozoic iron-formations of the Lake Superior region, Canadian Journal of Earth Sciences, v. 39, p. 999–1012, doi: 10.1139/E02-016. Schneider, D.A., Holm, D.K., O’Boyle, C., Hamilton, M., and, Jercinovic, M., 2004, Paleoproterozoic development of a gneiss dome corridor in the southern Lake Superior region, USA, in Whitney, D.L., Teyssier, C., and Siddoway, C.S., eds., Gneiss domes in orogeny, Geological Society of America Special Paper 380, p. 339–357. Schulz, K.J., and Cannon W.F., 2007, The Penokean orogeny in the Lake Superior region, Precambrian Research, v. 157, p. 4-25. Ueng, W.C., and Larue, D.K., 1988, The early Proterozoic structural and tectonic history of the south central Lake Superior region, Tectonics, v. 7, p. 369-388. Vallini, D.A., Cannon, W.F., and Schulz, K.J., 2006, Age constraints for Paleoproterozoic glaciation in the Lake Superior Region: detrital zircon and hydrothermal xenotime ages for the Chocolay Group, Marquette Range Supergroup, Canadian Journal of Earth Sciences, v. 43, p. 571-591. 38 �FIELD TRIP 2 Tuesday May 15, 2018 GEOLOGY OF THE HEMLOCK FORMATION Tomas Waggoner, Consulting Geolo Email: thomaswaggoner@hotmail.comgist INTRODUCTION The one day field trip will make eleven stops to examine the major rock types that make up or impact the 1,874 Ma Paleoproterozoic Hemlock Formation (Figure 1). The ~30,000 foot thickness of the primarily tholeiitic basalt is particularly rich in iron oxides and could easily have provided both the iron and silica incorporated in major portions of the Lake Superior type iron formations. The first stop will examine the Lake Ellen Kimberlite which is the only easily accessible kimberlite in the Upper Peninsula Kimberlite District and it also intrudes the Hemlock Formation. The differentiated West Kiernan sill will be visited including the ultramafic lower unit and the base metal rich differentiate near the base of the gabbro unit. The upper transition will be examined where plagioclase levels approach 80% and contains significant titaniferous magnetite, stilpnomelane and apatite. Other stops will illustrate rhyolite, volcanic conglomerates, amygdaloidal and pillowed basalts. Also, one of the stops will examine the Mansfield iron member which is one of several areal restricted iron formations present in the Hemlock. 39 �General Geology The ~2.7 Ma Margeson Creek Gneiss in the center of the Amasa Uplift is overlain by the Randville Dolomite of the Chocolay Group and is equivalent to the Kona Dolomite on the Marquette Range and the Badriver Dolomite on the Gogebic Range. Dating by Vallini and others (2006) indicates the age of the Sturgeon quartzite under the Randville dolomite in the Iron Mountain area is 2.32.2 Ga. The Randville dolomite directly underlies the Hemlock Formation and exhibits several lithologies. The most pervasive lithologic type is a sandy dolomite while a coarse orange arkose, not present in the Randville equivalents elsewhere, is common at the southern portion of the Amasa Uplift. The age of the Randville is at least 300 million years older than the 1874 Ma Hemlock Formation (Schneider and others, 2002) . The basal conglomerate and quartzite units present in the Chocolay Group of the Marquette Range and Menominee Range are absent around the Amasa Uplift. At the south end of the Amasa Uplift the Randville is approximately 1800 feet thick and consists of sandy dolomite, feldspathic quartzite, arkose and sericite argillite. Some authors have postulated that the Randville has pinched out on the north end (Cannon and others, 1976; Foose, 1981). However, subsequent drilling in section 28, T. 46 N., R. 33 W indicates the dolomite is present on the north end, but that same drilling was not extensive enough to identify lithologic types or thickness. Overburden depths on the north end of the Uplift can vary from 50 feet to over 560 feet with no reported outcrops. During the early Penokian orogeny the Pembine-Wausau terrane (Figure 2) was already accreted to the Superior Provence by 1875 Ma (Figure 2). Back-arc extensional basins south of the Niagara fault contains numerous volcanogenic massive sulfide deposits (VMS). Rifting on the continental margin produced a basin(s) into which banded iron formations formed coeval with volcanism. The age of banded iron formation deposits and the intra-arc rift massive sulfide deposits in the 40 �Pembine-Wausau terrane are both ~1874 Ma. One of the largest volcanic centers is the Hemlock Formation, located in Iron County, Michigan. The volcanic pile achieved a thickness in excess of 30,000 feet west of the Amasa Uplift and thins away in all directions (Figure 3). It is estimated the footprint exceeds 2,800 square miles. Geophysics A total field positive magnetic anomaly around the west side of the Amasa Uplift is caused by increased magnetite content in the upper 6000 feet of the Hemlock basalts (Figure 3). The Amasa Iron-formation is essentially non-magnetic as primary iron minerals have all been oxidized to hematite and in some cases enriched to iron ore that was mined early in the last century. Major conductivity zones are associated with graphitic slates internal to the volcanics and the overlying graphitic Michigamme slates. The presence of disseminated sulfides in the lower portion of the normal gabbro of the West Kiernan sill produce minor EM anomalies. Figure 3. Increase in iron content in basalt near the top of the Hemlock Formation 41 �Figure 4. Plan geology map showing the sub crop of the Hemlock Formation (Xm2) and the later thrust faulting. Field trip area is defined by the red box. (After USGS Map I-2356) Generally the exposed Archean portion of the Amasa Uplift exhibits low gravity readings. Most of the Hemlock sub crop area exhibits a muted gravity contrast. Positive gravity values are associated with the Riverton Iron-formation in the Iron River-Crystal Falls allochthon and an area northwest of the Amasa Uplift. A Bouguer ground gravity map with limited field stations of the Amasa Uplift area show an increase in the gravity field toward areas of thicker basalts of the Hemlock Formation, especially where the increased magnetite content increased the specific gravity of the rock. After deposition of the Michigamme Formation the continued northern push by the Wisconsin Magmatic terrane produced a number of east-west thrust fault panels replicating the stratigraphy in each panel (Figure 4). Hemlock Formation The Hemlock Formation was named (Clements, 1899) for volcanic rocks found near the Hemlock River west of the Amasa Uplift. The Hemlock belong to the Menominee Group of the Marquette Range Supergroup. The volcanics and sediments were deposited subaqueously and are believed to be terrigenous sedimentary sourced (Johnson, 1975; Dann, 1978; Ueng, 1987; and Beck, 1991). Beck (1991) described the volcanics as continental flood basalts. Principle lithologies listed in order of decreasing abundance are: basalt, vocaniclastics (referred to as agglomerates, hyaloclastites and breccias), rhyolite, graphitic slates and small iron formation members (one, the Bird, is an iron oxide chert and the other, the Mansfield, is a carbonate chert). 42 �The base of the Hemlock Formation rests unconformably on the Randville Dolomite except in the vicinity of Michigamme Mountain where the base of the Hemlock rests on a small deposit of unique quartzite composed of angular quartz and chert clasts along with minor massive chert. Portions of the quartzite have been replaced by magnetite and specularite while some of the chert contains secondary specularite that replaced the chert yielding a crude banding appearance. The Hemlock tholeiitic basalt magma (Figure 5) by its reduced nature, concentrated iron in the residual magma, principally as magnetite. Replacement iron oxides are found in the small clastic unit located on the southeast portion of the Amasa Uplift at the base of the Hemlock Formation. Iron oxide concentrations, principally magnetite, are found sporadically in the basaltic portion of the Hemlock and in the upper ~6000 feet of the basalts. Following active volcanism banded iron formations were formed (Amasa [west side of Uplift] and Fence River [east side of the Uplift]). There are a number of intrusive sills found within the Hemlock Formation. One, the West Kiernan sill, is approximately thirteen miles long and one mile thick as measured at outcrop. It is differentiated into four distinct units plus a chilled diabase contact zone. Relatively thin, but internally layered sills like the West Kiernan sill, consisting of basal peridotite overlain by pyroxenite, gabbro, and granophyre, are unusual in that injection of relatively crystal-free magma does not appear to form well layered, well-differentiated intrusions (March, 2006). For example, the large Sudbury impact melt sheet (3 km x 200 km; volume of ~30,000 km3) shows no sign of layering and very little sign of differentiation while the more than 300 m thick Palisades sill consists only of an olivine-enriched lower layer (~3 m thick), diabase, ferrodiabase, and granophyre (Walker, 1969). However, similarly layered and differentiated sills like the West Kiernan sill have been described from other Precambrian terranes including the Archean Vermilion district in Minnesota (Schulz, 1982), the Abitibi greenstone belt in Ontario (MacRae, 1969), the Eastern Goldfields region in Western Australia (Williams and Halberg, 1973) and the 43 �Barberton Mountain Land in South Africa (Anhaeusser, 1985). For the Archean examples, the layered sills appear to reflect crystallization from an iron-rich picrite magma (Schulz, 1982). Whether the West Kiernan sill is also related to a high-iron picrite magma is unclear. Until the late1980’s the volcanic extrusives (Badwater Greenstone) around the Iron River-Crystal Falls District were considered to be a distinct and separate mafic extrusive unit from the Hemlock. Several papers suggested that Badwater Greenstone was actually the same unit as the Hemlock based on similarity of rock types, textures and geochemistry (Dann, 1978, 1979). In the mid-1980’s several USGS personnel were looking south across the Paint River and noticed a pink colored rock beneath the greenstone and identified it as the Saunders Formation (Randville equivalent). This suggested that the entire Iron River-Crystal Falls district could be an allocthon (detachment fault) that had been forced northward up and over the underlying younger Michigamme Formation. In addition it was recognized that the Paint River Group was not a younger sequence than the Baraga Group but rather a fault repetition with the slates over the Riverton Iron Formation equivalent to the Michigamme slates (Figure 6). Regional metamorphism produced during the Penokean orogeny in the area of the field trip falls either in the chlorite or biotite isograd (Weir, 1986). Hemlock Formation as a Flux Source for portions of Iron Formations in the Lake Superior District Intermittingly, during and at cessation of active volcanism the magma source continued to yield significant quantities of iron, silica, phosphorus and carbon dioxide forming a large volume super plume (Figure 7) containing both instantly crystalized iron oxides and amorphous silica along with significant soluble Fe++ that spread in the ocean over thousands of square kilometers. Both 44 �magnetite and specularite were produced at the vent site(s) under equilibrium conditions that allowed instant crystallization, much like the sulfide particulate matter associated with black smoker vents. Very fine specular hematite is found at the core of much of the magnetite in banded iron formations worldwide (Han, 1966, 1978, 1988). The iron oxides (i.e. microplaty-specular hematite and magnetite) iron carbonate and amorphous silica settled to the sea floor (possibly on seasonal changes) forming alternating bands of chert and iron minerals with thickness based on flux available from the source, current directions and saturation conditions at the deposition site. Near shore banded iron formations exhibit granular textures, cross bedding and occasional reef development of stromatolites while deeper deposits are lithic without oolitic or granular layers or significant traces of biogenic activity. Ancient VMS (e.g. New Brunswick #3 and 6; Austin Brook; Manitowadge, Ontario; Lokken, Norway; Bending Lake, Ontario) and SED-EX (e.g. Little Commonwealth and Dunkel exploration, Wisconsin) deposits, with associated banded iron formations, have modern day analogues (without banded iron formations) at spreading plate centers where iron is precipitated as an oxyhydroxide ostensibly converted to hematite and magnetite at a later period. Oxyhydroxides do not convert to specularite or magnetite under normal natural conditions. Current sea floor discharge systems yield small sea floor deposits from small plumes under predominantly oxidizing conditions. Large systems capable of producing sufficient silica and iron for the creation of banded iron formations would require a supersized plume extending over large areas. Also, these large systems would create a significant reduced environment where early formed iron compound particles can scavenge additional iron. Sea floor ventings studied since discovery in the 1970’s are useful in observing the mechanisms in play in the formation of iron oxides and silica near the vent sites. Lake Nyos in Cameroon (Ozawa, 2016) illustrates the formation of crystalline siderite from a vent source high in Fe+2 and CO2 under a slightly acid environment. This would suggest that a plume environment at the bottom of the sea can quickly 45 �precipitate iron carbonate along with amorphous silica which can separate into chert and iron rich layers (Krapez, 2006). Conditions impacting the type and nature of the banded iron deposits produced include:  Plume buoyancy.  Hydrothermal water temperature, pH and Eh.  Amount of iron and silica flux including CH4, CO2, H2S and fluorine.  Discharge conditions including sea water temperature, salinity and pressure (depth).  Discharge and plume environment, oxidation and reduction conditions, plume size/thickness.  Size of the igneous flux source  Definable presence of iron and silica available to produce enough flux to form iron formations. FIELD TRIP STOPS Ten of the eleven stops are shown in Figure 8 below. Figure 8. Hemlock field trip stops 1-7 and 9-11 Kiernan and Lake Mary Quadrangle 46 �Stop 1. Lake Ellen Kimberlite UTM: 46o 10.478 N 88o 10.587 W, SWSW Sec. 27, 44-31 Gair (1956) in USGS Bulletin 1044, plate 2 noted a small magnetic anomaly in the SWSW of Section 27, T. 44N, R. 31 W. and made an observation that the magnetic high was caused by magnetite in the glacial till. In 1971 William Spence and Klaus Schulz discovered the Lake Ellen kimberlite (also referred to as Site 10) under thin soil cover while conducting base metal exploration for industry. Cannon (1981) wrote a report on the occurrence. Crystal Exploration acquired control of the property and conducted geophysics (Figure 9), drilling, trenching, bench testing and analytical work. This activity spurred other companies to conduct air and ground magnetics, soil and stream surveys, land leasing and follow-up drilling. Shapes of kimberlites are usually round or elliptical and can cover an area from a few acres and up to 200+ acres. The Lake Ellen kimberlite is approximately 600 feet in diameter (Figure 9). Most of the kimberlites in the district have a similar physical appearance (Figure 10). Depth of burial and degree of weathering can make geophysical prospecting difficult. The Lake Ellen pipe has a variable magnetic signature. Unweathered kimberlites contain minor magnetite which can be oxidized to hematite near surface and lose their magnetic properties. Where weathering produced clays or where epiclastic units are present the electromagnetic techniques can be an effective location tool. Inhomogeneity can mask certain parts of a single pipe. In addition to geophysical techniques, soil or stream sampling can be an effective method of discovery because the Lake Ellen is overlain by a thin soil cover. An excavation pit used for both bulk sampling and woods road construction constitutes the focus of the first stop. 47 �Figure 10. Kimberlite V-28-c-1 in Sec. 2, T. 38 N., R. 27 W. showing lapilli, xenocrysts and xenoliths of Ordovician limestone. The Lake Ellen kimberlite intruded the steeply dipping Hemlock rhyolite and Fence River Ironformation (Figure 11). Dating of inclusions found within diamonds mined in Australia, South Africa and Botswana (Kirkley, 1992) found the age of diamond formation was generally older than the kimberlites or lamproites themselves. It can be postulated that any diamonds present in the Upper Peninsula kimberlite district are also older. 48 �The known kimberlite district extends from the Michigamme Reservoir in the northwest to Powers in the southeast, a distance of 62 miles. The oval outline of the kimberlite district has a width of approximately 19 miles (Figure 12). Figure 12. Plan map of the Michigan Upper Peninsula Kimberlite District. A subset of purple chromium rich garnet xenocrysts can be used to project the potential of any kimberlite to contain diamonds. Orange, red-orange and light orange garnets are from eclogite xenoliths while purple, red and pink garnets fall within peridotite affinity. The Lake Ellen kimberlite has very few indicator garnets, 3 of a population of 178 (McGee, 1988) or 1.7% indicating a lower probability of containing diamonds. A 180 short ton bulk sample produced four diamonds which in quality and number make the Lake Ellen kimberlite uneconomic. None of the discovered kimberlites in the field have economic concentrations of diamonds. Xenoliths and xenocryst compositions indicate the diatreme originated in the upper mantle. Based on data obtained from garnets the calculated equilibrium temperatures range from 950o-1100o C. The kimberlite originated from approximately 140-160 km below the surface (Griffin, 2004). Age dating of Site 73 Kimberlite emplacement yielded a zircon age of 155 Ma while a K-Ar determination on phlogopite yielded a 190 Ma age indicating a Jurassic Period (138-205 Ma) emplacement. Granulite whole rock data suggest two age groups with affinities to tholeiitic basalts. Trace element analyses suggest one group is of Archean derivation while a second group is aligned with the Keweenawan extrusive rocks of the Mid-Continent Rift (Zartman and others, 2012). 49 �Stop 2a. Basal Hemlock Formation mineralized quartzite (Figure 13) UTM: 46o 9.545N 88o 11.087W, NENE Sec. 4, 43-31 Stop 2 will concentrate on Michigamme Mountain, a local topographic high (Figure 13). The principle rock at this location is a unique mineralized quartzite of limited areal extent. Gair (1956) named the quartzite Goodrich with the younger Hemlock overlying the quartzite. We now know that the age of the Hemlock (1.84 Ga) matches the age of the major iron formations in the Lake Superior region. The quartzite is composed of sub-angular quartz and granular fragments of chert. The quartzite overlies a thin massive chert. A later hydrothermal influx of specularite (Figure 14) and magnetite (Figure 15) associated with potassic and silica enrichment replacing some of the quartz and chert in the quartzite. Subsequent surface supergene oxidation converted some of the magnetite to martite. Soluble iron values are generally below 20% but can exceed 50%. At this location we are standing on an east-west structural anticline that may possibly be a chert dome. Stop 2b. Specularite chert from test pits. UTM: 46o 9.583N 88o 11.115W, SESE Sec. 33, 44-31 In the saddle between the two mineralized quartzite topo highs is a massive chert with secondary specularite (Fig. 13). You will note quite a bit of cross cutting specular hematite suggesting the iron oxide was emplaced in the chert at a later time. From Stops 2a & b you can see the highest elevation on Michigamme Mountain 80 yards to the southwest. The topographic high outcrop is quartzite with only minor magnetite and the occasional quartz vein containing micaceous selvages along the contacts 50 �Figure 14. Specularite chert from saddle on Michigamme Mountain Figure 15. Martite after magnetite with quartz/ adularia vein from Michigamme Mountain. Stop 2c. Hydrothermally enhanced rhyolite UTM: 46o 9.581N 88o 11.116W, SWSE Sec. 33, 44-31 While ascending Michigamme Mountain note the occasional black rock both to the north and underfoot. This is the same magnetite, now martite, quartzite to be seen at Stop 2a. The quartzite is overlain by a rhyolite flow representing the basal Hemlock Formation present along the eastern outcrop area of the Amasa Uplift. The rhyolite has quartz eyes and is typically pink on outcrop and orange on fresh surface (Figure 16 and 17). The orange color is primarily due to the addition of potassium feldspar. Chemistry of most of the rhyolite in the Hemlock shows the difference in K-spar (Table 1). Slight additional silica is provided by numerous random quartz veins that can be found in both the rhyolite (Figure 17) and underlying quartzite. Some of the veins in the quartzite contain selvages of micaceous hematite. It is suspected the potassium and silica were introduced at the time both magnetite and specularite replaced portions of the chert and quartzite at the base of the Hemlock. Further support of this timing is noted by the absence of alteration of overlying tuff and agglomerate at this stop. Figure 16. Rhyolite outcrop showing quartz veins Figure 17. Enhanced alkali/silica rhyolite porphyry 51 �Oxide SiO2 Al2O3 Fe2O3 FeO MgO Ca0 Na2O K2O TiO2 P2O5 Alt rhyolite*1 73.5 12.2 2.3 .22 .31 .21 .23 9.85 .99 .08 99.9 Unaltered rhy.*2 72.7 10.1 1.3 3.1 l.8 1.4 1.0 3.9 .47 .08 95.9 Table 1. Comparison of altered and unaltered rhyolite within the Hemlock Formation. *1 USGS Bull. 1044 p. 54 *2 USGS Bull. 1226 p. 21 Potassium values are twice those of other Hemlock rhyolite flows (see Table 1). Gair and Wier (1956, p. 53) “The acid volcanic rocks in the western part of the quadrangle (Kiernan-Sec. 36, T. 44 N., R. 32 W; Sec. 5, T. 43 N., R. 31 W) is much fresher and poorer in feldspar than the rock in the vicinity of Michigamme Mountain”. Table 1 shows the chemical difference between fresh and altered rhyolite. Chemistry of most of the Hemlock rhyolites matches that of the granophyre of the West Kiernan Sill. Ueng (1988) suggested the rhyolite formed in the magma chamber by crystal fractionation similar to the West Kiernan Sill differentiation that produced the granophyre. It can be speculated that the iron oxide addition to the underlying quartzite and chert occurred after extrusion of the rhyolite and prior to the next basaltic eruption. The overlying mafic tuff and agglomerate do not show any alteration beyond the generation of chlorite common to most of the Hemlock Formation indicating the oxide event occurred before emplacement of the basaltic rocks. On our trip returning to the vehicle, we will examine test pit material that represents fine specularite replacing the host quartzite. Stop 2d. Agglomerate (optional) UTM: 46o 9.692N 88o 10.985W, SESE Sec. 33, 44-31 A significant portion of the Hemlock is composed of fragmental tholeiitic basalt variously referred to as agglomerate, breccia or conglomerates (Figures 18-19) depending on visual degree of sorting or angularity. This stop has been highly weathered and is covered by moss and lichens. A more photogenic opportunity will be afforded at Stop 8. Figure 18. Fine grained volcanic conglomerate. Figure 19. Coarse grained agglomerate (hyaloclastites) breccia. 52 �Stop 3. East Kiernan Sill (optional) UTM: 46o 8.828N 88o 12.568W, SWSE Sec. 5, 43-31 The East Kiernan sill is a much smaller intrusion than the Western Kiernan sill. It is primarily an undifferentiated gabbro containing 50-80% prismatic hornblende after augite, albite, titaniferous magnetite, and apatite with less than 5% quartz. The plagioclase has been saussuritized and some hornblende converted to chlorite while the oxides have been converted to titanite and rutile. A number of gabbro samples of both Kiernan sills have been analyzed for PGEs (Bornhorst, 1990). None of the samples show values above background. Stop 4a. Mansfield Mine Location Monument UTM: 46o 6.835 N 88o 13.076 W, NWNW Sec. 20, 43-31 In 1889 the Mansfield Mining Co. took a lease on the Mansfield natural iron ore deposit from J.M. Longyear. Bessemer iron ore (<.045% P) was mined from six levels down to 435 feet below surface. Most of the workings were under the Michigamme River. Figure 20. Sign and plaque commemorating the Mansfield Mine Site disaster of 1893 On the evening of Sept. 27, 1893 a lower level pillar gave way causing the rock above to cave to surface flooding the mine workings with water from the Michigamme River. Twenty seven miners lost their lives in an instant. In 1897 the DeSoto Iron Co. of Springfield, IL bought the mine, redirected the river channel and reopened the mine. The Oliver Mining Co. (now USX) operated the mine from 1911 until the exhaustion of iron ore in 1913 at which time the workings had reached the 17th level about 1480 feet below the collar. A total of 1,462,504 long tons (LT) were mined between 1890 and 1913. In 1983 the mine site was designated the Mansfield Mine Location Historic District. A plaque with the names of the 27 men who died in the disaster marks the site (Figure 20). A few restored buildings mark the site of the mining community. 53 �Stop 4b. Mansfield Iron bearing slate member and agglomerate/volcaniclastics. UTM: 46o 6.945’ N 88o 13.185’W, NWNW Sec. 20, 43-31 The vertical #2 shaft went through fine volcanic conglomerate of the Hemlock and can be found on the mine dump. The slate portion is also represented on the dump (Figure 21) along with the iron formation. The Mansfield iron member has a limited areal extent of about 4 miles on strike. It was originally a siderite chert with soluble irons ranging from 10-47.3% that averaged ~25%. Near the contact with the intrusive West Kiernan Sill the Mansfield iron member was metamorphosed to a magnetite stilpnomelane chert. At the Mansfield Mine location the original siderite chert was far enough from the sill contact to avoid metamorphic alteration, but it did experience supergene oxidation and enrichment (Figure 22) with ore grades averaging about 52.8% natural iron with less than .045% P making it Bessemer type ore. Figure 21. Graphitic slate with pyrite cubes Figure 22. Direct shipping hematite ore with gypsum(white arrow). Stop 5. Steeply dipping Hemlock pillowed and massive basalt-Hemlock Falls Dam Site. UTM: 46o 7.840 N 88o 13.504 W, SESE Sec. 7, 43-31 The bulk of the Hemlock Formation is composed of massive, amygdaloidal and pillowed basalts and volcaniclastic rocks of the same composition. On the western edge of the Hemlock Falls Dam abutment is a large outcrop of pillow basalt dipping steeply to the west. Directly overlying the pillowed basalt is massive basalt as seen on the left side of Figure 23. 54 �Figure 23. Steeply dipping tholeiitic pillowed basalts at the west abutment of the Hemlock Falls Dam site. Stop 6. Amygdaloidal basalt of the Hemlock Formation UTM: 46o 6.461 N 88o13.931 W, NWNW Sec. 20, 43-31 On the north side of the County Road is a low outcrop of broken amygdaloidal basalt porphyry (Figure 24) with abundant gas vesicles filled with secondary albite quartz and carbonate. Some parts of the basalt contain a few per cent sulfides, mostly pyrite. Amygdaloidal basalts are not as abundant as massive basalts or agglomerates. 55 �Figure 24. Polished surface of amygdaloidal basalt at Stop 6 DIFFERENTIATED WEST KIERNAN SILL The West Kiernan Sill is approximately 13 miles long and averages 1.1 miles in thickness. It is designated a sill in that most contacts are concordant with the Hemlock units. Regional metamorphism for the entire sill is in the greenschist metamorphic zone. The sill can be divided into four distinct rock units: a basal serpentinized peridotite (Stop 11), a thick normal gabbro, a transition gabbro and, occasionally, an upper granophyre (Figure 25). A chill zone in the contact area is usually a diabase. The basal peridotite is composed of serpentine with aggregates of tremolite, talc and magnetite with minor amounts of carbonate and chlorite. Overlying the serpentine is a 900-1200 m thick normal gabbro. An increase in size and amount of plagioclase is noted from the base to the top of the unit (Figure 25). Alteration stilpnomelane is noted in the upper portion of the unit and becomes common in the overlying transition gabbro. The transition (Stop 7b) zone contains spotty concentrations of titaniferous magnetite that show up as magnetic bullseyes on airborne magnetic survey maps. The granophyre looks like a medium grained granite with minor or missing mafic minerals. Minerals present include albite, oligoclase, microcline and orthoclase with occasional needle of magnetite. Calcite is also pervasive. Geochemical analyses of the sill units are given in Table 2 along with analyses of the Hemlock basalt. 56 �1 2 3 Transition Oxide Peridotite*1 Gabbro*1 Gabbro*1 SiO2 39.01 44.94 57.2 Al2O3 6.56 19.70 10.8 Fe2O3 11.49 1.72 8.8 FeO 5.30 7.40 7.5 MgO 23.84 8.91 2.7 CaO 3.57 9.22 4.8 Na2O -1.94 2.8 K2O .02 .36 1.1 TiO2 2.04 .73 2.9 P2O5 .19 .09 .76 MnO .11 .13 .26 CO2 .08 .45 .37 H2O 7.10 4.56 3.90 Total 99.97 100.26 98.90 4 5 High Iron*2 34.1 3.4 13.1 22.3 8.5 7.6 .13 .10 5.42 .38 .38 3.0 98.40 Granophyre 69.12 12.83 .78 5.85 1.38 .70 1.72 3.44 .73 .11 .05 .70 2.25 99.75 *1 Bayley, 1959, Amer. Jour. of Sci. v. 257, p. 428 (col. 1-3) *2 Wier, 1967, USGS Bull. 1226, p. 35 (NW ¼ Sec. 12, 43-32) *3 Foose, 1981, N=7 *4 Foose, 1981, N=10 agglomerate/volcaniclastics Table 2. West Kiernan Sill-Whole rock analyses 57 7 Hemlock Basalt*3 48.2 14.6 6.7 7.8 5.9 5.8 3.03 .92 2.21 .30 .17 .71 2.64 98.98 8 Hemlock Agglomerate*4 47.6 14.4 5.4 7.6 6.6 7.8 3.29 .70 2.15 .19 .17 .70 2.57 99.17 �Based on the similarities in geochemistry between the West Kiernan and the Hemlock extrusives (Ueng, 1987) concluded both major igneous rocks were co-magmatic in origin. He further stated crystal fractionation in the magma chamber was responsible for the occasional rhyolite flows. Xenoliths of Hemlock are present in both the East and West sills (Bayley, 1959; Wier, 1967). Stops 7a-b-c will allow us to examine the transition gabbro (7c) containing abundant magnetite and the overlying granophyre (7a). Whole rock assays for these outcrops are shown on Table 3. Note the major increase in silica, sodium and potassium coupled with a significant reduction in iron oxides and titanium in the granophyre. The two rock types represent a significant chemical change during late stage differentiation. OXIDE A*1 B*2 SiO2 34.1 51.0 Al2O3 3.4 16.3 Fe2O3 13.1 1.4 FeO 22.3 9.2 MgO 8.5 8.5 CaO 7.6 10.1 Na2O .13 1.83 K2O .10 .68 H2O 3.0 Ti02 5.42 .85 P2O5 .38 .05 CO2 <.05 MnO .38 .14 S .38 *1 test pit with magnetite in transition gabbro. USGS Bulletin 1226, p. 35 *2 Fox M.S. thesis sample A5P outcrop near railroad cut. Table 3. Whole rock analyses upper Kiernan Sill transition gabbro near granophyre contact (Stop 7). Stop 7a. West Kiernan Sill granophyre UTM: 46o 8.545 N 88o 15.742 W, NWNW Sec. 12, 43-32 The outcrop is typical of the granophyre phase of the West Kiernan Sill. ferromagnesium mineralization is present and it does resemble a granite. Very little Stop 7b. Coarse grained transition gabbro UTM: same as 7a. This location is very coarse grained and contains around 80% plagioclase, typical of the upper part of the transition gabbro. The apatite is of the fluorine variety. Stop 7c. Transition gabbro with magnetite UTM: 46o 8.508 N 88o 15.717 W, NWNW Sec.12, 43-32 These test pits are near the top of the transition gabbro where abundant titaniferous magnetite is common. Additional minerals include stilpnomelane/bannisterite, ilmenite, biotite, augite, 58 �feldspar, chlorite and ferrohornblende (Figure 26-27). Ueng (1988) speculated that magnetite, ilmenite and apatite crystalized during fractionation in the slowly cooled sill. Figure 26. Skeletal titaniferous magnetite in transition gabbro. Figure 27. Gabbro porphyry with high titaniferous magnetite content. Stop 8. Top of Hemlock with increased magnetite UTM: 46o 14.923’ N 88o 26.206’ W, Sec. 4, 44-33 Figure 28. Hemlock agglomerate with iron oxide values above 20% mostly as magnetite. Stop 8 represents the upper 6000 feet of the Hemlock. Fine grained pillow and fragmental tholeiitic basalts (Figure 28) predominate and contain up to 25% FeO with most of the iron over 12% FeO as magnetite. 59 �Stop 9. West Kiernan Sill-copper/nickel gabbro UTM: 46o 4.923 N 88o 12.644 W, SENW Sec. 32, 43-31 The majority of the sill consists of a normal gabbro. Plagioclase and augite were the major minerals in the rock. The augite has been altered to hornblende. At this stop the gabbro outcrop contains disseminated pyrrhotite, chalcopyrite and pentlandite (Figure 29). The outcrop exhibits a typical northern climate gossan with shallow oxidation representing low grade disseminated sulfides. A grab sample from this outcrop assayed 0.35% copper/nickel. Figure 29. Polished slab showing Chalcopyrite-pentlandite-pyrrhotite gabbro from the West Kiernan sill. Stop 10. Magnetite bombs in the Hemlock Formation UTM: 46o 5.191’ N 88o 12.015’ W In 2005 Prime Meridian Resources drilled a coincident TEM and magnetic anomaly target in the Hemlock Formation east of the Michigamme River and just north of highway M-69. The magnetic anomaly is due to concentrations of magnetite and the conductor anomaly is due to increased amounts of pyrite and chalcopyrite in some parts of the magnetite rich zone. The hole identified as KS-102-2 penetrated alternating chert, tuffs and amygdaloidal basalts bottoming in magnetite rich basalts resembling volcanic ejecta or bombs (Figure 30). 60 �Figure 31. Photo of the outcrop of magnetite bombs with the up direction to the top. Field trip participants will examine diamond drill core representing the siliceous zone and the magnetite rich zone at the bottom of the hole. The best outcrop example is approximately 1 mile round trip from this location, however, the outcrop is now encrusted in lichen obscuring the salient features seen in the Figure 31. We will examine core from DDH KS-102-2 which exhibits excellent features of the magnetite bombs (Figure 32). 61 �Figure 32. Diamond drill core showing magnetite bombs with albite and calcite filling amygdules. Figure 33. Polished bomb showing amygdules and groundmass magnetite and titanite. Individual fragments range in size from 5-7 inches long and 1-2 inches thick and contain variable magnetite concentrations in the groundmass ranging from 15-70% (Figure 33). The groundmass consists of albite, titanite and magnetite. The amygdules have been filled with albite, epidote, calcite and ilmenite. Both the magnetite and titanite are absent from the vesicles supporting the concept that the magnetite was primary and was not involved with later hydrothermal overprint. The size of the magnetite is bimodal with two distinct morphologies present. The first are the small euhedral individual crystals of magnetite while the second type can best be described as aggregates of anhedral magnetite. The distribution of the magnetite in the groundmass is not uniform. In some areas the amount of magnetite approaches 70%. It is suggested the crystallization had occurred by the time it was expelled from the vent. The magnetite was observed to increase in size and amount near the edges of vesicles and between large vesicles. The fine interstitial material between the bombs is devoid of magnetite and titanite indicating a different source for the fine grained interstitial material. It has been suggested that what appears to be bombs are in reality very small flattened pillows and not bombs at all. This interpretation is a possibility, however, the lack of magnetite in the fine grained material between fragments would be difficult to reconcile pillow creation. One more significant fact is the magnetite is rather pure and does not contain any titanium suggesting that element separation occurred in the magma chamber. Ueng (1988) also described rhyolite bombs in the central portion of the Hemlock Formation. Stop 11. Serpentinized peridotite at the base of the West Kiernan sill UTM: 46o 4.894 N 88o 11.832 W, SWNW Sec. 33, 43-31 Stop 11 represents the basal peridotite of the West Kiernan Sill that has undergone extensive serpentinization. The roadside outcrop on the north side of M-69 exhibits fuzzy outlines of altered phenocrysts of pyroxene. Ueng (1988) did not observe any relic olivine. The process of serpentinization creates abundant secondary magnetite as evidenced by the strong magnetic attraction exhibited by the outcrop. Additional alteration minerals include tremolite, talc, 62 �carbonate, chlorite and some actinolite. Chemistry of the ultramafic rocks are shown in Table 2 where the MgO values are about 23.8%. Acknowledgements I would like to thank Klaus Schulz and Bill Cannon for their constructive conversations and comments on the guide content. Appreciation is expressed to two other reviewers for their helpful review. I would also like to thank Stacy Saari for work on several illustrations. Appreciation is extended to the multiple private land owners for allowing access to several of the field trip stops. During the poster session a display of North American kimberlite samples from the Doug Duskin collection will be on display for comparison with the Lake Ellen kimberlite visited on this field trip. The bibliography at the end of the field trip guide contains many more references than cited in the text to aid anyone who wishes to delve further into the subject in greater detail. Bibliography-Hemlock Formation Aldrich, L.T., Davis, G.L., and James, H.L., 1965, Ages of minerals from metamorphic and igneous rocks near Iron Mountain, Michigan: Journal of Petrology, v. 6, p. 445-472. Anhaeusser, C.R., 1985, Archean layered ultramafic complexes in the Barberton Mountain Land, South Africa in Ayers, L.D., Thurston, P.C., Card, K.D., and Weber, W. eds., Evolution of Archean supracrustal sequences: Geological Association of Canada Special Paper 28, p. 281-301. Banks, P.O., and Van Schmus, W.R., 1971, Chronology of Precambrian rocks of Iron and Dickinson Counties, Michigan: Institute on Lake Superior Geology abs, v. 17, p. 9-10. Bartlett, W.A., et al, 1976, Distribution of sulfur in the West Kiernan Sill, Iron County, Michigan: Bowling Green State University MS thesis, 87p. Bartlett, W.A., Lougheed, M.S., Mancuso, J.J., and Walter, L.J., 1976, Distribution of sulfur in the West Kiernan Sill, Iron County, Michigan: Institute on Lake Superior Geology abs, v. 22, p. 6. Bartlett, W.A., 1976, Distribution of sulfur in the West Kiernan Sill, Iron County, Michigan: Bowling Green State University M.S. thesis, 87 p. Barovich, K.M., Patchett, P.J., and Peterman, Z.E., 1987, Origin of the 1.9-1.7 Ga Penokean continental crust of the Lake Superior region abs: Eos, v. 68, p. 1547. Bayley, R.W., 1959, A metamorphosed differentiated sill in Northern Michigan: American Journal of Science, v. 257, p. 408-430. Bayley, R.W., 1959, Geology of the Lake Mary Quadrangle Iron County, Michigan: U.S Geological Survey Bulletin 1077, 112 p. Bayley, R.W., Dutton, C.E., and Lamey, C.A., 1966, Geology of the Menominee Iron Bearing District, Dickinson County, Michigan and Florence and Marinette Counties, Wisconsin: U.S. Geological Survey Professional Paper 573, 96 p. Baxter, D.A., and Bornhorst, T.J., 1988, Multiple discrete mafic intrusion of Archean to Keweenawan age, western Upper Peninsula, Michigan: Institute on Lake Superior Geology abs, v.34, p. 6-8. Beck, J.W., 1984, Nd and Sm isotopic studies of the Quinnesec and Hemlock Formations in northeastern Wisconsin and adjacent Michigan: Lake Superior Geology abs, v. 30, p. 1-2. Beck, J.W., 1988, Implications for early Proterozoic tectonics and the origin of continental flood basalts based on combined trace element and neodymium/strontium isotopic studies of mafic igneous rocks of the Penokean Lake Superior belt, Minnesota, Wisconsin and Michigan: University of Minnesota Ph.D. thesis, 262 p. Beck, J.W., and Murthy V.R.., 1991, Evidence for continental crystal assimilation in the Hemlock Formation flood basalts of the early Proterozoic Penokean orogeny, Lake Superior region: U.S. 63 �Geological Survey Bulletin 1904 I, p. 101-125. Bornhorst, T.J., and Baxter, D.A., 1990, Reconnaissance evaluation of platinum group elements in selected Precambrian rocks of the western Upper Peninsula, Michigan: Michigan Department of Natural Resources Geological Survey Division: Geology Report 90-2, 39 p. Cambray, F.W., 1978, Plate tectonics as a model for the environment of deposition and deformation of the early Proterozoic (Precambrian X) of Northern Michigan: Geological Society of America abs, p. 7 Cannon, W.F., 1973, The Penokean orogeny in northern Michigan in Huronian stratigraphy and sedimentation: Geological Society of America Special Paper 12, p. 251-271. Cannon, W.F., and Klasner, J.S., 1976, Geologic map and geophysical interpretation of the Witch Lake Quadrangle Marquette, Iron and Baraga Counties, Michigan: U.S. Geological Survey Map I-987, scale 1:62,500. Cannon, W.F., 1983, Mineral resource assessment of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Circular 887, 21 p. Cannon, W.F., 1985, Mineral-resources map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-A, scale 1:250,000. Cannon, W.F., 1986, Bedrock geologic map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-B, scale 1:250,000. Cannon, W.F., 1986, Structural and tectonic map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-D, scale 1:250,000. Chartier, R., 1985, The texture and mineralogy of the Lake Ellen kimberlite, Crystal Falls, Michigan USA: Institute on Lake Superior Geology abs, v. 31, p. 10. Clements J.M., and Smyth, H.L., 1899, The Crystal Falls Iron-Bearing District of Michigan: U.S. Geological Survey Monograph 36, 512 p. Cogan, M.J., 1993, Primary and secondary Bouguer gravity trend analysis and structural implications for early Proterozoic Kiernan sills, Iron County, Michigan: Geological Society of America, v. 25, p. 13. Cudzilo, T.F., 1978, Geochemistry of early Proterozoic igneous rocks, northeastern Wisconsin and upper Michigan: University of Kansas Ph..D dissertation, 202 p. Cummings, M.L., 1978, Metamorphism and mineralization of the Quinnesec Formation, northeastern Wisconsin: University of Wisconsin Ph.D. thesis, 190 p. Dann, J.C., 1978, Major-element variation within the Emperor Igneous Complex and the Hemlock and Badwater volcanic Formations: Michigan Technological University M.S. thesis, 159 p. Dann, J.C., 1978, Major-element variation within the Emperor Igneous Complex and the Hemlock and Badwater volcanic Formations: Institute on Lake Superior Geology abs, p. 15. DeMatties, T.A., Rowell, W.F., Munroe, J.F., 2007, An evaluation of the Prime Meridian midcontinent nickel-copper exploration program: Technical Report: Prime Meridian Resources, 538 p. Dutton, C.E., and Linebaugh, R.E., 1967, Map showing Precambrian geology of the Menominee IronBearing District and vicinity Michigan and Wisconsin: U.S. Geological Survey Map I-466, scale 1:125,000. Dutton, C.E., 1971, Geology of the Florence area, Wisconsin and Michigan: U.S. Geological Survey Professional Paper 633, 54 p. Foose, M.P., 1981, Geology of the Ned Lake Quadrangle, Iron and Baraga Counties, Michigan: U.S. Geological Survey Map I-1284, scale 1:62,500. Fox, T.P., 1983, Geochemistry of the Hemlock metabasalt and Kiernan sills, Iron County, Michigan: Michigan State University M.S. thesis, 73 p. Gair, J.E., and Wier, K.L., 1956, Geology of the Kiernan Quadrangle Iron County, Michigan: U.S. Geological Survey Bulletin 1044, 88 p. Graff, C.W., 1982, Iron-enriched basaltic fragmental rocks erupted in a shallow subaqueous environment, the Hemlock Formation, Amasa Quadrangle, Michigan: Institute on Lake Superior abs, v. 28, p. 11. 64 �Greenberg, J.K., and Brown, B.A., 1983, Lower Proterozoic volcanic rocks and their setting in the southern Lake Superior district, in Medaris. L.G., Jr., ed., Early Proterozoic geology of the Lake Superior region: Geological Society of America Memoir 160, p. 67-84. Han, T.M., 1966, Textural relations of hematite and magnetite in some Precambrian metamorphosed oxide iron-formations: Economic Geology, v. 61, p. 1306-1310. Han, T.M., 1978, Microstructures of magnetite as guides to its origin in some Precambrian iron formations: Fortschr. Mineral, v. 56, p. 105-142. Han, T.M., 1988, Origin of magnetite in Precambrian iron-formations of low metamorphic grade in Proceeding of the Seventh Quadrennial IAGOD Symposium: E. Schweizerbart’sche Verlagsbuchhandlung, D-7000 Stuttgart 1: p. 641-656. Heran, W.D., and Smith B.D. 1980, Description and preliminary map of airborne electromagnetic survey of parts of Iron, Baraga, and Dickinson Counties Michigan. U.S. Geological Survey Open File Report 80-297, 8 p. Hoffman, J.D., 1984, Nickel distribution in B-horizon soils, Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-K, scale 1:250,000. Hoffman, P.F., 1987, Early Proterozoic foredeeps, foredeep magmatism, and superior type iron formations of the Canadian shield, in Kroner, A., ed., Proterozoic lithospheric evolution: American Geophysical Union Geodynamics Series, v. 17, p. 85-98 Hotz, P.E., 1953, Petrology of granophyre in diabase near Dillsburg, Pennsylvania: Geological Society of America Bulletin, v. 64, p. 675-704. James, H.L., 1955, Zones of regional metamorphism in the Precambrian of northern Michigan, Geological Society of America Bulletin, v. 66, p. 1455-1457. James, H.L., Dutton, C.E., Pettijohn, F.J., and Wier, K.L., 1968, Geology and ore deposits of the Iron River-Crystal Falls District, Iron County, Michigan: U.S. Geological Survey Professional Paper 570, 134 p. James, H.L., 1958, Stratigraphy of pre-Keweenawan rocks in parts of northern Michigan: U.S. Geological Survey Professional Paper 314-C, 44 p. Johnson, D.J., 1975, Petrology of a portion of the Hemlock Formation, Iron County, Michigan: Michigan Technological University MS thesis, 51 p. Johnson, D.J., 1975, Petrology and tectonic setting of the Hemlock Formation, Iron County, Michigan: Institute on Lake Superior Geology abs, v. 21, p. 3 King, E.R., 1987, Aeromagnetic map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-F, scale 1: 250,000. Kirkley, M.B., Gurney, J.J., and Levinson, A.A., 1992, Age, origin and emplacement of diamonds: a review of scientific advances in the last decade: CIM Bulletin, v. 84, p. 48-57. Klasner, J.S., Ojakangas, R.W., Schulz K.J., and Laberge, G.L., 1988, Evidence for development of an early Proterozoic overthrust-nappe system in the Penokean orogeny of Northern Michigan: Institute on Lake Superior Geology abs, v. 34, 56-57. Klasner, J.W., and Jones, W.J., 1989, Bouguer gravity anomaly map and geologic interpretation of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-E, scale 1:250,000. Krapez, B., Barley, M.E., and Pickard, A.L., 2003, Hydrothermal and resedimented origins of the precursor sediments to banded iron formation: sedimentological evidence from the early Paleoproterozoic Brockman Supersequence of Western Australia: Sedimentology, v. 50 p. 979-1011. Kruger, C.L., 1967, Aeromagnetic map of the Crystal Falls Quadrangle and part of the Florence Quadrangle, Iron County, Michigan: U.S. Geological Survey Map GQ-607, scale 1:62,500. Kruger, C.L., 1967, Aeromagnetic map of the Perch Lake Quadrangle, Houghton, Baraga and Iron Counties, Michigan: U.S. Geological Survey Map GP-600, scale 1:62,500. Kruger, C.L., 1967, Aeromagnetic map of the Ned Lake Quadrangle and part of the Witch Lake Quadrangle, Iron, Baraga and Marquette Counties, Michigan: U.S. Geological Survey 65 �Map GP-609, scale 1:62,500. Larue, D.K., and Sloss, L.L., 1980, Early Proterozoic sedimentary basins of the Lake Superior region: Geological Society of America Bulletin, pt. II, v. 91, p. 1836-1874. Larue, D.K., 1983, Early Proterozoic tectonics of the Lake Superior region-tectonostratigraphic terranes near the purported collision zone in Medaris, L.G., Jr., Early Proterozoic geology of the Lake Superior region: Geological Society of America Memoir 160, p. 33-47. Larue, D.K., and Ueng, W.L., 1985, Florence-Niagara terrane-an early Proterozoic accretionary complex, Lake Superior region: Geological Society of America Bulletin, v. 96, p. 1179-1187. MacRae, N.D., 1969, Ultramafic intrusions of the Abitibi area, Ontario: Canadian Journal of Earth Sciences, v. 6, p. 281-303. March, B.D., 2006, Dynamics of magmatic systems: Elements, v. 2, p. 287-292. Ozawa, A., Ueda, A., “Fantong, W.Y., Anazawa, K, Yoshida, Y, Kusakabe, M., Ohba, T., Tanyileke, G., and Hell, J.V., 2016, Rate of siderite precipitation in Lake Nyos, Cameroon, geochemistry and geophysics of active volcanic lakes: Ohba, Capaccioni and Caudron, eds., Geolgical Society, London, Special Publications 437. Paddock, D.R., 1982, A Gravity investigation of eastern Iron County, Michigan: Michigan State University M.S. thesis, 110 p. Peterson, W.L., 1985, Surficial geologic map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-C, scale 1:250,000. Rehfuss, I.L., 1912, The Bird mine section: University of Wisconsin B.A. thesis, 49 p. Ruotsala, A.P., 1974, Composition and tectonic setting of middle Precambrian lavas, Crystal Falls area, Michigan: Institute on Lake Superior Geology abs, v. 20, p. 28. Schneider, D.A., Bickford, M.E., Cannon, W.F., Schulz, K.J. and Hamilton, M.A., 2002, Age of volcanic rocks and syndepositional iron formations, Marquette Supergroup: implications for the tectonic setting of Paleoproterozoic iron formations of the Lake Superior region: Canadian Journal of Earth Science, v. 39, p. 999-1012. Schulz, K.J., 1982, Magnesian basalts from the Archean terrains of Minnesota, in Arndt, N.T., and Nisbet, E.G., eds, Komatiites: London, George Allen and Unwin, p. 171-186. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4-25. Schulz, K.J., and Cannon, W.F., 2008, Synchronous deposition of Paleoproterozoic superior-type banded iron formations and volcanogenic massive sulfides in the Lake Superior region: implications for the tectonic evolution of the Penokean orogeny: Geological Society of America, abs. w/programs, v. 40, p. 387. Schulz, K.J., 1984, Early Proterozoic Penokean rocks of the Lake Superior region: geochemistry and tectonic implications: Institute on Lake Superior Geology abs, v. 30, p. 65-66. Sims, P.K., 1992, Geologic map of Precambrian rocks, southern Lake Superior region, Wisconsin and northern Michigan: U.S. Geological Survey Map I-2185, scale 1:500,000. Stahl, S.D., et al, 1993, Primary and secondary Bouguer gravity trend analysis of the Kiernan sills area, Iron County, Michigan: implications for early Penokean tectonics: The Geology of Michigan and its Geological Resources Symposium III, MI DNR. Ueng, W.C., Larue, D.K., and Sedlock, R.L., 1984, The early Proterozoic tectonic history of the southcentral Lake Superior region: Institute on Lake Superior Geology field trip guide, v. 30, p. 1-22. Ueng, W.C., Larue, D.K., and Sedlock, R.L., 1988, Geochemistry and petrogenesis of the early Proterozoic Hemlock volcanic rocks and the Kiernan sills, southern Lake Superior region: Canadian Journal of Earth Science, v. 25, p. 528-546. Vallini, D.A., Cannon W, F, and Schulz, K, J., 2006, Age constraints for Paleoproterozoic glaciation in the Lake Superior region: detrital zircon and hydrothermal xenotime ages for the Chocolay Group, Marquette Range Supergroup: Canadian Journal of Earth Science, v. 43, p. 571-591. Waggoner, T. D., Duskin, D., Karakus, M., and Gartner, J., 2011, Pyroclastic magnetite bombs in 66 �the Hemlock Formation, Iron County, Michigan: Institute on Lake Superior abs, v. 57, p 93-94. Walker, F., 1969, The Palisades Sill, New Jersey-a reinvestigation: Geological Society of America Special Paper 111, 178 p. Wier, K.L., 1967, Geology of the Kelso Junction Quadrangle Iron County, Michigan: U.S. Geological Survey Bulletin 1226, 47 p. Wier, K., 1986, Metamorphic map of the Iron River 1o x 2o Quadrangle, Michigan and Wisconsin: U.S. Geological Survey Map I-1360-G, scale 1:250,000. Williams, D.A.C. and Halberg, J.A., 1973, Archean layered intrusions of the eastern Goldfields region, Western Australia: Contribution to Mineralogy and Petrology, v. 38, p. 45-70 Zeitz, I., and Kirby, J.R., 1971, Aeromagnetic map of the western part of the northern peninsula, Michigan, and part of northern Wisconsin: U.S. Geological Survey Map GP-750, scale 1:250,000. __________, 1978, Aerial gamma ray and magnetic survey peninsula portion, Hancock Quadrangle Wisconsin, Minnesota, Michigan, Iron River Quadrangle, Michigan Wisconsin and Marquette Quadrangle: Michigan, v. 1 Final DOE Report. __________, 1980, Airborne electromagnetic map and magnetic survey of parts of the upper Peninsula Of Michigan and Northern Wisconsin: U.S. Geological Survey Open File Report 81-577, Scale 1:24,000. Bibliography-Lake Ellen Kimberlite. Cannon, W.F., and Mudrey, M.G., 1981, The potential for diamond-bearing kimberlite in northern Michigan and Wisconsin: USGS Circular 842, 15 p. Carlson, S.M., and Floodstrand, W., 1994, Michigan kimberlites and diamond exploration techniques: Institute on Lake Superior Geology field trip guide, Part 4, p. 1-15. Chartier, T., 1985, The texture and mineralogy of the Lake Ellen kimberlite: Institute on Lake Superior Geology abs, v, 31, p. 10-11. Clements, B., 2017, The Canadian diamond business: 25 years and going strong: SEG Newsletter, p. 1 and 12-17. Griffin, W.L., O’Reilly, S.Y., Doyle, B,J., Pearson, N.J., Coopersmith, H., Kivi, K., Malkovets, V. and Pokhilenko, N., 2004, Lithosphere Mapping beneath the North American Plate: Lithos, v. 77, p. 873-922. Hausel, W.D., 1998, Diamonds and mantle source rocks in the Wyoming craton with a discussion of other U.S. occurrences: Wyoming State Geological Survey Report of Investigations 53, 93 p. Head, J.W., and Wilson, L., 2002, Diatremes and kimberlites I: definition, geological characteristics and association: Micro Symposium 36 (MSO32), 2 p. Jarvis, W., and Kalliokoski, J, 1988, Michigan kimberlite province: Institute on Lake Superior Geology abs, v. 34, p. 46-48. Jarvis, W., 1993, Michigan kimberlites: an update: abs, 61swt PDAC, Paper M-10. Kirkley, M.B., Gurney, J.J., 1992, Age, origin and emplacement of diamonds: a review of scientific advances in the last decade: CIM Journal, v. 84, p. 48-57. Kjarsgaard, B.S., 2007, Kimberlite pipe models: significance for exploration, ore deposits and exploration technology in Proceeding of Exploration 07: Fifth Decennial International Conference on Mineral Exploration, Milkereit, B., ed.: Paper 46 p. 667-677. McGee, E.S., and Hearn, B.C., 1983, Lake Ellen kimberlite, USA: U.S. Geological Survey Open File Report 83-156, 34 p. McGee, E.S., and Hearn, B. C., 1984, The Lake Ellen kimberlite, Michigan, U.S.A.: in Kimberlites. I: kimberlites and related rocks, Korprobst, J., ed., p. 143-154. 67 �McGee, E.S., 1987, Garnet xenocryst analysis potential for diamonds in Williams kimberlite, north central Montana and the Lake Ellen kimberlite, northern Michigan: U.S. Geological Survey Open File Report 87-418, 15p. McGee, E.S., 1988, Potential for diamonds in kimberlites from Michigan and Montana as indicate by garnet xenocryst composition: Economic Geology, v. 83, p. 428-432. Paces, J.B., and Taylor, L.A., 1990, Petrography, mineral chemistry and geothermobarometry of mafic granulite and eclogite nodules from upper Michigan kimberlites: Institute on Lake Superior Geology abs, v 36, p. 82-84. Quigley, P.O., 2007, Michigan kimberlites revisited: new mineral, chemical and petrographic analysis: Institute on Lake Superior Geology abs, v. 53, p. 63. Scully, K.R., Canil, D., and Schulze, D.J., 2004, The lithospheric mantle of the Archean Superior Provence as imaged by garnet xenocryst geochemistry: Chemical Geology, v. 207, p. 189-221. Skillings, D., 1995, Crystal Exploration Inc. continuing to evaluate diamond potential of Michigan’s upper peninsula and in Wisconsin: Skillings Mining Review, v. 84, p. 5-7. Zartman, R.E., Kempton, P.D., Paces, J.B., Downes, H., Williams, I.S., Dobosi, G., and Futa, K., 2013, Lower-Crustal xenoliths from Jurassic kimberlite diatremes, upper Michigan (USA): evidence for Proterozoic orogenesis and plume magmatism in the lower crust of the southern Superior Province: Journal of Petrology, v. 54, p. 575-608. ________, 1992, Petrography of the Crotch Lake kimberlite, Michigan: Exmin document on file at the Michigan DEQ Core Library. ________, 1993, Ashton Mining of Canada Inc. Annual Report, 21 p. 68 �FIELD TRIP 3 Friday, May 18, 2018 GEOLOGY AND IRON ORES OF THE MENOMINEE IRON RANGE, DICKINSON COUNTY, MICHIGAN Thomas H. Mroz, BSGE, MSPG, CPG William F. Cannon, Klaus J. Schulz, Robert A. Ayuso, U.S. Geological Survey INTRODUCTION The Menominee Iron Range was visited on a previous ILSG field trip in 2003. This trip differs substantially from that previous trip and visits mostly new localities with only three exceptions. The following introductory material is reproduced in large part from the 2003 Guidebook (LaBerge et al., 2003). The trip visits most of the stratigraphic units of the area including the Carney Lake Gneiss, the Archean basement on which Paleoproterozoic sedimentary rocks were deposited. The emphasis of this trip is the Paleoproterozoic section and most stratigraphic units will be seen. Several stops are focused on the Vulcan Iron-formation, the principal iron-bearing unit of the Range. Both the unaltered formation and the secondary ores that formed within it will be seen. The Menominee Iron Range, one of the principal iron producing districts of the Lake Superior region, produced about 260 million tons of high grade iron ore between 1873 and 1946, but has been inactive since. The ores were secondary concentrations of iron oxides and hydroxides within the Paleoproterozoic Vulcan Iron-formation. The ores are generally believed to be paleosupergene concentrations that formed on the Cambrian surface and were covered by late Cambrian sandstone of the Munising Formation. The range also lies very near the Niagara fault zone, the paleosuture between the Superior craton and the accreted Pembine-Wausau arc terrane of northern Wisconsin, and bears the imprint of the strong deformation produced during the accretion. Stratigraphy. The presently accepted stratigraphic terminology for the Menominee Iron Range was developed by Bayley et al. (1966) and modified only slightly since. The stratigraphic relationships are shown in Figure 2, which is modified from Bayley et al. (1966) to reflect changes in terminology and radiometric age determination since that publication. Precambrian rocks range in age from Archean to Mesoproterozoic. They are capped by Late Cambrian sandstones, which occur as numerous outliers and underlie most of the higher ridges along the Range 69 �Figure 1. Generalized geologic map of the Menominee Iron Range showing the location of the field trip stops. . Figure 2. Sequence of formations in the Menominee Iron Range. Modified from Bayley et al. (1966, table 6). 70 �The following summary draws heavily on previous descriptions by Bayley et al. (1966), who provided the most recent comprehensive study of the Range, and LaBerge et al. (2003), who prepared the 2001 ILSG fieldtrip guidebook for the Range. The Precambrian stratigraphic units can be divided into four principal sequences ranging from Archean to Mesoproterozoic. Radiometric age determinations conducted since the previous ILSG guidebook provide new clarity on absolute ages of these sequences. Archean- Archean rocks form the basement on which the Paleoproterozoic rocks of the Menominee Iron Range were deposited. They underlie a large area north of the Range and are known as the Carney Lake Gneiss. The Carney Lake is a complex, multilithic assemblage of mostly granitic and lesser mafic rocks, which are described more fully in the description of Stop 1. The most recent rock-forming period of the complex was at about 2.75 Ga, but recent age determinations (Ayuso et al., 2017; this volume) have identified cores of zircon grains as old as about 3.8 Ga indicating a very extended history within these rocks. Paleoproterozoic- The Paleoproterozoic rocks of the Menominee Iron Range are entirely sedimentary and together make up the Marquette Range Supergroup. Three individual groups are present; from oldest to youngest they are the Chocolay, Menominee, and Baraga Groups. Each group is separated by unconformities and the Supergroup, as well, lies unconformably on the Archean Carney Lake Gneiss. Chocolay Group- The Fern Creek Formation, Sturgeon Quartzite, and Randville Dolomite make up the Chocolay Group. The basal formation, the Fern Creek Formation, is a glaciogenic unit that is preserved only sporadically in the area. Stop 2 is one of the best localities to study it. The stop has been well described by R.J Ojakangas (LaBerge et al., 2003). The medial unit of the Chocolay Group, the Sturgeon Quartzite (see Stop 2 description), forms a continuous blanket of orthoquartzite throughout the area and lies directly on the Carney Lake Gneiss in areas devoid of the Fern Creek. A sericitic unit at the top of the Fern Creek (see Stop 2 description) may be a paleosol indicating a period of weathering between deposition of the Fern Creek and the Sturgeon. The Randville Dolomite, the uppermost unit of the Chocolay Group, is an extensively exposed shallow-water carbonate unit containing numerous stromatolite horizons and other indications of shallow or intertidal deposition. Radiometric ages of detrital zircons and authigenic xenotime have bracketed the depositional age of the Chocolay Group between 2.2 and 2.3 Ga and support its correlation with lithologically similar units in the Huronian Supergroup of Ontario (Vallini et al., 2006). Menominee Group- The Menominee Group is composed of the basal Felch Formation and overlying Vulcan Iron-formation, the major iron-bearing member of the Menominee Iron Range. The group lies unconformably on the Chocolay Group, although it is structurally concordant with the older rocks in most places. Bayley et al. (1966) describe several localities where there is an angular discordance between the Felch and Vulcan; where there is a basal conglomerate in the Felch Formation it is composed in large part of clasts of the Randville. The absolute age of the Menominee Group in this area has not been determined but regional relationships provide some constraints. To the northwest, the Hemlock Volcanics, part of the Menominee Group and interlayered with iron-formation probably approximately coeval with the Vulcan, have been dated at 1.87 Ga (Schneider et al., 2002). That date indicates that more than 300 million years likely separate the end of Chocolay Group deposition and the deposition of the Menominee Group. 71 �Felch Formation- The Felch Formation is a sericitic slate and quartzite unit that overlies the Randville Dolomite. It consists of thin-bedded sericitic slate and phyllite and intercalated thinbedded quartzite, with the quartzite layers being more prevalent near the top of the formation (Bayley et al., 1966). It is about 100 feet thick on the south range but is as much as 500 feet thick on the north range. Bayley et al. (1966) considered the Felch Formation to be correlative with the Ajibik Quartzite and Siamo Slate of the Marquette district and the Palms Formation of the Gogebic district. The Felch Formation is conformable and gradational with the overlying Vulcan Iron-formation. Vulcan Iron-formation- The Vulcan Iron-formation is the major iron-bearing unit of the Menominee district. It is well known from numerous mines and drill holes, but generally is not well exposed in natural outcrops. The iron-formation is divided into four units, two composed mainly of granular iron-formation and two composed of slate and slaty iron-formation. In succeeding order the units are the Traders Iron-bearing Member, the Brier Slate, the Curry Ironbearing Member, and the Loretto Slate. They are described in detail by Bayley et al. (1966). The Traders and Curry Members contain layers of granular jasper alternating with layers of magnetite and hematite. The Brier and Loretto Members are mainly laminated siliceous iron-rich slate, which locally contains laminae of detrital quartz, feldspar, micas, zircon, and tourmaline. According to Dutton (1958), the iron-formation is about 1,000 feet thick, of which about 730 feet is ferruginous slate (Brier Slate - 330 feet, Loretto Slate - 400 feet) and 270 feet is granular ironformation (Traders - 100 feet, Curry - 170 feet). The stratigraphy within the Vulcan is seen well at Stops 3, 7, and 8. A detailed stratigraphic section (from oldest to youngest) from the Curry Mine located between the towns of Vulcan and Norway shows that above the Randville Dolomite there is a 20 foot conglomerate consisting of novaculite (dense, fine-grained siliceous rock resembling chert) boulders and smaller angular fragments cemented by ferruginous silica. A similar rock is seen at Stop 8. A fault breccia occurs in two zones bordering a dolomitic slate, over a 25-foot interval. Above the fault breccia is 26-foot interval of vitreous quartzite and then 69 feet of Felch Formation that is divided into several horizons including quartz slate, blocky green slate, massive brown chert, blocky brown talcose slate, shaly brown talcose slate, and topped by the “Trader’s quartzite” (informal terminology). The Vulcan Iron-formation lies on top of the quartzite and is 111 feet thick with several horizons noted; ferruginous slaty iron-formation, wavy-bedded red cherty iron-formation, massive wavy iron-formation with brownish red chert lenses, even-bedded iron-formation with brown chert beds, and the uppermost unit is a massive brown granular chert horizon. The Brier slate is in fault contact with the Traders Iron-bearing Member. It is grey to brown (oxidized) laminated slate, 100 feet thick that is also in fault contact with the Curry Iron-bearing Member. The Curry is 158 feet thick with a thin basal slaty phase and thick even-bedded, dark reddish purple granular chert with specular hematite laminae and in cross fractures. The Loretto Slate is the next formation horizon at about 45 feet thick and bounded by sheared contacts. It is a dark brown, thinly laminated, blocky ferruginous slate. The “Hanbury slate” (Michigamme Formation) overlies the Loretto and in the 5th evel of the Curry mine consists 405 feet of ferruginous mottled red and white slate, then a greenish grey thinly laminated slate with a high chlorite content, and finally a soft pyritic black carbonaceous slate with graphite on shear planes. This stratigraphic section from the Curry Mine is the most complete sequence known for the Range and was developed by Penn Iron Mining Company geologists. 72 �Baraga Group- The Baraga Group consists of a single unit, the Michigamme Formation. The belts underlain by the Michigamme Formation are very poorly exposed, which accounts, at least in part, for the lack of detailed mapping of what may well be otherwise discernible map units. According to Bayley et al. (1966), the Michigamme Formation consists chiefly of slate, especially quartzose, micaceous, and graphitic varieties, subgraywacke, quartzite, conglomerate, dolomite, dolomitic quartzite, and some iron-formation. More recent exploration drilling also has identified units of mafic volcanic rocks. An unconformity between the Michigamme and underlying Vulcan Iron-formation is indicated by the presence of basal a conglomerate, reported from a few localities, that contains clasts of iron-formation and other Menominee and Chocolay Group lithologies, and by regional truncation of pre-Baraga Group units beneath the basal Michigamme units. Although the Michigamme Formation lies on the Vulcan Iron-formation along both the north and south ranges, the Vulcan is largely absent to the north. The stratigraphic section bounding the Archean Carney Lake Gneiss consists of only the Chocolay and Baraga Groups, with the Menominee Group absent except for the extreme eastern end of the area. These relationships suggest that a topographic high existed to the north of the Menominee Range during or shortly after the time of Menominee Group deposition. Mesoproterozoic- The only rocks of Mesoproterozoic age are thin dikes of unmetamorphosed and undeformed diabase of probable Keweenawan age. They are known mostly where they cut the Carney Lake Gneiss. Typical dikes are only a meter or two wide and commonly have chilled margins against the rocks into which they are intruded. Structure. The Menominee iron district (Figure 3) is a south-facing homocline of Paleoproterozoic strata in which stratigraphic repetitions are created by two major faults and by folding internal to fault slices (Bayley et al., 1966). The faults cut the folds longitudinally, approximately along the fold axes, repeating the Paleoproterozoic sequence three times. The structural elements of the Range are shown in Figure 3, reproduced from Bayley et al. (1966, figure 22). The 3-D geometry is shown in Figure 4, reproduced from Bayley et al. (1966, figure 23). On the north, the Carney Lake Gneiss forms the core of a broad anticlinal structure (Figure 3). The Paleoproterozoic strata lie unconformably on the gneiss and dip steeply to the south or are overturned (as at Stop 2) and dip steeply north and face south. Interestingly, on this northernmost sequence of strata the Menominee Group, including the Vulcan Iron-formation, is absent and the Michigamme Formation lies directly on the upper unit of the Chocolay Group. This suggests that there was uplift in the area of the Carney Lake Gneiss concurrent with or shortly after deposition of the Menominee Group, creating a topographic high. In the south, the Paleoproterozoic strata are repeated twice by major faults to form the two ranges of the district. These faults were named the North Range fault and South Range fault by Bayley et al. (1966). The faults have steep dips at the present level of exposure and consistently show southside-up displacement. More recent interpretations (e.g., Sims and Schulz, 1993) consider them to have been thrust faults, which were steepened by continued shortening of the thrust panels. The rocks in the hanging wall (south side) of these faults have no indications of Archean basement rocks, in contrast to the area immediately to the north where the Carney Lake Gneiss is an integral part of the structure. The north range and south range panels may be allochthons detached from basement and thrust northward over the more autochthonous sequence of the northern part of the district. The Menominee Range is bounded on the south by the Niagara fault, along which it is in contact with volcanic rocks of the Wisconsin magmatic terranes. 73 �Figure 3. Structural elements of the Menominee Iron Range from Bayley et al. (1966, figure 22). The “south fault” is now referred to as the Niagara fault and is recognized as the suture between continental margin assemblages to the north and the accreted Wisconsin Magmatic Terranes to the south. 74 �Figure 4. Block diagram showing distribution of stratigraphic units of the Menominee Iron Range, from Bayley et al. (1966, figure 23). Iron deposits Iron ore was discovered in the Menominee district in 1848 by two explorers, J.W. Foster and S.W. Hill, according to Winchell (1895). However, iron mining did not begin until 1870, when N.P. Saxton started digging pits and trenches on the site of the Breene Mine, with the first ore being shipped in 1873 (Bayley et al., 1966). All the major mines had been opened by 1878. Production continued until 1946, with a total production from the district of approximately 85,000,000 tons (Bayley et al., 1966). Seven mines produced nearly 77,000,000 tons of ore, with a majority of the production from the district coming from three mines, the Chapin (27,500,000 tons), the Penn (21,700,000 tons) and the Aragon (11,200,000 tons) (Dutton, 1958). Production from the Chapin Mine ended in 1934 with a major collapse of the workings. The subsidence from this collapse formed the lake on the north side of Iron Mountain. A causeway across the lake now carries the traffic on Highways US 2 and US 141. Ore from the district was hauled by rail to Escanaba, Michigan; from there it was carried by boat to steel mills on the lower Great Lakes. The majority of the ore shipped from the district was high-grade (>50 % Fe) natural iron ore. Some ‘siliceous hematite’ ore (40-50% Fe) was produced from the Millie Pit and the Traders Pit. The Traders ore was used at one time for an experimental project at the Ardis Furnace where an attempt was made to high grade the ore utilizing high temperature roasting. The ruins of the Ardis Furnace are now on the National Register of Historic Places and can be visited near downtown Ironwood. The last of the mines in the area was the Groveland Mine in the Felch trough that operated until the early 1980’s using beneficiation methods to process both hematite and magnetite ore with complex iron silicates to produce a concentrate. 75 �Although the iron-formation in the Menominee district was studied as a possible source of beneficiating ore ("taconite ore"), no commercial operation has been undertaken. In the early 1950’s the Oliver Mining Company completed extensive diamond drilling and evaluation of underground ore reserves in the Vulcan – Norway portion of the Range and had designed open pit operations to extract ore from the Curry and Traders Iron-bearing Members. The area included the Curry, Brier Hill, and Aragon shafts which were the deepest mines on the Range at over 2000 feet. FIELD TRIP STOPS Stop 1: Carney Lake Gneiss (45.873°N, 87.86°W) The Carney Lake Gneiss occurs north of the Menominee Iron Range and forms the Archean basement on which the Paleoproterozoic strata of the Range were deposited. The Carney Lake Gneiss was defined and described by Bayley et al. (1966) who mapped the unit in some detail and published maps and lithologic descriptions of it, but did not attempt to decipher its obviously highly complex internal history. The general descriptions of the Carney Lake below are extracted from that publication. According to Bayley et al. (1966, p. 20-29) “Granitic gneiss constitutes about 85 percent of the Carney Lake Gneiss; of the remainder, about 5 percent is granodiorite and syenite dikes, and about 10 percent is inclusions of older rock. The gneiss is not uniform in composition or appearance, but varies from a gray plagioclase-biotite gneiss to red microcline-biotite gneiss. For the purpose of discussion these types will be designated gray gneiss, composite gneiss, and red gneiss, respectively.” “The gray gneiss probably constitutes about 25 percent of the complex collectively called the Carney Lake Gneiss. It is most abundant in the northern half of the complex where it contains many amphibolite inclusions. In thin sections the gray gneiss shows abundant plagioclase, quartz, and biotite. The foliation is well shown by aligned biotite, and also by the plagioclase and quartz, which are arranged in subparallel elongated grains and in lenticles. Cataclastic structures are common.” “The composite gneiss constitutes at least 70 percent of the complex. It is present almost everywhere, but is more abundant in the southern half of the area, where it contains minor patches of red gneiss and many inclusions of biotite schist. … The grain size of the composite gneiss ranges from medium to coarse. The gneiss is streaky and consists of red and gray elements, the red parts composed of pink microcline and quartz, the gray parts chiefly of plagioclase and biotite, which are the same minerals that constitute the gray gneiss. At some places the red part forms patches and streaks within the gray, at others the gray is enveloped by the red, and at still others the two elements form alternating layers. The red part commonly occurs as veins or layers of coarse pegmatite which cut across the foliation or bifurcate and join other layers. Here and there veins and layers of the red material swell and form large pods of pegmatite which fade transitionally into the gray rock. Pegmatite pods may also pinch and swell along the strike of the foliation of the gneiss. Locally they stop abruptly against the gray rock, only to appear again further along the strike. … The red gneiss is medium grained and weakly 76 �foliated. Fresh specimens are pink or red, whereas weathered specimens are brownish pink. Like the composite gneiss, the red gneiss consists of two components of different age, an older part composed of extensively altered plagioclase, and a younger part that consists of quartz, microcline, muscovite, and minor amounts of albite, but the red gneiss generally contains more quartz and microcline and less biotite and plagioclase than the other gneisses.” “Granodiorite occurs as rare dikes that are most abundant in the southern half of the complex and are more likely to be closely associated with the composite gneiss than the other types. The granodiorite is massive, equigranular, pink, medium to fine grained, and brownish-pink weathering.” “Inclusions in the Carney Lake Gneiss constitute less than 10 percent of the complex, but bear importantly on the character and origin of the gneiss. They consist of amphibolite, biotite schist, and metasedimentary rocks. The inclusions are clearly older than the gneiss and may represent ( 1) engulfed parts of the pre-gneiss Dickinson Group which occurs to the north of the complex consists, in part, of a metamorphosed series of basic tuffs, graywacke-type deposits, and basaltic flows (James et al., 1961), or ( 2) engulfed parts of the Quinnesec Formation which occurs to the south of the complex and consists of metavolcanic rocks, greenstone, amphibolite, and schist.” (Note: Radiometric ages determined since the Bayley et al. (1966) report indicate that both the Dickinson Group and Quinnesec Formation are younger than the Carney Lake Gneiss.) “The field relations show that (1) the gray gneiss grades into composite gneiss, into inclusions of amphibolite, and, more rarely, into inclusions of biotite schist, (2) this gneiss exhibits sharp contacts against the inclusions and appears as dikes in them, (3) the composite gneiss grades into red gneiss and into biotite schist, and (4) both the composite gneiss and the red gneiss contain inclusions of biotite schist and occur as dikes and stringers in some of the inclusions. Further, the gneisses are cut by red granodiorite dikes; one of these dikes, in turn, is cut by a late middle Precambrian metadiabase dike, and another metadiabase dike contains an inclusion of granodiorite. The relations of the syenite, which is known only in the southeast corner of the complex, and the grandiorite dikes are not clear, but the lack of foliation of the granodiorite dikes and the slight foliation of the syenite may indicate that the syenite was emplaced before the granodiorite dikes.” The descriptions in Bayley et al. (1966) and the relatively cursory examination that we have so far conducted in the Carney Lake Gneiss make it obvious that these rocks contain a very extended and complex history. In particular, our recent documentation of zircon grains with cores as old as 3.8 Ga (Ayuso et al, 2017; this volume) indicates that these rocks are undoubtedly part of the Gneiss Terrane defined by Sims et al. (1980) and contain vestiges of Eoarchean crust. Geochronology of samples of the Carney Lake Gneiss done using the USGS/Stanford Sensitive High-Resolution Ion Microprobe (SHRIMP) produced U-Pb data on zircons that confirms an Archean age (Ayuso et al., 2017; this volume). Two samples were collected for radiometric dating from the southern half of the complex: 1) sample 1 is from a granitic K-feldspar-bearing gneiss that is locally pegmatitic; 2) sample 2 is from a banded and folded gray to red granitic gneiss. Abundant zircons (70-200) were obtained from sample 1 that range from anhedral to subhedral, contain complex igneous and irregular growth zoning, and multiple growth rims; 77 �these zircons have irregular to pyramidal overgrowths. The zircons from sample 2 range from slightly rounded to subhedral and are otherwise mostly similar to zircons from sample 1. One hundred and twenty nine analyses of cores and rims were obtained. Individual zircons have older ages near their cores (mostly discordant) and younger ages near their rims. On a concordia diagram (Figure 5), U-Pb data plot as clusters of data points ranging from concordant to discordant and suggest several chords and intercepts that are common to both samples from the Carney Lake Gneiss (Figure 5). That study identified cores of individual zircons as old as 3.8 Ga. The most common age for individual zircons and for rims on older grains is about 2.75 Ga and records a younger major event in the late Archean (Figure 6). Figure 5. A-BSE (back scatter electron) image of a zircon from the Carney Lake Gneiss showing ages of four analyzed spots. B- Concordia diagram for 129 spot analyses from zircons in the Carney Lake Gneiss. 78 �Figure 6. Histogram showing the distribution of individual SHRIMP spot analyses of zircons from the Carney Lake Gneiss. Shaded areas are number of analyses. Solid line is relative probability. The gross structure within the Carney Lake Gneiss as mapped by Bayley et al. (1966) is a dome elongated in an east-west direction as defined by foliation and compositional layering of the gneisses. However, the age of the doming event is uncertain. Basal Paleoproterozic strata surrounding the dome are generally steeply dipping, in part overturned, and mostly concordant with the contact with the Carney Lake, indicating that much of the doming post-dates deposition of those strata that are as young as about 1850 Ma. Thus, much of the internal structure of the Carney Lake likely has a strong Paleoproterozic imprint superimposed on a complex set of Archean structures. A series of outcrops along a powerline east of Norway Truck Road provides a good example of various lithologies and structural complexity of the Carney Lake Gneiss. Figure 7 shows locations for outcrops and locations of photos in Figure 8. In general, the gneisses are more amphibolitic to the west and become progressively more granitic to the east, although a great deal of finer-scale complexity also is seen here. A few thin (1-3 m) mafic dikes that are younger than the complex gneissic structure can also be seen. 79 �Figure 7. Map showing location of outcrops of Carney Lake Gneiss along a powerline and location of photographs shown in Figure 8. 80 �Figure 8. Photographs of Carney Lake Gneiss along powerline. Locations shown on Figure 7. A-Amphibolitic gneiss cut by weakly deformed pegmatites. B-Contorted granitic gneiss with amphibolite inclusions. C-Amphibolitic gneiss with granitic stringers. D-Granitic gneiss with 81 �moderately dipping foliation. E-Foliated amphibolitic gneiss cut by undeformed pegmatite. FBanded gneiss with intrafolial isoclinal folds. Stop 2: Sturgeon River locality. Archean basement, Fern Creek Formation, and Sturgeon Quartzite. (45.784°N, 87.789°W) (This description is modified slightly from a previous field trip guide (LaBerge et al., 2003) that was written by Richard Ojakangas. This locality also has been modified in recent years by removal of a previous hydroelectric dam and draining of the impoundment. As a result, some additional outcrops have been created, especially of the Carney Lake Gneiss, but these are not the focus at this stop. The following description refers to the dam for locational purposes and is still useful in that vestiges of the dam can still be seen.) The Fern Creek Formation and Sturgeon Quartzite are the lower two formations of the Chocolay Group. The group is discontinuous and has been recognized in parts of the Marquette Iron Range and Gogebic Iron Range as well as here in the Menominee Range. The age of the group is bracketed between about 2.2 and 2.3 Ga based on ages of detrital zircon grains and hydrothermal xenotime (Vallini et al., 2006). The group appears to be equivalent to lithologically similar formations in the Huronian Supergroup in Ontario. Although the Sturgeon Quartzite is essentially continuous along the Menominee Range, the Fern Creek is preserved only locally, one of the best exposures being at this stop. Sericitic sediments near the top of the Fern Creek have been interpreted to be reworked paleosols formed prior to deposition of the Sturgeon Quartzite as discussed below (originally in Ojakangas’ stop description). If true, there is a disconformity between the Fern Creek and Sturgeon, which may account for the very limited preservation of the Fern Creek. Here the Sturgeon River has cut a deep gorge through the Sturgeon Quartzite; the formation was named for this locality. This small area has been well studied, especially because of the presence of the Archean-Paleoproterozoic contact at the dam. The area has been described by Credner (1869), Brooks (1873), Rominger (1881), Irving (1890), Bayley (1904), Lamey (1937), Pettijohn (1943), and Trow (1948). Substop 1. Walk past the gate to the end of the road at the powerhouse and dam. We will traverse back up the road to the vehicles, thus observing the rock units in stratigraphic sequence. The dam was constructed on Sturgeon River Falls, which was held up by a thick mafic dike that can be seen in the woods off the east end of the dam. The unconformity between the Archean Carney Lake Gneiss and the Paleoproterozoic Fern Creek Formation can be seen in a small ground-level exposure adjacent to the dam (Figure 9). The lowest bed in the Fern Creek is a diamictite at this spot, whereas a short distance to the west on the river bottom by the power station, the lowest unit is arkosic sandstone with rare oversized stones. 82 �Figure 9. Unconformity at Sturgeon Dam. Hammer head rests on Archean Carney Lake Gneiss and hammer handle is on basal diamictite of the Fern Creek Formation. Nearby in the river bottom, the basal unit of the Fern Creek is arkosic sandstone with rare dropstones, illustrated in Figure 11. Figure 10. Stratigraphic column at the Sturgeon River locality. SQ at the top of the column designates Sturgeon Quartzite. 83 �Figure 11. Granitic dropstone in lowest sandstone of the stratigraphic column. Note that the stone has pierced and bowed down the underlying strata. Figure 10 is a measured column of the Fern Creek Formation. The lower 25 m are well exposed when there is no water in the channel. Note that this portion of the formation consists of five beds of diamictite (matrix-supported conglomerate) as thick as 2.5 m, three arkosic sandstone beds as thick as 2.6 m with rare oversized stones, stacked arkosic sandstone beds with minor intercalated siltstone and argillite laminae, an argillite bed 4.5 cm thick, and a 15 cm conglomerate. Interestingly, the well-exposed section seen in the river bottom is not found on the west bank of the river; only 1½ m of conglomeratic rock is present there. Apparently, the more complete section is preserved in a topographic low on the Archean surface. However, faulting may be a factor as well, for weathered pyrite is present along a fault between the Archean basement and the Fern Creek west of the powerhouse. The middle 25 m of the Fern Creek Formation is relatively poorly exposed; Figure 10 shows this portion consisting of conglomerate, graywacke sandstone with oversized stones, and arkosic sandstone with oversized stones. Interpretation: This is a glaciogenic sequence. The diamictites may be thin tills deposited beneath glacial ice, but more likely are debris flow deposits as suggested by one diamictite bed that grades upward into sandstone. Some of the conglomeratic beds are difficult to clearly classify as 84 �either matrix-supported or clast-supported. One 20 cm bed at the 15 m level in the section is graded from medium sand to clay, suggestive of a turbidity current mechanism. Several of the oversized stones in the sandstone and greywacke beds show either a bowing down of the underlying laminae or an actual penetration, indicating that the stones were dropped into the basin from above and are indeed dropstones. Other lonestones may be dropstones, too, but clear evidence is lacking. The likely mechanism for deposition of dropstones is release from melting icebergs or from a floating glacier. Substop 2: The 25 m section between 50 and 75 m on Figure 10 is intermittently exposed on the west bank of the river, but this area is usually inaccessible because of high water. It includes beds of sericitic quartzite interbedded with sericite schist. The sericitic nature of this interval is illustrated by a small road-level outcrop between the road and the river just north of the quartzite ridge. This is a sericitic quartz pebble conglomerate with sericite clay chips, some reddish rather than yellow-green in color. Interpretation: This sericitic portion of the column is interpreted as a reworked paleosol that formed on the Fern Creek Formation during a warm climatic period that followed glaciation. Trow (1948) first suggested that this might be a paleosol. Substop 3: Sturgeon Quartzite ridge. Note that the bedding is slightly overturned towards the south, and that cross-bedding indicates that stratigraphic tops are to the south. Cross-bedding is of both trough and planar types. According to Trow (1948), the general cross-bedding indicates a paleocurrent trend from the northwest toward the southeast. Since the original field trip guidebook was prepared (LaBerge et al., 2003), geochronological studies (Vallini et al., 2006) have constrained the age of the Sturgeon Quartzite, and by inference of the underlying Fern Creek Formation. Most detrital zircons have ages between 2.5 and 2.7 Ga, but there is also a well-defined cluster of ages at about 2.3 Ga, thus providing a maximum age of deposition. Xenotime overgrowths on zircon grains are as young as 2.1 Ga and define a minimum age. These ages are consistent with age ranges determined for equivalent units in the Marquette and Gogebic Iron Ranges. Interpretation: Abundant asymmetrical ripple marks have low ripple indices (wave length/ripple height) indicative of deposition by water rather than by wind. The beds are generally of even thickness, indicative of a shallow marine rather than a fluvial environment of deposition. Stop 3. Underground tour of the Iron Mountain Iron Mine. (45.782°N, 87.864°W) The Iron Mountain Iron Mine has been in operation as a tourist locality for 60 years and thousands of people have enjoyed this historic site (Figure 12). The #2 adit is one of three exploration tunnels driven perpendicular to the strike of the Vulcan Iron-formation in search of high-grade (>50% Fe) ore in the early 1870’s along the south side of Brier Hill (Figure 13). The #1 adit was driven north to intercept at depth high grade ore that cropped out at the surface to the west of #2 adit. The first iron ore production came from this operation which was a combination of open pit and underground mining. The exploration then turned east along strike to explore the Traders Ironbearing Member of the Vulcan Iron-formation, but lost the member due to longitudinal faulting. The #2 adit was completed and crossed several faults and duplication of beds to find the Traders Iron-bearing Member 1,000 feet into the hillside. (Figures 14, 15). Further exploration sited the #3 adit a half a mile to the east and intercepted the complete section of “Hanbury Slate” 85 �(Michigamme Formation of present usage), Curry Iron-bearing Member, Brier Slate Member, and Traders Iron-bearing Member. The #3 East Vulcan Shaft was located based on finding high grade ore at this locality. Figure 12. The gateway to paradise. Portal to thousand-foot-long adit of the Iron Mountain Iron Mine. The tour will allow for the observation of a complete stratigraphic section from the “Hanbury Slate” (Michigamme Formation) at the entrance of the adit, through the Loretto Slate, Curry Iron-bearing Member into the Brier Slate Member (duplicated by folding and faulting), then through the Traders Iron-bearing Member and terminating at the “Footwall Slate”, which here is a breccia. These explorations were all drilled with hand-held drill bits and sledge hammers. Black powder was used for blasting and all material was hand loaded into tram cars and pulled with mules. The tunnel next was turned west to follow the contact of the Traders Iron-bearing Member and the footwall, but only goes about 70 feet before it intercepted another fault. The iron formation is brecciated and it is believed that the tunnel is near the subcrop of the formation with the overlying Cambrian sandstone because much of it is filled with friable sand through this section. The tunnel meanders a little and gains some elevation while intercepting brecciated and shallow dipping broken banded iron-formation as it comes to an intersection where tunnels go off in three directions (Figure 14). 86 �The tunnels to the north and west explore a wedge of Traders Iron-bearing Member and outlined a block showing low angle dips to the south. At this point a small ore body was found and extracted leaving a stope with broken slabs of rock. Along the south edge is a decline filled with spring water that wraps around a pillar. The stope connects to a shaft about 30 feet deep that was used to determine if more ore occurred at depth. The west tunnel explores the Traders Ironbearing Member about 500 feet along the East Vulcan longitudinal fault and terminates at the zone of caving to the west. The south tunnel crosses the strike of the Traders Iron-bearing Member and Brier Slate Member again and intercepts another fault where high grade ore is located. This passage was developed from the Old Central Shaft which was sited near the #1 adit. The extraction of ore at this location using sublevel caving methods left this large void (stope) as the ore was removed from below and the waste rock allowed to settle and partially prop up the workings. Several items should be noted: In the large stope, the unconformity between the iron formation, which strikes N 75° W and dips 70° SW, and the horizontal sandstone is striking. The ground water rain can be heard when it is quiet. The mined bottom of this ore block is about 600 feet to the southeast where it is cut by the large longitudinal fault and offset to the south. At that point in the operations, a ‘New’ Central shaft (Figure 18) was completed and used to extract ore to the 1200 foot level and develop ore to 1,600 feet toward East Vulcan #4 shaft. That was the extent of mining at the end of WWII when it became too costly to mine these orebodies without significant upgrades to equipment. Mines to the west: the West Vulcan, Curry, Brier Hill, and Aragon, were connected by tunnels to this property and went to depths of 2,400 feet. Figure 13. Plat of the Central Vulcan are in 1938. Geological interpretation and mine sites including subsidence area. 87 �Figure 14. Geological interpretation of the Iron Mountain Iron Mine adit area at the ‘Tunnel Level”, which correlates to the 1st Level of the old Central Vulcan Mine. Based on a blueprint from the Penn Iron Company, circa 1900 (see Figure 15). Walking tour, shown in heavy black line begins at the portal. Figure 15. Blueprint of Penn Mining Company adit #2. The initial straight section of the adit trending N 20o E, is 1,000 feet long for a scale reference. Note the offset of the Traders IronBearing Member on this image and Figure 14 and the location of the ore relative to fault geometry. When compared to other mines in the Range, this structure displays east dipping fault control where all mines to the west have west-dipping ore bodies. 88 �Figure 16. Traders Iron-bearing Member, shattered martite-jasper displaying 30o dip to the south into the wall. A carbide light for scale. Figure 17. Quartz druze on Jasper Iron Formation from the Central Vulcan Mine. Specimen was collected from a retaining wall on the north side of the west parking lot at the shaft. The wall contained several specimens of trace minerals that cement brecciated martite ore including calcite and pyrite. Only the quartz specimens survived the weathering since closure in 1946. The minerals occurred in “water courses”, in the ore bodies, usually in vertical channels as described by local miners. 89 �Figure 18. The New Central Shaft, Vulcan, Michigan. View looking west. Old Central shaft, circa 1877, was due north 1,500 feet’ (500m) and the underground workings connected at the 6th and 9th levels between the shafts. Stop 4. Randville Dolomite. (45.806°N, 87.951°W) The Randville Dolomite, the youngest formation of the Chocolay Group, is exposed extensively in the region. It is well described by Bayley et al. (1966, p. 35) from which the following description is excerpted. “Massive clastic dolomite makes up a large part of the Randville Dolomite and is closely associated with thick- and thin-bedded sandy dolomite, dolomitic and quartzose slate and phyllite, and pebbly dolomite conglomerate. Thick beds of nearly pure crystalline dolomite are present in some areas and probably make up an important part of the formation. A most distinctive rock type in the formation shows algal structures (stromatolites). These are domical, 1-3 inches high, 3-12 inches in diameter, and composed of nested laminae of pure dolomite. The algal structures occur nearly every place in the district where the dolomite is exposed. They form reefs as much as 50 feet thick and of great but undetermined linear extent. They are also present in the Randville Dolomite of central Dickinson County (James et al., 1961) and in the Kona Dolomite of the Marquette district. As pointed out by James, stromatolite structures are also reported in nearly all dolomite of late Precambrian age-in the western United 90 �States and Canada, Australia, South Africa, and Fennoscandia-and most geologists now accept the view that they represent fossil algal colonies. In the mapped area the algal dolomite is usually associated with thin-bedded sandy and conglomeratic dolomite of shallow-water deposition. This general association may be best observed in the outcrop area southeast of Lake Antoine, where algal dolomite, ripple-marked sandy dolomite, and thin dolomite beds showing mud cracks occur together.” At Stop 4, an active gravel/stone operation, recent activities have exposed conglomeratic dolomite. The rock is a poorly sorted, thick-bedded, intraformational conglomerate composed almost entirely of clasts of dolomite as much as about 10 cm diameter. Clasts visible in hand specimen range down to sand-sized grains. All clasts are composed of very fine-grained gray to pinkish dolomite. The matrix is somewhat darker, coarser-grained dolomite. The rock appears to be an intraformational conglomerate and we have seen no exotic clasts that would indicate clastic input for a distant source. We also have not seen any clearly biogenic features here. The total thickness of this conglomeratic unit was not exposed in September, 2017, but it appears to be at least 10 meters thick. Figure 19 illustrates the typical lithology. Figure 19. Conglomerate composed entirely of clasts of Randville Dolomite. Stop 5. Michigamme Slate. (45.777° N, 87.889° W) Brickyard Road, Norway Michigan. Exposures are on private property to which we have been granted access for this field trip. The bedrock ridge south of Hanbury Lake in sections 15 and 16, T. 39 N., R 29 W. contains the most extensive exposures of the Michigamme Slate on the Menominee Range. These rocks have been informally referred to as the Hanbury Slate in some early reports, but were renamed the Michigamme Slate by Bayley et al. (1966). 91 �Figure 20. Geologic map of the Michigamme Slate in the Hanbury Lake area. The figure shows parts of plates 2 and 3 of USGS Professional Paper 513 (Bayley et al., 1966). The area is entirely underlain by Michigamme Slate except for a few bodies of intrusive metagabbro (pCmg). Areas of outcrop are shown by darker shade. Various lithologies of the Michigamme are indicated as sl-slate, qtz-quartzite, dolo-dolomite, gw-graywacke. Stop 5 is near the west end of the ridge where lithologies change from highly sheared greywacke slate to dolomitic slate intruded by metagabbro. Folds indicate a duplication of strata that impacts true thickness estimates of the Michigamme Slate in the southern portion of the Range. Complex folding of the sediments is evident in outcrop along with a significant change in lithology which includes greywacke, quartzite, carbonate, and pyritic carbonaceous shale. Folds have vertical to steep southward-dipping axial planes as indicated by the prominent foliation, and plunge from 30-40° to the east. In USGS Monograph XLVI (Bayley, 1904), a detailed description of the area describes both large (meters) to small (centimeters) scale folding that exhibits strike and dips normal to the regional strike of the range (N 75° W). On the west and north ends the area the slate is cut by metadiabase dikes. These dikes are likely the same age as those identified in the mine workings at the Penn Mines Central Shaft, and at the Cyclops and Norway mines open pits on Norway Hill. Descriptions and discussion of the area from USGS Professional Paper 513 (Bayley et al., 1966, p. 60) follows: Dolomitic rock.- Dolomitic quartzite, dolomitic shale, and dolomite occur chiefly in the broad belt of outcrops south of Hanbury Lake. Dolomitic quartzite occurs south of Hanbury Lake only, where it is associated with dolomitic slate and dolomite and with intrusive metagabbro. The quartzite beds appear to be confined to the eastern three-fourths of the group of outcrops south of the lake, probably because the quartzite beds lens out to the west or are doubled back in a fold. Numerous minor folds in the slate show small areas where the beds strike north across the overall northwest foliation, and folding is thus indicated as the more likely cause of the limited distribution of the quartzite. 92 �The dolomitic quartzite is dark grey or, if encrusted with limonite, brown. The beds are 1-10 feet thick and commonly show quartz-filled cross-fractures that do not enter the adjacent slate beds. A distinctive characteristic of the quartzite is the presence of chips of black slate as much as 6 inches long in most beds. The rock is made up of about equal parts of well-rounded and wellsorted quartz grains and dolomite, and trace amounts of carbonaceous dust. The quartz grains all show undulatory extinction when viewed under the microscope, a feature probably inherited from the source rocks inasmuch as the quartzite does not appear to be deformed internally. The only exposed dolomite in the formation is confined to a belt of outcrops trending northwest from south of Hanbury Lake. The northwestern most rocks on the belt are dolomitic slates which outcrop in secs. 4 and 5, T. 39 N., R. 30 W. South of Hanbury Lake the dolomite is light colored, banded, and somewhat slaty: it occupies the north part of the group of outcrops. The best exposed rock is at the lakeshore. The beds are folded, and in the northern-most outcrop the general strike is nearly normal to the trend of the outcrop belt; The dips are low to the southeast, because these beds are on the crest or in thee trough of a minor fold. The lesser plications and folds on the beds plunge at low angles, less than 30 degrees SE. West of Hanbury Lake, in the south parts of secs. 7 and 8, T. 39 N., R. 29 W., are outcrops of siliceous and dolomitic grey slate and rather thick-bedded siliceous gray dolomite. A rind of limonite that coats the exposed surfaces of the dolomite indicates that the carbonate is probably ferruginous, an observation previously made by W. S. Bayley, who reported the chemical analysis shown in Table 29. On the assumption that all the iron , magnesia, and lime form carbonates, W. S. Bayley gives the composition of the carbonate as about 9 percent FeCO3, 41 percent MgCO3, and 50 percent CaCO3. Stop 6. Niagara fault splay at Piers Gorge. (45.759°N, 87.942°W) (text reproduced from 2003 ILSG guidebook. Laberge et al., 2003) Rocks exposed along the Menominee River at Piers Gorge are almost certainly a branch of the Niagara fault zone and represent one of the few exposures of the fault zone. This location is about one kilometer south of the mapped trace of the Niagara fault. The hill lying north of the gorge, but still south of the mapped fault, is underlain by metagabbro that is much less deformed than the rocks in the gorge. These relationships indicate that strain along the fault zone was distributed very heterogeneously and concentrated in discrete zones of very high strain surrounding islands of weakly deformed rocks. The rocks in the gorge are highly foliated and lineated quartz-sericite schists and chloritic schists, probably developed from felsic and mafic volcanic rocks. Felsic and mafic volcanic rocks with only weak foliation, along with mafic sills with little internal deformation, are exposed on both sides of this strongly foliated zone. Metagraywacke of the Marquette Range Supergroup is exposed in Norway, about 2 miles north of this locality, and volcanic and plutonic rocks of the Wisconsin magmatic terranes are exposed along the Menominee River in this area. The foliation here strikes N 80-85° W and dips 80-85° N. and has a stretch lineation that plunges 60-65°, N 85° W. As the recognized boundary between the dominantly sedimentary rocks of the Marquette Range Supergroup to the north and the Wisconsin magmatic terranes to the south, the Niagara fault zone is commonly referred to as a suture. However, it lacks some features 93 �(such as a mélange) that are typical of suture zones. Geophysical evidence (Attoh and Klasner, 1989; and LaBerge and Klasner, 2001) suggests that thinned continental crust of the Superior craton has been overridden by the Wisconsin magmatic terranes, and extends in the subsurface for 10-50 miles south of the Niagara fault zone. If this is the case, the Niagara fault zone may be the frontal thrust on which oceanic rocks of the Wisconsin magmatic terranes overrode the continent margin assemblage of the Marquette Range Supergroup. Continued compression of the suture zone resulted in the steepening of the thrust surfaces into their present, nearly vertical orientation. Figure 21. Schist in Niagara fault zone at Piers Gorge. A- Highly foliated schist along north bank of the Menominee River. B- Nearly vertical foliation with modern slump toward river producing a spurious shallower foliation. Both photos from 2003 ILSG field trip. Stop 7. Quinnesec Mine (45.810°N, 87.991°W) The abandoned workings of the Quinnesec mine (known locally as the Devil’s Icebox) are mainly in the Traders Iron-bearing Member of the Vulcan Iron-formation. The property is fenced and accessible by arrangement with the property owner. The mine lies on the overturned north limb of a second-order syncline (Figure 22). The Precambrian strata at the mine dip about 60° north, but face southward, inasmuch as the Brier Slate Member of the Vulcan is along the south side of the excavated approach to the mine, and the Felch Formation is along the north wall of the workings. Cross-sections through the Vivian Mine immediately to the west of Stop 7 show the geometry of the westward plunging folds and the overturned structure we see at the mine exposure. The ore is specular hematite and jasper and the brecciated subcrop was considered an ore where it has been reworked in a shoreline environment during the Cambrian Period. 94 �The mine workings provide an exceptional view of the unconformity and basal Cambrian sandstone overlying the mine workings along the north side of the hill (Figure 24 A, B). The basal portion of the sandstone contains numerous angular slabs of oxidized iron-formation, iron ore, and slate in a sandy matrix (Figure 24 C). Clearly, this area was a small island as the Cambrian sea advanced over the area. Cross sections (such as Figure 23) also show the steep local relief that existed on the Precambrian erosional surface. The complex folding and duplication of beds made for a more resistant area of iron-formation that likely led to the development of a topographic high. The clasts of iron ore in the basal conglomerate also indicate that the ore here was formed before the Cambrian sea covered the area. Figure 22. Geologic map of the Quinnesec Mine and vicinity showing that the workings were developed in the overturned northern limb of a small syncline. From LaBerge et al. (2003, based on mapping by Bayley et al. (1966)). 95 �Figure 23. A portion of plate XXX from U.S. Geological Survey Monograph XLVI (Bayley, 1904). 96 �Figure 24. Photographs of the Quinnesec Mine workings. A- View looking west into the mine workings. The Traders Iron-bearing Member of the Vulcan Iron-formation dips steeply north and faces south. The unconformity with the overlying Munising Sandstone (Late Cambrian) is well exposed and shows steep topography that existed on the Precambrian units during Cambrian marine transgression. Photo by Thomas Waggoner. B- Close-up view of the unconformity. C- Basal breccia of the Munising Sandstone, probably talus deposited at the base of the paleoescarpment formed by the Vulcan. Large clasts are entirely iron-formation, many of which show secondary iron enrichment. Lighter matrix is quartzose sand. 97 �Stop 8. Keel Ridge area (45.810°N, 88.028°W) The area of the previous Keel Ridge mine is now operated for crushed stone as well as being excavated for future business development. The Keel Ridge mine was one of the earliest mines opened on the Menominee Range in 1880, but only produced until 1899 with a total production of 93,101 tons. The mine was located just to the northwest of the large stripped area we will examine at this locality. The stratigraphic section exposed is from the Randville Dolomite on the north to the Michigamme Slate (“Hanbury Slate”) on the south and includes an excellent cross section of various members of the Vulcan Iron-formation. Although the area is easily accessible from U.S. 2 and the various units of the Vulcan Iron-formation can be observed and sampled, it is private property and should not be entered without permission of the owner. Figure 25. The Keel Ridge mine area. Geologic units are as indicated in USGS Professional Paper 513, Plate 1 (Bayley et al., 1966). Excavation reveals the northwest-striking formations that include the upper part of the Felch Formation (“Traders Slate”), Traders Iron-bearing Member, Brier Slate, and Michigamme Slate (“Hanbury Slate”). The formations face to the south and dip to the north at 80o- 90o. 98 �Figure 26. View looking east on east side of exposure showing the upper part of the Felch Formation (“Traders Slate”) and conformable contact with the Traders Iron-bearing Member of the Vulcan Iron-formation. Dips are vertical to overturned toward the south. The northernmost exposures are of the Randville Dolomite, which here consists of a breccia of angular siliceous material (chert?) in a siliceous carbonate matrix. This may be a residual accumulation of chert nodules and beds formed by solution of the carbonates of the Randville on the long-lived erosion surface, now expressed as the unconformity between the Randville and the Felch Formation. The next outcrop to the south is the Felch Formation (“Traders Slate”) which is a rare exposure of this unit in the entire range. Bayley (1904) referred to the distinctive sericitic slate or quartzite found at the top of the Felch Formation as either the “Traders Quartzite” or the “Traders Slate” depending on the predominant lithology. The formation is described in USGS Professional Paper 513 (Bayley et al., 1966, p. 38) as: “Lithology of the Felch Formation is remarkably uniform throughout the length of the south iron range, but variable along the north range. On the south range the formation is about 100 feet thick and consists of thin-bedded sericite slate and phyllite, and intercalated thin-bedded quartzite. The quartzite layers appear to be prevalent in the upper part of the formation, and a thin (4 in. to 3 ft.) key bed of dark ferruginous quartzite, the so-called “Traders quartzite,” is commonly present near the top of the formation. The fine-grained clastic rocks which make up the major part of the formation on the south iron range include slate, phyllite, siltstone, and schist. All these rocks show minor differences imposed during deposition and modifications imposed by later deformation and low-rank metamorphism, but they bear a close outward resemblance to one another and show a common mineralogy. 99 �They are predominantly thin-bedded rocks, the layers commonly less than 1 cm thick, and most show bedding-plane fissility. Cleavage surfaces of phyllite are lustrous and spangled with tiny plates of white mica. On fresh surfaces the rocks are gray to greenish gray, but where weathered they may be pale green, red, or light buff, or almost white and mottled with red; color banding is not conspicuous. The chief mineral components of all fine-grained types are pale green sericite and quartz; minor components are feldspar, chlorite, biotite, hematite, and magnetite. Most of the rock layers are composed of about equal parts of the two chief components, but layers composed predominately of one or the other are common. Medium to coarse well-rounded grains of quartz and potassic feldspar, commonly visible to the unaided eye, are scattered throughout many specimens of the slaty rock and form wafer-thin discontinuous quartzite stringers between the slaty layers. These latter characteristics of the slate are useful but not infallible criteria for identifying Felch strata in the field. The prevailing texture of the rocks is micro-schistose. In some specimens the quartz grains as well as the sericitic groundmass are elongated in the plane of schistosity, which at most places parallels the bedding.” Several tens of feet are exposed at this location including the sharp conformable contact with the overlying Traders Iron-bearing Member. The two members of the Vulcan Iron-formation that were most significant economically are the Traders and Curry Iron-bearing Members. They were described in USGS Professional Paper 513 (Bayley et al., 1966, p. 43) as follows: “The rocks of the Traders and Curry members are iron formation, which has been defined by James (1954, p. 239) as “a chemical sediment, typically in bedded or laminated, containing 15% or more of a layer of sedimentary origin, commonly but not necessarily containing layers of chert.” The iron-formation of the Vulcan is thin bedded in commonly laminated, but it does not display uniformity in the thickness of the beds. Individual beds generally range from 1 mm to 30 cm in thickness. As a rule, beds of granular jasper alternate with beds composed chiefly of oxides of iron, principally hematite Fe2O3 (69.94% iron), and a lesser amount of magnetite Fe3O4 (72.4% iron). Almost all of the iron-rich layers contain a small amount of crystalline quartz, and at some places dolomitic carbonate and chlorite as well. The iron-formation usually is dark. Viewed from a distance it commonly appears dark gray or reddish-brown, but at close range it appears as a medley of deep red or maroon, metallic gray, and black. If much oxidized, hues of orange and red are dominant. Most jasper beds are maroon (liver colored) or red. They are generally thicker than adjacent iron rich beds and most are uniformly straight bedded, but irregular beds and lenticular beds are common. The jasper beds are composed chiefly of red jasper granules, specular hematite, magnetite, and metachert (a fine-grained mosaic of crystalline quartz). The granular character of most jasper beds can be seen by the unaided eye, but a wetted surface and a hand lens are helpful. In their primary state the jasper granules are a mixture of amorphous silica and red iron oxide. In their primary state the jasper granules are a mixture of amorphous silica and red iron oxide (fig. 13). In their characteristic crystallized state, the iron oxide is specular hematite, magnetite, or both, and the silica is crystalline quartz (fig 14). Most jasper beds contain, in addition to jasper granules, ooliths which are made up of concentric layers of red amorphous hematite and silica 100 �about a nucleus of quartz or jasper. Re-crystallized ooliths form the same products as the granules. The granules, in shape, size, and appearance, resemble the greenalite granules, that are so characteristic of the Biwabic Iron-Formation of the Mesabi Range. They may represent the analog of the greenalite granules, formed under oxidizing conditions.” Figure 27. Figure 13 from Bayley et al., 1966, p. 44. “Photomicrograph showing jasper granules in a metachert matrix. The very dense granules are bright-red noncrystalline jasper. The gray (salt-and-pepper) granules show an early stage of crystallization-segregation of iron oxide as hematite or magnetite from silica. Ordinary light.” 101 �Figure 28. Figure 14 from Bayley et al., 1966, p. 45. “Photomicrograph showing several types of crystallization of japer granules. Quartz, white; specular hematite and magnetite, gray or black. Note the wide variation in the iron to silica ratios from one granule to another. The black spicules are specular hematite, and the square to rectangular sections are magnetite. Most granules pictured contain both iron minerals. The mottled granule (top center) represents an early stage of crystallization and segregation. The shapes of the iron oxide segregations suggest incipient specular hematite. Plane-polarized light.” The Brier Slate is in fault contact with the Traders Iron-bearing Member and is oxidized at this location. The Curry Iron-bearing Member is missing either because of faulting or nondeposition. This interpretation is based on the formations exposed in the underground workings of the Keel Ridge mine. The Michigamme Slate (“Hanbury Slate”) is in fault contact with the Brier Slate and is best exposed on the west side of the excavation. The outcrop shows significant shearing and oxidation of the slate with probable duplication of section to the south Presented next are detailed descriptions of the Traders Iron-bearing Member, the Curry Ironbearing Member, and the Brier Slate by Oliver Mining Company geologists from crosscuts in the mines east of this locality. Comments in their archived records state that the direct shipping ores (+50% Fe) from the various mines could be identified by their physical and mineralogical characteristics. 102 �Traders Iron-bearing Member. A detailed description of the Traders Iron-bearing Member from the Penn Mining Co. to the east shows a measured section of 112’, less than the average 132’ measured in sections in the south range by the USGS. An abrupt conformable contact with the underlying “Traders Quartzite” occurs in the mine. Basal 3 feet: Ferruginous slate or slaty iron-formation consisting of a micaceous, medium-grained, even-bedded unit, with a high silica content slate. Iron content, 2632%. Earthy hematite occurs with specular hematite scattered throughout. Next 12 feet: Wavy-bedded jasper iron-information with fine-grained red (liver colored) chert bands. Beds consist of thin even slaty specular hematite laminae (not readily cleavable), and long narrow (1/4” – 4”), lenses of fine grained to saccharoidal red (liver colored) chert. Grades gradually to a granular chert phase (next layer), by change of color and finer grain size. Iron content 33-40%. Next 18 feet: Massive wavy bedded iron formation with reddish brown granular chert lenses. A fairly massively, bedded jasper. Thin laminae of slaty specular hematite (not readily cleavable) occur with heavy lenses of reddish-brown saccharoidal to granular chert. Top grades into overlying even-bedded iron-formation by decrease in the amount of granular chert. Bottom grades into red chert phase by a gradual change of color and becomes finer grained while the specular hematite background remains the same. Iron content is 34-40%. Next 64 feet: Even-bedded iron-formation with dark brown, fine-grained chert. (and occasional granular chert lenses). Even-bedded cherty iron-formation composed of thin (1/64” to 1/4”) laminae of slaty specularite with narrow (1/4”) lenticular bands of darkbrown, fine-grained chert. Near the base occur heavier bands of the dark brown, finegrained chert up to 2 inches in thickness which disappear toward the top and bedding there becomes uniformly thin and quite even. One section shows one of these heavy finegrained chert bands. Another section shows a typical thin, even-bedding of the higher phase. Occasional heavy lenses of dark, reddish-brown, granular chert occur frequently near the base where this unit grades into the granular chert below but they become rarer in the upper portion. Iron content is 34-43%. Next 15 feet: Massive dark granular chert. Massive, irregularly bedded, lean, granular chert, dark-brownish-purple in carbide light. Contains very little slaty ferrugenous matter and the chert is all granular with many red jasper granules. Specularite occurs in thin irregular veinlets through the body of the chert. Iron content 25-33%. Curry Iron-bearing Member. A measured section from this locality is 158 feet in fault contact with Brier Slate. Basal 14 feet: Slaty basal phase: Even-bedded blocky, dark-brown, siliceous, slaty ferruginous rock, containing very little free chert. Laminations are 1/8 inch- to 1/2 inchthick and consist of slaty brown, hematite, becoming bluer with increase of specularite toward the top. What free chert exists is purplish granular Curry-type. This horizon seems favored for ore concentration. Iron content is 32-33%. Next 144 feet: Cherty phase: It is a heavy bedded, straight-bedded blocky specular cherty iron-formation with groups of thin, even, rich specular laminae alternating with 103 �irregular lenses (1/4 inch to 4 inches thick) of dark reddish-purple granular chert, shot through with the regular veinlets and mottles of specularite. The chert is invariably granular and dark reddish purple in color. Iron content is 32-43%. Brier Slate. The Brier Slate separates the Traders and Curry Iron-bearing Members in this mine cross-section. The slate is 104 feet thick and displays contacts that are faulted with both ironformations. The Brier Slate is a soft, fine to medium grained, thinly laminated, blocky, ferruginous slate. The color is very dark with a white streak where it is unoxidized, but it is generally oxidized to a chocolate brown color with a dull red streak. Bedding laminae are thin but prominent, especially near the bottom, and grain size varies by laminae, with the coarser grains in the heavier laminae. A coarse phase occurs near the middle. Concentration may produce higher iron analysis locally, but never an ore body. Iron content averages 23%, varying between 15 to 30%. References Attoh, K., and Klasner, J.S., 1989, Tectonic implications of metamorphism and gravity field in the Penokean orogen of northern Michigan, Tectonics, v. 8, p. 911-933. Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vasquez, J.A., and Jackson, J., 2017, Evidence for the presence of Eoarchean crust in northern Michigan, Institute on Lake Superior Geology, Proceedings of 63rd annual meeting, Part 1: Program and abstracts, p. 910. Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., and Jackson, J., 2018, New U-Pb zircon ages for rocks from the granite-gness terrane in northern Michigan, Institute on Lake Superior Geology, Proceedings of 64th annual meeting, Part 1: Program and abstracts. Bayley, W.S., 1904, The Menominee iron-bearing district of Michigan, U.S. Geological Survey Monograph XLVI, 513 p. Bayley, R.W., Dutton, C.E., and Lamey, C.A., 1966, Geology of the Menominee iron-bearing district, Dickinson County, Michigan, and Florence and Marinette Counties, Wisconsin, U.S. Geological Survey Professional Paper 513, 96 p. Brooks, T.B., 1873, Iron-bearing Rocks (Economic), Michigan Geological Survey, v. 1, Pt. 1, Chapters 1-4, 319 p. Credner, H., 1869, Die vorsilurischen Gebilde der "obern Halbinsel von Michigan" in Nord Amerika, Deutsche Geologische Gesellschaft, v. XXI, p. 51 6-554. Dutton, C.E., 1958, Precambrian geology of parts of Dickinson and Iron Counties, Michigan, Field Guide for Michigan Basin Society, 44 p. Irving, R.D., 1890, The greenstone schist area of the Menominee and Marquette regions of Michigan, explanation and historical notes, U.S. Geological Survey Bulletin 62, 241 p. 104 �LaBerge, G.L., and Klasner, J.S., 2001, Geology and tectonic significance of Early Proterozoic rocks in the Monico area, northern Wisconsin, U.S. Geological Survey Miscellaneous Investigations Series Map 1-2739, scale 1:24,000. LaBerge, G.L., Cannon, W.F., Schulz, K.J., Klasner, J.S., and Ojakangas, R.W., 2003, Paleoproterozoic stratigraphy and tectonics along the Niagara suture zone, Michigan and Wisconsin, 49th Annual Meeting of the Institute on Lake Superior Geology, Part 2 Field Trip Guidebook, 107 p. Lamey, C.A., 1937, Republic Granite or basement complex, Journal of Geology, v. 46, p. 48751. Pettijohn, F.J., 1943, Basal Huronian conglomerates of Menominee and Calumet districts, Michigan, Journal of Geology, v. 51, p. 387-397. Rominger, C., 1881, Menominee Iron Region: Michigan Geological Survey, v. IV, p. 190-192. Schneider, D.A., Bickford, M.E., Cannon, W.F., Schulz, K.J., and Hamilton, M.A., 2002, Age of volcanic rocks and syndepositional iron formations, Marquette Range Supergroup: implications for tectonic setting of Paleoproterozoic iron formations of the Lake Superior region, Canadian Journal of Earth Sciences, v. 39, p. 999-1012. Sims. P.K., Card, K.D., Morey, G.B., and Peterman, Z.E., 1980, The Great Lakes tectonic zone-a major crustal structure in central North America, Geological Society of America Bulletin, v. 91, p. 690-698. Sims, P.K., and Schulz, K.J., 1993, Geologic map of Precambrian rocks in parts of Iron Mountain and Escanaba 30' X 60' quadrangles, northeastern Wisconsin and adjacent Michigan, U.S. Geological Survey Miscellaneous Investigations Series Map 1-2356, scale 1:100,000. Trow, J.W., 1948, The Sturgeon Quartzite of the Menominee district, Michigan, Ph.D. thesis, Chicago, Illinois, University of Chicago, 60 p. Vallini, D.A., Cannon, W.F., and Schulz, K.J., 2006, Age constraints for Paleoproterozoic glaciation in the Lake Superior Region: detrital zircon and hydrothermal xenotime ages for the Chocolay Group, Marquette Range Supergroup, Canadian Journal of Earth Sciences, v. 43, p 571-591. Winchell, H.V., 1895, Historical sketch of the discovery of mineral deposits in the Lake Superior region, Geological and Natural History Survey of Minnesota, 23rd Ann. Report, p. 116-155. 105 �FIELD TRIP 4 Friday May 18, 2018 GRANITOID ROCKS OF THE PEMBINE-WAUSAU TERRANE IN NORTHEASTERN WISCONSIN Klaus J. Schulz, U.S. Geological Survey With a contribution from Marcia Bjornerud, Lawrence University INTRODUCTION This trip examines granitoid rocks of the Pembine-Wausau terrane that are exposed in northeastern Wisconsin. The Pembine-Wausau terrane is one of the Paleoproterozoic magmatic arcs that comprise the internal domain of the Penokean orogen (Figure 1; Schulz and Cannon, 2007). The terrane consists of mafic to felsic volcanic rocks ranging from tholeiitic to calcalkaline in composition, subordinate sedimentary rocks, and granitoid intrusive rocks of largely calc-alkaline affinity. It was accreted to the southern margin of the Archean Superior craton beginning about 1,875 Ma along a paleosuture now marked by a major ductile deformation zone, the Niagara fault zone (Sims et al., 1985). The rocks in northeastern Wisconsin have been key to understanding the stratigraphic and tectonic evolution of the Pembine-Wausau terrane and the nature of the Penokean orogeny. Rocks in the area are fairly well exposed, especially compared to other areas in northern Wisconsin. Granitoid rocks constitute nearly half of the outcropping rocks of the area and are mainly granodiorite and tonalite, but include gabbro, diorite, and granite (Sims and Schulz, 1993). An older suite, ranging in age from about 1,890–1,870 Ma, is dominantly calcic to calcalkaline and appears to be cogenetic with the volcanic arc magmatism, while younger, 1,860– 1,840 Ma, calc-alkaline to alkaline plutons are broadly contemporaneous with collision of the Pembine-Wausau terrane with the Superior craton margin (Sims et al., 1992). Younger posttectonic intrusions, emplaced at about 1,835 and 1,760 Ma, consist of alkali-feldspar granite suites (Sims et al., 1993). In the area of the field trip (Figure 2) most of the exposed granitoid rocks are part of the Dunbar dome, an irregular, asymmetrical structure ~470 km2 in area, composed of gneiss, migmatite, amphibolite, and foliated to unfoliated granitoid rocks mantled by steeply dipping sedimentary and volcanic rocks of Paleoproterozoic age (Sims et al., 1992). The Dunbar dome is one of several roughly correlative domes in northern Wisconsin (Morey et al., 1982) which have less well-exposed gneiss and granitoid rocks with comparable isotopic ages (Sims and Peterman, 1980; Sims et al., 1989) and chemical compositions (Sims et al., 1993). In the Dunbar dome, compositionally varied gneisses, assigned by Cain (1964) to the Dunbar Gneiss, are intruded by five major plutons named the Marinette Quartz Diorite, Newingham 106 �Tonalite, Hoskin Lake Granite, Spikehorn Creek Granite, and Bush Lake Granite (Sims et al., 1992). This field trip will examine exposures of each of the major units of the Dunbar dome except the Bush Lake Granite. Detailed descriptions of these major units including their petrography, structure, geochemistry, age, and petrogenesis are given in Sims et al., 1984; 1985; and 1992 (the 1984 reference is available for free download from the ILSG website http://www.lakesuperiorgeology.org/; the 1992 reference is available for free download from the USGS website https://www.usgs.gov/products/publications/official-usgs-publications). In addition to the granitoid units of the Dunbar dome, the field trip will examine exposures of the Twelve Foot Falls Quartz Diorite, a subvolcanic intrusion that is comagmatic with calc-alkaline volcanic rocks in the Quinnesec Formation (Sims et al., 1992) as well as the post-tectonic (~1,835 Ma) Athelstane Quartz Monzonite and Yavapai-age (~1,750 Ma) Amberg Granite which occur south of the Dunbar dome (Sims, 1990; Sims et al., 1993). Sims et al. (1992) concluded that the granitoid rocks within and outside the Dunbar dome were derived from different sources based on their contrasting chemistry, and presumably were developed in different tectonic environments, and were subsequently superposed tectonically. They attributed the Newingham Tonalite and Twelve Foot Falls Quartz Diorite to the subduction processes that formed the volcanic arc represented by the Quinnesec volcanic rocks. In contrast, the intrusions within the dome, which range from syn- to post-tectonic, were attributed to melting of continental lithosphere during collision of the arc with the continental margin of the Superior craton. However, these conclusions need to be reevaluated in light of acquired Nd isotope data (Van Wyck and Johnson, 1997; Schulz and Ayuso, 1998) and new understanding of the processes that produce granitoid rocks in orogenic belts (Hildebrand and Whalen, 2017). Although the granitoid rocks of the Dunbar dome range in composition from calcic tonalite to calc-alkaline granodiorite to alkali-calcic quartz diorite, they have a surprisingly small range of enriched ɛNd values centered on 0 and depleted mantle model ages of ~2.0 to 2.2 Ga. The Dunbar Gneiss has the most negative ɛNd(1,860) values of -2.1 to -3.4 and the Hoskin Lake Granite the most positive ɛNd(1,835) value of +1.71; the Marinette Quartz Diorite and Newingham Tonalite have similar ɛNd(1,860) values near 0 (+0.12 and +0.39, respectively). Only the Twelve Foot Falls Quartz Diorite, which is a subvolcanic intrusion, has a strongly positive ɛNd(1,900) value of +4.54 indicating derivation from a long-term light rare earth element (REE) depleted source. The narrow range of enriched ɛNd values for the Dunbar dome granitoids is unlikely to be the result of crustal contamination as it would be highly fortuitous for granitoids of such varying chemistry to all have similar degrees of crustal contamination. Instead, the narrow range in ɛNd values is more likely a characteristic of the source from which the granitoids were derived. In addition, the data indicate that the Newingham Tonalite is likely not a syn-volcanic intrusion, as suggested by Sims et al. (1992), but rather is a syn-collisional intrusion and part of the Dunbar dome suite. A characteristic feature of the Dunbar dome granitoids is that they are relatively enriched in Ba, K, Nb, Rb, Sr, Ta, and Th, and have steep, light REE-enriched patterns (Sims et al., 1992). Recently, Hildebrand and Whalen (2017) examined the geochemistry of a number of major 107 �Cordilleran-type batholiths including the Sierra Nevada, Peninsular Range, Idaho-Montana, and Cascades-Coast Plutonic Complex of North America among others. They noted a clear compositional distinction between plutons generated as syn-volcanic intrusions during subduction and those emplaced as syn- to post-tectonic intrusions during slab failure (breakoff). In particular, they showed that magmas generated during slab failure have relatively high Nb/Y, Sr/Y, and Sm/Yb ratios (Figure 3) as well as evolved radiogenic isotopes. These characteristics are shown by intrusive rocks ranging from gabbro to granite and calcic to alkaline in composition (Hildebrand and Whalen, 2017). They concluded that the distinctive whole-rock geochemistry, as well as radiogenic and stable isotope compositions, of slab failure magmas involve only minor amounts of crustal contamination and are derived mainly from plagioclaseabsent melting of garnet-bearing rocks in the mantle (for example, garnet pyroxenite, eclogite, and/or subcrustal lithosphere). As seen in Figure 3, the Dunbar dome granitoids, including the Newingham Tonalite, plot in the fields defined by Hildebrand and Whalen (2017) for slab failure magmas. Figure 1. Generalized geologic map of the Penokean orogen in the Lake Superior region showing approximate location of Figure 2 (from Schulz and Cannon, 2007). 108 �Figure 2. Geologic map for a portion of northeast Wisconsin showing the locations of the field trip stops (from Sims and Schulz, 1993). 109 �Figure 3. Plots of Nb vs Y (A), La/Sm vs Sm/Yb (B), Nb/Y vs Sr/Y (C), and La/Yb vs Sr/Y (D) for samples from the Dunbar dome and Twelve Foot Falls Quartz Diorite. Fields after Hildebrand and Whalen (2017). FIELD TRIP STOPS Stop 1. Hoskin Lake Granite (Outcrop on north side of County road N; 45.764° N., 88.071° W.; Note that the outcrop is on private property and permission from the owner is required) The Hoskin Lake Granite is an arcuate, convex-northward body of granite on the north margin of the Dunbar dome characterized by (1) pink to gray, medium- to coarse-grained inequigranular granite with large, oriented, tabular potassium feldspar crystals (Figure 4), (2) abundant inclusions of mafic-intermediate volcanic rocks of the Quinnesec Formation, and (3) late, euhedral crystals of potassium feldspar that lie athwart to an older foliation and, at least locally, transect centimeter-thin quartz veins. As noted by Cain (1964), the southern margin of the granite appears to be gradational into rocks assigned to the Marinette Quartz Diorite and evidence for K-metasomatism along the border of the two units is compelling. To the east, the Hoskin Lake Granite appears to grade into the post-tectonic Spikehorn Creek Granite. Although different in appearance, the two granites have similar compositions (Sims et al., 1992). Excellent descriptions of the Hoskin Lake Granite are given in Bayley et al. (1966) and Sims et al. (1992). An Nd isotope analysis of one sample of the Hoskin Lake Granite gave a ɛNd (1,835) = +1.71 110 �with a depleted mantle model age of 1.99 Ga (Schulz and Ayuso, 1998), suggesting derivation from a slightly light REE depleted source. Figure 4. Hoskin Lake Granite. Stop 2. Marinette Quartz Diorite (Railroad cut on County road O; 45.747° N., 88.033° W.) Note: Access to this railroad grade is strictly prohibited without prior approval of the owner. The railroad cut shows metamorphosed Marinette Quartz Diorite cut by dikes of Hoskin Lake Granite and leucogranite. The Marinette Quartz Diorite is a large, layered sill-like intrusive body that was emplaced in the contact zone between the Dunbar Gneiss and the Quinnesec Formation and the intrusive Newingham Tonalite. It is dominantly composed of quartz diorite and diorite with moderately high biotite (~10 to 20%), hornblende (trace to 30%), and sphene (trace to 6%) contents. In the north-central part of the Dunbar dome, the rocks contain variable amounts of potassium feldspar and are interpreted as hybrid rocks reflecting post-crystallization Kmetasomatism (Sims et al., 1992). In the eastern part of the body, which is relatively unmetamorphosed, the Marinette Quartz Diorite is a dark gray to black, medium-grained, hypidiomorphic granular rock with well- to ill-defined layering and generally lacks a penetrative foliation; rare clinopyroxene occurs as relicts in the cores of some hornblende crystals. South of Dunbar, the Marinette Quartz Diorite is medium to dark gray, mesocratic, layered quartz diorite and diorite cut by abundant granite pegmatite and aplite dikes (Figure 5). Based on the presence of mineralogical layering and geochemistry, the Marinette Quartz Diorite is interpreted as dominantly cumulate rocks derived from an alkaline mafic to intermediate magma with within plate–syn-collisional compositional characteristics (Sims et al., 1992). Uranium-lead zircon dating of the Marinette Quartz Diorite gives an age of 1,862±15 Ma, which overlaps the age of the Dunbar Gneiss that it intrudes (Sims et al., 1992). Neodymium isotope analyses of the Marinette Quartz Diorite show ɛNd(1,860) = +0.7 to +0.12 with a depleted mantle model age of ~2.1 Ga (Barovich et al., 1989; Schulz and Ayuso, 1998), suggesting derivation from an enriched source. 111 �Figure 5. Marinette Quartz Diorite south of Dunbar cut by granite pegmatite dikes. Stop 3. Marinette Quartz Diorite (Brown Spur road going east from County road O; 45.713° N., 88.011° W.) Although there appears to be no actual outcrop of Marinette Quartz Diorite at this stop, there are numerous angular boulders of black, medium- to coarse-grained diorite to quartz diorite typical of the eastern, less metamorphosed part of the intrusion. Stop 4. Newingham Tonalite (At the intersection of Highway 8 and 1 Mile Road, north side of road; 45.628° N., 88.096° W.) The Newingham Tonalite forms a large body, ~75 km2 in area, that intrudes the volcanic rocks of the Quinnesec Fomation along the southeast margin of the Dunbar dome (Figure 2). The contact zone is at least 100 m wide, and consists of interlayered tonalite bodies and generally angular amphibolite (metabasalt) inclusions. Near the contact with the Marinette Quartz Diorite, the Newingham Tonalite has been largely converted to granodiorite or granite by post-crystallization addition of potassium feldspar forming a hybrid, megacrystic facies of the Newingham Tonalite (Sims et al., 1992). The Newingham Tonalite is mostly a uniform light gray, medium-grained, slightly porphyritic rock (Figure 6) that generally has a good secondary foliation except in the eastern portion north of Pembine. It is locally cut by dikes of slightly porphyritic tonalite and, occasionally, granite pegmatite. The Newingham Tonalite has the compositional characteristics of high Al2O3-type tonalite-trondhjemite suites including high Al2O3 (>15 wt.%) and Sr (>600 112 �ppm), low K and Rb, and steep REE patterns depleted in heavy REE (La/Yb = 60-90) (Sims et al., 1992). Uranium-lead zircon dating of the Newingham Tonalite gives an imprecise age of 1,861±40 Ma (Sims et al., 1992). Neodymium isotope analysis of one sample of the Newingham Tonalite gave a ɛNd(1,860 Ma) = +0.39 with a depleted mantle model age of 2.09 Ga (Schulz and Ayuso, 1998), suggesting derivation from an enriched source or contamination by older crustal rocks. Figure 6. Newingham Tonalite. Stop 5. Dunbar Gneiss (Intersection of Highway 8 with County road U, west side of road; 45.655° N., 88.199° W.) Exposed here is a low outcrop of mainly megacrystic granite gneiss that contains rafts of amphibolite and is intruded by granite pegmatite. The gneiss has a pervasive steeply dipping N. 50° W. foliation. The Dunbar Gneiss generally consists predominantly of gray biotite gneiss, which is layered at scales ranging from a few centimeters to several meters reflecting differences in the amount and kind of major minerals as well as differences in grain size (Figure 7). Granite pegmatite and aplite intrude the Dunbar Gneiss particularly in the western part of the dome and can compose more than 50 percent of the outcrop. As described by Sims et al. (1992, p. 7)….”The biotite gneisses are mylonitic rocks. They have a dominant xenomorphic granular (granoblastic) texture and a penetrative foliation expressed by oriented biotite and, less commonly, by elongate and flattened aggregates of quartz and plagioclase that are generally subparallel to compositional layering….The biotite gneisses have a moderate range in composition from layer to layer; their average and modal composition is granodiorite. Plagioclase (calcic oligoclase-andesine) and quartz are the principal minerals, potassium feldspar varies from 0 to about 30 percent, and biotite generally composes from 10 to 20 percent 113 �of the rocks. Myrmekite and myrmekitic plagioclase are abundant….Sphene (titanite) is the principal accessary mineral and comprises as much as 2 percent of the rock.” The Dunbar Gneiss is calc-alkaline in composition with intermediate SiO2 contents (~62 to 72 wt.%), moderately high Al2O3 (~14 to 17.5 wt.%), and high K2O (2.4 to 5.1 wt.%). It also is enriched in Ba, Nb, Rb, light REE, Ta, and Th, and has steep REE patterns with depleted heavy REE (Sims et al., 1992). A sample of Dunbar Gneiss just north of this outcrop gave a U-Pb zircon concordia upper intercept age of 1,862±5 Ma (Sims et al., 1992). Neodymium isotope analyses by Barovich et al. (1989), Schulz and Ayuso (1998), and Van Wyck and Johnson (1997) gave very similar results with ɛNd(1,860) = -2.1 to -3.4 and depleted mantle model ages of 2.20 to 2.41 Ga. The isotope data suggest the protolith of the Dunbar gneiss was derived from an enriched source. Figure 7. Dunbar Gneiss and granite pegmatite. Stop 6. Dunbar Gneiss (Spur Lake Road going west from County road U; 45.684° N., 88.232° W.) The exposures on the east side of the road consist of compositionally layered biotite gneiss and lesser amphibolite intruded by megacrystic biotite gneiss, granite pegmatite, and aplite. All rocks are deformed and have a vertical N. 50-55° W. foliation. Stop 7. Twelve Foot Falls Quartz Diorite and mylonite, ultramylonite, and pseudotachylyte along the Twelve Foot Falls shear zone The Twelve Foot Falls Quartz Diorite is an elongate, east-west trending body that intrudes and locally contains inclusions of metavolcanic rocks of the Quinnesec Formation south of the Dunbar dome (Figure 2). The quartz diorite is generally massive in the eastern part, but becomes foliated towards the west. At the type locality at Twelve Foot Falls on the north branch of the Pike River, the quartz diorite has been intensely sheared by the Twelve Foot Falls shear zone and is mainly a mylonitic gneiss. As described by Sims et al. (1992, p. 43)….”Outside the shear zone, the quartz diorite is medium to coarse grained and is characterized by subhedral plagioclase (sodic andesine) crystals as much as 1 cm long, smaller subhedral hornblende 114 �crystals, in part pseudomorphic after pyroxene, and anhedral crystals of blue quartz as much as 1 cm in diameter. Microcline locally occurs as a late interstitial mineral. The primary texture is hypidiomorphic granular, but finer grained secondary textures are superposed on it at many places. Characteristically, the rock is considerably retrograded: plagioclase is partly to largely altered to epidote and albite, and hornblende is partly altered to biotite, epidote, and chlorite. Other alteration minerals are sphene, opaque oxides, and calcite.” The quartz diorite has a calc-alkaline andesite (SiO2 = 57 wt.%) composition with low TiO2 (0.44 wt.%) and high field strength element contents (Sims et al., 1992). It is similar in composition to andesite volcanic rocks in the Quinnesec Formation and is interpreted to be a subvolcanic intrusion. A sample dated by Schulz and Schneider (2005) gave a U-Pb zircon concordia upper intercept age of 1,889±6 Ma age. This age shows that the volcanic rocks of the Quinnesec Formation are significantly older than the intrusive rocks of the Dunbar dome and places a minimum age on the Pembine ophiolite present within the Quinnesec Formation (Schulz and Schneider, 2005). An Nd isotope analysis of the quartz diorite gave a ɛNd(1,900) = +4.54 with a depleted mantle model age of 1.87 Ga (Schulz and Ayuso, 1998) indicating a depleted source and no crustal contamination. The strongly positive ɛNd value for the quartz diorite is similar to that determined for the mafic volcanic rocks of the Quinnesec Formation (Beck and Murthy, 1991). (Material below and Stops 7a and 7b contributed by Marcia Bjornerud, Lawrence University) The Twelve Foot Falls shear zone (Sims, 1990) can be traced for at least 20 km along strike in northwest Marinette County, Wisconsin, from Twelve Foot Falls County Park on the north branch of the Pike River to just south of Kidd Lake. The timing of displacement on the shear zone is only broadly constrained; the shear zone transects the Twelve Foot Falls Quartz Diorite (1,889±6 Ma) as well as the metavolcanic Quinnesec Formation, and it lies immediately south of the 1,862±5 Ma Dunbar Gneiss (Sims et al., 1992). The vertical to steeply northeast-dipping foliation and mylonite bands in the Twelve Foot Falls Quartz Diorite are broadly parallel to the foliation in the southern part of the Dunbar dome, but the Twelve Foot Falls shear zone does not cut through the dome itself. Sims (1990) suggested that the northern part of the Amberg Granite (U-Pb zircon age 1,752±8 Ma) also is transected by the shear zone; if so, the zone would have developed during or after the Yavapai orogenic cycle. The sense of slip also is poorly constrained; a weak down-dip lineation points to dip-slip motion, but it is not clear whether the slip sense was reverse – which would suggest activity synchronous and sympathetic with convergence on the Niagara Fault – or normal, which would indicate slip related to late-orogenic relaxation. Stop 7a: Twelve Foot Falls County Park (45.579° N., 88.137° W.) Note that Marinette County parks require an entrance fee; use self-service registration box in main parking area. The main falls are visible from the parking area across a small pool in the Pike River. Follow the narrow foot path north of the picnic area to reach the outcrop adjacent to the falls, where the Twelve Foot Falls Quartz Diorite is well-exposed. The rock there is strongly foliated and locally mylonitized, and both the foliation and mylonitic fabric are defined by bands of quartz and feldspar alternating with aligned hornblende crystals (partly regressed to chlorite), indicating that 115 �the overall schistosity and localized zones of high strain formed at peak metamorphic (amphibolite facies) conditions. Depending on water levels, there is another area of exposed rock about 220 m downstream from Twelve Foot Falls, just above Eight Foot Falls. There, dark, branching discordant veins 0.3-0.5 cm wide and 10-15 cm long cut across the foliation in the host rock (Figure 8). In thin section, the veins are found to contain a mesh of fine retrograded hornblende (?) crystals with high aspect ratio, arranged with no preferred orientation in a non-crystalline matrix that is dark in plane polarized light. These macro- and micro-scale characteristics suggest that the veins represent devitrified pseudotachylyte injection veins – frictional melt generated on a fault plane during seismic slip and injected as ‘hydro’-fractures into the surrounding rock (Nadziejka and Bjørnerud, 2014; Larson and Bjørnerud, 2017). Significantly, the pseudotachylyte material can be seen in both outcrop and thin section to have been cut by, and in places incorporated into, the mylonitic bands. This indicates that brittle seismic failure occurred at least once while the rocks were still at depths and temperatures where crystal plastic deformation was predominant. We have also found small amounts of pseudotachylyte at Eighteen Foot Falls, about 1 km upstream from Twelve Foot Falls, and at Dave’s Falls near Amberg. Thin sections of specimens from Eight Foot Falls also show mutually cross-cutting relationships between plastically deformed quartz veins and pseudotachylyte (Figure 9). This indicates that brittle tensile fracture and fluid flow occurred in alternation with seismic failure and ductile deformation. Some of the quartz veins contain significant amounts of pyrite. In addition, the foliation is in places transected by discontinuous cm-long veins in which hornblende and quartz occur as fibrous crystals perpendicular to the walls. The crystals have growth bands and fluid inclusion planes oriented parallel to the vein walls. These features suggest that the veins formed by the ‘crack-seal’ mechanism, in which cyclic fluid pressure variations cause hydrofracturing and incremental mineral growth. Crack-seal veins are most commonly found in the shallow upper crust; the fact that hornblende is one of the vein-filling minerals indicates that in this case, the process occurred at greater depth and higher temperatures. In combination, these observations provide an exceptional glimpse into the complex interplay of deformation mechanisms and fluid flow in the middle crust during an orogenic event. Large earthquake ruptures apparently penetrated downward into rocks that were otherwise at temperatures high enough for full crystal plasticity. Such mutually cross-cutting relationships between mylonites and pseudotachylytes have been reported from only a small number of sites around the world (Sibson and Toy, 2006). Strain incompatibilities related to these ruptures may have caused dilatancy and large fluid pressure gradients that led to the formation of quartz-pyrite and quartz-hornblende veins. 116 �Figure 8. Outcrop photos of pseudotachylyte and mylonite in the Twelve Foot Falls Quartz Diorite at Eight Foot Falls. Brunton compass and pencil indicate scale. White arrows show places where pseudotachylyte has been offset along the mylonitic foliation. Although the apparent offset is left lateral, lack of three-dimensional exposure makes true slip vector difficult to determine. 117 �Figure 9. Photomicrograph (cross-polarized light) of finely recrystallized quartz diorite host rock cut by a quartz vein that was in turn intruded by pseudotachylyte (altered entirely to clinochlore). Scalloped edge of quartz grain in middle of image is consistent with melting. Vein quartz shows undulatory extinction, indicating that plastic deformation followed or alternated with brittle fracture. Stop 7b: Powerline exposure of Twelve Foot Falls Shear Zone (From the intersection of Twelve Foot Falls Road and Forest Road 513, drive 1.3 km (0.8 miles) west on 513 to an open area where a major powerline crosses the road; 45.584° N., 88.156° W.) This site provides a glimpse of the strain heterogeneity typical of the Twelve Foot Falls shear zone. The textural character of the Twelve Foot Falls Quartz Diorite ranges from igneous to ultramylonitic, in some cases over distances of centimeters. The foliation and mylonitic bands dip about 80° NE. In places, a weak down-dip mineral lineation is discernible. In thin section, feldspar porphyroclasts show extremely long tails – suggesting very high shear strains – but no consistent asymmetry that would allow the sense of shear to be determined unambiguously. Stop 8. Athelstane Quartz Monzonite and Amberg Granite (U.S. Highway 141 and Black Sam Road just north of Amberg; 45.517° N., 87.996° W.) This pavement outcrop consists of Athelstane Quartz Monzonite cut by dikes of Amberg Granite (Figure 10; Medaris et al., 1973). The Athelstane Quartz Monzonite intrudes felsic volcanic rocks (Beecher Formation) north of this stop and extends for considerable distance both to the west and south (Sims, 1990). It is a pink medium- to coarse-grained granite to granodiorite with 118 �allotriomorphic granular texture and 5 to 10 percent biotite and(or) hornblende (Sims et.al., 1993). Typically, the quartz monzonite has a clotty appearance due to the interstitial nature of the mafic minerals (Figure 10). Small metavolcanic inclusions are present locally. The Athelstane Quartz Monzonite is mildly peraluminous and has high SiO2 (68–77 wt.%), intermediate Al2O3 (12–15 wt.%), K2O greater than Na2O, and enrichment in iron (FeOt/(FeOt + MgO) = ~0.9) (Sims et al., 1993). Samples plot in within-plate and syn-collisional fields on trace element tectonic discriminant diagrams (Sims et al., 1993). A U-Pb zircon age of 1,835±15 My was determined on a sample from a large quarry south of this stop (Sims, 1990). This age overlaps with that determined for the Spikehorn Creek Granite on the northeast side of the Dunbar dome (Sims et al., 1992). Barovich et al. (1989) determined a ɛNd(1,835) = +1.1 with a depleted mantle model age of 2.07 Ga on a sample of the quartz monzonite. The positive ɛNd value is similar to that determined for the Hoskin Lake Granite in the Dunbar dome (see Stop 1) and suggests derivation from a similarly light REE depleted source. The Amberg Granite, seen here as dikes cutting the Athelstane Quartz Monzonite, is gray, fineto medium-grained, with a hypidiomorphic granular texture and has biotite as the major ferromagnesian phase. It also occurs in at least three intrusive bodies within the Athelstane Quartz Monzonite (Sims, 1990). It has a U-Pb zircon age of 1,752±8 Ma (Sims, 1990); it is one of a number of small plutons of this age found across northern Wisconsin (Sims et.al., 1993). These ~1,750 Ma plutons are coeval with the anorogenic granite-rhyolite terrane in south-central Wisconsin (Anderson et al., 1980; Smith, 1983). A sample of the Amberg Granite has an ɛNd(1,750 Ma) = -0.91 and a depleted mantle model age of 2.17 Ga (Schulz and Ayuso, 1998), suggesting derivation from an enriched source. Figure 10. Athelstane Quartz Monzonite cut by dikes of Amberg Granite. 119 �Stop 9. Athelstane Quartz Monzonite and mafic dikes (Dave’s Falls County Park just south of Amberg; 45.497° N., 87.989° W.) Excellent exposures of the Athelstane Quartz Monzonite occur on both sides of the Pike River. The Athelstane is cut by mafic dikes (~3 to 20 m wide) that strike about north-south. The dikes, at least three of which are exposed in the park, weather recessively relative to the Athelstane (Figure 11). The dikes have an andesitic composition (SiO2 ~53 to 55 wt.%) with low MgO (~2.5 to 3.5 wt.%) and TiO2 (~1.8 wt.%) contents, enriched light REE chondrite normalized patterns, and negative Nb-Ta and Ti anomalies on a primitive mantle normalized trace element plot (Figure 12). Dikes with similar composition have been observed cutting outcrops of the Twelve Foot Falls Quartz Diorite and in drill core cutting the Back Forty massive sulfide deposit in Michigan (Schulz, unpublished data). The dikes are post-1,835 Ma and pre-Keweenawan in age, but their actual age is not known. Figure 11. Mafic dike (in valley, looking north) cutting the Athelstane Quartz Monzonite. 120 �Figure 12. Chondrite normalized rare earth element plot (A) and primitive mantle normalized trace element plot (B) for andesite dikes from northeastern Wisconsin (Schulz, unpublished data). Stop 10. Spikehorn Creek Granite (East side of U.S. Highway 8 at intersection with Morgan Park Road; 45.703° N., 87.981° W.) The road cut shows the Spikehorn Creek Granite with metabasalt inclusions (Figure 13). As described in Sims et al. (1992, p. 27-28)…”The Spikehorn Creek Granite is a gray to pinkishgray, fine- to medium-grained rock containing scattered anhedral potassium feldspar grains as much as 2 cm in diameter. It is generally massive, but locally (especially near the margins of the body), it bears a mylonitic foliation expressed mainly by recrystallized quartz leaves and oriented biotite….The granite has sharp intrusive contacts against the Quinnesec volcanic rocks and the Marinette Quartz Diorite, and small ramifying dikes intrude these rocks for distances as much as 400 m from the contact….The Spikehorn Creek Granite (of the Niagara lobe) and the Hoskin Lake Granite are compositionally similar except that the Hoskin Lake has slightly higher K2O content.” 121 �Figure 13. Spikehorn Creek Granite with angular metabasalt inclusions. References Anderson, J.L., Cullers, R.L., and Van Schmus, W.R., 1980, Anorogenic metaluminous and peraluminous granite plutonism in the Mid-Proterozoic of Wisconsin, USA, Contributions to Mineralogy and Petrology, v. 74, p. 311–328. Barovich, K.M., Patchett, J.R., Peterman, Z.E., and Sims, P.K., 1989, Origin of 1.9 - 1.7 Ga Penokean continental crust of the Lake Superior region, Geological Society of America Bulletin, v. 101, p. 333–338. Bayley, R.W., Dutten, C.E., and Lamey, C.A., 1966, Geology of the Menominee iron-bearing district Dickinson County, Michigan and Florence and Marinette Counties, Wisconsin, U.S. Geological Survey Professional Paper 513, 96 p. Beck, W., and Murthy, V. R., 1991, Evidence for continental crustal assimilation in the Hemlock Formation flood basalts of the Early Proterozoic Penokean Orogen, Lake Superior region, U.S. Geological Survey Bulletin 1904-I, 28 pp. Cain, J.A., 1964, Precambrian geology of the Pembine area, northeastern Wisconsin Papers of Michigan Academy of Science, Art, and Letters, v. 49, p. 81–103. Hildebrand, R.S., and Whalen, J.B., 2017, The tectonic setting and origin of Cretaceous batholiths within the North American Cordillera–the case for slab failure magmatism and its significance for crustal growth, Geological Society of America Special Paper 523, 113 p. 122 �Larson, M., and Bjørnerud, M., 2017, Seismic slip, mylonitization and fluid flow along the Penokean Twelve Foot Falls shear zone, Marinette Country, northeastern Wisconsin, Institute on Lake Superior Geology, Part 1: Program and Abstracts, v. 63, p. 56–57. Medaris, L.G., Van Schmus, W.R., Lahr, M.M., Myles, J.R., and Anderson, J.L., 1973, Guidebook to the Precambrian geology of northeastern and northcentral Wisconsin, 19th Institute on Lake Superior Geology, Wisconsin Geological and Natural History Survey, 86 p. Morey, G.B., Sims, P.K., Cannon, W.F., Mudrey, M.G., Jr., and Southwick, D.L., 1982, Geolgical map of the Lake Superior region, Minnesota, Wisconsin, and northern Michigan,Minnesota Geological Survey State Map Series S-13, scale 1:1,000,000. Nadziejka, B., and Bjørnerud, M., 2014, Petrographic characterization of the Penokean Twelve Foot Falls shear zone, Marinette County, WI, Evidence for coeval ductile and seismic behavior: Institute on Lake Superior Geology, Part 1: Program and Abstracts, v. 60, p. 91– 92. Schulz, K.J., and Ayuso, R.A., 1998, Crustal recycling in the evolution of the Penokean orogen: Isotopic evidence for Archean contributions to crustal growth in the Pembine-Wausau terrane, northern Wisconsin, Institute on Lake Superior Geology, Part 1: Program and Abstracts, v. 44, p. 113–114. Schulz, K.J., and Schneider, D.A., 2005, Age constraints on the Paleoproterozoic Pembine ophiolite-island arc complex and implications for the evolution of the Penokean orogeny, Geological Society of America Abstracts and Program, v. 37, no. 5, p. 4. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region, Precambrian Research, v. 157, p. 4–25. Sibson, R.H., and Toy, V.G., 2006, The habitat of fault‐generated pseudotachylyte: Presence vs. absence of friction‐melt, in Abercrombie, Rachel, McGarr, Art, Di Torro, Giulio, and Kanamori, Hiroo, eds., Earthquakes: Radiated Energy and the Physics of Faulting: American Geophysical Union Monograph 170, p. 153–166. Sims, P.K., 1990, Geologic map of Precambrian rocks of the Iron Mountain and Escanaba 1° x 2° quadrangles, northeastern Wisconsin and northwestern Michigan, U.S. Geological Survey Miscellaneous Investigations Series, Map I–2056, scale 1:250,000. Sims, P.K., and Peterman, Z.E., 1980, Geology and Rb-Sr age of lower Proterozoic granitic rocks, northern Wisconsin, in Morey, G.B., and Hansen, G.N., eds., Selected studies of Archean gneisses and lower Proterozoic rocks, Southern Canadian Shield: Geological Society of America Special Paper 182, p. 139–146. Sims, P.K., and Schulz, K.J., 1993, Geologic map of Precambrian rocks of parts of Iron Mountain and Escanaba 30’ x 60’ quadrangles, northeastern Wisconsin and adjacent Michigan, U.S. Geological Survey Miscellaneous Investigations Series, Map I–2356, scale 1:100,000. Sims, P.K., Schulz, K.J., and Peterman, Z.E., 1984, Guide to the geology of the Early Proterozoic rocks in northeastern Wisconsin, Institute on Lake Superior Geology, Field Trip 1, 93 p. Sims, P.K., Peterman, Z.E., and Schulz, K.J., 1985, The Dunbar Gneiss-granitoid domeimplications for Early Proterozoic tectonic evolution of northern Wisconsin, Geological Society of America Bulletin, v. 96, p. 1101–1112. 123 �Sims, P.K., Van Schmus, W.R., Schulz, K.J., and Peterman, Z.E., 1989, Tectono-stratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean Orogen, Canadian Journal of Earth Sciences, v. 26, p. 2145–2158. Sims, P.K., Schulz, K.J., and Peterman, Z.E., 1992, Geology and geochemistry of Early Proterozoic rocks in the Dunbar area, northeastern Wisconsin,: U.S. Geological Survey Professional Paper 1517, 65 p. Sims, P.K., Schulz, K.J., DeWitt, E., and Brasaemle, B., 1993, Petrography and geochemistry of Early Proterozoic granitoid rocks in Wisconsin magmatic terranes of the Penokean orogen, northern Wisconsin—a reconnaissance study, U.S. Geological Survey Bulletin 1904-J, 31 p. Smith, E.I., 1983, Geochemistry and evolution of the early Proterozoic, post-Penokean rhyolites, granites, and related rocks of south-central Wisconsin, U.S.A., in Medaris, L.G., Jr., ed., Early Proterozoic geology of the Great Lakes region: Geological Society of America Memoir 160, p. 113–128. Van Wyck, Nicholas, and Johnson, C.M., 1997, Common lead, Sm-Nd, and U-Pb constraints on petrogenesis, crustal architecture, and tectonic setting of the Penokean orogeny (Paleoproterozoic) in Wisconsin, Geological Society of America Bulletin, v. 109, p. 799– 808. 124 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2018 Description An account of the resource Institute on Lake Superior Geology. Iron Mountain, Michigan. May 15-18, 2018. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2018 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/a3135b02ae42629dc9f91152dd8734c3.pdf ee8b7508494e14facf4b73ce8c98ac45 PDF Text Text 65th Annual Meeting Terrace Bay, Ontario - May 8-9, 2019 Institute on Lake Superior Geology Part 1 – Program and Abstracts �Thank you to our sponsors! Individual contributors to student travel scholarship: Al MacTavish, Mary Kay Arthur, L. Gordon Medaris, Jr., Nick Swanson-Hysell �65th Annual Meeting Institute on Lake Superior Geology May 8-9, 2019 Terrace Bay, Ontario HOSTED BY: Mark Smyk and Pete Hollings Co-Chairs Ontario Geological Survey and Lakehead University Proceedings - Volume 65 Part 1 – Program and Abstracts Compiled and edited by Mark Puumala Cover Photos: Left - Little Pic River Breccia zone, Coldwell Complex, Middle - Toe of pahoehoe flow, Slate Islands. Right - Glacially polished syenite, Coldwell Complex. ��65th Institute on Lake Superior Geology Volume 65 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trip 1: The Slate Islands Trip 2: Midcontinent Rift-Related Carbonatites and Diatremes Trip 3: Geology of the Western Schreiber-Hemlo Greenstone Belt Trip 4: Geology of the Nipigon Area Trip 5: A stratigraphic transect across the Northern flank of the Midcontinent Rift near Rossport Trip 6: Geology of the Coldwell alkaline complex Trip 7: Building and ornamental stone sites of the Marathon Area, Ontario Trip 8: Geology of the past-producing Winston Lake Cu-Zn Mine Reference to material in Part 1 should follow the example below: Bedrosian, P., 2019. Multi-scale AEM and MT mapping of the Precambrian in Upper Michigan, Northern Wisconsin, and Eastern Minnesota. In; Puumala, M., (Ed.), Institute on Lake Superior Geology Proceedings, 51st Annual Meeting, Nipigon, Ontario, Part 1 - Abstracts and Proceedings. v.65, part 1, 5-6. Published by the 65th Institute on Lake Superior Geology and distributed by the ILSG Secretary: Pete Hollings - ILSG Secretary Department of Geology Lakehead University 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Email: peter.hollings@lakeheadu.ca ��Part 1: Program and Abstracts Table of Contents Institutes on Lake Superior Geology, 1955-2019 Sam Goldich and the Goldich Medal iii v Goldich Medal Guidelines vii Goldich Medalists and Goldich Medal Committee ix Citation for Goldich Medal Award to Mark Severson x Honoring the Pioneers of Lake Superior Geology xii Memoriam to Gene L. LaBerge xiii Eisenbrey Student Travel Awards xv Joe Mancuso Student Research Awards xvi Doug Duskin Student Paper Awards and Award Committee xvii Board of Directors and Session Chairs xviii Field Trip Leaders and Guidebook Authors xix Report of the 64th Annual Meeting xx Technical Program xxiii Poster Presentations xxx Abstracts 1-103 ii �Institutes on Lake Superior Geology, 1955-2019 # Date Place Chairs 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 1955 1956 1957 1958 1959 1960 1961 1962 1963 1964 1965 1966 1967 1968 1969 1970 1971 1972 1973 1974 1975 1976 1977 Minneapolis, Minnesota Houghton, Michigan East Lansing, Michigan Duluth, Minnesota Minneapolis, Minnesota Madison, Wisconsin Port Arthur, Ontario Houghton, Michigan Duluth, Minnesota Ishpeming, Michigan St. Paul, Minnesota Sault Ste. Marie, Michigan East Lansing, Michigan Superior, Wisconsin Oshkosh, Wisconsin Thunder Bay, Ontario Duluth, Minnesota Houghton, Michigan Madison, Wisconsin Sault Ste. Marie, Ontario Marquette, Michigan St. Paul, Minnesota Thunder Bay, Ontario C.E. Dutton A.K. Snelgrove B.T. Sandefur R.W. Marsden G.M. Schwartz & C. Craddock E.N. Cameron E.G. Pye A.K. Snelgrove H. Lepp A.T. Broderick P.K. Sims & R.K. Hogberg R.W. White W.J. Hinze A.B. Dickas G.L. LaBerge M.W. Bartley & E. Mercy D.M. Davidson J. Kalliokoski M.E. Ostrom P.E. Giblin J.D. Hughes M. Walton M.M. Kehlenbeck iii �# 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 Date 1978 1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 Place Milwaukee, Wisconsin Duluth, Minnesota Eau Claire, Wisconsin East Lansing, Michigan International Falls, Minnesota Houghton, Michigan Wausau, Wisconsin Kenora, Ontario Wisconsin Rapids, Wisconsin Wawa, Ontario Marquette, Michigan Duluth, Minnesota Thunder Bay, Ontario Eau Claire, Wisconsin Hurley, Wisconsin Eveleth, Minnesota Houghton, Michigan Marathon, Ontario Cable, Wisconsin Sudbury, Ontario Minneapolis, Minnesota Marquette, Michigan Thunder Bay, Ontario Madison, Wisconsin Kenora, Ontario Iron Mountain, Michigan Duluth, Minnesota Nipigon, Ontario Sault Ste. Marie, Ontario Lutsen, Minnesota Marquette, Michigan Ely, Minnesota 56 2010 International Falls, Minnesota 57 58 59 60 61 62 2011 2012 2013 2014 2015 2016 Ashland, Wisconsin Thunder Bay, Ontario Houghton, Michigan Hibbing, Minnesota Dryden, Ontario Duluth, Minnesota 63 2017 Wawa, Ontario 64 2018 Iron Mountain, Michigan 65 2019 Terrace Bay, Ontario iv Chairs G. Mursky D.M. Davidson P.E. Myers W.C. Cambray D.L. Southwick T.J. Bornhorst G.L. LaBerge C.E. Blackburn J.K. Greenberg E.D. Frey & R.P. Sage J. S. Klasner J.C. Green M.M. Kehlenbeck P.E. Myers A.B. Dickas D.L. Southwick T.J. Bornhorst M.C. Smyk L.G. Woodruff R.P. Sage & W. Meyer J.D. Miller & M.A. Jirsa T.J. Bornhorst & R.S. Regis S.A. Kissin & P. Fralick M.G. Mudrey & Jr., B.A. Brown P. Hinz & R.C. Beard L. Woodruff & W.F. Cannon S. Hauck & M. Severson M. Smyk & P. Hollings A. Wilson & R. Sage L. Woodruff & J. Miller T.J. Bornhorst & J. Klasner J. Miller, G. Hudak, & D. Peterson M. Jirsa, P. Hollings, & T. Boerboom, P. Hinz & M.Smyk T. Fitz P. Hollings T.J. Bornhorst & A. Blaske J. Miller & M. Jirsa R. Cundari & P. Hinz J. Miller, C. Schardt, & D. Peterson A. Pace, A. Wilson, & T.J. Bornhorst L. Woodruff, W. Cannon, & E.K. Stewart P. Hollings & M.C. Smyk �Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970’s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 v �INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL vi �Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. vii �Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. viii �Goldich Medalists 1979 Samuel S. Goldich 1998 Zell Peterman 2016 Mark A. Jirsa 1980 not awarded 1999 Tsu-Ming Han 2017 Philip Fralick 1981 Carl E. Dutton, Jr 2000 John C. Green 2018 Val W. Chandler 1982 Ralph W. Marsden 2001 John S. Klasner 1983 Burton Boyum 2002 Ernest K. Lehmann 1984 Richard W. Ojakangas 2003 Klaus J. Schulz 1985 Paul K. Sims 2004 Paul Weiblen 1986 G.B. Morey 2005 Mark Smyk 1987 Henry H. Halls 2006 Michael G. Mudrey 1988 Walter S. White 2007 Joseph Mancuso 1989 Jorma Kalliokoski 2008 Theodore J. Bornhorst 1990 Kenneth C. Card 2009 L. Gordon Medaris, Jr 1991 William Hinze 2010 William D. Addison & 1992 William F. Cannon Gregory R. Brumpton 1993 Donald W. Davis 2011 Dean M. Rossell 1994 Cedric Iverson 2012 James D. Miller 1995 Gene La Berge 2013 Tom Waggoner 1996 David L. Southwick 2014 Laurel Woodruff 1997 Ronald P. Sage 2015 Rodney J. Ikola 2019 GOLDICH MEDAL RECIPIENT Mark Severson Goldich Medal Committee Serving through the meeting year shown in parentheses. Klaus Schultz (2016-2019) U. S. Geological Survey Dan England (2017-2020) Eveleth Fee Office Steve Kissin (2018-2021) Lakehead University ix �Citation for the Goldich Medal Recipient to Mark Severson ILSG Members, Goldich Medal recipients and guests, it is my honor to present the citation for this year’s recipient of the Goldich Medal, Mark J. Severson. Mark J. Severson has made significant contributions to understanding a vast number or topics associated with the geology of the Lake Superior region during his 30+ year career. Few can say they have contributed to the ILSG as a student, as an industry geologist, as an academic, as a thesis advisor, and as a teacher. In fact, Mark’s contributions to understanding Lake Superior Geology fill at least 15 pages of a Google Scholar search! He exemplifies the essence of an “Institute on Lake Superior Geology” geologist, possessing both exceptional field skills and extraordinary lab skills which have enabled him to conduct comprehensive, high quality scientific research. Mark also possesses the rare skills that allow him to communicate complicated geological features, models, and stories to professionals, students and the public in a way that teaches them (and even more importantly, gets them excited about) the amazing geology of the Lake Superior region and the countless geological wonders in their backyards. Mark’s research in the Lake Superior region began in the mid-1970s after obtaining his Bachelor of Sciences degree in Geology at Western Illinois University. In 1978, he was awarded his Master’s Degree in Geology at the University of Minnesota Duluth, studying the “Petrology and Sedimentation of Early Precambrian Graywackes in the Eastern Vermilion District, Northeastern Minnesota” under the advisement of (at that time) future Goldich Medal awardee Dr. Richard Ojakangas. After stints as an exploration geologist searching for base metals, gold and uranium with US Steel and Santa Fe Pacific Mining, Mark started a distinguished 25-year long career with the Economic Geology Group at the Natural Resources Research Institute (NRRI) at the University of Minnesota Duluth (UMD). While at the NRRI, Mark established himself as one of the leading economic geologists in the Lake Superior region, producing nearly 40 NRRI technical reports, eight geologic maps, and numerous peer-reviewed journal and public poster presentations. The geologic topics covered in this work are diverse, and include: • • • Performing a wide variety of research associated with the igneous stratigraphy, coppernickel-platinum group element and titanium mineralization in the Duluth Complex (which included logging of over 1 million feet of Duluth Complex drill core and the production of 10 NRRI Technical Reports, 1 NRRI geologic map, and 11 peer reviewed journal publications): Completing substantial research evaluating sedimentary environments, mineralization, and the stratigraphy of the Biwabik Iron Formation, culminating in NRRI Technical Report NRRI/TR-2009/09, where he established the “Rosetta Stone” for interpreting the stratigraphy of the Biwabik Iron Formation; Writing numerous technical reports describing SEDEX-type mineralization in Carleton County and the Cuyuna District of eastern and east-central Minnesota, respectively; x �• • • • • • Producing an NRRI technical report describing the history of gold exploration in Minnesota; Completing an NRRI technical report explaining metallic exploration, mining, and processing permits in Minnesota; Co-authoring a significant NRRI technical report which describes rare earth element (REE) mineral potential across Minnesota; Developing technical reports describing clay deposits in the Minnesota River Valley; Producing the most detailed heat flow maps available for the State of Minnesota; and Co-authoring a federally-funded report describing possibilities for the development of pumped-hydro energy storage systems in legacy iron-mining landscapes in northeastern Minnesota; During his time at the NRRI, Mark contributed to the education of undergraduate and graduate students, as well as teachers through his efforts as an Adjunct Professor in the Department of Geology at the University of Minnesota Duluth, as an instructor for the Precambrian Research Center geology field camp, and via the Minnesota Minerals Education Workshop. Throughout his career, Mark collaborated on numerous projects with the Minnesota Geological Survey (MGS). This included co-authoring three Open File Reports (maps and reports) about Duluth Complex mineralization, as well as a significant Report of Investigation which describes the geology and mineral potential of the Duluth Complex and related rocks. It is important to note that these MGS publications were co-authored with Goldich Medal awardees John Green, Jim Miller, Mark Jirsa, and Val Chandler. Since 2013, Mark has worked (and is now “semi-retired”) as a Senior Geologist for Teck American, where he continued to define Cu-Ni resources at the Mesaba Deposit in NE Minnesota. Despite his “semi-retired” status, Mark continues to make significant contributions to understanding Lake Superior geology in his role as Vice President for the Mesabi Range Geological Society. Mark’s contributions to the ILSG since 1989 include authoring or co-authoring 22 abstracts and seven field trip guidebooks, serving as a session chair, and serving on student paper committees over the course of at least 20 ILSG meetings since 1989. It is worth noting that Mark served as the co-chair with Steve Hauck for the 50th Annual ILSG meeting that took place in Duluth in 2004. As well, Mark has undoubtedly increased the knowledge of those attending the many ILSG field trips that he participated in over the past 29 years. All of us who have known and worked with Mark know of his passion for the geology. His significant contributions to understand and teach about the spectacular and diverse geology of the Lake Superior region, as well as his significant contributions to the Institute on Lake Superior Geology have all been accomplished with the highest level of professionalism and distinction. Please join me in congratulating Mark J. Severson as the 2019 recipient of the Goldich Medal from the Institute on Lake Superior Geology. Submitted by George J. Hudak Director, Minerals-Metallurgy-Mining Initiative Natural Resources Research Institute, UMD xi �Honoring the Pioneers of Lake Superior Geology (Adopted by the Board of Directors, 2016) Preamble At the suggestion of Gene LaBerge, the 2016 executive board agreed to implement a program to recognize historic pioneers in the understanding of geology in the Lake Superior region. Beginning with the 2017 annual meeting, nominations will be accepted from the membership for geologists whose work was conducted primarily before inception of the institute in 1955. Biographical sketches of those pioneers will be presented at future annual meetings so that all might appreciate the value of their contributions. Selection of nominees will be decided in part by the organizing committee of each year's annual meeting, in consultation with the Board, to ensure equitable geographic representation in the selection process. Award Guidelines 1) Nominations from the membership will be submitted via the Institute web site and forwarded to the Chair of the next Annual Meeting. The nominations will be no more than half a page in length and will summarise the contribution of the nominee. 2) The Organising Committee will select one or two individuals to be highlighted at the next Annual meeting and submit those names to the Board for approval. 3) The nominator will be requested to prepare a brief presentation to be given during the next annual meeting with a summary to be included in the Proceedings volume. 4) Unsuccessful nominations will be kept by the Secretary for two years and forwarded to the next meeting Chair; these nominations may be resubmitted at a later date. The Board will review this award every five years. Pioneers of Lake Superior Geology 2017 Douglass Houghton (1809-1845) 2018-19 not presented xii �In Memoriam Gene L. LaBerge This winter the Institute on Lake Superior Geology, its members, and countless others lost a dedicated geologist, and outstanding teacher, mentor, colleague, and friend-Gene LaBerge. Gene’s contributions to the geology of the Lake Superior region and the study of iron-formation have been many and impactful. He will be greatly missed, not only for his scientific contributions, but also for his good humor, wise council, and generous nature. Gene was a product of the Northwoods, born and raised in Ladysmith, Wisconsin, the eventual home of the Flambeau copper-gold mine in the mid-1990s. After serving a stint in the U.S, Marine Corp during the Korean War, he went on to study geology, obtaining his B.S., M.S., and Ph.D. degrees all at the University of Wisconsin-Madison. While a graduate student in 1957, he was hired by Ralph Marsden, head of exploration for U.S. Steel, to explore for iron-formation in northern Michigan. This, along with field trips around the region led by Ralph and Stan Tyler of UW-Madison, sparked his interest in iron-formation and the geology of the Lake Superior region. While he was working to finish his Ph.D. research, his advisor, Stan Tyler, suddenly died. Fortunately, Ralph Marsden was able to step in, allowing Gene to finish and receive his Ph.D. His dissertation was on the origin of magnetite in iron-formation. After graduate school, he continued his study of iron-formation accepting a post-doctoral fellowship in Adelaide, Australia where his new bride, Sally, had received a Fulbright Scholarship. During this post-doc, he also spent several months in South Africa. After a year in Australia, Gene accepted a second postdoctoral fellowship, this time with the Geological Survey of Canada in Ottawa. Some of the samples of iron-formation he collected for his studies he subsequently used to build a beautiful fireplace in his house in Omro, Wisconsin. In 1965, Gene joined the faculty at UW-Oshkosh as the third member of the then expanding Geology Department. It would remain his home for all his career. Early on, he and Sally would spend weekends driving the back roads of the Lake Superior region looking for outcrops and planning field trips. Gene’s passion for geology was clearly expressed through the many (>100) overnight field trips he led during his 33-year teaching career. While still a graduate student, he had helped conduct a pebble survey in northern Wisconsin for U.S. Steel, identifying thousands and thousands of pebbles in gravel pits across the region. This experience led him to devise one of his classic (or infamous) student exams-the pebble test- where students had to identify the rock type and mineralogy of small pebbles using only a hand lens. Gene often said that “a rock or mineral was much easier to identify if you had seen it before”. To reinforce that, students in his mineralogy and lithology classes learned to identify not only the most common minerals and rocks but also many less common ones, particularly those important in exploration for certain types of mineral deposits. His classes were always rigorous, comprehensive, and taught with infectious enthusiasm. Gene retired from teaching in 1998. Over his teaching career, Gene xiii �received all the teaching and research awards offered by UW-Oshkosh, the only faculty member to have done so. In the late 1960s, on the advice of Carl Dutton, Gene began mapping the geology around Wausau in Marathon County, Wisconsin for the Wisconsin Geological and Natural History Survey (WGNHS). The project eventually grew to include mapping all of Marathon County in collaboration with Paul Myers of UW-Eau Claire. In 1983, Gene began working part-time for the U.S. Geological Survey (USGS), an association he would maintain both formally and informally for the rest of his career. Much of the work he did for the USGS was done collaboratively with his good friend and colleague John Klasner of Western Illinois University. During his work for the WGNHS and USGS, Gene probably walked over more of Wisconsin and northern Michigan than anyone ever has. Along with authoring many technical journal articles, book chapters, and WGNHS and USGS publications, he used his in-depth knowledge of the geology of the Lake Superior region to write a book for non-specialists, Geology of the Lake Superior Region, first published in 1994. Gene was an active and long-time member of the Institute on Lake Superior Geology (ILSG), giving his first presentation at the1958 meeting in Duluth, only the fourth meeting of the Institute. He went on to give presentations at many more ILSG meetings, served as Chair for two meetings (1969 and 1984), and led field trips for several meetings. Gene received the Goldich Award in 1995 for his significant contributions to the geology of the Lake Superior region and the ILSG. Gene also had other interests and pursuits. He was an avid mineral collector, building a worldclass collection whose centerpiece was gem-quality tourmalines from locations around the world. The collection was effectively displayed in his house in cases he designed and build himself. The collection was eventually sold (a hard decision resulting from having to move and downsize), but not before he had his three daughters select their favorite specimens. In mid-life, when some go out and buy a new sports car or take up sky diving, Gene took up playing the guitar, a talent he found useful for all those nights sitting around a campfire on field trips. In 1999, he published a book with the help of his daughter Michelle, Travels with Sophie, chronicling the experiences of his mother, Louise, while she served as the first supervising teacher of Rusk County, Wisconsin in 1917-1918. Sophie was the name of the Model T Ford his mother used to travel between more than 100 one-room schools in the county at the time. Gene also recently completed a book with George Robinson, curator of the Seaman Mineral Museum at Michigan Technological University, on the minerals in iron ores, Minerals in the Iron Ores of the Lake Superior Region. Gene led an active and productive life that touched and influenced many people through his research, teaching, writing, and public out-reach. He is survived by his wife, Sally, and daughters Michelle, Rene, and Laura and their families. Gene’s passing leaves a hole in the fabric of the lives of all who knew him and called him a friend. Klaus J. Schulz xiv �Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xv �Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive one-half of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. In fall 2018, the ILSG Board of Directors selected two students to be granted research funding of $750.00 each from the Joe Mancuso Student Research Fund. The awardees were: Jacqueline L. Drazan University of Minnesota-Duluth, MSc, Department of Earth and Environmental Sciences, draza004@d.umn.edu TOPIC: Morphological and Geochemical Comparison between Archean Marine Peperites (Fivemile Lake, MN) and Pleistocene Freshwater Peperites (Sveifluhals, Iceland) Thomas Bodden Michigan Technological University, MSc, Department of Geological and Mining Engineering and Sciences, tjbodden@mtu.edu TOPIC: Stable isotopic composition of calcite precipitated with native copper and other minerals of the Keweenaw Peninsula, Michigan xvi �Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and longtime friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. 2019 Student Paper Awards Committee Katarina Bjorkman – Bjorkman Prospecting George Hudak – Natural Resources Resarch Institute–UMD David Good – Western University xvii �Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Esther Stewart (2018-2021) – Wisconsin Geological & Natural History Survey Anthony Pace (2017-2020) – Ontario Geological Survey Christian Schardt (2016-2019) – University of Minnesota Duluth Pete Hollings - Secretary (2016-2019) – Lakehead University Mark Jirsa – Treasurer (2017-2020) – Minnesota Geological Survey Session Chairs Ben Drenth- United States Geological Survey Dan England – Eveleth Fee Office Mary Louse Hill - Lakehead University Amy Radakovich- Minnesota Geological Survey Nicholas Swanson-Hysell – University of California, Berkeley Laurel Woodruff – United States Geological Survey Michael Zieg – Slippery Rock University Shannon Zurevinski – Lakehead University xviii �Field Trip Leaders and Guidebook Authors Field trips have been the mainstay of the ILSG since its inception 65 years ago. We want to give a special thanks to the field trip leaders and guidebook authors who volunteered their time and talent in carrying that tradition forward. 1) The Slate Islands Pete Hollings – Lakehead University Mark Smyk – Ontario Geological Survey Bill Addison and Philip Fralick – Lakehead University 2) Midcontinent Rift-related carbonatites and diatremes Shannon Zurevinski – Lakehead University Dorothy Campbell and Mark Puumala – Ontario Geological Survey 3) Geology of the western Schreiber-Hemlo greenstone belt Seamus Magnus – Ontario Geological Survey 4) Geology of the Nipigon area Philip Fralick – Lakehead University Robert Cundari – Ontario Geological Survey 5) A stratigraphic transect across the northern flank of the Midcontinent Rift near Rossport Pete Hollings and Philip Fralick – Lakehead University 6) Geology of the Coldwell alkaline complex Allan MacTavish – Panoramic PGMs (Canada) Limited Mark Smyk – Ontario Geological Survey David Good – Western University John McBride – Stillwater Canada Inc. 7) Building and ornamental stone sites of the Marathon area, Ontario Peter Hinz – Ministry of Energy, Northern Development and Mines 8) Geology of the past-producing Winston Lake Cu-Zn Mine Robert Lodge – University of Wisconsin-Eau Claire Mark Smyk and Mark Puumala – Ontario Geological Survey xix �REPORT OF THE 64th ANNUAL MEETING OF THE INSTITUTE ON LAKE SUPERIOR GEOLOGY IRON MOUNTAIN, MICHIGAN The U.S. Geological Survey with assistance from the Wisconsin Geological and Natural History Survey hosted the 64th Annual Institute on Lake Superior Geology on May 15 – 18, 2018 at the Pine Mountain Resort in Iron Mountain, Michigan. The meeting consisted of two days of technical sessions with pre- and post-technical session field trips. Laurel Woodruff (USGS), Bill Cannon (USGS), and Esther Stewart (WGNHS) were co-chairs for the 2018 meeting. Tom Mroz and Tom Waggoner helped with pre-meeting logistics. Darlene Comfort and Ted Bornhorst (A.E. Seaman Mineral Museum, Michigan Technological University) handled all pre-meeting registration and printing needs. Ted also supplied the poster boards and helped with many aspects of the meeting. Mary Kay Arthur and Dave Wilhelm (Geological Society of Minnesota) provided valuable logistical assistance on-site at Pine Mountain during the technical sessions. Connie Dicken (USGS) was the media czar for the technical sessions, keeping all presentations on track with fewer glitches than normal. Generous contributions to the ILSG general fund and in support of 2018 student travel scholarships came from Lundin Mining, the Geological Society of Minnesota, Ron Seavoy, Mary Kay Arthur, L. Gordon Medaris, Jr., and Steven Baumann. Total meeting registration was 189 (37 students), an excellent turn-out. Proceedings Volume 64 was published in two parts: Part 1 – Program and Abstracts, edited by Esther Stewart, contains 62 published abstracts for 34 oral and 28 poster presentations; Part 2 – Field Trip Guidebooks, compiled by Bill Cannon, contains descriptions of four field trips, two pre-meeting and two post-meeting. The 64th ILSG marked the second time in its long history that the annual meeting was held in Iron Mountain. The prior meeting was in 2003. Field trips visited two areas new to the ILSG and two trips that provided new stops in areas of prior trips. On Tuesday, May 15, Bill Cannon, Klaus Schulz, Robert Ayuso, and Tom Mroz led a field trip of 46 people to examine the stratigraphy, structure and economic geology of regional Precambrian rocks in the Felch District, Central Dickinson County, Michigan. Also, on Tuesday, Tom Waggoner was the leader of 37 people for a trip that looked at the Paleoproterozoic Hemlock Formation. Most stops of these two trips were to locations and geology new to ILSG attendees. On Friday, May 18 Tom Mroz and Bill Cannon shepherded a large crowd of 56 somewhat wandering individuals on a field trip that looked at the geology of the Menominee Iron Range. This trip had little overlap with a trip of a similar title that was given in 2003 as it included a visit to the Archean Carney Lake Gneiss, newly recognized as containing zircons with cores as old as 3.8 Ga, and an underground tour of the Iron Mountain Iron Mine. Another Friday trip led by Klaus Schulz, with a contribution from Marcia Bjørnerud, examined the granitoid rocks of the PembineWausau terrane in northern Wisconsin. In addition to examining granite, the 30 people on that trip had an additional task of keeping Klaus out of jail when he was caught looking at an outcrop along a railroad line. One hundred and fifteen participants attended the annual ILSG banquet on Wednesday night, cheerfully bringing chairs from the technical session room into the banquet room. After an excellent dessert, everyone moved their chairs back to the technical session room. The 2018 Homer award was given to Pete Hollings for his confidence in the appropriate vehicles one needs to travel around Iceland. Al McTavish, fresh off the success of last year’s Iceland trip, despite the Land xx �Rovers, gave a short promotional presentation for a proposed 2019 trip to another volcanic island, Hawaii. As always, a highlight of the post-banquet activities was presentation of the 2018 Goldich Medal. This year’s very deserving recipient was Val Chandler of the Minnesota Geological Survey (MGS). Val’s wife and three adult children were all able to attend the banquet and award ceremony. The Goldich was presented to Val by David Southwick, Director Emeritus of the MGS and his colleague for many years. Dave’s citation described Val’s many professional contributions to the geophysical mapping and interpretation of Minnesota’s mostly hidden geology. Val’s long history with the ILSG started with a field trip when he was fresh out of Purdue, which serendipitously led to his distinguished career with the MGS. This year’s banquet speaker was Nancy Langston, a professor in the Department of Social Sciences at Michigan Technological University. Dr. Langston (or Nancy, as we all called her) gave a presentation that drew on her recent book titled Sustaining Lake Superior: An extraordinary lake in a changing world. The presentation described past, present, and future environmental challenges to Lake Superior, such as logging, Reserve Mining taconite disposal, and climate change. We all were encouraged by Nancy’s final optimism that with responsible stewardship, the largest freshwater lake in the world will endure. In 2018, the student paper committee had its usual difficult job of selecting the best among 7 excellent oral presentations and 16 excellent poster presentations for the Doug Duskin Student Paper Awards. This year’s committee included Robert Cundari (Ontario Geological Survey), Esther Stewart (WGHNS), performing double duty along with her co-chair responsibilities, and Latisha Brengman (University of Minnesota – Duluth). In the end, there was a three-way tie for first place. Poster awards ($300 each) were awarded to Samuel Hone (Slippery Rock University) for his poster titled: Olivine crystal size distribution in the Black Sturgeon Sill, Nipigon, Ontario, and William Fitzpatrick (University of Wisconsin- Eau Claire) for his poster titled: Mineral chemistries of the Tower Mountain Intrusive Complex Au-deposit, Ontario. Kira Arnold (Lakehead University) was recognized for her oral presentation titled: Geology and geochemistry of the Terrace Bay Batholith, N. Ontario ($400). Eisenbrey Student Travel Grants were given to 19 students: Daniel Wilkes, Emily Gorner, Kira Arnold, Vittoria Smith, and Simon Dolega – all from Lakehead University; Schuyler Borages, Erica Craddock, Ryan Leonard, Walter Johnson-Geis, Lily Atkinson, and Juliana Olsen-Valdez, all from Lawrence University; Jacqueline Drazan, Margaret Upton, and Matthew Matko, from the University of Minnesota-Duluth; Victoria Stinson, University of Saskatoon; Dustin Liikane, University of Toronto, Katharine Rose and Kevin Rupp, both from Western Michigan University, and Joseph Rasmussen, University of Wisconsin-Platteville. The Institute’s Board of Directors met on May 16, 2018 and a brief overview of the meeting is provided below: 1. Accepted the Report of the Chair for the 63rd ILSG from Ted Bornhorst and minutes of the last Board meeting from ILSG secretary, Pete Hollings. 2. Accepted the 2017-2018 ILSG Financial Summary from ILSG treasurer, Mark Jirsa. 3. Approved one co-chair from the 64th annual meeting, Esther Stewart, as the on-going board member. xxi �4. Nominated Steve Kissin from Lakehead University to replace Shannon Zurevinski on the Goldich Committee. 5. Approved Terrace Bay, Ontario as the location for the 2019 ILSG annual meeting with cochairs Pete Hollings and Mark Smyk. The 64th ILSG meeting was a great success and we wish to thank all the people who contributed to that success. The staff of Pine Mountain was professional and responsive to the needs of a large group – plenty of excellent donuts. The weather was perfect, not too hot, not too cold, not rainy, not buggy. The field trips this year had many participants, and thanks are due to field trip leaders, intrepid bus drivers, those who drove support vehicles on field trips and handled each trip’s logistics, as well as everyone else who stepped up when needed. As always, everyone who attended the 64th ILSG was willing to help as necessary and to adapt to any situation that developed. The meeting this year was well attended, and we are heartened by the excellent student participation and attendance, a trend we hope continues. Your co-chairs are very pleased with the final outcomes of the 64th ILSG. Organizing a meeting and compiling the two Proceeding volumes requires a significant time commitment from the cochairs and others, and we thank our respective organizations for their recognition of the importance of the ILSG. We also thank the ILSG community and members who make the experiences of the co-chairs almost fun, especially once the meeting is over, and we encourage others to take on the task. Laurel Woodruff, Bill Cannon, and Esther Stewart Co-Chairs, 64th Institute on Lake Superior Geology xxii �TECHNICAL PROGRAM TUESDAY MAY 7, 2019 All field trips begin and end at the Terrace Bay Cultural Centre 8:00 am - 5:00 pm PRE-MEETING FIELD TRIPS 1) The Slate Islands Pete Hollings – Lakehead University 2) Midcontinent Rift-related carbonatites and diatremes Shannon Zurevinski – Lakehead University 3) Geology of the western Schreiber-Hemlo greenstone belt Seamus Magnus – Ontario Geological Survey 4) Geology of the Nipigon area Philip Fralick – Lakehead University 4:00 pm - 10:00 pm Registration (Terrace Bay Cultural Centre) 7:00 pm - 10:00 pm Welcoming Reception and Poster Session (Terrace Bay Cultural Centre) xxiii �WEDNESDAY MAY 8, 2019 8:00 am – 11:30 am Registration (Terrace Bay Cultural Centre) 8:30 OPENING REMARKS (Terrace Bay Cultural Centre) Pete Hollings and Mark Smyk, Co-Chairs, 2019 ILSG TECHNICAL SESSION I Session Chairs: Shannon Zurevinski – Lakehead University Michael Zieg – Slippery Rock University * denotes a student eligible for Best Student Paper Award. To be eligible students must have graduated no more than one month before the ILSG meeting, be first author, and present the paper at the meeting. + denotes author that will present abstract, if different than the first author. 8:40 Brigitte Gelinas, +Pete Hollings and Richard Friedman Geology and geochemistry of the Laird Lake property and associated gold mineralization, Red Lake greenstone belt, Ontario 9:00 *Munira Afroz, Philip Fralick, Brian Killingsworth, Martin Homann, Pierre Sansjofre and Stefan Lalonde Sulfur, Carbon, and Oxygen Isotope Geochemistry of ~2.93 Ga Mesoarchean Chemical Sedimentary rocks in the Red Lake Area, Ontario 9:20 *Brittany Ramsay, Philip Fralick, Paul Bielski, Martin Homann, Pierre Sansjofre and Stefan Lalonde Mesoarchean chemical sedimentary rocks of northwestern Ontario: Implications for hydrosphere composition in deep time 9:40 Brad Gottschalk and Caroline Rose Recent efforts to curate and provide access to the historical documents of the E.K. Lehmann and Associates Exploration Company 10:00 COFFEE BREAK 10:20 William F. Cannon, Klaus J. Schulz and Benjamin J. Drenth The Dickinson Group in the Central Upper Peninsula of Michigan: Part 1- Age and tectonic setting based on new geophysical, geochronological, and geochemical data xxiv �10:40 Benjamin J. Drenth, William F. Cannon and Klaus J. Schulz The Dickinson Group in the central Upper Peninsula of Michigan: Part 2Geophysical expression and a preliminary interpretation of its eastward extent under Paleozoic cover 11:00 Ryan Clark, David Peate, Alison Kusick, Kenny Horkley and Raymond Anderson Reexamining the Osborne core for new insights into the age and petrology of the Northeast Iowa Intrusive Complex (NEIIC) 11:20 Wouter Bleeker, Michael Hamilton, Sandra Kamo, Dustin Liikane, Jennifer Smith, Pete Hollings, Robert Cundari, Michael Easton and Don Davis High-resolution dating of the magmatic plumbing system of the Midcontinent Rift System—Insights into rift evolution and mineralization processes 11:40 End of Technical Session I 11:40 LUNCH BREAK – BUFFET PROVIDED ILSG BOARD OF DIRECTORS MEETING TECHNICAL SESSION II Session Chairs: Dan England – Eveleth Fee Office Laurel Woodruff – United States Geological Survey 1:00 Mark Puumala Using graphitic sedimentary rock geochemistry as an indicator of gold potential in the Shebandowan greenstone belt, northwestern Ontario 1:20 *Chanelle Boucher and Pete Hollings Geology and geochemistry of ultramafic rocks in the Lake of the Woods area 1:40 *Kira Arnold, Pete Hollings, Seamus Magnus, Shannon Zurevinski and Robert Creaser Geology and geochemistry of the Terrace Bay Batholith, N. Ontario 2:00 David Holder, Francois Robert and John Hay Geological characteristics and structural controls of Au mineralisation at the enigmatic Hemlo deposit 2:20 COFFEE BREAK xxv �2:40 Paul A. Bedrosian Multi-scale AEM and MT mapping of the Precambrian in Upper Michigan, northern Wisconsin, and eastern Minnesota 3:00 John McBride, David Good, D. Hollis and N. Arndt Pilot study: Using ambient noise passive seismic surveys for Ni-Cu-PGE mineral exploration at the Marathon PGM-Cu deposit, Marathon, Ontario 3:20 Dave Good, Pete Hollings and Andrew Jedemann Recognizing MCR magmas generated by partial melting in the SCLM: Lessons from mafic magmas in the Coldwell Complex 3:40 Ross Sherlock and Kate Rubingh Geologic architecture and precious metal mineralization in the southern Abitibi; new insights from the Larder Lake area 4:00 POSTER VIEWING - AUTHORS WILL BE PRESENT AT THEIR POSTERS 5:00 END OF TECHNICAL SESSION II 6:00 RECEPTION AND CASH BAR (Terrace Bay Cultural Centre) 7:00 ANNUAL BANQUET (Terrace Bay Cultural Centre) • Announcement of 66th Annual Meeting Location • 2019 Goldich Award Presentation to Mark Severson xxvi �THURSDAY MAY 9, 2019 8:30 INTRODUCTORY REMARKS AND UPDATES (Terrace Bay Cultural Centre) Pete Hollings and Mark Smyk, Co-Chairs, 2019 ILSG TECHNICAL SESSION III Session Chairs: Mary Louise Hill – Lakehead University Nicholas Swanson-Hysell – University of California-Berkeley 8:40 Wouter Bleeker, +Sandra Kamo, Michael Hamilton and K. Chamberlain New age data and insights into the ca. 1887-1870 Ma Circum-Superior Belt, with startling implications for the Lake Superior area geology 9:00 Robert Michael Easton What do detrital zircon studies of the Huronian Supergroup tell us? an analysis of all published data 9:20 *Sophie Kurucz, Philip Fralick, Stefan Lalonde and Martin Homann Paleoproterozoic snowball earth? Sedimentology and geochemistry of a Huronian glacial cycle 9:40 L.G. Medaris Jr., D.H. Malone, G.C. Hill, B.S. Singer, B.R. Jicha, A. Van Lankvelt, M.L. Williams and P.W. Reiners The Wolf River Orogeny: Geon 14 magmatism, sedimentation, and deformation in the southern Lake Superior region 10:00 COFFEE BREAK 10:20 Jim Miller The importance of “tablesetting” intrusions in creating economic Ni-Cu-PGE deposits in the Midcontinent Rift 10:40 Robert Nowak, Espree Essig and Robert Mahin Geochemical vectoring towards a serpentinized peridotite chonolith, Eagle East NiCu-Co-PGE deposit, Upper Peninsula, Michigan 11:00 Jennifer Smith, Wouter Bleeker, Mike Hamilton, and Duane Petts An investigation into the distribution of chalcophile elements and timing of mineralization within the Crystal Lake intrusion: A U-Pb geochronology and LAICP-MS study 11:20 Jack Gibbons, Tamara Diedrich and Thomas Quigley Petrography of several cobalt-enriched samples from the Atikokan River Intrusions, Atikokan, Ontario xxvii �11:40 End of Technical Session III 11:40 LUNCH BREAK – BUFFET PROVIDED TECHNICAL SESSION IV Session Chairs: Amy Radakovich – Minnesota Geological Survey Ben Drenth – United States Geological Survey 1:00 Thomas W. Buchholz, Alexander U. Falster and Wm. B. Simmons Updated mineralogy of a roadside pegmatite in the Stettin Complex, Wausau Syenite Complex, Marathon County, Wisconsin 1:20 *Paul Bielski and Philip Fralick LA-ICP-MS micro-sampling of iron formation: what it can tell us 1:40 Tamara Diedrich and Stephen Day Neutralization of proton acidity with sequestration of atmospheric CO2 during experimental weathering of intrusive rocks from the Midcontinent Rift System 2:00 Carson G. Prichard, +James J. Student, Jory L. Jonas, Nicole M. Watson and Kevin M. Pangle Catchment geology correlation with fish otolith microchemistry across disparate glacial till depths in the Lake Michigan basin 2:20 COFFEE BREAK 2:40 J.M. DeGraff, C.W. Tyrell, G.E. Hubbell and B.T. Carter Keweenaw Fault system along Bête Grise Bay, Michigan: geometry, kinematics, and tectonic significance 3:00 Esther K. Stewart, V.J.S. Grauch, Laurel G. Woodruff and Samuel Heller Seismic stratigraphy of the 1.1 Ga Midcontinent Rift beneath western Lake Superior shows evidence of differing subsidence histories for syn-magmatic sub-basins 3:20 V.J.S Grauch, Esther K. Stewart, Laurel G. Woodruff and Samuel Heller Evaluating Alternate Geophysical Models along the Isle Royale-Superior Shoal Aeromagnetic Anomaly, Central Lake Superior 3:40 Nicholas L. Swanson-Hysell Insights into Midcontinent Rift development resulting from a strengthened chronostratigraphic framework xxviii �4:20 BEST STUDENT PAPER AWARDS STUDENT TRAVEL AWARDS CLOSING REMARKS 4:40 END OF TECHNICAL SESSIONS FRIDAY MAY 10, 2019 8:00am – 5:00pm POST-MEETING FIELD TRIPS Field trips 5 to 8 begin and end at the Terrace Bay Cultural Centre 5) A stratigraphic transect across the northern flank of the Midcontinent Rift near Rossport Pete Hollings – Lakehead University 6) Geology of the Coldwell alkaline complex Allan MacTavish – Panoramic PGMs (Canada) Limited David Good – Western University 7) Building and ornamental stone sites of the Marathon Area, Ontario Peter Hinz – Ministry of Energy, Northern Development and Mines 8) Geology of the past-producing Winston Lake Cu-Zn Mine Mark Puumala and Mark Smyk – Ontario Geological Survey xxix �POSTER PRESENTATIONS *Thomas J. Bodden, Theodore J. Bornhorst, Florence Begue and Chad Deering Stable isotopic composition of calcite precipitated with native copper and other minerals of the Keweenaw Peninsula, Michigan Terrence J. Boerboom Recognition of probable distal ejecta from the 1850 Ma Sudbury meteorite impact event along the southern edge of the Animikie basin in Minnesota J.M. DeGraff and I.S. DeGraff Southwest Margin of the Midcontinent Rift System in Eastern Lake Superior: Review and Preliminary Interpretation *Jacqueline L. Drazan, George Hudak and Howard Mooers Morphology, mineralogy, texture, and genesis of peperite, Fivemile Lake, Vermilion District, Minnesota: Comparison with Pleistocene peperite, Iceland Benjamin J. Drenth, William F. Cannon and Klaus J. Schulz High-resolution aeromagnetic survey, central Upper Peninsula, Michigan Don Elsenheimer, Cari Deyell-Wurst and Lionel C. Fonteneau Hyperspectral Imaging of Bedrock Core from the Minnesota DNR Drill Core Library: A New Tool for Archival Preservation and Mineral Exploration V.J.S. Grauch and K.J. Schulz Superior Shoal revisited: Evidence for Keweenawan basalts with reversed- and normalpolarity remanent magnetization and early magma chemistry, central Lake Superior Linnea L. Johnson, David H. Malone and John P. Craddock Detrital Zircon Geochronology of Keweenaw Interflow Sediments within the North Shore Volcanic Group, Minnesota, U.S.A. Seamus Magnus Precambrian Geology of the Western Schreiber–Hemlo Greenstone Belt Amy Radakovich, Val Chandler and Mark Jirsa Wawa, undercover: Bedrock geologic and bedrock topographic mapping in north-central Minnesota Laura Ratcliffe Precambrian Geology of the Eastern Shebandowan Greenstone Belt - Insights into Stratigraphy and Structural History xxx �Christian Schardt and Mady David High-technology metals in ore-forming environments and their signature in volcanic-hosted sulfide mineralization in northern Minnesota and Wisconsin K.J. Schulz, W.F. Cannon, L.G. Woodruff and R.A. Ayuso Geochemistry of Archean Gneisses in Dickinson County, Northern Michigan Clarence Surette and Jill Taylor-Hollings Towards understanding geoarchaeological contexts in Northwestern Ontario: The newly formed lithic material comparative collection at Lakehead University Nicholas L. Swanson-Hysell, Sonia M. Tikoo and L.M. Fairchild New paleomagnetic constraints on the formation of the Slate Islands impact structure Nicholas L. Swanson-Hysell, Sarah P. Slotznick and L.M. Fairchild An oxygenated Paleolake Nonesuch and primary detrital hematite in the Freda river system Shiwei Wang, Pete Hollings and Ben Kuzmich Petrological and geochemical characteristics of the granitic rocks from the Dog Lake Granite Chain: Implications for the genesis of Quetico Basin Laurel G. Woodruff, Suzanne W. Nicholson, Connie L. Dicken and Klaus J. Schulz Mineral deposits of the Midcontinent Rift System - A new space/time classification *Jackie Wrage, Adrian Fiege, Brian Konecke, Adam Simon, Philipp Ruprecht and Harald Behrens Sulfur mobility in arc magma systems: Implications for porphyry ore deposits Michael J. Zieg Multiscale Layering in the Black Sturgeon Sill, Nipigon, Ontario * denotes a student eligible for Best Student Paper Award. To be eligible students must have graduated no more than one month before the ILSG meeting, be first author, and present the paper at the meeting. xxxi �ABSTRACTS 1 �Sulfur, Carbon, and Oxygen Isotope Geochemistry of ~2.93 Ga Mesoarchean Chemical Sedimentary rocks in the Red Lake Area, Ontario AFROZ, Munira1, FRALICK, Philip1, KILLINGSWORTH, Bryan2,3, HOMANN, Martin3 SANSJOFRE, Pierre3 and LALONDE, Stefan3 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada. Institut de Physique du Globe de Paris, 1 Rue Jussieu, Paris, France. 3European Institute for Marine Studies, CNRS-UMR6538 Laboratoire Géosciences Océan, Technopôle Brest-Iroise, Plouzané, France. 2 Isotope geochemistry provides important insight into ancient marine carbon and sulfur sources and their role in evolving biologic activity. This research studied ~2.93Ga Mesoarchean chemical sedimentary rocks and carbonaceous slate directly underlying the Red Lake carbonate platform to explore these interactions through the analysis of sulfur, carbon and oxygen isotopes. Core samples of sulfidic iron formation, black slate, and carbonate rocks from 11 drill holes through the carbonate platform were analyzed using mass spectrometry. Multiple isotopes of sulfur (i.e. δ32S, δ33S, δ34S, and δ36S) were measured from sulfidic iron formation samples, while δ13C was examined from carbonaceous slate and δ13C and δ18O from inorganic carbonates. δ34S (‰ VCDT) -15 -10 -5 0 5 10 15 EBL-27 PB-32 PB-33 PB-34 PB-35 Figure 1: δ34S plot of samples from different drill-holes. Figure 2: δ34S vs. ∆33S plot of Red Lake samples with additional literature data (After Johnston, 2011) The analysis showed that sulfur in pyrite was derived from multiple sources as evident from the δ34S values in Fig. 1. Near zero values of δ34S indicate sulfur leached from primary sources due to high-temperature hydrothermal fluids (Thode et al., 1961), whereas δ34S values of >5‰ indicate that some of the sulfur was derived from Archean seawater (Ono et al., 2003). Finally, the lower negative values are indicative of bacterial sulfate reduction in sediments (Seal, 2006). In addition, the δ34S vs. ∆33S plot (Fig. 2) reveals that mass-independent fractionation of sulfur (diagonal array of samples) as well as microbial processing of sulfur (horizontal trend of samples) was active in the Mesoarchean sulfidic iron formation (Ono et al., 2003). The organic δ13C isotope plot (Fig. 3) has lighter δ13C values (~ -30‰) near the bottom of the stratigraphy while heavier δ13C values (~ -17‰) are exhibited towards the carbonate platform. This trend, especially less fractionated values of C, indicates that purple sulfur bacteria might be present in the shallow water carbonate platform along with cyanobacteria as these bacteria fractionate carbon isotopes differently (Posth et al., 2017). Furthermore, the dolostone samples have lighter δ18O isotope values (Fig. 4) which suggests dolomitization was not confined to the marine 2 �environment, instead, it was influenced by fresh water that produces lighter isotopic signatures (Wright and Tucker, 1990). Figure 3: δ13C plot with stratigraphy Figure 4: Mg/Ca vs. δ18O plot of carbonates (After Jaffrés et al., 2007) Based on the results, it is concluded that the source of sulfur was varied in the sediments below the Red Lake carbonate platform and was fractionated by both mass-dependent and massindependent processes. The δ13C trend of organic carbon hints that different bacterial communities were living on the carbonate platform. The δ18O signature indicates that dolostones were precipitated from a mixed water environment. References Jaffrés, J. B. D., Shields, G. A., & Wallmann, K. (2007). The oxygen isotope evolution of seawater: A critical review of a long-standing controversy and an improved geological water cycle model for the past 3.4 billion years. Earth-Science Reviews, 83(1–2), 83–122. Johnston, D. T. (2011). Multiple sulfur isotopes and the evolution of Earth’s surface sulfur cycle. Earth-Science Reviews, 106(1–2), 161–183. Ono, S., Eigenbrode, J. L., Pavlov, A. A., Kharecha, P., Rumble, D., Kasting, J. F., & Freeman, K. H. (2003). New insights into Archean sulfur cycle from mass-independent sulfur isotope records from the Hamersley Basin, Australia. Earth and Planetary Science Letters, 213(1), 15–30. Posth, N. R., Bristow, L. A., Cox, R. P., Habicht, K. S., Danza, F., Tonolla, M., Canfield, D. E. (2017). Carbon isotope fractionation by anoxygenic phototrophic bacteria in euxinic Lake Cadagno. Geobiology, 15(6), 798–816. Seal, R. R. (2006). Sulfur Isotope Geochemistry of Sulfide Minerals. Reviews in Mineralogy and Geochemistry, 61(1), 633–677. Thode, H. G., Monster, J., & Dunford, H. B. (1961). Sulphur isotope geochemistry. Geochimica et Cosmochimica Acta, 25(3), 159–174. Wright, V. P., & Tucker, M. E. (1990). Carbonate sedimentology. Blackwell scientific publications. 3 �Geology and Geochemistry of the Terrace Bay Batholith, N. Ontario ARNOLD, Kira1, HOLLINGS, Pete1, MAGNUS, Seamus2, ZUREVINSKI, Shannon1, CREASER, Robert3 1 Department of Geology, Lakehead University, Thunder Bay, Ontario P7B5E1 Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, Earth Resources and Geoscience Mapping Section, 933 Ramsey Lake Road, Sudbury, ON, P3E 6B5, Canada 3 Department of Earth & Atmospheric Sciences, University of Alberta, 126 ESB Edmonton, Alberta, T6G2R3, Canada 2 The Terrace Bay Batholith is a 25 km long oval shaped granitoid intrusion located in the western portion of the Schreiber-Hemlo greenstone belt, part of the larger Wawa-Abitibi terrane (Fig. 1). The pluton, emplaced at 2689±1.1 Ma (Kamo 2016) intrudes circa 2720 Ma metavolcanic rocks, and a nearby pluton of equivalent age intrudes circa 2698-2693 Ma clastic metasedimentary rocks (Kamo 2016; Davis and Sutcliffe 2017). Younger plutonism in the region occurred between 2673 and 2667 Ma (Kamo 2016; Kamo and Hamilton 2017). The purpose of this study was to classify the Terrace Bay Batholith petrographically and geochemically in order to investigate the petrogenesis and tectonic setting in which the pluton formed, and to characterize the gold and base metal mineralization associated with the intrusion. Detailed mapping showed that the pluton can be separated into three mineralogically distinct lithologies (Fig. 1): granodiorite (typically composed of medium to coarse quartz and feldspar phenocrysts in a groundmass of fine-grained amphibole, biotite, disseminated magnetite, and sulphide minerals), monzogranite (composed of medium-grained quartz and feldspar with increased amounts of potassium feldspar and amphibole relative to the granodiorite), and diorite (composed of medium-grained amphibole and plagioclase with little to no quartz or potassium feldspar present). Two types of hydrothermal alteration are present: chlorite-epidote alteration and a pervasive hematite alteration. The faults and shears in the pluton likely acted as pathways for the hydrothermal fluids. Geochemically, the pluton is a homogenous calc-alkaline pluton, with minimal variation between lithologies. The pluton exhibits trace element signatures that are characteristic of suprasubduction zone magmas, including: fractionated heavy rare earth elements, negative high field strength element anomalies, enrichment of Th over light rare earth elements and enrichment of light rare earth elements. The fractionated heavy rare earth elements and the Th-Nb-La systematics are consistent with formation in a subduction zone at depths where garnet is stable. The Sr/Y and La/Yb signatures support formation within the garnet stability field and suggest small amounts of slab-derived melt were incorporated into the mantle wedge. The εNd values ranging from +2.16 to +2.49 suggest that the pluton underwent minimal crustal contamination during melting and emplacement. The emplacement of the pluton was determined to be through multiple injections derived from a single source. Prolonged fractional crystallization may have resulted in the formation of subtle mineralogical variation but no geochemical differences. Molybdenum mineralization in the pluton is spatially associated with gold mineralization, which suggests it was deposited during the same hydrothermal event. Gold and molybdenum mineralization is generally disseminated throughout the pluton at low concentrations, with higher concentrations of the metals hosted in sulphide mineralized quartz veins. Rhenium-Osmium isotopes from samples of molybdenum from these sulphide-mineralized quartz veins yielded an age of 2671 ±12 Ma, as well as postdating the emplacement of the pluton. Candela (1991) suggests that in plutons emplaced at greater depths, aqueous phases will remain dispersed 4 �throughout the magma, resulting in disseminated mineralization such as that in the Terrace Bay pluton. Figure 1. Simplified bedrock geology map of the Terrace Bay batholith and surrounding greenstone belt in Priske, Strey and Syine townships. Modified from Arnold et al. (2017). References Arnold, K.A., Hollings, P., Magnus, S.J. 2017. Geology and mineral potential of the Terrace Bay pluton, western Schneider-Hemlo greenstone belt; in Summary of Fieldwork and Other Activities, 2017, Ontario Geological Survey, Open File Report 6333, p.12-1 to 12. Candela, P. A. 1991. Physics of aqueous phase evolution in plutonic environments. American Mineralogist, p. 76. Davis, D.W. and Sutcliffe, C.N. 2017. U-Pb geochronology by LA-ICPMS in samples from northern Ontario; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 131p. Kamo, S.L. 2016. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 1: 2015-2016; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. Kamo, S.L. and Hamilton, M.A. 2017. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 2: 2016-2017; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 72p. Kamo, S.L. 2018. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 3: 2017-2018; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 44p. 5 �Multi-scale AEM and MT mapping of the Precambrian in Upper Michigan, Northern Wisconsin, and Eastern Minnesota BEDROSIAN, Paul A. U.S. Geological Survey, MS 964, Federal Center, Denver, CO, 80225 The U.S. Geological Survey is conducting airborne electromagnetic (AEM) and magnetotelluric (MT) surveys over parts of Minnesota, Upper Michigan, and Wisconsin to map Precambrian geology and inform mineral resource assessments in this complex region. A total of 2,700 line-km of AEM data were collected along 16 regional transects over an area of 75,000 km2. These transects range from 100 to 300 km in length and cross numerous structural boundaries and more than 2 billion years of geology. An additional 100+ MT stations have been collected along some of these transects to refine regional resistivity models (Bedrosian, 2016) based on broadly-spaced EarthScope MT stations (Fig. 1). Pronounced contrasts in electrical resistivity exist between conductive sedimentary/metasedimentary rocks and resistive volcanic/intrusive rocks. Archean rocks of the Superior and Minnesota River Valley provinces are imaged as monolithic resistors (Figure 2), whereas strong conductors are linked to metamorphic graphite and metallic sulfides within Paleoproterozoic (PP) rocks of the Penokean orogen (most notably the Michigamme Formation). These conductors are evident in regional-scale resistivity models (Fig. 1b) and extend well into the lower-crust beneath the Penokean orogen (Bedrosian, 2016). Their upper-crustal geometry is being refined by ongoing MT investigations, while in the near-surface, AEM models image narrow (100s of meters wide) sub-vertical conductors (Fig. 2) extending for tens of kms along mapped or inferred faults and shear zones. Laboratory measurements on core samples confirm the presence of graphite in several of these conductive zones. Additional conductors mapped by the AEM data are preferentially located along the periphery of Archean gneiss domes in the region, suggesting that either the oldest PP units are anomalously conductive, and/or that locallyenhanced metamorphic grade is required to form conductive minerals. MT and AEM models of conductive PP rocks further constrain and refine structural details, such as a southern dip on the Niagara fault and a northward extension of PP rocks beneath younger rocks as far north as the Keweenaw fault. Within rocks of the Mesoproterozoic Midcontinent Rift System (MRS), the primary electrical contrast is between resistive volcanic and intrusive rift rocks and the conductive sedimentary successions of the Oronto Group, the Bayfield Group, and the Jacobsville Sandstone. East of the Keweenaw Peninsula, a thick succession of the latter exhibits a similar resistivity signature to that of the Freda Formation on the other side of the peninsula. Relative to lab measurements on these rocks, conductivity in the AEM models for both Freda and Jacobsville is elevated, possibly indicating elevated salinity in the pore waters of these Precambrian aquifers. Together with well control, the AEM models refine structure in several locations, including recognition of the Bayfield Group as more spatially limited than previously recognized and models that suggest the concealed edge of the MRS crosses the central U.P. along a linear magnetic boundary (near 46°15´N, 86°30´W). Imaged younger features include paleochannels cut into Precambrian sediments beneath Lake Superior (Fig. 2), an eastward-thickening wedge of Paleozoic cover in the Eastern survey area (the Northwestern edge of the Michigan basin), and a veneer of glacial sediments, the variable thickness of which can be mapped along each of the AEM profiles. The latter presents modeling 6 �challenges, currently under investigation, due to strong induced-polarization effects in clay-rich glacial tills. References Bedrosian, P.A. 2016. Making it and breaking it in the Midwest: Continental assembly and rifting from modeling of EarthScope magnetotelluric data, Precambrian Res., 278, 337-361, doi: 10.1016/j.precamres.2016.03.009. Figure 1: (a) AEM flight lines (black) and MT stations (white and green circles) atop magnetic anomaly map. Magnetic highs (red) are primarily due to thick volcanic successions (V), iron formations (IF) and intrusive complexes (IC). (b) Regional electrical resistivity model at 5 km depth. Conductors (red) correlate with PP metasedimentary rocks (shaded). White lines denote regional faults; yellow line indicates profile shown in Figure 2. Figure 2: Interpreted resistivity cross-section derived from AEM data. Profile location highlighted in Figure 1. Vertical exaggeration 10:1. 7 �LA-ICP-MS Micro-Sampling of Iron Formation: What it Can Tell Us BIELSKI, Paul and FRALICK, Philip Department of Geology, Lakehead University, Thunder Bay, ON, Canada The occurrence of iron formation during the Archean is well documented, however the mechanisms of their genesis are poorly understood within shallow waters and even less so within the deep-ocean. At the same time our understanding of Archean deep-ocean chemistry is also limited and poorly constrained. To address these issues, a new method for analysis of sulphide facies iron formation geochemistry is being conducted. This method involves a geochemical analysis of deep-ocean iron formation facies at a sub-lamination scale with attention to possible indicators of deposition rate and changes in water chemistry due to mixing of ambient seawater with a hydrothermal plume. Thus, changes in water chemistry during individual cycles of deposition can be measured. This method is conducted using Laser Ablation Inductively Coupled Mass Spectrometry (LA-ICP-MS) alongside Scanning X-Ray Fluorescence (XRF). To test this new application of small scale geochemical analysis, an investigation of the Morley Occurrence was conducted. The Morley Occurrence is a deep-water Neoarchean (~2.7 Ga) sulphide-facies iron formation sitting upon intermediate flows and pyroclastic rocks and overlain by mafic flows and minor turbidites (Fralick et al., 1989). The occurrence itself is about 3 km south-east of Schrieber, Ontario. What makes this site interesting for this application is that oxides replace pyrite at the top of some thin colloform laminations (Fig. 1). These colloform structures are composed of sub-millimeter to millimeter thick pyrite laminations with increasing chert, carbon, and detrital minerals toward their tops. Applying LA-ICP-MS to these pyrite laminations at a sub-laminae level has provided information on geochemical changes of the depositional waters (Fig. 2) in addition to being proof of concept for this application to be used on other facies of iron formation. Laser ablation data (Fig. 2) shows a decrease in Ti upwards through pyrite laminations while Zr increases before resetting at each new lamination. Comparison with data on other hydrothermally sourced metals, such as Ni, Mn, Zn, and Pb, indicates that Ti is of hydrothermal origin while Zr is unrelated to venting fluids. This agrees with the pattern generated when plotting the series of laser ablation shots against Ti and Zr. The source of Zr could be thought to be from detrital origin, however the Zr concentrations are quite low for detrital sediments (Fig. 2). In addition, Zr has a positive relation to Y, Hf, U, and Th (high-field strength elements) along with samples having variable Zr/Hf ratios comfortably below and occasionally significantly above both chondrite and continental values which points to possible preferential scavenging of Hf from seawater by non-detrital sediment leading to fractionation between the two (Bau and Alexander, 2009). This explanation agrees with the data: a resetting Zr value at the beginning of each new pyrite lamination which increases with an assumed decrease in deposition rate with the general rate of deposition based off of chert and detrital sediments increasing upwards. 8 �The LA-ICP-MS data from the Morley Occurrence indicates that hydrothermal influence decreased upwards through each laminae, while seawater influence increased upward. A decreasing deposition rate upward through a laminae resulted in increased scavenging time for elements such as Zr and possibly increased concentration of rainout detritus containing Zr. Figure 1. Left: A photomicrograph of colloform laminations. Right: An example of LA-ICP-MS shots through colloform laminations (Yellow scale bar is 500 um). Figure 2. Plots of LA-ICP-MS Ti and Zr data taken through a set of colloform laminations. Square points represent where each of the three new pyrite laminations begin. References Bau, M. and Alexander, B.W., 2009. Distribution of high field strength elements (Y, Zr, REE, Hf, Ta, Th, U) in adjacent magnetite and chert bands and in reference standards FeR-3 and FeR-4 from the Temagami iron-formation, Canada, and the redox level of the Neoarchean ocean. Precambrian Research, 174(3-4), pp.337-346 Fralick, P.W., Barrett, T.J., Jarvis, K.E., Jarvis, I., Schnieders, B.R. and Vande Kemp, R., 1989. Sulfidefacies iron formation at the Archean Morley occurrence, northwestern Ontario; contrasts with oceanic hydrothermal deposits. The Canadian Mineralogist, 27(4), pp.601-616. 9 �High-resolution dating of the magmatic plumbing system of the Midcontinent Rift System—Insights into rift evolution and mineralization processes BLEEKER, Wouter1, HAMILTON, Michael2, KAMO, Sandra2, LIIKANE, Dustin2,3, SMITH, Jennifer1, HOLLINGS, Pete4, CUNDARI, Robert5, EASTON, Michael6, and DAVIS, Don2 1 Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario K1A 0E8 2 Jack Satterly Geochronology Laboratory, U. of Toronto, 22 Russell St., Toronto, Ontario M5S 3B1 3 Dept. of Earth Sciences, University of Toronto, U. of Toronto, 22 Russell St., Toronto, Ontario M5S 3B1 4 Department of Geology, Lakehead University, 955 Oliver Rd, Thunder Bay, Ontario P7B 5E1 5 Ontario Geological Survey, 435 James Street South, Thunder Bay, Ontario P7E 6S7 Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, Ontario P3E 6B5 Emails: wouter.bleeker@canada.ca; mahamilton@es.utoronto.ca; dustin.liikane@mail.utoronto.ca 6 North America’s Midcontinent Rift System is one of the best preserved and most accessible Proterozoic failed intra-cratonic rifts in the world, and therefore a pre-eminent natural laboratory for understanding the evolution of complex rift systems in cratonic settings, what generates them, what makes them fail, and the myriad of processes associated with their magmatic, sedimentary, and structural evolution. A key dataset that is fundamental to any deeper understanding of this rift system, including its endowment of mineralization, consists of precise and accurate ages on all the different components that make up this rift system. Already, there is a rich literature on dating (mostly U-Pb) of the rift system (see Bleeker, 2018, for a recent summary). Much recent progress has focused on improving the age resolution of volcanic rocks that fill the rift, in conjunction with detailed paleomagnetic investigations, to resolve in more detail the rapidly evolving apparent polar wander path (e.g., Fairchild et al., 2017). Nevertheless, many key components of the rift system, including a wide variety of intrusions that are part of the complex plumbing system of the rift, remain undated or have ages that require refinement, or have dates that are clearly puzzling outliers in the evolving temporal framework of U-Pb ages. Some of the published ages were obtained on limited amounts of small baddeleyite crystals and suffer from associated complications (Pb loss and variable discordance, elevated common Pb and associated corrections, ambiguity in choice of regression line and upper intercept, subtly different systematics between baddeleyite and zircon, etc.). In some cases, there exists doubt on the exact provenance or sample location of dated samples. Our aim is to revisit many, if not all, of the intrusions, and particularly those associated with mineralization, to improve and complete the U-Pb age framework, ideally at ~1 Myr resolution; and to link the intrusive record to the better resolved volcanic record as well as the overall tectono-magmatic evolution of the rift. Initially we have focused on some of the “outliers” such as the Inspiration Sill (Nipigon area, with a published age of 1159±33 Ma), or the iconic Logan Sills overlooking Thunder Bay (Fig. 1), with an age interpretation of 1114.7±1.1 Ma based on limited and discordant baddeleyites (Heaman et al., 2007). These problems can be tackled by searching for more optimum samples in the field, and applying ever improving U-Pb analytical techniques (lower blanks, new and better calibrated spike solutions, chemical abrasion of zircons, etc.). Searching for zircon-bearing samples is the key for ultra-high precision ages. 10 �Figure 1: Above: the iconic Logan Sills (s.s.) overlooking the Kaministiquia River and the city of Thunder Bay. Two sills are visible, having intruded the mudstones and thinly bedded turbiditic sedimentary rocks of the ca. 1.85 Ga Rove Formation, Animikie Basin; an upper main sill capping the mesas, and a thin lower sill forming a minor ledge in the trees. Right: our optimum sample of evolved, late-stage, varitextured and in part pegmatitic gabbro from near the top of the main upper sill, at Mount McKay. Figure 2: Left: Cross-cutting relationship of younger NNE-trending Pigeon River dyke cutting across, and chilled against, older coarse-grained and sparsely porphyritic diabase of one of the main Cloud River dykes. New age data are available for both the Pigeon River and Cloud River dykes. Together with searching for more optimum samples in the field, or in drill core, a key aspect of our study also involves resolving cross-cutting relationships in the field, where they exist, to help guide overall interpretation (Fig. 2). Already we have new and more robust age data on ~10 key units, including previously undated mineralized intrusions, which will be discussed at the meeting. Among those are: the Inspiration Sill, the main Logan Sill (Fig. 1), Pigeon River dykes, Cloud River dyke, Sunday Lake and Current Lake intrusions, Crystal Lake intrusion, Mount. Mollie dyke, Bovine Igneous Complex, and several others. Acknowledgements: we thank numerous industry partners and colleagues at the USGS for their keen interest in this study, their scientific input, and their generous cooperation. References Bleeker, W., Liikane, D.A., Smith, J., Hamilton, M., Kamo, S.L., Cundari, R., Easton, M., and Hollings, P., 2018, Controls on the localization and timing of mineralized intrusions in intra-continental rift systems, with a specific focus on the ca. 1.1 Ga Mid-continent Rift system. Geological Survey of Canada, Open File 8373, p. 15–27. https://doi.org/10.4095/306594. Fairchild, L.M., Swanson-Hysell, N.L., Ramezani, J., Sprain, C.J., and Bowring, S.A., 2017. The end of Midcontinent Rift magmatism and the paleogeography of Laurentia. Lithosphere, vol. 9, p. 117–133. Heaman, L.M., Easton, R.M., Hart, T.R., Hollings, P., Mac-Donald, C.A., and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences, vol. 44, p. 1055–1086. 11 �New age data and insights into the ca. 1887-1870 Ma Circum-Superior Belt, with startling implications for the Lake Superior area geology BLEEKER, W.1, KAMO, S.2, HAMILTON, M.2, and CHAMBERLAIN, K.3 1 Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario K1A 0E8, wouter.bleeker@canada.ca Jack Satterly Geochronology Laboratory, U. of Toronto, 22 Russell St., Toronto, Ontario M5S 3B1 3 Department of Geology and Geophysics, University of Wyoming, Laramie, USA 2 Introduction: The ca. 1887-1870 Ma Circum-Superior Belt (Baragar and Scoates, 1981) has long been recognized as one of Canada’s major metallogenic belts, principally because of its world-class Ni-Cu-PGE deposits at Thompson and Raglan, as well as a number of other significant prospects elsewhere along the belt (e.g., Labrador Trough). Discontinuous outcrop along the margin of the Superior craton, remoteness, and the great extent of the belt has hampered detailed correlations between different segments. Although first-order correlations were suspected, and included volcanic rock in the Lake Superior area, U-Pb geochronology has only recently advanced to a point where we can demonstrate that peak mafic-ultramafic magmatism was coeval between localities such as Thompson and Raglan, with large volumes of Mg-rich ultramafic rocks being emplaced at 1882 Ma, both as flows, sills, and feeder dykes. Similar age mafic-ultramafic magmatism is now known from around the Superior craton, from northern Quebec to Minnesota, while ca. 1882 Ma dyke swarms intrude far into the cratonic hinterland. Carbonatites and kimberlites are coeval, within age uncertainty, with the maficultramafic magmatism or precede peak magmatism by a few million years, an age pattern also seen in other large igneous provinces (e.g., Bushveld Complex and slightly earlier carbonatites, Phalaborwa). A model that best explains the present observations is that of a mantle plume impinging on the base of thick cratonic lithosphere of supercraton Superia (Bleeker, 2003), with hot, low-viscosity, plume mantle then flowing laterally into multiple thin spots and incipient rifts, localized along the present margins of the Superior cratonic fragment, where large-scale and nearly synchronous decompression melting ensued at 1882±1 Ma (Fig. 1; see Bleeker and Kamo, 2018; and references therein). Figure 1: Interpreted geodynamic setting of the Circum-Superior Belt: plume ascent, interaction with cratonic lithosphere, and continental breakup. a) Ascending mantle plume impinging on thick lithosphere of the ancestral Superior craton, i.e. supercraton Superia. b) Flattening and rapid lateral flow of hot, buoyant plume mantle to thin spots, leading to nearly synchronous large-volume mafic-ultramafic magmatism (after Bleeker and Kamo, 2018). 12 �Continenal breakup: The emerging picture, with nearly synchronous mafic-ultramafic magmatism around what are now the margins of the Superior craton, clearly indicates a geological setting of continental breakup rather than that of accreting arcs at 1882 Ma. It is thus important to think about the Superior craton, and the geology of the Lake Superior area, in the context of progressive continental breakup. There is growing evidence that the Kaapvaal craton of South Africa, and thus also supercraton Vaalbara, was attached to the southwestern corner of the Superior craton (Bleeker et al., 2016; see also Gumsley et al., 2017), as part of an ancient terrane (also comprising much of Wyoming craton and Karelia) that collided with the growing Superia landmass at ca. 2650 Ma, and remained there until the ca. 1880 Ma breakup event introduced above. Indeed there are ca. 1880 Ma dykes in the eastern Kaapvaal craton. Furthermore, the proposed reconstruction is paleomagnetically viable. On final breakup, the Minnesota River Valley terrane, a piece of the ancient crust of the eastern Kaapvaal craton, was left stranded as an exotic terrane on the southern breakup margin of the Superior craton (Bleeker et al., 2016). Predictions: The startling conclusion must be that Kaapvaal craton, indeed entire Vaalbara, Wyoming and Karelia, were contiguous with the southern Superior craton from ca. 2650 Ma until progressive breakup from ca. 2000 Ma to 1880 Ma, within the context of supercraton Superia. Indeed dykes of exact Bushveld age (2056 Ma) have been identified in the Western Superior craton (Bleeker et al., 2016), and 2167 Ma Biscotasing dykes have been identified in the eastern Kaapvaal craton (with matching trend!), two independent and exact age matches of short-lived mafic magmatic events that demonstrate, without any doubt, a “nearest neighbour” relationship (Bleeker and Ernst, 2006) of these cratonic fragments over this time interval. This leads to another startling prediction: a potential plume track that starts with the LIP-scale Marathon magmatic event in the eastern Lake Superior area, at ca. 2130 Ma, can be traced to the southwest, with pulsed 2125-2100 Ma mafic magmatism along the southern margin of the Superior craton; it then was responsible for the Fort Frances giant mafic dyke swarm at ca. 2070 Ma; it can then be traced into the easternmost Kaapvaal craton where it is fist manifested by the 2060 Ma Phalaborwa and related carbonatites; and, finally, further west, it eroded the lithosphere and produced Earth’s largest layered mafic intrusion at 2056 Ma, the Bushveld Complex. In conclusion: two of the best-known Archean cratons in the world, Superior and Kaapvaal, shared a common history from ca. 2650 Ma terrane collision until ca. 1880 Ma breakup. On breakup, a piece of ancient Kaapvaal crust, Minnesota River Valley terrane, was left behind. Kaapvaal and Superior were joined and contiguous all through the lead-up to the Bushveld Complex, and the magmatism that culminated with the Bushveld Complex started with the ca. 2130 Ma Marathon event, tracing a continuous plume track from the southern Superior craton into the Kaapvaal craton. References Baragar, W.R.A. and Scoates, R.F.J., 1981. In: Developments in Precambrian Geology, v. 4: p. 297–330. Bleeker, W., 2003. Lithos, v. 71(2): p. 99-134; Bleeker, W., Chamberlain, K.R., Kamo, S.L., Hamilton, M., Kilian, T.M. and Buchan, K.L., 2016. 35th IGC, Cape Town, South Africa, Paper Number 5222. Bleeker, W. and Ernst, R., 2006. In: Dyke Swarms—Time Markers of Crustal Evolution. Balkema, Rotterdam, p. 326. Bleeker, W. and Kamo, S.L., 2018. In: GSC Open File 8373, p. 5–14, https://doi.org/10.4095/306592. Gumsley, A.P., Chamberlain, K.R., Bleeker, W., Söderlund, U., de Kock, M.O., Larsson, E.R. and Bekker, A., 2017. PNAS, v. 114(8): p. 1811-1816. 13 �Stable isotopic composition of calcite precipitated with native copper and other minerals of the Keweenaw Peninsula, Michigan BODDEN, Thomas J.1, BORNHORST, Theodore J.2, BÉGUÉ, Florence3, and DEERING, Chad1 1 Department of Geological and Mining Engineering and Sciences, Michigan Tech, Houghton, MI 49931 2 A. E. Seaman Mineral Museum, Michigan Tech, Houghton, MI 49931 3 Institute of Earth Sciences, University of Lausanne, Lausanne, Switzerland Hydrothermal native copper deposits are hosted by Midcontinent Rift-filling volcanic and sedimentary rocks in Michigan’s Keweenaw Peninsula. Butler and Burbank’s (1929) classic U.S.G.S. Professional Paper documented the district-wide paragenesis of hydrothermal mineral precipitation and the relative age of minerals with respect to the precipitation of native copper, the principal ore mineral. Puschner (2002) subdivided hydrothermal mineral paragenesis into three Stages: Stage 1 (pre-native copper), Stage 2 (syn-native copper), and Stage 3 (post-native copper). Since Stage 1 and 2 minerals represent a single continuous episode of mineral deposition, they will be combined for data presentation below. Stage 3 is a distinct later episode (veins that cross-cut native copper deposits) and thus, will be considered separately. Stages 1-2 formed at a significantly higher temperature than Stage 3 based on mineral equilibria, stable isotope pairs, chlorite geothermometry, and fluid inclusions (Puschner, 2002; Livnat, 1983). In addition to temporal (paragenetic) variation of mineral precipitation within the native copper district, both within the district and the broader Keweenaw Peninsula, there is spatial variation in the suite of minerals in a particular locality (Stoiber and Davidson, 1959). The spatial mineral variation corresponds to a regular variation in the temperature of precipitation of minerals from the hydrothermal fluids. This research is an extension of Bornhorst and Woodruff (1997) who proposed fluidmixing was an important mechanism facilitating native copper precipitation on the basis of the variability of stable isotope data derived from Stage 1-2 calcite from the Kearsarge deposit; the largest basalt-hosted deposit in the district. Calcite is good to study the evolution of the hydrothermal fluids as it precipitates in all three stages and with native copper in Stage 2. The purpose of this study was to test the hypothesis of fluid-mixing proposed by Bornhorst and Woodruff (1997) using a geographically broader data set as well as using secondary ion mass spectrometry (SIMS) to obtain in-situ stable isotope values for calcite. We have compiled 159 published oxygen-carbon stable isotope pairs for calcite from Livnat (1983; 88 pairs), Puschner (2002; 31 pairs), and Bornhorst and Woodruff (1997; 40 pairs) determined by traditional bulk mineral analysis. We have added an additional 101 pairs determined by SIMS on selected spots from three samples. Each of the pairs have been grouped according to paragenetic Stage when possible based on sample description, geographic location, and textural observation from cathodoluminescence imaging; those not able to be grouped are not included in the discussion below. Puschner’s (2002) data was only obtained for Stage 1-2 calcite, as was the case for Bornhorst and Woodruff (1997) with the exception of one Stage 3 calcite. The new SIMS data are from: Stage 1-2 calcite from the Quincy deposit, Stage 2 calcite and Stage 3 calcite from the Kearsarge deposit. The variability of the SIMS spot data from only 14 �three samples is similar to the entire range of the 159 bulk samples as a result of averaging by bulk sampling. The oxygen and carbon isotopic composition of the hydrothermal fluids in equilibrium with the calcite has been calculated considering both temperature variation among paragenetic stages and differences in geographic location. For Stage 1-2 calcite a temperature of 250°C +/50°C was used and for Stage 3 calcite a temperature of 125°C +/- 25°C (Puschner, 2002; Livnat, 1983). The variation of Stage 1 and 2 water in equilibrium with calcite is widely scattered between δ18OH2O of about +22 to +4 ‰ and δ13CCO2 of about +3 to -8 ‰ (using midpoint temperatures). The total variability in Stage 1-2 water stable isotopic compositions can only partly be explained by considering paragenetic, spatial, and local temperature variation. Thus, the larger data set compiled for this study, representing Stage 1-2 water in equilibrium with calcite, is consistent with the observations of Bornhorst and Woodruff (1997). To explain the oxygen isotopic data in his limited data set, Puschner (2002) proposed that ore fluids were derived through metamorphism and mixed with a meteoric water at shallow depths during formation of the native copper deposits. The range in δ18OH2O and δ13CCO2 from our data supports this conclusion. Stage 3 ranges in δ18OH2O from about +8 to -2 ‰ and in δ13CCO2 from about 0 to -10 ‰ using midpoint temperatures. In δ18OH2O and δ13CCO2 space Stage 3 calcite is generally different than Stage 1-2, but overlaps with Stage 1-2 at the higher values of δ18OH2O. Stage 3 calcite isotopic composition can only partly be explained by local temperature variation. The range of δ18OH2O for Stage 3 calcite overlaps the expected range of values for meteoric and metamorphic waters and is consistent with a tentative interpretation of fluid-mixing. Further interpretations are in progress. This study was partially supported by an ILSG Student Research Grant. References Bornhorst, T.J., and Woodruff, L.G., 1997, Native Copper Precipitation by Fluid-Mixing, Keweenaw Peninsula, Michigan: Institute on Lake Superior Geology Proceedings, 43rd Annual Meeting, v. 43, part 1, p. 9-10. Butler, B.S., and Burbank, W.S., 1929, The copper deposits of Michigan: U.S. Geological Survey Professional Paper 144, 238 p. Livnat, A., 1983, Metamorphism and copper mineralization of the Portage Lake Lava Series, northern Michigan: Ph.D. Dissertation, University of Michigan, Ann Arbor, 292p. Puschner, U.R., 2002, Very low-grade metamorphism in the Portage Lake Volcanics on the Keweenaw Peninsula, Michigan, USA: Ph.D. Dissertation, University of Basel, Basel, Switzerland, 82p. and appendices Stoiber, R.E., and Davidson, E.S., 1959, Amygdule mineral zoning in the Portage Lake Lava Series, Michigan copper district: Econ. Geol., v. 54, p. 1250-1277, 1444-1460. 15 �Recognition of probable distal ejecta from the 1850 Ma Sudbury meteorite impact event along the southern edge of the Animikie basin in Minnesota BOERBOOM, Terrence J., Minnesota Geological Survey Petrographic examination of a drill core (LM-13-4) obtained in 2013 by Minerals Processing Corporation, located at the southeastern margin of the Animikie basin (Fig. 1), has revealed the presence of an approximately 5 m thick interval with features that can be attributed to distal Sudbury ejecta. These include sphere-in-sphere structures, vesiculated devitrified glass, zoned accretionary lapilli, anatase, and possible (albeit questionable) rare and poorly preserved decorated PDF lamellae in some quartz grains. These features are similar to those described from other locations, including Michigan (Cannon and others, 2010), and Ontario and Minnesota (Addison and others, 2005), among others. This location is approximately 950 km from Sudbury. This core is from a belt that has historically been mapped as part of the southward-adjacent Mille Lacs Group which is thought to predate deposition in the main Animikie basin. However, a more recent reinterpretation (Boerboom, 2009) places this belt in the lower part of the Animikie basin, as part of the Thomson Formation, an interpretation supported by the presence of this ejecta layer. The vertical drill core intersects bedrock at 137’/42m depth (beneath glacial drift) and ends at 437’/133m depth. The ejecta horizon occurs in the Figure 1. Cartoon map of the 291-309’ interval. Bedding is upright and (most likely) dips north an average of Animikie basin showing 60 degrees. Major folds appear to be lacking in this core, despite the presence location of drill core LM13-4. of a weak, nearly vertical cleavage. The approximately 5.5m thick interval attributed to the ejecta horizon lies within a thick, low-grade turbidite sequence. The core above the ejecta horizon is gray and orange-ochre banded ‘ferruginous slate’, likely a more weathered and oxidized version of the gray carbonaceous argillite and graywacke below the ejecta horizon which contains numerous thin beds of brownish carbonate beds and pyrite. The ejecta horizon can be divided into three distinct portions (Fig. 2) – a lower lapilli-rich layer (20cm, not all shown in Fig. 2), a middle fragmental and brecciated layer (10cm), and an upper sandy layer that Figure 2. Lower portion of the eject interval. Top of core is to the left. Bottom 6 inches/13cm is not shown. Drill core is 3.5 cm in width. continues upward for several meters and appears to grade into the overlying turbidites. The lower portion contains abundant dark gray vertically flattened, weakly and concentrically zoned accretionary lapilli that increase in size and abundance upward, in a matrix of small pale green and dark gray, angular shard – like clasts. The middle layer contains angular chert clasts at the base and larger elongate pale green shard-like clasts at the top. The upper layer is composed of a grayish-green sandy graywacke with 1-3mm lapilli in the lower part that are weakly concentrated along bedding planes. This interval may represent an influx of debris ultimately derived from the Sudbury impact crater, possibly a submarine debris flow slumped downslope from its original depositional source, within an otherwise unbroken turbidite sequence. Despite thorough replacement by secondary phyllosilicate minerals, there are many well-preserved features (Fig. 3) that compare to those elsewhere attributed to ejecta fallout. To date positive identification of shocked quartz has been unsuccessful. However, some grains bear parallel arrays of linear bubble trains which may represent decorated quartz planar deformation features (Fig. 3c). Nonetheless there are many other features that can be attributed to ejecta fallout, and the location along the southern margin of the Animikie basin is where it logically would be expected. 16 �Huber and others (2014) describe spherules (their term) from drill cores near Coleraine that contain microcrystalline rutile and anatase in the outer rims. They state that because the transition from rutile to anatase at low pressures is in the 500-600 C range, and because the rocks at Coleraine were only subject to low T metamorphism, that anatase formed in the spherulitic melt droplets as they cooled. Thin sections from core LM13-4 contain abundant anatase as confirmed by SEM and by optical properties. However, the anatase does not occur as fine granular masses, but rather as prismatic crystals up to 0.8mm in length concentrated in microscopically dark-opaque zones interpreted to be deformed devitrified glass shards. The anatase laths are commonly rimmed by zones that are not opaque, implying they may have formed by some secondary mechanism such as diagenesis, hydrothermal alteration or lowgrade metamorphism. In contrast, ilmenite is the dominant Ti-phase within the accretionary lapilli. The significance of this is currently unknown. Ongoing petrographic work will attempt to more positively identify shocked quartz, and SEM and XRD work will be conducted to better characterize the secondary mineralogical assemblages. If further analytical and petrographic data conclude the material is ejecta-bearing, it will be the first such occurrence along the southern Animikie basin in Minnesota. A B C D E Figure 3. A. Sphere-in-sphere structures interpreted as melt droplets, with chloritic cores and sericitic rims. B. Clast of devitrified vesicular glass with internal spherical structures. C. Straight bubble trains in quartz – possible relict deformation lamellae? D. Zoned accretionary lapilli most visible at thin edge of thin section. E. Reflected light image of showing ilmenite (Ilm) in accretionary lapilli, and anatase (An) in dark semiopaque zones (in transmitted light) that are interpreted as possible deformed pumice-like fragments. References Boerboom, T.J., 2009, Bedrock geologic map of Carlton County, Minnesota; Minnesota Geological Survey County Atlas Series C-19, Plate 2; scale 1:100,000. Huber, M.S., McDonald, I., and Koeberl, C., 2014, Petrography and geochemistry of ejecta from the Sudbury impact event: Meteoritics and Planetary Science 49, No. 10, P. 1749-1768 Cannon, W.F., Schulz, K.J., Horton, J.W., Jr., and Kring, D.A., 2010: The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA, GSA Bulletin v. 122; no. 1/2; p. 50–75. Addison, W.D., Brumpton, G. R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W., and Hammond, A.L., 2005: Discovery of distal ejecta from the 1850 Ma Sudbury impact event, Geology; March 2005; v. 33; no. 3; p. 193–196. 17 �Geology and Geochemistry of Ultramafic rocks in the Lake of the Woods Area BOUCHER, Chanelle and HOLLINGS, Pete Department of Geology, Lakehead University, Thunder Bay, Ontario P7B5E1 The Archean komatiites of the Lake of the Woods greenstone belt in Kenora, Ontario formed on the western extension of the Superior Province southern margin and have not been studied using modern methods. Although Archean plate tectonic processes have been the subject of decades of research, the nature of these processes remains the subject of considerable debate. Recent work has investigated the link between komatiites and Archean subduction zones. Komatiites are widespread in Archean terranes and together with spatially associated tholeiitic basalts form an important part of many Late Archean greenstone belts, therefore a better understanding of Archean geodynamic processes and comparison to modern day processes is required. The Lake of the Woods greenstone belt (LWBG) is located in the Western Wabigoon Terrane which is composed dominantly by mafic volcanic rocks with large tonalite-granodiorite plutons. The LWGB is situated along the northwestern margin of the Western Wabigoon Terrane, bounded to the north by the Winnipeg River and English River terranes and to the south by the Quetico terrane (Ayer and Davis, 1997). It consists of a northeast trending metavolcanic plutonic belt that extends for 900km and is about 150km wide (Ayer and Davis, 1997). The LWGB is divided into three supracrustal assemblages: a lowermost mafic volcanic Lower Keewatin assemblage; a compositionally diverse, predominantly volcanic middle Upper Keewatin assemblage; and a predominantly sedimentary uppermost Electrum assemblage (Ayer and Davis, 1997). The Upper Keewatin assemblage consists of 1) mafic to felsic metavolcanic rocks of calc-alkalic affinity, 2) ultramafic to mafic metavolcanic rocks of komatiitic to tholeiitic affinity, and 3) turbiditic metasedimentary rocks (Ayer and Davis, 1997). Detailed mapping in the Upper Keewatin Assemblage identified komatiites on the southern margin of the Long Bay Group. The komatiites are typically metamorphosed to upper greenschist facies and include a variety of schists that do not show any preserved primary textures or mineralogy. Polyhedrally jointed flow tops were observed in rare locations. Mineral assemblages include dominantly anthophyllite-tremolite-chlorite (Fig. 1A) and serpentinetremolite-chlorite (Fig. 1B) schists, as well as lesser talc-tremolite-chlorite schists. These units are moderately to intensely foliated with chlorite and lesser amphibole defining the foliation and also include randomly oriented bladed amphibole grains that typically have tremolite cores and anthophyllite rims. The amphiboles show a chemical transition from core to rim with a loss in Ca as anthophyllite appears. Accessory phases include chromite, magnetite, ilmenite and apatite. Ultramafic rocks are very fine-grained, and mineralogy has been described using a compilation of petrography, XRD (x-ray diffraction) and SEM (scanning electron microscope) analysis. Whole-rock geochemical analyses were conducted on 110 samples collected during field work in 2017 and 2018. The Upper Keewatin Assemblage is composed of dominantly mafic to intermediate volcanic rocks that are typically of tholeiitic affinity with rare calc-alkalic units. A total of 41 samples were determined to be ultramafic. The komatiite units are Al-undepleted rocks that display primitive mantle normalized patterns, as well as major and trace element concentrations consistent with melts derived from outside the garnet stability zone. They can be 18 �subdivided into three suites with primitive mantle patterns that display strong Th and Nb depletions with flat HREEs (heavy rare earth elements), weak Th and Nb depletions with flat HREEs and enriched Th with moderate Nb depletions and flat HREEs. Neodymium isotope analyses, in conjunction with trace element geochemistry, suggests that some units have been weakly to moderately contaminated. Mafic tholeiitic units have low- and high-Ti varieties, in which most units are dark grey to black amphibolites and rare chlorite-tremolite schists. The geochemistry of the mafic units shows similar contamination trends to the ultramafic units. Figure 1. Field photographs of grey to dark green foliated ultramafic metavolcanic rocks. A (LOW17CB19 15U 368614 5467778): Pervasive chlorite creates deep green color and moderate foliation. B (LOW17CB92 15U 368861 5465424): Talc alteration with red staining along strong foliation planes. References Ayer, J.A., and Davis, D.W. 1997. Late Archean evolution of differing convergent margin assemblages in the Wabigoon Subprovince: Geochemical and geochronological evidence from the Lake of the Woods greenstone belt, Superior Province, northwestern Ontario; Precambrian Research, 81:155-178. 19 �Updated mineralogy of a roadside pegmatite in the Stettin Complex, Wausau Syenite Complex, Marathon County, Wisconsin. BUCHHOLZ, Thomas W.1, FALSTER, Alexander. U. 2, and SIMMONS, Wm. B. 2 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494; 2Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217. In 2017 we presented an initial report regarding a recently re-exposed pegmatite along 120 Avenue in the SW 1/4 NW 1/4 of Sec. 22, T.29, R. 6E near the western margin of the intrusion. It has become apparent that this aplite/pegmatite is the same roadside pegmatite described in Weidman (1907), having been obscured by slumped soil, rock and vegetation for the intervening 110 years. The Stettin Complex is the oldest (1565 +3-5 Ma, Van Wyck 1994) and most alkalic of the four intrusions that comprise the Wausau Syenite Complex, and is primarily composed of amphibole, pyroxene, tabular and nepheline syenites, and syenite aplite. th In 2017 we reported on the occurrence of albite, arfvedsonite, aegirine, microcline, pyrochlore, monazite-(Ce), bastnäsite-(Ce), cerianite, xenotime-(Y), zinnwaldite, zircon, goethite/hematite replacements after siderite, pyrite, a TiO2 phase, columbite-(Fe), bismuthinite, astrophyllite and fluorite. Several other species were included in the poster presentation but not described in the abstract, hence these are included in the below descriptions. Minor additional xenotime has been identified as sheaves of pale blue crystals in albiterich aplite, associated with aegirine, while further analysis of pyrochlore indicates they are largely fluorcalciopyrochlore, although due to strong, ubiquitous zoning three or more pyrochlore species may be present in various zones in one crystal. Graphite is not uncommon as thin, black crystals in late quartz pods in and near pegmatitic portions of the dike but is easily missed, and several yellow-brown grains of thorbastnäsite have been found in microcline in core zone material. Euxenite-(Y) forms rare small, brown, elongated crystals in pegmatite, and probable thorite and grayite are sparse. Careful visual examination of samples has revealed the first Be-bearing minerals in the Stettin complex in albite-rich aplite, located close to the transition to pegmatite. Phenakite is found as patches of clear, colorless phenakite poikilitically including albite crystals and as isolated grains in pegmatite, and true to its name (from Greek phenas for “deceiver) is difficult to distinguish from similar quartz without the use of optical methods. Bertrandite was found as very pale blue platy crystals in patches in albite near phenakite, and feathery, pale yellow bavenite near bertrandite. Heavy mineral separates have revealed a suite of unusual inconspicuous phases, some present in very small amounts. These include sparse grains of galena and sphalerite, native bismuth with small amounts of a Ca-Bi phase (perhaps kettnerite or beyerite) one small grain of akanthite in native bismuth, and an Ag-Bi-S phase, also in bismuth. The Ag-Bi-S phase remains unidentified as the stoichiometry does not match benjaminite, dantopaite, matildite or pavonite (the known Ag-Bi-S minerals), and paucity of material precludes further investigation. 20 �Cassiterite is not uncommon in some samples, but is difficult to visually distinguish from abundant zircon. Worldwide, cassiterite is very rare from alkalic complexes, as a review of applicable literature and Mindat listings revealed very few occurrences worldwide. Apparently, Sn is not commonly enriched in alkalic environments, although several additional occurrences of cassiterite have been noted in Stettin Complex pegmatites. The occurrence of graphite in and near the core zone suggests a reducing environment during crystallization of those portions of the dike, while the late crystallization of cerianite (requiring oxidation of Ce3+ to Ce4+), replacement of siderite by goethite and hematite, and common partial replacements of aegirine and arfvedsonite by Fe-oxide phases suggests a late transition to an oxidizing environment. References Van Wyck, N. (1994) The Wolf River A-type magmatic event in Wisconsin: U/Pb and Sm/Nd constraints on timing and petrogenesis. Institute on Lake Superior Geology, 40th Annual Meeting, Part 1, Program and Abstracts, p. 81-82. Weidman, Samuel (1907). The Geology of North Central Wisconsin. Wisconsin Geological and Natural History Survey Bulletin No. XVI, Scientific Series No. 4, 697 pp. 21 �The Dickinson Group in the Central Upper Peninsula of Michigan: Part 1- Age and tectonic setting based on new geophysical, geochronological, and geochemical data CANNON, W.F.1, SCHULZ, K.J.1 and DRENTH, Benjamin J.2 1 U.S. Geological Survey, Reston, VA 20192 U.S. Geological Survey, Denver, CO 80225 2 A unique sequence of metasedimentary and metavolcanic rocks is exposed in a ~100 km2 area of central Dickinson County, Michigan. James (1958) divided these rocks into three formations that comprise the Dickinson Group. James et al. (1961) provided additional details of structure and stratigraphy. Our recent geophysical, geochemical, and geochronological studies shed new light on this group and suggest substantial changes to previous interpretations. The basal East Branch Arkose (arkose, conglomerate, and minor basalt flows) grades upward to the Solberg Schist (finer grained-clastics with probable metavolcanic interbeds and a medial banded iron-formation, the Skunk Creek Member). The uppermost formation, the Six Mile Lake Amphibolite (massive to banded hornblende-plagioclase rock), is presumed to be mafic metavolcanics. James et al. (1961) concluded that the Six Mile Lake Amphibolite grades southward into Archean gneiss, so ascribed an Archean age to the entire Dickinson Group. The exposed Dickinson Group lies in a vertical, south-facing monocline about 5 km wide. James et al. (1961) considered this the approximate stratigraphic thickness of the group because they found no indication of internal folding or faulting across that distance. The lack of internal folding is especially well documented for the East Branch Arkose where abundant cross beds all indicate south-facing strata. In the Solberg Schist the Skunk Creek Member can be traced by its strong aeromagnetic anomaly (as much as 2000 nT) as a single horizon for 50 km without indication of structural repetition. This apparent structural simplicity is belied by a ubiquitous penetrative foliation that is steeply dipping and essentially beddingparallel in the East Branch and Solberg as shown by oriented micas, stretched quartz grains, and, in the East Branch Arkose, flattened and elongated pebbles. In the Six Mile Lake Amphibolite oriented hornblende grains define a gently-plunging lineation (fold axes?). Development of these penetrative structures appears to have been synchronous with metamorphism that peaked at about 1.83 Ga (Holm et al., 2007) and thus records deformation during the Penokean orogeny. The exposed Dickinson Group may be the north limb of a large syncline whose southern limb is truncated by a fault against the Archean rocks to the south. East of the exposed area our new aeromagnetic data indicates that the belt of Dickinson Group rocks widens and is more structurally complex (Drenth et al., 2019). The East Branch Arkose contains a significant population of detrital zircons with 2.1 Ga ages (Craddock, et al., 2013) and is clearly Paleoproterozoic rather than Archean. The most abundant clast type in the East Branch conglomerates is orthoquartzite likely derived from the older 2.2-2.3 Ga Sturgeon Quartzite. An age of 2.1 Ga has been determined for the “porphyritic red granite”(prg) (Ayuso et al., 2018), which is surrounded by Dickinson Group strata. The prg likely was an important source of detritus, including zircons, for the East Branch Arkose. Correlation of the Dickinson Group with other Paleoproterozoic sequences of the region is not fully resolved. It is clearly younger than the 2.2-2.3 Ga Chocolay Group and its metamorphism at 1.83 Ga provides an upper age limit. Within current age constraints it could be equivalent to parts of the Menominee and/or Baraga Groups. But, another possibility is that the Dickinson Group is a vestige of a unique sequence deposited during the long hiatus between about 2.1 Ga (prg) and 1.9 Ga (Menominee Group). and provides a record of the final separation of the Superior and Wyoming cratons. The Six Mile Lake amphibolite and a metadiabase sill in the Solberg Schist have distinctive trace element chemistry consistent with a mantle plume source. Within the Lake Superior region that composition is known only in mafic dikes north of Lake Superior (Schulz et al., 2018) that were intruded between 2126 and 2067 Ma and mark a long-lived mantle plume event during separation of the two cratons (Halls et al., 2008). The coarse, locally derived fluvial sediments of the East Branch Arkose are consistent with extensional uplift during which much of the Chocolay Group was stripped from it basement and erosion unroofed 2.1 Ga granite plutons. The ensuing transition of fine-grained clastic sediments and banded iron-formation of the 22 �Solberg Schist marks the transition to marine sedimentation culminating in plume-related mafic volcanism (Six Mile Lake Amphibolite) marking the final continental separation. Geologic map of the Dickinson Group and surrounding units (modified from James et al., 1961). References Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., and Jackson, J., 2018, New U-Pb zircon ages for rocks from the granite-gneiss terrane in northern Michigan: Evidence for events at ~3750, 2750, and 1850 Ma: Institute on Lake Superior Geology, Proceedings of 64th Annual meeting, Part 1: Program and Abstracts. p. 7-8. Craddock, J.P., Rainbird, R.H., Davis, W.J., Davidson, C., Vervoort, J.D., Konstantinou, A., Boerboom, T., Vorhies, S., Kerber, L., and Lundquist, B., 2013, Detrital zircon geochronology and provenance of the Paleoproterozoic Huron (~2.4-2.2 Ga) and Animikie (~2.2-1.8 Ga basins, southern Superior Province, Journal of Geology, v. 121, p. 623-644. Drenth, Benjamin J., Cannon, W.F., and Schulz, K.J., 2019, The Dickinson Group in the Central Upper Peninsula of Michigan: Part 2- Geophysical expression and a preliminary interpretation of its eastward extent under Paleozoic cover, Institute on Lake Superior Geology, Proceedings of 65 th Annual meeting, Part 1: Program and Abstracts. Halls, H.C., Davis, D.W., Stott, G.M., Ernst, R.E., and Hamilton, M.A., 2008, The Paleoproterozoic Marathon large igneous province: New evidence for a 2.1 Ga long-lived mantle plume event along the southern Superior Province, {Precambrian Research, v. 162, p. 327-353. Holm, D.K., Schneider, D.A., Rose, S., Mancuso, C., McKenzie, M., Foland, K.A., and Hodges, K.V., 2007, Proterozoic metamorphism and cooling in the southern Lake Superior region, North American and its bearing on crustal evolution, Precambrian Research, v. 157, p. 106-126. James, H.L., 1958, Stratigraphy of pre-Keweenawan rocks in parts Northern Michigan, U.S. Geological Survey Professional Paper 314-C, 24 p. James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., 1961, Geology of Central Dickinson County, Michigan, U.S. Geological Survey Professional Paper 310, 176 p. Schulz. K, J., Cannon, W.F., and Woodruff, L.G., 2018, Geochemistry of mafic rocks in Dickinson County, Michigan: Evidence for 2.1 Ga rifting: Institute on Lake Superior Geology, Proceedings of 64 th Annual meeting, Part 1: Program and Abstracts. p. 93-94. 23 �Reexamining the Osborne core for new insights into the age and petrology of the Northeast Iowa Intrusive Complex (NEIIC) CLARK, Ryan1, PEATE, David2, KUSICK, Alison2, HORKLEY, Kenny2, and ANDERSON, Raymond2 1 Iowa Geological Survey, University of Iowa, Iowa City, IA 52242 USA 2 Department of Earth and Environmental Sciences, University of Iowa, Iowa City, IA 52242 USA The Keweenawan Midcontinent Rift System (MRS) has been the focus of decades of research for its enigmatic geologic history and its wealth of economic minerals. The latter has been concentrated in the Lake Superior region where the MRS is exposed at or near the land surface. Copper-nickel sulfide and platinum group element deposits have been identified along the north shore in Ontario (Coldwell Complex) and along the western shore in Minnesota (Duluth Complex). These magmatic deposits are related to the MRS and are geophysically distinct, with high amplitude magnetic anomalies and associated gravity highs (Drenth et al., 2015 and Drenth & Brown, 2016). Since 2012, the U.S. Geological Survey (USGS) has conducted two major high-resolution geophysical surveys in northeastern Iowa and southeastern Minnesota in an attempt to better understand the nature of the Precambrian basement geology concealed beneath at least 1,000 feet (300 m) of Paleozoic sedimentary rocks. The surveys, both magnetic and gravity, have succeeded in refining the area previously identified as the Northeast Iowa Plutonic Complex (Anderson, 2006), now called the Northeast Iowa Intrusive Complex (NEIIC). The NEIIC has an aerial extent of over 6,000 mi2 (15,500 km), including several large ring/horseshoe shaped anomalies and associated linear features. Some of these features have been characterized using geophysical techniques, yet with a limited number of boreholes that reach the NEIIC, accurate lithologic and geochronologic data has remained elusive. One iron exploration core, the Osborne core, drilled in 1963 intersected a dike extending northeastward from the main part of the NEIIC and encountered more than 700 feet (213 m) of ultramafic olivine-plagioclase cumulate. The Iowa Geological Survey (IGS) and the University of Iowa Department of Earth and Environmental Sciences are reexamining the Osborne core to identify and characterize datable minerals. A systematic survey of compositional variations was done using a handheld portable XRay Fluorescence (pXRF) analyzer, with replicate analyses made on individual cores pieces, at an average sampling interval of 2 m along the core. These data show there are two distinct zones within the core that have elevated zirconium (Fig. 1), together with high K, P and Rb, indicative of trapped residual liquid. X-ray element mapping and backscatter images have identified baddeleyite and zirconolite minerals in these zones (Fig. 2). Samples have been selected and are being processed for geochronologic analyses. Obtaining a reliable age date from the Osborne core could provide a missing piece to the NEIIC puzzle and help answer the question of whether it is in fact related to the MRS and other economic mineral deposits in the Lake Superior region. 24 �Osborne Core pXRF Results Zr (ppm) 0 200 400 600 800 1000 1800 1900 2000 Depth (ft) 2100 2200 2300 2400 2500 2600 Figure 1: PXRF results for zirconium through the Precambrian sequence encountered in the Osborne core. Figure 2: Backscatter image of a zirconolite crystal from the Osborne core at 2,416' depth. References Anderson, R.R. 2006. Geology of the Precambrian surface of Iowa and surrounding area. Iowa Geological Survey, Open File Map OFM-06-7. Drenth, B.J., Anderson, R.R., Schulz, K.J., Feinberg, J.M., Chandler, V.W., and Cannon, W.F. 2015. What lies beneath: geophysical mapping of a concealed Precambrian intrusive complex along the IowaMinnesota border. Canadian Journal of Earth Science, v. 52, p. 1-15. Drenth, B.J., and Brown, P.J. 2016. Airborne magnetic total-field survey, Manchester region, Iowa, USA. U.S. Geological Survey data release, https://doi.org/10.5066/F7416V52. 25 �Keweenaw Fault System along Bête Grise Bay, Michigan: Geometry, Kinematics, and Tectonic Significance DEGRAFF, J.M. 1, TYRELL, C.W.1, HUBBELL, G.E. 1, and CARTER, B.T. 2 1 Michigan Technological University, Houghton, MI 49931 Structural Geology Consultant (now at Repsol) Houston, TX 77027 2 The Keweenaw Fault (KF) extends along the southern margin of the Midcontinent Rift System from northwest Wisconsin to near Keweenaw Point in Michigan. Reverse movement on the fault has thrust Portage Lake Volcanics (PLV, 1.1 Ga) over younger, mostly flat-lying Jacobsville Sandstone (JS) (Fig. 1), imparting a regional northerly tilt to PLV strata (1). The KF near Keweenaw Point is of interest because 1950s USGS maps (2-3) show five coastal areas with juxtaposed PLV and JS strata connected by an anomalously sinuous fault trace (Fig. 2a). Based on geophysical data, some have proposed that the KF continues offshore beyond Keweenaw Point in an arc curving over 90° to a southeasterly direction (4-5). These geometries seem incompatible with a simple thrust system. Furthermore, a lack of reported slip indicators prevents defining the ratio of dip to strike slip and estimating principal stress directions responsible for fault motion. New mapping of the KF system along Bête Grise Bay reveals that the oddly sinuous fault trace of the 1950s oversimplifies important geologic relationships in the area (Fig. 2a-b). Three to perhaps four of seven PLV-JS contacts previously mapped as faulted instead have an unconformity between PLV lava flows and basal JS strata. Unconformable contacts to the west show fractured, locally saprolitic PLV basalt below moderately dipping JS strata of alternating muddy siltstone, lithic to quartzose sandstone, and pebble conglomerate with angular basalt fragments in a muddy matrix. An unconformable contact to the east shows slightly deformed JS strata overlying steeply dipping, intensely faulted and brecciated PLV strata, indicating major slip on this KF segment before local JS deposition. At other shoreline locations, deformed JS strata truncated on fault contacts with PLV lavas provide evidence of a second period of slip on the KF system after some or all JS deposition. Recognition of unconformable PLV-JS contacts, combined with mapping both onshore and offshore, breaks the sinuous single fault trace of the 1950s into at least six segments generally striking ESE and forming a left-stepping, en echelon pattern. Well exposed fault surfaces near the shoreline have provided many opportunities to measure orientation of slip indicators and to infer slip sense. Analysis of such measurements at 36 sites indicates that the last period of activity on this part of the KF system was dominated by strike slip, with a 2:1 ratio of dextral strike slip to reverse dip slip (N side up). Geologic relationships across major fault segments are consistent with their north sides sliding to the right and upward relative to opposing sides. Inversion of fault-slip data further confirms a mostly strike-slip regime and indicates a maximum shortening direction of N80°W during the last period of fault motion. South of the Bare Hill rhyolite, a major ENE-trending fault appears to link two ESE-trending, en echelon fault segments (Fig. 2b). The linking fault follows the core of a tight upright anticline in PLV strata with an interlimb angle of 30° or less. PLV strata on the SE flank of the anticline dip steeply to moderately SE (counter-regional) for at least 3.5 km along the shore. Poles to bedding on both flanks of the anticline define a fold axis plunging 21° at N82°E. The tightly folded nature of the faulted anticline in relatively rigid strata implies that the fold formed during dextral strike slip on the linking fault or was modified afterward by such shearing. The trace of the KF system changes direction from NNE near Houghton to ESE at Bête Grise Bay (> 70⁰), which mimics the change in strike of PLV layers over the same distance (Fig. 1). Large crustalscale faults often curve and split into segments near their terminations. The new mapping results thus imply that the KF system terminates near the end of the peninsula in a series of fault splays, possibly transferring slip to other faults farther southeast. Based on these results and regional information, we suggest that slip on the KF system changes from mostly reverse dip-slip along its NNE-trending portion 26 �near Houghton to mostly dextral strike-slip near the tip of the Keweenaw Peninsula, and that total slip magnitude decreases over this same distance. Acknowledgements: We appreciate primary funding by the USGS EDMAP program, additional funding by the Keweenaw Community Forest Company, field and GIS support from D. Lizzadro-McPherson, and discussions with USGS geologists W. Cannon, K. Schulz, and L. Woodruff. Figure 1 (left): Keweenaw Peninsula where Portage Lake Volcanics are thrust over Jacobsville Sandstone. Black rectangle near tip of the peninsula marks area of Figure 2. (adapted from 1). Figure 2 (below): Study area along the Keweenaw Fault east from Bête Grise Bay. Main units: PLV mafic = greens; PLV felsic = reds; JS = pink-A / yellow-B. A) USGS maps from 1950s (2-3). B) New map highlighting fault pattern and PLV-JS unconformity. References 1. 2. 3. 4. 5. Cannon, W.F. and Nicholson, S.W., 2001, Geologic Map of the Keweenaw Peninsula and Adjacent Area, Michigan: United States Geological Survey, Map I-2696, Scale = 1:100,000. Cornwall, H. R., 1954, Bedrock Geology of the Lake Medora Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-52, scale 1:24,000. Cornwall, H.R., 1955, Bedrock Geology of the Fort Wilkins Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-74, scale 1:24,000. Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332. Hinze, W.J., Allen, D.J., Braile, L.W., and Mariano, J., 1997, The Midcontinent Rift System: a major Proterozoic continental rift: in Ojakangas, R.W., Dickas, A.B., and Green, J.C. (eds.), Middle Proterozoic to Cambrian Rifting, Central North America: Boulder, Colorado, Geological Society of America Special Paper 312, p. 7-35. 27 �Southwest Margin of the Midcontinent Rift System in Eastern Lake Superior: Review and Preliminary Interpretation DEGRAFF, J.M.1 and DEGRAFF, I.S.2 1 Michigan Technological University, Houghton, MI 49931 Geologic Consultant, Houston, TX 77042 2 The relatively well-defined southwest branch of the Midcontinent Rift System (MRS) transitions to the less well-defined southeast branch near Keweenaw Point and the postulated Thiel Fault zone (1-3; Fig. 1). The southwest branch has abundant outcrops of rift-related rocks around Lake Superior and has been extended farther southwest beneath Paleozoic strata with geophysics and widely spaced deep drill holes. In contrast, rift-related rocks of the southeast branch mostly lie beneath lakes Superior and Michigan, post-rift Jacobsville Sandstone (JS), or Phanerozoic strata of the Michigan Basin. Current understanding of the southeast branch of the MRS mostly comes from geophysical data and rare deep boreholes that penetrate Precambrian basement. To better define the transition between the two branches of the MRS, we initiated research on the offshore geology between the Keweenaw Peninsula and the south shore of Lake Superior east of Marquette, Michigan (Fig. 2). Data to be obtained and used include new shoreline and underwater outcrop descriptions and existing seismic and potential field data. This poster provides some initial perceptions and thoughts in the context of prior investigations. Outcrops of rift-related rocks in the area are mainly along the Keweenaw Peninsula and Manitou Island, but probably occur at Stannard Rock shoal and perhaps at other shallow areas along an arc from Keweenaw Point southeastward to Munising (Fig. 2). Information about Stannard Rock is limited to sketchy early reports and one rock sample described as “quartzless porphyry” or rhyolite (4-5). If Stannard Rock shoal proves to have rhyolitic rocks, this outcome together with the rhyolitic flows at the bottom of the Amoco St. Amour 1-29R borehole (6) could indicate a significant area of felsic volcanism along the southwest margin of the MRS in eastern Lake Superior. Other relevant outcrops in the area are Jacobsville strata that rim Keweenaw Bay and extend eastward along the south shore of Lake Superior to Sault Ste. Marie. Along the shore near Munising, aerial imagery available through Google Earth shows JS strata on the rift margin generally striking NS and dipping eastward toward the rift axis defined by geophysical data. Qualitative review of available geophysical data provides additional insight into the nature of the rift margin north of Munising and structural trends in pre-rift basement between the two MRS branches. Five seismic reflection lines in eastern Lake Superior define: (1) an uplifted rift flank to the southwest with sub-horizontal strata, and (2) a rift margin-slope with strata dipping moderately northeast toward the rift basin. Some lines show evidence of a component of reverse faulting along the rift margin (7), but others may be interpreted as having only a flexure without obvious faulting. The NNW-trending rift margin has two jogs, a southern one near Munising and a northern one near Stannard Rock, implying that rift-margin faults are not continuous along the entire margin. The orientation of such faults is more than 90° off trend of the Keweenaw Fault, and so their slip direction must differ from that of the Keweenaw Fault. Therefore, faults along this rift margin are best regarded as distinct from the Keweenaw Fault and should have different names (e.g., Munising, Au Train). Broad trends in potential field data are consistent with rift margin trends interpreted on seismic data, including the two jogs. In addition, aeromagnetic data define several circular to arcuate anomalies up to 16 km in diameter that cluster on the rift flank near the jogs in the rift margin (Fig. 2). Each cluster of circular anomalies appears to lie along ENE-trending zones that 28 �may represent crustal-scale fracture zones. It is possible that these anomalies are caused by eruptive centers that sourced volcanic rocks in their immediate surroundings. Further work is required to test these ideas and to improve understanding of this less studied sector of the MRS. Acknowledgements: We appreciate the helpful and encouraging comments by Bill Hinze (Purdue University) during a review of this abstract. Figure 1 (left): Major rock units and faults in the Lake Superior area; KF-Keweenaw Fault, DF-Douglas Fault, IRF-Isle Royale Fault, TF-Thiel Fault (1). Inset map shows extent of Midcontinent Rift System (MRS) from Lake Superior southwest to Kansas (K) and southeast to Detroit (D). Black rectangle is area of Figure 2. Figure 2 (below): Main structural elements between the Keweenaw Peninsula and Munising. H-Houghton, Ma-Marquette, MuMunising, MI-Manitou Island, SR-Stannard Rock; KF-Keweenaw Fault, TF-Thiel Fault, F-unnamed rift-margin faults. Dark red “faults” inferred from geophysical data. Purple and blue features interpreted from aeromagnetic data and explained in poster. References 1. Miller, Jr., J.D., 2007, The Midcontinent Rift in the Lake Superior region: a 1.1 Ga Large Igneous Province: IAVCEI Large Igneous Provinces Commission, p. 1-18. 2. Hinze, W.J., Allen, D.J., Braile, L.W., and Mariano, J., 1997, The Midcontinent Rift System: a major Proterozoic continental rift: in Ojakangas, R.W., Dickas, A.B., and Green, J.C. (eds.), Middle Proterozoic to Cambrian Rifting, Central North America: Boulder, Colorado, Geological Society of America Special Paper 312, p. 7-35. 3. Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R., and Spencer, C., 1989, The North American Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332. 4. Irving, R.D., 1883, The copper-bearing rocks of Lake Superior: U.S.G.S. Mono., v. 5, p. 360-361. 5. Hubbard, L.L., 1898, Keweenaw Point with particular reference to the felsites and their associated rocks: Geol. Survey Michigan, v. 6, part 2, 155 p. 6. Ojakangas, R.W. and Dickas, A.B., 2002, The 1.1-Ga Midcontinent Rift System, central North America: sedimentology of two deep boreholes, Lake Superior region, Sediment. Geol., v. 147(1-2), pp. 13-36. 7. Mariano, J. and Hinze, W.J., 1994, Structural interpretation of the Midcontinent Rift in eastern Lake Superior from seismic reflection and potential-field studies: Canadian Journal of Earth Sciences, v. 30, p. 619-628. 29 �Neutralization of proton acidity with sequestration of atmospheric CO2 during experimental weathering of intrusive rocks from the Midcontinent Rift System DIEDRICH1, Tamara, DAY2, Stephen 1 MineraLogic LLC, 306 W. Superior St., Alworth Building, Suite 408, Duluth, MN 55802 USA 2 SRK Consulting (Canada) Inc., 1066 West Hastings St., Vancouver, BC, V6E 3XS Canada Intrusive rocks associated with the Mesaba Deposit1, contained predominantly in the Bathtub Intrusion of the Duluth Complex2, have been the subject of a comprehensive geochemical characterization program, initiated in 2010, to inform plans for managing water and waste rock on any potential future mining project. This program includes multiple experimental components to characterize subaerial weathering reactions, rates, and products; including, but not limited to, laboratory testwork using an ASTM standard method under the “humidity cell test” configuration, laboratory testing on columns of rock, and a field-based, larger scale barrel test program. Duluth Complex rocks tend to contain abundant olivine and/or plagioclase, both of which are relatively reactive acid-neutralizing and, potentially, carbonate-forming silicate minerals. Experimental weathering outcomes confirm the effectiveness of silicate dissolution in neutralizing proton acidity through three distinct mechanisms: 1) consumption of protons as reactants in silicate mineral dissolution reactions; 2) reaction with dissolved alkalinity formed during dissolution of silicate minerals in the presence of atmospheric CO2; and, 3) as reactants during dissolution of secondary carbonate minerals, which were precipitated as weathering products of primary silicate phases. Furthermore, as suggested by the latter two of the above numerated mechanisms, silicate weathering reactions in the presence of atmospheric CO2, represents a well-established net sink for atmospheric CO2 in the form of carbonate mineral weathering products. Weathering reactions for relatively reactive silicate minerals that are abundant in Duluth Complex rock include those shown below for An50 and olivine, respectively: Na0.5Ca0.5Al1.5Si2.5O8(s) + 1.5 H+ + 6.5 H2O ↔ 0.5 Ca2+ + 0.5 Na+ + 1.5 Al(OH)3 + 2.5 H4SiO2 (Mg,Fe)2SiO4(s) + 4 H+ ↔ 2 (Mg2+, Fe2+) + H4SiO4 Subsequent oxidation and hydrolysis of iron from the olivine breakdown reaction releases hydrogen through the following reaction: Fe2+ + 1/4O2 + 5/2H2O ↔ Fe(OH)3 + 2H+ Every cationic charge unit3 added to solution corresponds to a proton being removed. 1 The Mesaba Deposit is a magmatic copper-nickel-PGM deposit described by >800,000 feet of diamond drilling that is owned by Teck American Inc. a wholly owned subsidiary of Teck Resources Limited. 2 Severson, M J, Hauck, S A, 2008. Finish Logging of Duluth Complex Drill Core (And a Reinterpretation of the Geology at the Mesaba (Babbitt) deposit). Natural Resources Research Institute. 3 Charge unit concentration is equal to molar concentration times charge. Release of Fe2+ during dissolution consumes protons, which are re-released upon oxidation and hydrolysis of iron. Therefore, release of iron is overall proton-neutral and not included. 30 �The relationship between molar concentrations of cations and sulfate in weathering test leachate is a robust indicator of leachate pH across all experimental configurations. Figure 1 shows data from 3,053 individual leachate samples from over 40 different tests. The y-axis represents the relative rates of proton consumption and production during weathering, as indicated by the “charge unit balance” (defined in the figure) of the leachate sample. When the composition of the leachate indicates that protons are being consumed by silicate dissolution faster than they are being produced during sulfide mineral oxidation, the leachate pH is higher than the blank; i.e., there is a net decrease in proton concentration. Conversely, pH of the leachate becomes acidic when the composition of the leachate indicates that protons are being released faster than they are consumed. The clear relationship between rates of proton production and consumption, and drainage pH is an indication of the effectiveness of silicate mineral dissolution in neutralization of proton acidity in the weathering tests. Figure 1. Leachate data from weathering tests (n=3053) showing relationship between charge unit balance and pH. “Charge unit balance”, defined as the molar ratio of cationic charge unit concentration to sulfate charge unit concentration (oxidation of one mol of sulfur in pyrrhotite to sulfate releases two protons). When charge unit balance is equal to one (shown as dashed line), the rate of consumption and production of protons during weathering is equal. Dotted line shows lowest pH observed in leachate from blank tests. In addition to sulfide mineral oxidation, dissolution of atmospheric CO2 into rainwater can provide protons for silicate dissolution, through equilibria between dissolved CO2 and carbonic acid (H2CO3), and the subsequent dissociation of carbonic acid to bicarbonate alkalinity and protons. Therefore, in the presence of CO2, consumption of protons during silicate dissolution would continue to drive this reaction toward the reaction products, resulting in accumulation of bicarbonate alkalinity in associated waters. Under select conditions, this carbonate builds up and eventually reacts with the calcium and magnesium released during silicate dissolution to precipitate secondary carbonate minerals. While secondary carbonate minerals have not, yet, been directly detected as experimental products, leachate chemistry suggests that, as the ratio of the rock to water increases for different experimental configurations, the leachate becomes more concentrated in calcium, magnesium, and bicarbonate alkalinity, until, eventually, calcium/magnesium ratios decrease, as calcium carbonate is presumably being preferentially precipitated out of solution. 31 �Morphology, mineralogy, texture, and genesis of peperite, Fivemile Lake, Vermilion District, Minnesota: Comparison with Pleistocene peperite, Iceland. DRAZAN, Jacqueline L.1, HUDAK, George2, MOOERS, Howard1 1 Department of Earth and Environmental Sciences, University of Minnesota - Duluth, 1114 Kirby Dr., 229 Heller Hall, Duluth, MN, 55812; 2Natural Resources Research Institute, 5013 Miller Trunk Hwy, Hermantown, MN, 55811. Peperites are defined as a “rock formed essentially in situ by disintegration of magma intruding and mingling with unconsolidated or poorly consolidated, typically wet sediments” (White et al., 2000, p. 65). Pillowed dikes and associated peperite is well exposed at Fivemile Lake in the Vermilion District of northeastern Minnesota (Hudak et al., 2002; Hudak et al., 2003; Hudak et al., 2004). The rocks are Neoarchean in age (~2.7 billion years, Peterson et al., 2001), and contain well-preserved and well-studied volcanic facies (e.g. Hudak et al., 2002, 2003, 2004). The sequence at Fivemile Lake has been interpreted as recording a series of northeasttrending mafic dikes which have intruded wet volcaniclastic sediments to produce peperite deposits at different levels within the seafloor in a relatively shallow (<1500 m) submarine volcanic system (Hudak et al., 2004). In the current study, outcrops were mapped at a scale of 1:39 with field work focused on extremely detailed mapping to evaluate peperite deposit morphology, mineralogy, and textures. The igneous component is pillowed to massive, dominated by amygdules, and grades into the host sediment (Fig. 1, right). Outside the margins of the pillowed dikes, both globular and blocky peperite comprising isolated igneous clasts floating in the host volcaniclastic sedimentary rock are present. The volcaniclastic sedimentary rocks are moderately- to highly-vesiculated, and locally contain 1cm wide highly vesiculated zones that are parallel- to sub-parallel to igneous component fragments within the peperite deposits. Although textures are well-preserved both at the outcrop and thin-section scales at Fivemile Lake, the original minerals/glass have been postdepositionally altered to quartz, epidote, and carbonate minerals. As an analog for the Archean Fivemile Lake peperite, Pleistocene peperites from three locations in Iceland were described, sampled, and analyzed. Locations include three sites in móberg, two near Sveifluháls, Iceland, and a site at Reynisfyara Beach near Vik, Iceland. Kagy (2011) identified peperites along pillowed dyke margins in móberg near Sveifluháls, Reykjanes Peninsula, Iceland. Both blocky and fluidal types were described along with anomalous igneous clasts in a host rock of hydrothermally altered lapilli tuff (palagonite formation). The rocks are relatively unaltered and contain abundant glass with some alteration to palagonite. Host sediment is highly vesiculated and glassy, with broken, jagged hyaloclastite fragments making up the matrix (Fig. 1, left). Some fragments have phenocrysts, while other fragments are separated by a junky, opaque matrix, likely a result of surficial weathering at the outcrop. Evaluation of the peperite at Fivemile Lake with comparison to Pleistocene peperites aids in identification of primary peperite textures and morphologies. Documentation and mapping of peperites is useful in determining and understanding magma-water interactions and hydrovolcanic processes like magma explosions in wet sediment. Formation of peperites at Fivemile Lake is spatially associated with synvolcanic faults and occurred near the paleo- 32 �seafloor, which is a prospective geologic setting for volcanogenic massive sulfide deposits (Gibson et al., 1999; Rosa et al., 2016). Figure 1: Field images on left from Iceland (Site 2L; Kagy, 2011) indicate pillowed dyke with peperite next to host sediment. On right, field image shows pillow lava intruding and budding off into host sediment on Peperite Point. References Gibson, H. L., Morton, R. L., and Hudak, G. J., 1999. Submarine volcanic processes, deposits, and environments favorable for the location of volcanic-associated massive sulfide deposits: Reviews in Economic Geology, v. 8, p. 13-48. Hudak, G. J., Newkirk, T. T., Odette, J., and Hauck, S., 2002. Comparative Geology, Stratigraphy, and Lithogeochemistry of the Fivemile Lake, Quartz Hill, and Skeleton Lake VMS Occurrences, Vermilion District, NE Minnesota: Natural Resources Research Institute Technical Report NRRI/TR-2002/03, 390 p. Hudak, G. J., Newkirk, T. T., Odette, J., and Hauck, S., 2003. Comparative Geology, Stratigraphy, and Lithogeochemistry of the Fivemile Lake, Quartz Hill, and Skeleton Lake VMS Occurrences, Vermilion District, NE Minnesota: Natural Resources Research Institute Report of Investigation NRRI/RI-2003/18, 390 p. Hudak, G. J., Newkirk, T. T., Drexler, H., Odette, J. D., and Hocker, S. M., 2004. Neoarchean Peperites in the Vicinity of Fivemile Lake, Vermilion District, NE Minnesota: Institute on Lake Superior Geology, V. 50, Part 1- Proceedings and Abstracts, p. 84-85Kagy, H.M. 2011. Interaction Of Basaltic Dikes And Wet Lapilli Tuff At Glaciovolcanic Centers: A Case Study Of Sveifluháls, Iceland As A Terrestrial Analog For Dike-cryosphere Interaction On Mars, Master’s thesis. University of Pittsburgh, Department of Geology and Planetary Science. Mercurio, E. C. 2011. Processes, Products and Depositional Environments of Ice-Confined Basaltic Fissure Eruptions: A Case Study of the Sveifluháls Volcanic Complex, SW Iceland, Ph.D. dissertation, University of Pittsburgh, Department of Geology and Planetary Science. Peterson, D.M., Gallup, C., Jirsa, M.A., and Davis, D.W. 2001. Development of Archean lode-gold and massive sulfide deposit exploration models using geographic information system applications: targeting mineral exploration in northeastern Minnesota from analysis of analog Canadian Mining camps: unpublished Ph. D. dissertation, University of Minnesota, Duluth, Minnesota, 503 p. Rosa, C.J.P., McPhie, J., Relvas, J.M.R.S. 2016. Distinguishing peperite from other sediment-matrix igneous breccias: Lessons from the Iberian Pyrite Belt. Journal of Volcanology and Geothermal Research: 315, p. 28-39. White, J.D.L., McPhie. J., Skilling, L. 2000. Peperite: a useful genetic term. Bulletin of Volcanology: 2, p. 65-66. 33 �The Dickinson Group in the Central Upper Peninsula of Michigan: Part 2 - Geophysical expression and a preliminary interpretation of its eastward extent under Paleozoic cover DRENTH, Benjamin J.1, CANNON, William F.2, and SCHULZ, Klaus J.2 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver Federal Center, Denver, CO, 80225 2 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA, 20192 The Dickinson Group crops out in central Dickinson County, Michigan, and includes three formations (described in detail by James et al., 1961) that may contain a unique metasedimentary and volcanic record of the final breakup of the Superior and Wyoming cratons (Cannon et al., this volume). The basal East Branch Arkose is made up of arkose, conglomerate, and basalt flows. The overlying Solberg Schist consists of finer clastic rocks, metavolcanics rocks, and an iron-formation, the Skunk Creek Member. The uppermost member, the Six Mile Lake Amphibolite, consists of mafic metavolcanic rocks. The contact between the East Branch Arkose and Solberg Schist is gradational. The contact between the Solberg Schist and Six Mile Lake Amphibolite is not exposed, but was interpreted to be conformable (James et al., 1961). Where exposed west of the edge of Paleozoic cover, the Dickinson Group forms a nearly vertical, south-facing monocline extending more than 20 km with consistent east-west strike (Fig. 1). The Dickinson Group was originally interpreted as Archean, based on an apparent gradational contact between the Six Mile Lake Amphibolite and Archean granite to the south (James et al. 1961). However, various lines of evidence establish an apparent age range of ~2.1 to 1.83 Ga for the entire Dickinson Group (Holm et al., 2007; Craddock et al., 2013; Ayuso et al., 2018; Schulz et al., 2018; Cannon et al., 2018b; Cannon et al., this volume). Cleary, the nature of the Six Mile Lake Amphibolite-Archean contact (queried on Fig. 1) is critical to the interpretation of the Dickinson Group and merits further study. Parts of the Dickinson Group have geophysically distinctive features compared to surrounding Precambrian rocks. The high-density Six Mile Lake Amphibolite is the dominant source of an east-west elongated, ~13 mGal gravity high that extends 10s of km over both the area of exposure and Paleozoic cover to the east (Drenth et al., 2018). The ~2.1 Ga (Ayuso et al., 2018) “porphyritic red granite” (prg, Fig. 1), a probable source of detritus for sedimentary parts of the Dickinson Group, produces a ~4 mGal gravity low and a zone of mostly quiet aeromagnetic anomalies. Geophysical data show that it is a larger body than shown by previous mapping. Numerous narrow, strike-parallel elongated aeromagnetic highs lie over all units of the Dickinson Group, including the following examples. Aeromagnetic highs with amplitudes up to 600 nT lie over the East Branch Arkose, interpreted to reflect interbedded basalt flows (James et al., 1961). The Skunk Creek Member iron-formation of the Solberg Schist produces an aeromagnetic high with a maximum amplitude of 2000 nT, distinguishing it from other anomaly sources in the area. Other aeromagnetic highs with amplitudes <500 nT do not have confirmed sources, but have been generally ascribed to diabase dikes, gabbroic intrusions, and other magnetic layers within the Dickinson Group (James et al. 1961). A preliminary interpretation of the eastward subcrop extension (under Paleozoic cover) of the Dickinson Group (Fig. 1) is based on 3D inverse gravity modeling of the geometry of the Six Mile Lake Amphibolite, tracing the distinctive aeromagnetic signature of the Skunk Creek Member, and following the strikes of other aeromagnetic anomalies. The volume of the Six Mile Lake Amphibolite is interpreted to increase dramatically to the east of where it is exposed, and the Skunk Creek Member is interpreted to be complexly folded east of the Paleozoic contact, in 34 �contrast to the monoclinal structure to the west. At least two, and perhaps three folds are indicated by aeromagnetic patterns. Collectively, these interpretive observations may be best reconciled by a model that involves complexly faulted and folded Solberg Schist and Six Mile Lake Amphibolite, including a possible thrust sheet (Fig. 1). The broader tectonic significance of this model hinges on the true nature of the Six Mile Lake Amphibolite-Archean contact. Figure 1: Preliminary interpretation of the full extent of the Dickinson Group, modified from James et al. (1961), Craddock et al. (2013), Cannon et al. (2018a,b), and Cannon et al. (this volume). References Ayuso, R. A., Schulz, K. J., Cannon, W. F., Woodruff, L. G., Vasquez, J. A., Foley, N. K., and Jackson, J., 2018, New U-Pb zircon ages for rocks from the granite-gneiss terrane in northern Michigan: evidence for events at ~3750, 2750, and 1850 Ma: Institute on Lake Superior Geology 64th Annual Meeting Proceedings, Part 1: Program and Abstracts, p. 7-8. Cannon, W. F., Schulte, R., and Bickerstaff, D., 2018a, Exposed Precambrian bedrock in part of Dickinson County, Michigan, and Marinette and Florence Counties, Wisconsin: U.S. Geological Survey data release: https://www.sciencebase.gov/catalog/item/59a5b942e4b075bb795913e1. Cannon, W. F., Schulz, K. J., Ayuso, R. A., and Mroz, T. H., 2018b, Field Trip 1: Archean and Paleoproterozoic geology of the Felch District, Central Dickinson County, Michigan, in Cannon, W. F., ed., Institute on Lake Superior Geology 64th Annual Meeting Proceedings Volume 2: Field Trip Guidebooks, p. 1-38. Cannon, W.F., Schulz, K.J., Drenth, B.J., this volume, The Dickinson Group of Dickinson County, Michigan: Part 1- age and tectonic setting based on new geophysical, geochemical, and geochronologic data: Institute on Lake Superior Geology, Proceedings of 65 th Annual Meeting, Part 1: Program and Abstracts. Craddock, J.P., Rainbird, R.H., Davis, W.J., Davidson, C., Vervoort, J.D., Konstantinou, A., Boerboom, T., Vorhies, S., Kerber, L., and Lundquist, B., 2013, Detrital zircon geochronology and provenance of the Paleoproterozoic Huron (~2.4-2.2 Ga) and Animikie (~2.2-1.8 Ga basins, southern Superior Province, Journal of Geology, v. 121, p. 623-644. Drenth, B.J., Woodruff, L.G., Schulz, K.J., Cannon, W.F., and Ayuso, R.A., 2018, On the source(s) of the FelchArnold gravity anomaly, Upper Peninsula, Michigan: Institute on Lake Superior Geology, Proceedings of 64 th Annual Meeting, Part 1: Program and Abstracts, p. 27-28. Holm, D. K., et al. (2007). "Reinterpretation of Paleoproterozoic accretionary boundaries of the north-central United States based on a new aeromagnetic-geologic compilation." Precambrian Research, v. 157, p. 71-79. James, H. L., Clark, L. D., Lamey, C. A., and Pettijohn, F. J., 1961, Geology of Central Dickinson County, Michigan: U.S. Geological Survey Professional Paper 310, 176 p. Schulz, K.J., Cannon, W.F., and Woodruff, L.G., 2018, Geochemistry of mafic rocks in Dickinson County, Michigan: evidence for ~2.1 Ga rifting: Institute on Lake Superior Geology, Proceedings of 64 th Annual Meeting, Part 1: Program and Abstracts, p. 93-94. 35 �High-resolution aeromagnetic survey, central Upper Peninsula, Michigan DRENTH, Benjamin J.1, CANNON, William F.2, and SCHULZ, Klaus J.2 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver Federal Center, Denver, CO, 80225 2 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA, 20192 We present a new aeromagnetic dataset from a high-resolution (150 m line spacing, 80 m nominal terrain clearance) regional fixed-wing survey (~37,000 line km) flown over portions of the central Upper Peninsula of Michigan in 2018. The survey footprint includes areas with Precambrian bedrock between Marquette and Iron Mountain and extends eastward over a large area with weakly magnetized Paleozoic sedimentary cover (Fig. 1), which will allow interpretation of Precambrian subcrop. Archean rocks of the gneiss terrane south of the Great Lakes Tectonic Zone (GLTZ), a Neoarchean suture, are generally weakly magnetized. A swarm of north-northeast trending magnetic dikes are imaged cutting the gneiss terrane between the Bush Lake fault and the GLTZ. These dikes are not detected north of the GLTZ, indicating the swarm predates the suture. Magnetic highs north of the GLTZ lie over exposures of the Archean greenstone-granite terrane and trend subparallel to the GLTZ trend. Metasedimentary rocks of the Paleoproterozoic Chocolay and Baraga Groups are generally weakly magnetized. Iron formations within the Menominee Group (i.e., the Vulcan Iron-formation) produce very large amplitude positive anomalies. Anomaly amplitudes in the Felch and Calumet troughs reach ~15,000 nT. Several other very large amplitude anomalies (up to ~35,000 nT) lie over the Paleozoic sedimentary cover to the east and are produced by very strongly magnetized iron formations in the Precambrian subcrop that have been drilled by the private sector (Waggoner, 2007). The Dickinson Group, once thought to be Archean (James et al. 1961) but now considered to be at least partly Paleoproterozoic (e.g., Cannon et al., 2018), is characterized by numerous east-west elongated, narrow magnetic highs. Some of these highs have been interpreted to reflect mafic volcanic rocks and an iron formation, but the sources of others are not explicitly known (James et al., 1961). Multiple generations of likely Proterozoic dikes are expressed in the aeromagnetic data. Numerous reversely polarized dikes interpreted to be Keweenawan (i.e., related to the ~1.1 Ga Midcontinent Rift System) trend east-northeast. Normally polarized dikes that are also likely Keweenawan trend west-northwest. A swarm of northwest-trending dikes of unknown age trends subparallel to the GLTZ. References Cannon, W.F., and Ottke, D., 1999, Preliminary digital geologic map of the Penokean (early Proterozoic) continental margin in northern Michigan and Wisconsin: U.S. Geological Survey Open-File Report 99-547: http://pubs.usgs.gov/of/1999/of99-547/. Cannon, W. F., Schulz, K. J., Ayuso, R. A., and Mroz, T. H., 2018, Field Trip 1: Archean and Paleoproterozoic geology of the Felch District, Central Dickinson County, Michigan, in Cannon, W. F., ed., Institute on Lake Superior Geology 64th Annual Meeting Proceedings Volume 2: Field Trip Guidebooks, p. 1-38. James, H. L., Clark, L. D., Lamey, C. A., and Pettijohn, F. J., 1961, Geology of Central Dickinson County, Michigan: U.S. Geological Survey Professional Paper 310, 176 p. Waggoner, T. D., 2007, Definition of the Proterozoic terrain under the Paleozoic -- central U.P., Michigan: Institute on Lake Superior Geology 53rd Annual Meeting, p. 85-86. 36 �Figure 1: Simplified bedrock geology of the aeromagnetic survey region, modified from Cannon and Ottke (1999) and Cannon et al. (2018). 37 �What do detrital zircon studies of the Huronian Supergroup tell us? an analysis of all published data EASTON, Robert Michael1 1 Adjunct Professor, Department of Earth Sciences, Carleton University, Ottawa, Ontario Since the publication of the first detrital zircon analyses from the Huronian Supergroup in 2006 (Rainbird and Davis 2006), detrital zircon work has been completed on more than 25 samples of the supergroup, from almost every unit (except for the Pecors, Espanola and Bruce formations) (Craddock et al. 2013; Davis et al. 2018; Easton and Heaman 2008, 2011; Hill et al. 2018; Kenny et al. 2018; Long et al. 2011; Ménard 2017; Petrus et al. 2016; Rasmussen et al. 2013). Most of this work occurred in the area between Sudbury and Sault Ste. Marie, all north of the Murray fault, with only 2 samples studied so far from the Cobalt basin northwest of Sudbury. These data are summarized in Table 1, with age ranges and averages based on grains that are < 5% discordant, a lower cutoff than used in most studies. Key observations are: • Zircons between circa 2450 and 2490 Ma, likely derived from either Huronian Supergroup volcanic rocks and/or related mafic and felsic intrusions, so far have been reported only from the Matinenda or the Mississagi formations, generally from sample sites near the base of the supergroup. • Samples from the lower Huronian Sgp (Elliot Lk and Hough Lk groups) are dominated by Geon 26 detritus, consistent with provenance dominated by local sources characteristic of the RamsayAlgoma granitoid complex. Where detailed stratigraphic sampling has occurred, the lowermost units have unimodal populations, becoming more diverse with increasing stratigraphic height (e.g., Easton and Heaman 2011). The only exceptions are the 2 samples from the Cobalt basin, which are dominated by Geon 27 populations, consistent with more >2.7Ga basement in that area. • Above the Mississagi Formation, Geon 27 populations are dominant, but Geon 28, 29 and Geon 30 grains are also commonplace. This may reflect a change in sedimentation style, and/or increased erosion of the hinterland resulting in a wider range of source material becoming available. • The uppermost Huronian Sgp units have ages of circa 2310 Ma (Hill et al. 2018; Rasmussen et al. 2013), meaning deposition of the entire supergroup took place between circa 2460 to 2310 Ma. • Persistent throughout the sequence are occasional Geon 25 grains, typically with ages of 2550-2590; these grains become somewhat more abundant in the upper two groups. These grains have no known local source, and as suggested by Bleeker (pers. comm. 2019). may have a source region to the south, such as the Kaapvall craton, that was subsequently rifted away from North America. • Currently it is not possible to determine if the detrital zircon populations differ between glaciogenic (e.g., Ramsay Lake, Gowganda) units and the non-glaciogenic sandstone units. • Grains >3.0 Ga occur sporadically throughout the supergroup, mainly in the Matinenda and Mississagi formations, and could be sourced locally from Michigan (see Ayuso et al. 2017). More difficult to explain is the population of 29 ancient grains, 3.0-3.6 Ga, in the Gowganda Formation sample from Cobalt. Is this sourced locally in the Cobalt area, or have these grains been transported from sources currently exposed on the northeast shore of Hudson’s Bay? It is unclear if the sampled unit is glaciogenic or not, as the sampled rock type was not specified by Kenny et al. (2017). References Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A. and Jackson, J. 2017. Evidence for the presence of Eoarchean crust in northern Michigan; in 63rd Institute on Lake Superior Geology Annual Meeting, Wawa, ON, Proceedings v.63, pt.1, .9-10. Craddock, J.P., Rainbird, R.H., Davis, W.J., Davidson, C., Vervoort, J.D., Konstantinou, A., Boerboom, T., Vorhies, S., Kerber, L., and Lundquist, B. 2013. Detrital zircon geochronology and provenance of the Paleoproterozoic Huron (∼2.4–2.2 Ga) and Animikie (∼2.2– 1.8 Ga) Basins, southern Superior Province; Journal of Geology, v.121, 623-644. Davis, D.W., Ménard, J. and Sutcliffe, C.N. 2018. U-Pb geochronology by LA-ICP-MS in samples from northern Ontario; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 94p. Easton, R.M. and Heaman, L.M. 2008. Detrital zircon geochronology of Huronian Supergroup sandstones located within the Vernon structure, north of Espanola, Ontario; 54th Institute on Lake Superior Geology, Proceedings, v.54, pt.1, 21-22. 38 �Table 1. Summary of data for all Huronian Supergroup samples based on grains ≤ 5% discordant, in most studies many more grains were analyzed. For samples with significant discordance, the lower numbers shown are for grains ≤ 10% discordant. Also indicated are grains per Geon. All samples are sandstones unless otherwise noted. Samples from he Cobalt Basin are in italics. Abbreviations: cong, conglomerate; EL, Elliot Lake area; MCB, main conglomerate bed; S, Sudbury area; TH, Thessalon area. Formation Bar River mudstone Bar River EL Gordon Lake EL Gordon Lake EL Lorrain EL Gowganda Number n=16 n=62 n=57 n=30 n=172 Range (Ma) 2279-2745 2523-3074 2284-2840 3 sites 2684-2890 2520-3614 Serpent EL Serpent EL-S n=46 n=10 n=19 n=63 n=22 n=130 n=117 n=72 n=65 n=25 n=37 n=36 n=210 n=39 n=27 n=30 n=47 n=36 n=5 n=15 n=28 2549-3576 2531-3317 2531-3317 2443-3617 2591-2832 2388-3286 2414-2978 2544-2949 2656-2887 2526-2719 2607-2821 2533-2752 2366-2906 2505-3774 2451-2714 2650-2742 2620-2897 2617-2776 2634-2651 2621-2684 2546-2838 Main Peak (Ma) 2344 2706 (27>>26) 2317, 2702 (26≈27) 2308, 2308, 2311 2713 (27>26) 2705, 2857, 2965, 3076, 3316 (27>26) 2719 (27>>26) 2688 (5%) 2688 (10%) 2466, 2692 (26>27) 2663 (26>>27) 2477, 2697 (26≈27) 2490, 2560, 2689 2683 (26>>27) 2697 (26>27) 2659 (26>>27) 2677 (26>>27) 2670 (26>>27) 2459, 2703, 2771 2557, 2661 (26>>27) 2457, 2671 (26>>27) 2680 (26>>>27) 2664 (26>>>27) 2649 (26>>>27) 2641 (5%) 2643 (10%) 2641 (26>>>27) n=37 2507-2890 2698 (26≈27) Mississagi EL Mississagi EL-S Mississagi (upper) S Mississagi S Mississagi S Ramsay Lake EL-S Ramsay Lake S Ramsay Lake S cong McKim S Mississagi cong Matinenda S Matinenda EL-S Matinenda EL Matinenda (upper) EL Matinenda EL Matinenda above MCB EL Matinenda below MCB EL Livingstone Creek TH 24 2 4 2 3 25 5 5 3 26 1 18 20 27 2 30 22 28 29 >3.0 2 1 4 3 2 6 34 18 51 6 38 18 29 22 4 8 13 1 57 39 22 28 3 8 7 122 5 4 5 3 3 11 1 3 1 7 9 3 9 4 8 24 18 57 47 44 35 19 27 26 78 22 20 25 47 33 5 15 24 1 17 16 1 1 2 5 2 2 10 4 3 1 3 3 11 3 4 1 9 11 2 2 2 7 2 4 1 1 1 1 3 Easton, R.M. and Heaman, L.M. 2011. Detrital zircon geochronology of Matinenda Formation sandstones (Huronian Supergroup) at Elliot Lake, Ontario: Implications for uranium mineralization; 57th Institute on Lake Superior Geology, Proceedings, v.57, pt.1, 31-32. Hill, C.M., Davis, D.W. and Corcoran, P.L. 2018. New U-Pb geochronology evidence for 2.3 Ga detrital zircon grains in the youngest Huronian Supergroup formations, Canada; Precambrian Research, v.314, 428-433. Kenny, C.G., Petrus, J.A., Whitehouse, M.J., Daly, J.S., and Kamber, B.S. 2017. Hf isotope evidence for effective melt homogenisation at the Sudbury impact crater, Ontario, Canada; Geochimica et Cosmochimica Acta, v.215, 317-336. Long, D.G.F., Ulrich, T. and Kamber, B.S. 2011. Laterally extensive modified placer gold deposits in the Paleoproterozoic Mississagi Formation, Clement and Pardo Townships, Ontario; Canadian Journal of Earth Sciences, v.48, 779-792. Ménard. J.A. 2017. Sedimentary provenance of the Elliot Lake and Hough Lake groups, Huronian Supergroup, Sudbury area; in Summary of Field Work and Other Activities, 2017; Ontario Geological Survey, Open File Report 6333, 17-1 to 17-7. Petrus, J.A., Kenny, G.G., Ayer, J.A., Lightfoot, P.C. and Kamber, B.S. 2016. Uranium-lead zircon systematics in the Sudbury impact crater-fill: implications for target lithologies and crater evolution; Journal of the Geological Society; v.173, 59-75. Rainbird, R.H. and Davis, W.J. 2006. Detrital zircon geochronology of the western Huronian Basin; in 52nd Institute on Lake Superior Geology Annual Meeting, Sault Ste. Marie, ON, Proceedings v.52, pt.1, 55-56. Rasmussen, B., Bekker, A. and Fletcher, I.R. 2013. Correlation of Paleoproterozoic glaciations based on U–Pb zircon ages for tuff beds in the Transvaal and Huronian Supergroups; Earth and Planetary Science Letters, v.382, 173-180. 39 �Hyperspectral Imaging of Bedrock Core from the Minnesota DNR Drill Core Library: A New Tool for Archival Preservation and Mineral Exploration ELSENHEIMER, Don1, DEYELL-WURST, Cari2, and FONTENEAU, Lionel C.3 1 Minnesota Department of Natural Resources, 500 Lafayette Rd, St. Paul, MN 55155 USA Corescan Pty Ltd, 22033 Boul Gouin Ouest, Montreal, QC, CANADA 3 Corescan Pty Ltd, 1/127 Grandstand Road, Ascot WA 6104, AUSTRALIA 2 The Minnesota Department of Natural Resources (DNR) hired Corescan Inc. to scan 4900m of bedrock core from the DNR Drill Core Library (DCL) using Corescan’s hyperspectral core imaging system (Martini et al, 2017). The technique integrates both Visible Near InfraRed (VNIR) and Shortwave Infrared (SWIR) reflectance spectroscopy with high-resolution photography (50 µm) and 3-d laser profiling (200 µm) to identify minerals, estimate mineral abundances and create textural maps at 500 µm resolution. Hyperspectral imaging is a non-destructive analytical technique that supports the archival preservation of limited core material. Project results support DNR land management decisions on state mineral rights and promote mineral exploration and development. This project for the first time will provide public access to hyperspectral imaging data archived within the Coreshed® Virtual Core Library. DNR anticipates public release of project data and public access to Coreshed by summer, 2019. The DNR selected project core from thirty-two (32) drill holes located in five areas in Northern and Central Minnesota with distinct mineral deposits and/or high mineral potential. Initial project results are from an Archean Wabigoon Subprovince greenstone terrane near International Falls (Seine Group) and Biwabik Iron Formation core from the Mesabi Range. The Seine Group of greenschist-facies, metasedimentary and metavolcanic rocks sits at the contact between the Wabigoon and Quetico Subprovinces of the Archean Superior Province (Jirsa et al., 2014). Gold exploration in the region included an active period of drilling in the late 1980’s. Frey (2012) re-logged and re-sampled several of the DCL-archived Seine Group cores, and identified alteration patterns and features favorable for gold mineralization, including greater abundances of porphyoblastic and vein tourmaline. Hyperspectral imaging of twelve archived DCL cores from the area extends Frey’s tourmaline observations to drill cores that (due to active exploration) were not available at the time of his study. There is a positive correlation between gold concentrations and hyperspectral mineral identification of under-recognized tourmaline. Variations in the 2350nm feature position (Bierwirth, 2008) suggest tourmaline compositions within the dravite-schorl series (Figure 1). Complete or near complete transects of the Biwabik Iron Formation (BIF) were imaged in six Mesabi Range drill cores (LWD99-1, LWD99-2, MDDP-2, -5, -7, and -8). Hyperspectral imaging of core from LWD99-2 is able to differentiate microplaty hematite banding from more martite-rich bands. Two chlorite types are also recognized within this same core based on absorption features; an Mg-Fe intermediate composition that occurs in the Virginia Formation and its contact with the underlying Upper Slaty Unit, and a more iron-rich chamosite found in the Lower Cherty Unit and its contact with the underlying Pokegama Quartzite. Average albedo in the visible spectral range (448-740nm) highlights variation within the heavily sampled contact between the BIF and overlying Virginia Formation, where Addison et al. (2005) identified an ~25 to ~58cm thick ejecta layer associated with the 1850Ma Sudbury impact event. White mica is recognized based on absorption features within an ~ 2.6m interval of LWD99-2 core at the transition from BIF to Virginia Formation. Within this occurrence interval, a much smaller ~ 38cm interval with ammonium-rich white mica (feature around 2010nm, Canet et al. (2015)) is recognized in a thin layer of cherty carbonate. The discovery of relatively rare ammonium-rich white mica in association with an identified ejecta layer, if confirmed, would be significant. 40 �References Addison W.D., Brumpton G.R., Vallini D.A., McNaughton N.J., Davis D.W., Kissin S.A., Fralick P.W., and Hammond A.L. (2005) Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology 33:193-196. Bierwirth, P.N. (2008) Laboratory and imaging spectroscopy of tourmaline - a tool for mineral exploration. 14th Australasian Remote Sensing and Photogrammetry Conference, Darwin. Canet C., Hernández-Cruz B., Jiménez-Franco A., Pi T., Peláez B., Villanueva-Estrada R.E., Alfonso P., González-Partida E., Salinas S. (2015) Combining ammonium mapping and short-wave infrared (SWIR) reflectance spectroscopy to constrain a model of hydrothermal alteration for the Acoculco geothermal zone, Eastern Mexico. Geothermics 53:154-65. Frey B.A. (2012) International Falls Drill Core Descriptions and Chemistry, Koochiching County, Minnesota. Project 378 Open-File Report, Minnesota Department of Natural Resources, Division of Lands and Minerals, 39p. Jirsa M.A., Boerboom T.J., and Chandler V.W. (2014) M-197 Bedrock Geology of the International Falls and LittleFork 30’x60’ Quadrangles, northern Minnesota. Minnesota Geological Survey, Retrieved from the University of Minnesota Digital Conservancy, http://hdl.handle.net/11299/166157. Martini B.A., Harris A.C., Carey R., Goodey N., Honey F., and Tufilli N. (2017) Automated Hyperspectral Core Imaging – A Revolutionary New Tool for Exploration, Mining and Research. in “Proceedings of Exploration 17: Sixth Decennial International Conference on Mineral Exploration” edited by V. Tschirhart V. and M.D. Thomas, p. 911-922. Figure 1: Hyperspectral imaging of tourmaline within an 8cm-long section of quarter-core from DDH TC35-1. This section is within a larger 4 foot (1.22m) core interval that assayed at 4020ppb Au. Variations in the 2350nm feature position (Bierwirth, 2008) suggest compositions within the dravite-schorl series. 41 �Geology and Geochemistry of the Laird Lake Property and Associated Gold Mineralization, Red Lake Greenstone Belt, Ontario GÉLINAS, Brigitte, HOLLINGS, Pete1, FRIEDMAN, Richard2 1 Department of Geology, Lakehead University, Thunder Bay, Ontario P7B5E1 Pacific Centre for Isotopic and Geochemical Research, University of British Columbia 2 The Red Lake greenstone belt (RLGB) is one of world’s best endowed gold districts and like many other gold-rich regions, the individual deposits are closely associated with regional contacts, in part unconformable (Robert et al., 2005). A regional break in Red Lake separates the Mesoarchean and Neoarchean assemblages and hosts 94% of all gold (production, reserves, and resources; Dubé et al., 2003) in the greenstone belt, yet, the relationship between the two Archean packages is still disputed in terms of tectonic history (Stott, 1996; Stott and Corfu, 1991; Hollings and Kerrich, 2000; Roger et al., 2000; Sanborn-Barrie et al., 2001; 2004; Hollings and Kerrich, 2006). The Laird Lake property encompasses the regional break between the Balmer (2.99 to 2.96 Ga) and the Confederation (2.74 to 2.73 Ga) assemblages on the south-western end of the Red Lake greenstone belt, Northwestern Ontario. Multiple gold occurrences on the Laird Lake property generally occur within 200 m of the regional break and could represent the continuation of a similar gold system as seen at the Madsen Mine. The purpose of this study was to determine the tectonic setting in which the assemblages formed, and to characterize the controls on and nature of the gold mineralization associated with the tectonic contact between the Balmer and Confederation assemblages. Only 10 km east of the study area is the past-producing Madsen Mine, which lies on the north side of the regional break between the Balmer and Confederation assemblages. The ore is locally defined by the Austin and McVeigh ore zone, which displays a characteristic mineral banding (Dubé et al., 2000). Detailed mapping of the Laird Lake area highlighted major differences between the two assemblages (Gélinas, 2018). The Balmer assemblage is typically composed of fine-grained, aphyric, locally pillowed mafic volcanic rocks, ultramafic intrusive and volcanic rocks with flow-breccia textures and local spinifex-bearing clasts, and banded-iron formations. In contrast, the Confederation assemblage consists of porphyritic (feldspar) or poikiloblastic (amphibole) mafic volcanic rocks intercalated with intermediate to felsic volcanic rocks that include crystal lapilli tuffs, crystal tuffs and tuffs. Syn-volcanic and syn- to post-D2 intrusions commonly cross-cut the volcanic packages. A regional foliation (~Etrending) is present throughout the volcanic rocks and increases in intensity at the tectonic contact between the two assemblages where a deformation zone no thicker than 100 m is present within the Balmer assemblage. Whole-rock geochemical analyses were undertaken on 161 samples from the Laird Lake area. The Balmer assemblage is composed of tholeiitic mafic volcanic rocks with minor Al-undepleted komatiites, whereas the Confederation assemblage is composed of transitional mafic and calc-alkalic intermediate to felsic volcanic rocks, which display FI, FII, and FIIIb rhyolite trends. Neodymium isotope analyses, in conjunction with trace element geochemistry, suggests that parts of the Balmer assemblage were weakly contaminated by an older intermediate basement. The data suggests both arc and back arc volcanism within the Confederation assemblage, with the arc rocks showing stronger a crustal component than the back-arc rocks. U-Pb geochronology of volcanic and intrusive Confederation units yielded ages of 2741 ± 19 Ma (FI quartz-feldspar porphyritic crystal tuff) and 2737.68 ± 0.79 Ma (diorite). The geochemistry and age of the tuff correlates within error to the Heyson sequence of the Confederation, whereas the diorite is likely a syn-volcanic intrusion. 42 �The Balmer assemblage is interpreted to represent an oceanic plateau formed by plume magmatism on the margins of the North Caribou Terrane whereas the Confederation assemblage was likely built in an oceanic arc setting where both arc and back arc volcanism were occuring simultaneously. The presence of xenocrystic zircons within the 2741 Ma quartz-feldspar porphyritic crystal tuff suggest that melts within the main arc incorporated xenocrystic zircons during ascent through a thin Mesoarchean crustal fragment. Juxtaposition of the Confederation assemblage onto the Mesoarchean assemblages likely occurred between 2739-2733 Ma. Gold mineralization at the Laird Lake property is controlled by a D2 deformation zone within the Balmer assemblage at the tectonic contact between the Balmer and Confederation assemblages. The mineralization is commonly found associated with a mineral banded parallel to the main D2 fabric, accompanied by disseminated arsenopyrite, pyrrhotite, pyrite ± chalcopyrite, similar to the features observed at the nearby Madsen Mine. The Laird Lake property likely represents the continuation of the same mineralized structure found at both the Madsen and Starrat-Olsen mines and was later displaced as far as 10 km west by the dextral Laird Lake fault post-2704 Ma. References Dubé B, Balmer W, Sanborn-Barrie M, Skulski T, Parker J (2000). A preliminary report on amphibolite-facies, disseminated-replacement-style mineralization at the Madsen gold mine, Red Lake, Ontario. Geological Survey of Canada, Current Research 2000-C17, 14 p. Dubé B, Williamson K., and Malo, M., 2003. Gold mineralization from the Red Lake mine trend: Example from the Cochenour-Willans mine area, Red Lake, Ontario, with new key information from the Red Lake Mine and potential analogy with the Timmins camp. Geological Survey of Canada Current Research 2003-C21, 15 p. Gélinas, B., 2018. Geology and Geochemistry of the Laird Lake Property and Associated Gold Mineralization, Red Lake Greenstone Belt, Northwestern Ontario. Unpublished MSc thesis, Lakehead University, 360 p. Hollings P., and Kerrich R., 2000. An Archean arc basalt – Nb-enriched basalt – adakite association: The 2.7 Ga Confederation assemblage of the Birch-Uchi greenstone belt, Superior Province. Contributions to Mineralogy and Petrology, vol. 139, p. 208-226. Hollings P., and Kerrich R., 2006. Light rare earth element depleted to enriched basaltic flows from 2.8 to 2.7 Ga greenstone belts of the Uchi Subprovince, Ontario, Canada. Chemical Geology, vol. 227, p. 133-153. Robert F, Poulsen HK, Cassidy KF, Hodgson CJ (2005) Gold Metallogeny of the Superior and Yilgarn Cratons. Economic Geology 100th Anniversary volume p. 1001-1033. Rogers N., McNicoll V., van Staal C.R., and Tomlinson K.Y., 2000. Lithogeochemical studies in the UchiConfederation greenstone belt, northwestern Ontario: implications for Archean tectonics; Geological Survey of Canada, Current Research 2000-C16, 11 p. Sanborn-Barrie M, Skulski T, Parker J (2001) Three hundred million years of tectonic history recorded by the Red Lake greenstone belt, Ontario. Geological Survey of Canada, Open File 4594, 30 p. Sanborn-Barrie M, Rogers N, Skulski T, Parker J, McNicoll V, Devaney J (2004) Geology and tectonostratigraphic assemblages, east Uchi Subprovince, Red Lake and Birch–Uchi belts, Ontario. Geological Survey of Canada, Open File 4256; Ontario Geological Survey, Preliminary Map P.3460, scale 1:250 000. Stott G. M., 1996. The geology and tectonic history of the central Uchi Subprovince; Ontario Geological Survey, Open File Report 5952, 178 p. Stott, G. M., and Corfu, F., 1991. Uchi subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p. 145-238. 43 �Petrography of several Co-enriched samples from the Atikokan River Intrusions, Atikokan, Ontario GIBBONS1, Jack, DIEDRICH1, Tamara, QUIGLEY, Thomas2 1 MineraLogic LLC, 306 W. Superior St., Alworth Building, Suite 408, Duluth, MN 55802 USA 2 Great Lakes Exploration Inc., Menominee, MI 49858 USA The Atikokan River Intrusions (ARIs) consist of five or more sulfide and oxide-rich mafic intrusive bodies that have been emplaced along a 28-km section of the Quetico Fault Zone (QFZ) east of Atikokan, ON. Sulfide and oxide mineralization within these intrusions was historically explored for iron ore, and locally developed into at least one small-scale, open pit and underground mine in the late 19th century. The intrusions and associated mineralization are also variably enriched in copper, nickel, and cobalt. Great Lakes Exploration, Inc. (GLE) currently controls an approximately 12-km long stretch of the ARIs, which includes a significant portion of the mineralized intrusions. GLE is currently evaluating the potential for these intrusions to contain cobalt, copper, and nickel at concentrations and in mineral phases that are economically recoverable. Optical petrography (reflected and transmitted light) observations, and bulk geochemical data, from seven ARI hand samples, constrain the nature of Co-mineralization and provide information on textural relationships between minerals, as described here. Petrographic characterization indicates that sulfide mineral assemblages includes pyrrhotite, pyrite, and chalcopyrite. The presence of trace sphalerite was previously identified in other samples by QEMSCAN (conducted by XPS Consulting and Testwork Services). Pyrite occurs as 100- to 200-µm sized grains, while pyrrhotite and chalcopyrite occur as smaller 20- to 50-µm sized grains that compose larger aggregates that partially encompass pyrite grains. None of the observed sulfides display complex intergrowth or exsolution textures in the samples evaluated. Magnetite occurs with most sulfide assemblages, contains both chalcopyrite and pyrrhotite inclusions, appears to be roughly positively correlated with pyrrhotite abundance, and can locally replace pyrite. The abundance (15 to 20 volume percent) and textural relationship (e.g., contains sulfide inclusions and crosscuts/replaces igneous phenocrysts) suggest that at least a portion of the observed magnetite is secondary. Though no stoichiometric cobalt phase was definitively identified via optical petrography, cobalt assay results correlate well with pyrite abundance, consistent with the presence of cobaltiferous pyrite. Textural relationships suggest that the observed sulfide assemblage evolved from an early pyrite- to a late pyrrhotite-dominant assemblage, with chalcopyrite present in both early and late assemblages but likely increased in abundance in the latter assemblage. Examples of sulfide and oxide mineral occurrences are provided in Figure 1A-B. Observations on silicate mineralogy help to establish potential peak metamorphic conditions, understand origin and composition of mineralizing hydrothermal fluids, and provide guidance in constraining timing of mineralization. Coarse-grained chlorite intergrowths with pyrrhotite appear to suggest that a portion of the observed mineralization possibly occurred during metamorphism. The lack of primary igneous minerals, in most samples, suggests that secondary alteration was intense, at least locally, and that peak metamorphic conditions reached upper 44 �greenschist facies; historic reports (MacTavish, 1999) of rarely preserved garnet suggest that metamorphic conditions could have reached lower-most amphibolite facies at other locations within the ARIs. Examples of alteration products and textures are shown in Figure 2A-B. A B Figure 1. Examples of typical ARI sulfide assemblages. Both images taken in plane-polarized, reflected light. A) Coarsegrained pyrite and magnetite locally supported by a pyrrhotite-rich matrix. Trace chalcopyrite occurs between pyrite grains. Magnetite locally replaced several pyrite grains near the center portion of image. Pyrrhotite displays a slender reaction rind. B) . Typical pyrite, pyrrhotite, and chalcopyrite assemblage. Several of the pyrite grains contain a distinct pitted core (partially outlined by dashed line) surrounded by a broad growth zone lacking inclusions, which might possibly indicate pyrite growth occurred in two stages. The scale of the image is the same as Fig. 1A. A B Figure 2. Mineral textures that help to constrain the timing and origin of observed sulfide assemblage. Both images taken in plane-polarized, reflected light. Images have been edited to highlight contrast between mineral phases. A) Large igneous orthoclase phenocryst replaced by magnetite. Trace amounts of pyrite and chalcopyrite occur within the magnetite. Pyrrhotite is absent from the sample. B) Coarse-grained, secondary chlorite is intimately intergrown with pyrrhotite in lower left portion of image. Alteration rind on pyrrhotite is very well developed in this sample. Pyrite locally replaced by magnetite. Chalcopyrite and pyrrhotite exhibit strong spatial association that is typical of this sulfide assemblage. Reference MacTavish, A.D. 1999. The mafic-ultramafic intrusions of the Atikokan-Quetico area, northwestern Ontario; Ontario Geological Survey, Open File Report 5997, 127p. 45 �Recognizing MCR magmas generated by partial melting in the SCLM: Lessons from mafic magmas in the Coldwell Complex GOOD, Dave1, HOLLINGS, Pete2 and JEDEMANN, Andrew2 1Department of Earth Sciences, Western University, London, ON N6A 5B7 Canada 2Department of Geology, Lakehead University, Thunder Bay, ON P7B 5E1 Canada We present new interpretations of a comprehensive data set for basalt and intrusive mafic rocks from the Midcontinent Rift. These data display well-defined trends for trace element abundances that demonstrate variable degrees of partial melting in a plume-like source and subsequent fractional crystallization. For instance, diagrams that compare highly incompatible elements Zr and La, La and Yb, or Nb and Th in MCR rocks show the majority of data plot in a field that spans compositions from E-MORB to OIB along approximately linear trends with constant inter-element ratios. Data that deviates from these MCR trends are explained by interaction of the magma with continental crust during ascent, consistent with elevated Th contents or Rb-Sr and Sm-Nd isotope values that confirm contamination. However, there are cases where evidence such as relative Th or major element abundances contradict isotopic evidence that may or may not agree with crustal contamination. Indeed, multiple isotope systems (Pb-Pb and Nd-Sm) are sometimes in disagreement with respect to the degree of contamination. Another mechanism that might explain such irregularities is partial melting of a metasomatized SCLM source (Furman and Graham, 1999; Sgualdo et al., 2015). It has been well established that initial SCLM isotope values can be overprinted by metasomatism, and therefore isotope systematics, in particular Rb-Sr and Sm-Nd are ineffective for distinguishing between crustal contaminated plume magmas and SCLM-derived magmas. Establishing a set of geochemical criteria that could be used to distinguish between mafic rocks in the MCR that were generated by partial melting in the metasomatized SCLM from plume magmas that were contaminated in the crust is the subject of this presentation. In wide-ranging studies of mantle xenoliths from Africa, New Zealand and Europe, secondary minerals found in veins include phlogopite, clinopyroxene and pargasitic amphibole (Frezzotti et al., 2010; Scott et al., 2014). The trace element signatures of each phase exhibit distinguishing features that, since they are among the first minerals to disappear during a partial melting event, will impart distinctive trace element signatures to the resulting magma composition. For instance, the different compatibilities of Rb, Ba and Sr in amphibole compared to phlogopite, or Nb and Th in amphibole compared to clinopyroxene will result in decoupling of LILE abundances due to the relative proportions of each mineral in the source rock. As these minerals contain very high concentrations of incompatible elements relative to the depleted protolith SCLM rock, a relatively small amount (<1-2%) of each mineral will have a very large impact on the resultant magma composition enabling recognition of trace element signature. Magmas generated from Areas of SCLM that have been less impacted by metasomatism, and thus might have a very low proportion of secondary minerals, will have a depleted HFSE signature marked by sub-chondritic Zr/Y, Zr/Hf and very low La/Yb values, and possibly anomalous Sr and Ba. A key example of volcanic rocks that show contradictory isotopic and geochemical evidence for crustal contamination is Mamainse Point Volcanic Group 5b (Shirey et al., 1994), in 46 �which the εNd values of -3.5 and -6.3 indicate significant crustal contamination, but Pb-Pb data imply a maximum of 2% crustal material. The combination of sub-chondritic Zr/Y and Zr/Hf, low La/Yb, very low La, Th, and TiO2 abundances, and corresponding positive Ba and Sr anomalies is strong evidence for derivation from a weakly metasomatized but initially depleted SCLM source. Examples of MCR magmatism from the Nipigon embayment that exhibit SCLMlike signatures are presented to test the usefulness of key features identified in mafic rocks of the Coldwell Complex that distinguish them as originating from the SCLM. The geochemical characteristics of the Nipigon intrusions are examined as test cases to establish whether or not they were derived from the SCLM. References Beccaluva et al., 2001, J. Pet. 42, 173-187. Bodinier, J.L., Menzies, A.M., et al., 2004, J. Pet. 45, 299-320. Frezzotti, M.L., Ferrando, S. et al., 2010, Geochim. Cosmochim. Acta 74, 3023-3039. Furman, T., and Graham, D. 1999, Lithos 48, 237-262. Good D.J. and Lightfoot P.C., CJES, in press. Hollings, P., Hart, T., Richardson, A., MacDonald, C.A., 2007, CJES 44, 1087-1110. Scott, J.M., Hodgkinson, A. Palin, J.M., et al. 2014, Contrib Mineralogy Petrol., 167: 963. Sgualdo, P., Aviado, K., Beccaluva, L., et al., 2015, Tectonophysics, v. 650, p. 3-17. Shirey, S.B., Klewin K.W., Berg, J.H. and Carlson R.W., 1994, Geochim. Cosmochim. Acta, 58, 44754490. Lightfoot, P.C., Sage, R.P., Doherty, W., Naldrett, A.J. and Sutcliffe, R.H. 1999. OGS OFR 5998, 57p. 47 �Recent Efforts to Curate and Provide Access to the Historical Documents of the E.K. Lehmann and Associates Exploration Company GOTTSCHALK, Brad, and ROSE, Caroline Wisconsin Geological and Natural History Survey, 3817 Mineral Point Road, Madison, Wisconsin, 53705 In October of 2015, the Wisconsin Geological and Natural History Survey (WGNHS) received a large donation of documents from Kate Lehmann, the daughter of renowned exploration geologist Ernest K. Lehmann. While Lehmann worked primarily in Minnesota, his company, E.K. Lehmann and Associates, also worked in northern Wisconsin from the late seventies into the mid-nineties and, during that time, donated a large amount of core from their drilling projects to WGNHS. After Ernest Lehmann passed away, his family donated the records related to his work in Wisconsin to WGNHS, and the paperwork related to his work in Minnesota to the Minnesota DNR. With financial assistance from the Lehmann family, the Minnesota DNR scanned all of the documents donated to them and added them to their Drill Core Library and Mineral Exploration Collections’ interactive map to provide online access. Staff at WGNHS have provided something similar to our users, but with a more narrowly focused scope. The donation to WGNHS was quite large—31 record boxes of reports and other documents, and two cardboard cases of rolled maps and figures. As WGNHS possesses limited resources, we knew we would have to find a way to focus our curation efforts. Consulting with Tom Evans, an emeritus Survey staff member and an authority on mining in Wisconsin, we decided to concentrate on documents that directly related to rock core in our possession. From December of 2015 to February of 2017, Tom Evans and Brad Gottschalk, the WGNHS archivist, searched for documents that provided data for drillholes. Once these were identified, they matched the Lehmann drillholes to records in our geological database, Geobase. 351 individual drillholes in 67 exploration targets were identified in the Lehmann documents. Of these 351 holes, we had physical core from 289. Some of these Lehmann targets are still considered areas of interest for mineral development. The targets most extensively explored were Bend in Taylor County, Ritchie Creek in Price County, and Horse Shoe in Lincoln County. The documentation for drillholes in these and other targets include location maps, geological maps, logs, geological and geophysical cross-sections, chemical analyses of samples and assay reports. In 2017, we received a grant from the USGS’s National Geological and Geophysical Data Preservation Program (NGGDPP) to scan the Lehmann papers and put them online using an interactive map application. Selecting the documents for scanning was a complicated task. In the paperwork were monthly reports for many of the targets, as well as memos and final reports. The documentation for the drillholes was frequently duplicated in multiple monthly reports as well as in the final report. There were multiple cross-sections for the more widely explored targets, and, especially for the Bend target, which showed promise as a gold deposit, there was a great deal of assay data. Gottschalk and two student employees weeded out duplicates and scanned each unique document. In the end, some drillholes represented in the Lehmann papers were not included in the web application due to poor or incomplete data. After excluding these, we compiled data for 331 drillholes in 65 targets contained in 1153 individual documents. Of the 331 holes represented in the project, we have physical core samples from 288. 48 �As the scanning portion of the project neared completion, Caroline Rose, GIS specialist, began to construct the ArcGIS application that would provide online access to the documents. Rose used ArcGIS Online’s Storymaps templates and Web App Builder to present the Lehmann collection in two web maps: one organized by drillhole and one organized by exploration target. The first map features point locations and details of individual drillholes and links to all related documents. Document details can be followed to show all drillholes related to the document. Documents can be opened in PDF format from the map popup or from a table in the interface. The second web map shows exploration targets, which are collections of drillholes, as circular symbols sized according to the number of related documents. It is immediately apparent that three of the targets are related to more than fifty documents (Bend, Ritchie Creek, and Horse Shoe). Several other targets are related to more than ten documents. The targets are linked to their related documents. Again, documents in PDF format can be opened from the map or the table. Figure 1: The interactive map showing the exploration targets represented in the Lehmann papers. Rose configured the data using ArcGIS Pro to establish many-to-many relationships between the datasets, as one drillhole could be related to many documents, and one document could be related to many drillholes. She then used the ArcGIS Online WebApp Builder to create the interface, and a Storymaps template to create the tabbed layout. At the end of the project, metadata records for the 1153 documents scanned and put online were uploaded to the USGS National Digital Catalog. References Minnesota DNR, 2016, Lehmann Family fund collection of Mineral Exploration Documents (including the Polaris Joint Venture (https://www.dnr.state.mn.us/lands_minerals/polaris/index.html) 49 �Superior Shoal Revisited: Evidence for Keweenawan Basalts with Reversed- and Normalpolarity Remanent Magnetization and Early Magma Chemistry, Central Lake Superior GRAUCH, V.J.S. 1 and SCHULZ, K.J. 2 1 U.S. Geological Survey, MS 964, Federal Center, Denver, CO, 80225 U.S. Geological Survey, MS 954, National Center, Reston, VA, 20192 2 Superior Shoal is an easterly trending, ~20-km-long bathymetric bedrock high below the water’s surface near the center of Lake Superior. Being the only accessible bedrock within a radius of about 70 km, the Shoal can provide evidence critical to understanding the structure of the 1.1 Ga Midcontinent Rift in central Lake Superior, yet debates remain about its geology. Located at the intersection of two geophysically interpreted faults, the bathymetric high is composed of a series of ridges of Keweenawan basalts on the south and a broader ridge of sandstone on the north (Manson and Halls, 1991). Previous studies of Superior Shoal give conflicting results on the age of the basalts based on the polarity of remanent magnetization. Magnetic polarities are commonly used to recognize early (>1100 Ma) Keweenawan lavas (reversed-polarity) from younger (<1100 Ma) lavas (normal-polarity), while acknowledging a separate normal-polarity event between ca. 1101-1103 Ma (Swanson-Hysell et al., 2019). Manson and Halls (1991) concluded from paleomagnetic measurements that the basalts have normal-polarity remanence, whereas Teskey et al. (1991) concluded from analysis of aeromagnetic data that the basalts have reversed-polarity remanence. To resolve the apparent disagreement regarding magnetic polarity, we (1) reviewed the paleomagnetic results from the Manson and Halls study, (2) expanded on the aeromagnetic analysis of the Teskey et al. study, and (3) analyzed basalt samples collected during the paleomagnetic study of Manson and Halls to determine if they are chemically affiliated with typical early or late rift lavas (Nicholson et al., 1997). Review of Paleomagnetic Study of Manson and Halls A review of the methods, analyses, estimated errors, and results of the Manson and Halls (1991) study from their three basalt sites at Superior Shoal give confidence in their results. They found primary normal-polarity components, although orientations are somewhat dissimilar to those expected for typical normal-polarity Keweenawan basalts. They attributed the dissimilar directions to tectonic tilts that are nonuniform, but generally have northerly dip. Expansion of Aeromagnetic Analysis by Teskey et al. Teskey et al. (1991) analyzed the negative aeromagnetic anomaly at Superior Shoal using the principle that magnetic rocks forming rugged bathymetry should produce aeromagnetic anomalies that correspond to bathymetric shapes. In comparing the bathymetry of Superior Shoal to aeromagnetic anomalies along profiles, Teskey et al. noted an inverse correlation between bathymetric and aeromagnetic highs and lows, suggesting a reversed-polarity remanence. Expanding on this approach, a three-dimensional model of bathymetry was assigned magnetizations typical of normal versus reversed polarity for Keweenawan basalts. Comparisons of the magnetic fields computed from these models to the observed aeromagnetic anomaly show a good correspondence with the reversed-polarity model, supporting the conclusion that the bulk of the rock volume at Superior Shoal possesses very strong, reversed-polarity remanence. 50 �Chemical Analysis of Paleomagnetic Samples Recently, 10 samples from basalt sites 1 and 2 of Manson and Halls (1991) were analyzed for major and trace elements. The Superior Shoal basalt samples have similar geochemical characteristics with a limited range in MgO = 5.4 to 8.1 wt.%, TiO2 = 1.6 to 2.4 wt.%, and La/Yb = 6.5 to 7.2. They are most similar in composition to Siemens Creek Type II basalts and are comparable to the Central suite of the Osler Group (Fig. 1), both of which are composed of early, reversed-polarity lavas that are mostly older than ca. 1105 Ma (Nicholson et al., 1997; Swanson-Hysell et al., 2019). The results of the basalt analyses combined with the paleomagnetic results suggest that basalts with early magma chemistry but with normal-polarity remanence are present at Superior Shoal. Reconciliation of the Results A more detailed analysis of flight-line aeromagnetic data allows that basalts of both polarities likely exist at Superior Shoal. Low-amplitude positive anomalies are superposed on the broader, high-amplitude negative anomalies, suggesting that a large volume of reversedpolarity early lavas underlie normal-polarity lavas of smaller volume (and/or lower magnetization). The apparent conflict of normal-polarity, early magma chemistry may be due to (1) magma typical of early rift magmatism that continued erupting into one of the later normalpolarity times, or (2) a previously unrecorded normal polarity event that occurred sometime between 1105 Ma and 1103 Ma. Further study at Superior Shoal appears warranted. References Lightfoot, P.C., Sutcliffe, R.H., and Doherty, William, 1991, Crustal contamination identified in Keweenawan Osler Grop tholeiites, Ontario: A trace element perspective: Journal of Geology, v. 99, p. 739–760. Manson, M.L., and Halls, H.C., 1991, An investigation of Superior Shoal, central Lake Superior, with a manned submersible: Canadian Journal of Earth Sciences, v. 28, p. 145–150. Nicholson, S.W., Shirey, S.B., Schulz, K.J., and Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504–520. Swanson-Hysell, N.L., Ramezani, J., Fairchild, L.M., and Rose, I.R., 2019, Failed rifting and fast drifting: Midcontinent Rift development, Laurentia’s rapid motion and the driver of Grenvillian orogenesis: Geological Society of America Bulletin, 29 January 2019, https://doi.org/10.1130/B31944.1 Teskey, D.J., Thomas, M.D., Gibb, R.A., Dods, S.D., Kucks, R.P., Chandler, V.W., Fadaie, K., and Phillips, J.D., 1991, High resolution aeromagnetic survey of Lake Superior: Eos, v. 72, no. 8, p. 81, 85–86. 51 �Evaluating Alternate Geophysical Models along the Isle Royale-Superior Shoal Aeromagnetic Anomaly, Central Lake Superior GRAUCH, V.J.S. 1, STEWART, Esther Kingsbury 2, WOODRUFF, Laurel G. 3, and HELLER, Samuel 4 1 U.S. Geological Survey, MS 964, Federal Center, Denver, CO, 80225 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Rd., Madison, WI 53705 3 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 4 U.S. Geological Survey, MS 939, Federal Center, Denver, CO, 80225 2 As much as 3 km of Midcontinent rift basalts exposed on the NE-elongate island of Isle Royale (IR) in central Lake Superior dip SE and are commonly regarded as part of the upthrown block of a post-rift reverse fault just off the northern IR shore. A prominent, narrow, aeromagnetic high-low pair (IR-SS anomaly) emanates from the NE tip of IR, curving toward the SE to a strong negative anomaly at Superior Shoal (SS), a bathymetric high near the center of the lake (Fig. 1). The IR-SS anomaly is commonly interpreted as an extension of the IR reverse fault, involving younger (normal magnetic polarity) and possibly older (reversed magnetic polarity) rift basalts. Broad, linear to curvi-linear gravity highs parallel the IR-SS anomaly to the south (Fig. 1). A seismic-reflection line (GLIMPCE A) crosses the IR-SS anomaly at a complicated area of multiple linear magnetic anomalies (Fig. 1). The seismic section shows a 12-km-wide disrupted zone that extends vertically below the complicated area and divides packages of subhorizontal reflections that cannot be connected across the zone. Previous geophysical models of the IR-SS anomaly satisfy some of the data sets, but none integrate all of them satisfactorily. For example, a vertical reverse fault with ~2.5 km of throw has been interpreted and modeled from the seismic-reflection and gravity data (Thomas and Teskey, 1994), but does not account for the shallow basalts observed near SS. Conversely, a magnetic model along flight-line L11260 east of GLIMPCE A (Fig. 1) fits the sharp IR-SS anomaly using a <5-km-wide zone of igneous rock extending from the lake bottom to ~4 km depth (Teskey and Thomas, 1994), but does not fit the broader gravity anomaly. To develop models of the IR-SS anomaly that better integrate all the data and are constrained by new information that early lavas are preserved at SS (Grauch and Schulz, this volume), we tested a number of 2D gravity and magnetic models considering 3 conceptual models: (1) reverse fault with steeply dipping, south-facing basalt layers; (2) localized intrusions, such as volcanic feeder zones or dikes; and (3) remnants of an early-lava plateau extending northward from the IR-SS anomaly. We were unable to construct fully integrated models using the steeply dipping reverse-fault concept. Instead, models with moderately southward-dipping (<45°) basalt layers worked for profiles IRKP and IRKP2 (Fig. 1). The early-lava plateau remnant concept works well for profile L11070 across SS (Fig. 1). This model depicts strongly magnetic, reversed-polarity layers north of the IR-SS anomaly that abruptly terminate at the south side of the shoal (the linear positive anomaly there is the expression of this termination). Geologically, this model suggests that early lavas rest at shallow levels (~1 km depth) north of the IR-SS anomaly, possibly overlying a pre-rift sedimentary basin that has been pervasively intruded by a younger, rift-related mafic igneous complex. This is consistent with preliminary reinterpretations of GLIMPCE A, which suggest that a ~10-km thick, pre-rift sedimentary basin exists north of the IR-SS anomaly. Models that work best for profiles 8ext, 25ext, and 43ext include moderately dipping, south-facing basalts combined with concepts of both early-lava plateau remnants and localized intrusions. 52 �The 2D model testing suggests that (1) the IR-SS aeromagnetic anomaly is likely the product of multiple geologic causes; (2) a steeply dipping reverse fault model is the least favored; and (3) models involving localized intrusions and shallow, early-lava and/or pre-rift rocks at or on the north side of the IR-SS anomaly need to be considered further. References Anderson, E.D. and Grauch, V.J.S., 2018. Updated aeromagnetic and gravity anomaly compilations and elevationbathymetry models over Lake Superior: U.S. Geological Survey data release, https://doi.org/10.5066/F7F18X8S. Grauch, V.J.S. and Schulz, K.J., 2019. Superior Shoal Revisited: Evidence for early Keweenawan lavas with both reversed and normal-polarity remanent magnetization, central Lake Superior: Institute on Lake Superior Geology 65, Part 1 – Program and Abstracts Teskey, D.J. and Thomas, M.D., 1994. Three-dimensional magnetic modelling of the Midcontinent Rift beneath central Lake Superior: Canadian Journal of Earth Sciences, v. 31, p. 675–681. Thomas, M.D. and Teskey, D.J., 1994. An interpretation of gravity anomalies over the Midcontinent Rift, Lake Superior, constrained by GLIMPCE seismic and aeromagnetic data: Canadian Journal of Earth Sciences, v. 31, p. 682–697. Fig. 1. Aeromagnetic, gravity, and geology maps for the study area showing locations of seismicreflection lines and 2D model profiles. The IR-SS aeromagnetic anomaly is traced on all maps by the dashed yellow line. IR – Isle Royale; SS – Superior Shoal. Gravity and aeromagnetic compilations from Anderson and Grauch (2018). Geology generalized by E. Anderson from a USGS GIS compilation by C. Dicken (accessed January, 2015). 53 �Geological characteristics and structural controls of Au mineralisation at the enigmatic Hemlo deposit. HOLDER, David1, ROBERT, Francois1 and HAY, Jonathan1 1 Barrick Gold Corporation, Hemlo Operations, Marathon, Ontario, Canada. email:david.holder@barrick.com Hemlo is one of Canada’s largest and most well-known mines, producing ~23 Moz Au since discovery in 1981. The deposit is located in the Hemlo-Schreiber greenstone belt within the Wawa subprovince of the Superior craton. The Wawa sub-province, with ~40 Moz Au endowment (past production + reserves + resources) represents the western continuation of the highly auriferous southern Abitibi subprovince (~281 Moz Au). The sub-provinces are separated by the Kapuskasing structural zone, which is thought to have facilitated uplift and erosion of the Wawa block, exposing deep, high-grade metamorphic rocks ranging westward from granulite to amphibolite (e.g. Thompson 2006). Hemlo represents a rather unique deposit which is effectively isolated within the HemloSchreiber greenstone belt. Located along the Hemlo shear zone, Hemlo is hosted within amphibolite grade tectonites of volcanic (predominately volcanoclastics and hypabyssal intrusion) and sedimentary origin (e.g. Muir 1997). The mineralisation is characterised by an unusual metal assemblage with significant enrichments of Mo-As-Sb-Hg-Tl-V-Ba, associated with K-metasomatism and pervasive feldspathisation (Poulsen et al., in press). The unusual characteristics of Hemlo mean it has been the focus of many scientific studies over the past ~35 years. However there is still no consensus regarding the deposit genesis and its origins remain enigmatic. This is in part due to the effects of high-grade metamorphism and intense deformation, which have modified the original character of mineralization and geometry of the ore body. Historically, mining of the deposit has been carried out as 3 distinct operations; David Bell, Golden Giant (Main) and Williams (B- and C-zones) which has further hampered understanding of the system. Since unification of the mine by Barrick, a concerted effort has been made to determine the geological controls of mineralisation, focused primarily on the western-most C-zone, the main area of current operations. The deposit can be split into two distinct zones (Fig. 1); [1] the Williams B-zone and eastern extensions; Golden Giant Main zone and David Bell (referred to as B-zone herein) and [2] the Williams C-zone. The B-zone, which accounts for most of the gold, is a moderate to steeply NE-dipping tabular ore body developed on the contact of a series of felsic volcanic rocks known as the Moose Lake Volcanic complex (MLVC) and a heterolithic volcanoclastic unit locally referred to as the “fragmental” unit (Poulsen et al., submitted). The B-zone represents the “classic” Hemlo ore, characterised by textually destructive K-feldspar alteration (microcline) with abundant pyrite, molybdenite, barite, and a variety of As- and Hg-bearing sulfides and sulfosalts. The grade-thickness distribution on a longitudinal section across the deposit (Fig. 1) highlights the overall NW-plunge of the mineralisation in this zone, with a main shoot plunging ~30o and a number of steeper internal shoots plunging ~60o. The geologic controls of these plunges are poorly understood at present and are the focus of on-going study. 54 �Grade-Thickness Williams C-zone Figure 1: Interpolant gram.meter long sections (looking north) of the B-zone-David Bell (east) and Williams C-zone 100-series (west). The 300-series and B-zone footwall lodes not shown. Black-dashed lines highlight two apparent plunges to the mineralised system [60o-NW and 30o-NW]. Interpolant based on 0.5 g/t indicator grade shell. Williams B- / Golden Giant Main -zone David Bell The C-zone, located in the west part of the deposit comprises two sub-parallel Wstriking moderately (~60o) plunging shoots (Fig. 1) known as the 100- and 300- series lodes. The 100-series mineralisation is developed within a tight NW plunging fold closure of the “fragmental” unit, whereas the 300-series is situated within the MLVC. The mineralisation is characterised by pervasive, textually destructive K-feldspar alteration with Figure 2. Photograph of [A] molybdenite and pyrite disseminations and folded and transposed Kstringers (Fig. 2a), cross-cut by high-grade feldspar alteration and [B] rerecrystallized quartz veins and quartz-pyrite crystallised early quartz vein A with abundant Au visible. replacement zones (Fig. 2b). It is evident from underground exposure and drill-core that the bulk of mineralisation pre-dates metamorphism and deformation: feldspar and quartz-pyrite alteration zones are folded and transposed by the penetrative S2 foliation, which also transposes molybdenite and pyrite stringers (Fig. 2a). The early quartz veins and quartz-pyrite replacements display diffuse lobate contacts typical of recrystallised quartz, with sulfides and visible gold also transposed into the foliation planes (Fig. 2b). The current geometry of the C-zone mineralisation was evidently controlled by the development of F2 folds. The overall moderate to steep NW-plunge of the mineralisation corresponds with plunge measurements of F2 parasitic fold hinges and D2 stretching lineations (e.g. Muir 2003). A late, post-D2 mineralisation event is evident from a number of late crack-seal ribbon veins, oblique to and cross-cutting the S2 fabric, and cutting the earlier quartz-pyrite mineralisation. These distinct and superimposed styles of mineralization indicate a complex and multistage history of the Hemlo deposit, a characteristic common to many giant gold deposits. References Muir, T.L., 1997. Precambrian geology, Hemlo gold deposit area; Ontario Geological Survey, Report 289:1-219 Muir, T.L., 2003. Structural evolution of the Hemlo greenstone belt in the vicinity of the world-class Hemlo gold deposit; Canadian Journal of Earth Science. 40:395-430. Poulsen, H.K., Robert, F. & Barber, R., (submitted) Hemlo Gold System, Superior Province, Canada, Society of Economic Geologists Special Publication on Gold Deposits. Thompson, P.H. 2006. A new metamorphic framework for the Hemlo greenstone belt: Implications for deformation, plutonism, alteration and gold mineralization; Ontario Geological Survey, Open File Report 6190:1-80. 55 �Detrital Zircon Geochronology of Keweenaw Interflow Sediments within the North Shore Volcanic Group, Minnesota, U.S.A. JOHNSON, Linnea L.1, MALONE, David, H.1, CRADDOCK, John, P.2 1 Geography-Geology, Illinois State University, Normal, Illinois 61790 Geology, Macalester College, 1600 Grand Avenue, Saint Paul, Minnesota 55105 2 During the early stages of the Mesoproterozoic Midcontinent Rift, the North Shore Volcanic Group was deposited around 1100 Ma. This group of volcanic rocks, composed of rhyolite, basalt, and andesitic basalt, are interlaid with detrital sediments whose source zircon ages do not coincide with the age of the rift system. These interflow sediments vary in composition, comprised of quartz arenite, lithic arenite, conglomerate, and conglomeratic sandstone. Collection of samples took place at two locations along the north shore of Lake Superior in Minnesota, USA. Samples were collected from ~10 m thick conglomeratic sandstone at Caribou Creek , a ~1 m thick overturned lithic arenite entrained in a xenolith of the Beaver Bay Complex at milepost 61 on Highway 61, and cross bedded sandstones at Leif Ericson Park in Duluth. Zircon analysis using LA-ICPMS at the University of Arizona Laserchron Center, determine the provenance of both these sandstones. Milepost 61 sample set (n=102) contains zircons with a maximum deposition age of 1081 Ma in addition to zircon ages ranging from 1073.0-1879.9 Ma. Using an age probability plot, four peak ages are identified to be 1116, 1440, 1688, 1778 Ma. The Caribou Creek sample set (n=61) contains zircon ages ranging from 1051.2-3184.1 Ma, with three peak ages of 1109, 1377, and 1730 Ma. A total of 101 zircons were analyzed for the Leif Erickson sample. Zircons from this sample ranged in age from 1074-2707 Ma and has a maximum depositional age of 1081 Ma. Age peaks for this sample are 1111, 1446, 1690 and 1778 Ma. Prior notions that interflow sediments were sourced only from within the rift system cannot be entirely true. New data we collected suggests that some of the interflow sediment was derived from an external source outside of the Midcontinent Rift basin. Zircon ages coincide with Archean terranes to the south, and may also include the Midcontinent Granite-Rhyolite, Mazatzal and Yavapai provinces. Fluxes in high lands from reactivation of faults bounding these provinces may have uplifted these potential source areas. References Craddock, J.P., Konstantinou, A., Vervoort, J.D., Wirth, K.R., Davidson, C., Finley-Blasi, L., Juda, N.A., and Walker, E., 2013, Detrital zircon provenance of the Proterozoic Midcontinent Rift, Lake Superior region, USA: Journal of Geology, v. 121, p. 57-73. Davis, D.W., and Green, J.C. 1997, Geochronology of the North American Midcontinent Rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences, v. 34, p. 476-488. Fairchild, L.M., Swanson-Hysell, N.L., Ramezani, J., Sprain, C.J. and Bowring, S.A., 2017, The end of Midcontinent Rift magmatism and the paleogeography of Laurentia: Lithosphere, v. 9, p.117-133. Gehrels, G. and Pecha, M., 2014, Detrital zircon U-Pb geochronology and Hf isotope geochemistry of Paleozoic and Triassic passive margin strata of western North America: Geosphere, v. 10, p. 49-65. Gehrels, G.E., Valencia, V., Pullen, A., 2006, Detrital zircon geochronology by Laser-Ablation Multicollector ICPMS at the Arizona LaserChron Center, in Loszewski, T., and Huff, W., eds., Geochronology: Emerging Opportunities, Paleontology Society Short Course: Paleontology Society Papers, v. 11, 10 p. Gehrels, G.E., Valencia, V., Ruiz, J., 2008, Enhanced precision, accuracy, efficiency, and spatial resolution of U-Pb ages by laser ablation–multicollector–inductively coupled plasma–mass spectrometry: Geochemistry, Geophysics, Geosystems, v. 9, Q03017 Jirsa, M.A. 1984. Interflow sedimentary rocks in the Keweenawan North Shore Volcanic Group, northeastern Minnesota. Minn. Geol. Surv. Rep. Invest. 30, 20 p. Malone, D.H., Stein, C.A., Craddock, J.P., Kley, J., Stein, S., and Malone, J.E., 2016, Maximum depositional age of the Neoproterozoic Jacobsville Sandstone, Michigan: Implications for the evolution of the Midcontinent Rift: Geosphere, v. 12, p. 1–12. Whitmeyer, S.J., and Karlstrom, K E., 2007, Tectonic model for the Proterozoic growth of North America: Geosphere, v. 3, p. 220-259. 56 �Figure1: A regional tectonic map with the Midcontinent Rift and major geologic contacts (Craddock et al., 2013). Milepost 61 and Caribou Creek sample locations are marked with red stars. The location for the KP samples (Craddock et al. 2013) are marked with blue stars. Figure2: Stratigraphic column of the North Shore Volcanic Group strata in the Keweenaw Supergroup, showing the interflow sediments and sample localities (Craddock et al., 2013). Red indicates samples from the interflow sediments. Gray indicates samples from the Beaver Bay Complex and North Shore Volcanics. Figure3: Stacked probability density plot comparing Caribou Creek zircon ages to mile post 61 zircon ages. Age peaks in Ma. Figure 4: Cumulative probability plot comparison of current samples at Caribou Creek and Mile Post 61, with additional interflow sediment data set KP10 and KP 16 (acquired from Craddock et al., 2013). Figure 5: Stacked age probability plots comparing interflow sediment with overlying sandstone units (acquired from Craddock et al., 2013) data set. Age probability plots are categorized with orogeny events. 57 �Paleoproterozoic Snowball Earth? Sedimentology and Geochemistry of a Huronian Glacial Cycle KURUCZ, Sophie1, FRALICK, Philip1, LALONDE, Stefan2, HOMANN, Martin2 1 Department of Geology, Lakehead University, Thunder Bay, ON, skurucz@lakeheadu.ca 2 European Institute for Marine Studies, CNRS-UMR6538 Laboratoire Géosciences Océan, Brest, France The Paleoproterozoic Huronian Supergroup is a ~12km thick sequence of mostly sedimentary rocks that outcrops along the southern margin of the Superior craton and contains evidence for three complete glacial cycles within its stratigraphy. The second glacial event, represented in the Bruce Formation of the Quirke Lake Group is unique because of its overlying cap carbonate, the Espanola Formation, which is the only appreciable carbonate unit within the Huronian Supergroup. A cap carbonate overlying the glacial deposits of the Bruce Formation suggests that the Quirke Lake Group may record evidence for extreme climatic perturbations on the same scale as the later Neoproterozoic glacial cycles, where cap carbonates are ubiquitous overlying glacial deposits. The Neoproterozoic glaciations have been the source of much speculation regarding the cause of the formation of cap carbonates and the possibility of their representing the resulting effects of global ice cover during periods known as ‘Snowball Earth’ events (eg. Kirschvink, 1992). Thus, the presence of a cap carbonate overlying only the second of three glacial deposits in the Huronian Supergroup suggests that the conditions that led to its deposition were unique within the Paleoproterozoic and perhaps akin to those that prevailed during the Neoproterozoic glaciations. To assess the extent of the similarities between the Espanola Formation and the Neoproterozoic cap carbonates, the sedimentology, geochemistry, and isotopic composition of the Bruce glacial event was studied in its entirety. Some of the most interesting and useful results were uncovered through systematic sampling of drill hole E150-2 (Figure 1). Firstly, the presence of a hitherto unmentioned laminated dropstone facies occurs in the uppermost Bruce Formation. This unit is unique because it records evidence of both carbonate precipitation and glacial activity at the same time; a feature that is not recorded elsewhere in the Quirke Lake Group sedimentology. In this facies, 1-10cm thick carbonate-rich laminae occur in a clastpoor diamictite unit with dropstones occasionally punctuating the laminae. The laminated dropstone facies is also exceptional for its extremely negative δ13Ccarb values of ~-10‰, which is on the same order of magnitude as the Shuram-Wonoka anomaly, the most extreme anomaly recorded from the Neoproterozoic cap carbonates (Halverson et al., 2005). Even more perplexing, are the unique REE patterns associated with this extreme δ13Ccarb anomaly. This unit is characterised by REE patterns with consistent negative Eu anomalies, flat light (L) REE and highly variable heavy (H) REE that range from negatively to positively sloped. These patterns stand in stark contrast to REE patterns of samples from the overlying interlaminated carbonate and siltstone facies of the Espanola Formation. Carbonates from the overlying Espanola Formation have patterns with consistently depleted LREE and moderately enriched middle (M) REE, while HREE have a relatively flat pattern that transitions to a positive slope moving up stratigraphy. The relative depletion of LREE in these units that was not present in the underlying laminated dropstone facies indicates a stronger seawater signature, which may reflect a decrease in the influence of meltwater on the geochemical composition. Systematic sampling of the middle and upper Espanola Formation stratigraphy also produced a trend of upwards increasing δ13Ccarb values. Over approximately 110m of stratigraphy the δ13Ccarb values increase from ~4.5‰ to -2‰. This is another feature that has been noted from some Neoproterozoic cap carbonates and has been interpreted to be related to a marine regressive sequence (eg. Giddings and Wallace, 2009). 58 �Thus, the similarity between the Espanola Formation δ13Ccarb values and those of some Neoproterozoic cap carbonates supports the hypothesis that the Espanola Formation may have been formed under similar conditions as its Neoproterozoic counterparts. Figure 1: A ~75m section of stratigraphy sampled from drill hole E150-2 of the contact between the Bruce Formation and Espanola Formation. Red samples (lower REE plot) are from the laminated dropstone facies in the upper Bruce Formation. They have extreme negative δ13C values of approximately -10‰ and consistent negative Eu anomalies. The δ18O values do not show as anomalously low values but are noticeably lower than values further up stratigraphy and fall in the range of -21‰ to -20‰. The purple samples (upper REE plot) are from the interlaminated carbonate and siltstone facies of the lower Espanola Formation. These samples show a rapid trend upwards in δ13C values from ~-4.5‰ to -2‰ and they have REE patterns with consistent LREE depletion and moderate MREE enrichment. References Giddings, J.A., Wallace, M.W., 2009. Sedimentology and C-isotope geochemistry of the “Sturtian” cap carbonate, South Australia. Sediment. Geol. 216, 1–14. Halverson, G.P., Hoffman, P.F., Schrag, D.P., Maloof, A.C., Rice, A.H., 2005. Towards Neoproterozoic composite carbon-isotope record. Geol. Soc. Am. Bull. 117, 1181–1207. Kirschvink, J.L., 1992. Late Proterozoic low-latitude global glaciation - The Snowball Earth. In: Schopf, J.W., Klein, C. (Eds.), The Proterozoic Biosphere. Cambridge University Press, Cambridge, 51– 52. 59 �Precambrian Geology of the Western Schreiber–Hemlo Greenstone Belt MAGNUS, Seamus Ontario Geological Survey, 933 Ramsey Lake Road Sudbury, ON, P3E 6B5 Canada The Schreiber–Hemlo greenstone belt is located within the Wawa–Abitibi terrane of the Superior Province. The greenstone belt includes Neoarchean supracrustal and intrusive rocks that have been crosscut and unconformably overlain by Paleoproterozoic and Mesoproterozoic intrusive and supracrustal rocks of the Southern Province. Bedrock mapping in this area by the Ontario Geological Survey from 2015 to 2018 focussed on the Archean rocks of the western part of the Schreiber–Hemlo greenstone belt, with an emphasis on applying modern geochemical and geochronological techniques. The supracrustal rocks in the western Schreiber–Hemlo greenstone belt are arranged in an upright stratigraphy consisting of four distinct depositional packages, with chemical and clastic metasedimentary rocks along disconformable contacts (Figure 1). The oldest rocks in the greenstone belt are felsic and mafic metavolcanic rocks of Package A, deposited circa 2720 Ma (Davis and Sutcliffe 2017) in a volcanic arc environment. These are overlain by Package B, which is composed mainly of mafic metavolcanic rocks deposited in a “back-arc” volcanic environment. In the western part of the project area, Package B is overlain by Package C, which is composed mainly of mafic metavolcanic rocks deposited in an “oceanic plateau” volcanic environment. In the eastern part of the project area, Package B is overlain by Package D, which is composed of turbiditic wacke and mudstone deposited between 2696 and 2690 Ma (Fralick, Purdon and Davis 2006; Davis and Sutcliffe 2017). The chronostratigraphic relationship between packages C and D is unknown, as contacts between these packages have not been observed. The oldest felsic plutons that crosscut the supracrustal rocks are the circa 2690 Ma Terrace Bay and Steel River plutons (Kamo 2016). Regional ductile deformation likely started at this time, however, whether it began before or after emplacement of the plutons is uncertain. The circa 2667 Ma Santoy Lake pluton shows little evidence for ductile deformation along its margins, which suggests that regional ductile deformation ceased at approximately this time (Kamo 2016). Northwest ductile and brittle-ductile shear zones crosscut and displace all of the Archean rocks. Dikes of the Paleoproterozoic Matachewan, Biscotasing and Marathon dike swarms crosscut Archean rocks in the project area, and outliers of the base of the Paleoproterozoic Gunflint Formation unconformably overlie the Archean rocks at the west end of the project area, southwest of Schreiber. The Coldwell Alkalic Intrusive Complex intrudes the Archean rocks at the east end of the Schreiber–Hemlo greenstone belt. Alkalic diabase dikes crosscut the Archean rocks and the intrusive rocks of the Coldwell Alkalic Intrusive Complex and are believed to be related to volcanism during rifting associated with formation of the Keweenawan Midcontinent Rift. The Archean rocks host a variety of base metal and precious metal occurrences which have been the subject of exploration and limited mining activities for over a century. The circa 2720 Ma felsic metavolcanic rocks are correlative with rocks in the nearby Winston Lake and Manitouwadge areas that host past-producing Zn-Cu mines (Davis, Schandl and Wasteneys 1994; Zaleski, van Breemen and Peterson 1999). Gold mineralization is hosted in sheared and altered metavolcanic rocks and in veined and altered granitoid rocks. Proterozoic rocks in the north shore of Lake Superior region have potential to host magmatic sulphide and oxide mineralization including a variety of transitional metals and rare earth elements. 60 �Figure 1: Simplified geological map of the western Schreiber–Hemlo greenstone belt, highlighting the major Archean rock types, some of the stratigraphic younging indicators observed during this study, all of the U-Pb zircon geochronological data in the area, and the inferred fold axial traces. An inset figure outlines the inferred depositional packages A, B, C and D. Note that Proterozoic diabase dikes, which are abundant in the map area, are not shown for clarity. Abbreviations: DHR = Dead Horse Road, HWY 17 = Trans-Canada Highway 17, LLR = Long Lake Road. See references for ages. All UTM co-ordinates provided using NAD83 in Zone 16. References Davis, D.W., Schandl, E.S. and Wasteneys, H.A. 1994. U-Pb dating of minerals in alteration halos of Superior Province massive sulphide deposits: Syngenesis versus metamorphism; Contributions to Mineralogy and Petrology, v.115, p.427-437. Davis, D.W. and Sutcliffe, C.N. 2017. U-Pb geochronology by LA-ICPMS in samples from northern Ontario, internal report for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 131p. Fralick, P., Purdon, R.H. and Davis, D.W. 2006. Neo-Archean trans-subprovince sediment transport in southwestern Superior Province: sedimentological, geochemical and geochronological evidence; Canadian Journal of Earth Sciences, v.43, p.1055-1070. Kamo, S.L. 2016. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: Bedrock Mapping Projects, Ontario, Year 1: 2015-2016, internal report prepared for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. Zaleski, E., van Breemen, O. and Peterson, V.L. 1999. Geological evolution of the Manitouwadge greenstone belt and Wawa-Quetico subprovince boundary, Superior Province, Ontario, constrained by U-Pb zircon dates of supracrustal and plutonic rocks; Canadian Journal of Earth Sciences, v.36, p.945-966. 61 �Pilot study: Using ambient noise passive seismic surveys for Ni-Cu-PGE mineral exploration at the Marathon PGM-Cu deposit, Marathon, Ontario MCBRIDE, J.1, GOOD, D.2, HOLLIS D.3, and AARNDT, N.3 1 Stillwater Canada Inc. 90 Peninsula Rd. Marathon, ON P0T 2E0, Canada 2 Department of Earth Sciences, University of Western Ontario, London, ON N5A 5B7, Canada 3 Sisprobe, 38240 Maylan, France Active seismic surveys are a powerful geophysical tool for exploring to significant depth, and are commonly used in the oil and gas industry. However, because of the high cost and environmental impact associated with conducting a seismic survey, this method is rarely used for mineral exploration. Nevertheless, with the increased difficulty of finding economic mineral deposits, exploration companies continue to look deeper and there is a growing need to develop cheaper methods with less environmental impact to do so. Passive seismic methods currently being tested by SISPROBE Inc. at the Marathon deposit have the advantage of being a low impact and low-cost method for examining velocity contrast in geologic units to depths below surface approaching 1 km. Passive seismic methods use ambient noise generated from the natural environment. At Marathon, the dominant noise source is wave action in Lake Superior with a minor contribution from waves in the North Atlantic Ocean. Additional noise is generated by traffic on the nearby highway and railway. The use of autonomous seismic data recorders allows for flexibility when designing sensor arrays, which is necessary in remote or environmentally sensitive areas that include challenging topography. The Coldwell Complex is approximately 25 km in diameter and is composed of three centers of predominantly alkaline magmatism that intruded the Archean greenstone terrane (Mitchell and Platt, 1977) along the northern margin of the Midcontinent rift between 1108 and 1094 Ma (Heaman et al., 2007). Centre I is composed of augite syenite, quartz syenite and the Eastern Gabbro Suite. The Eastern Gabbro Suite outcrops along the eastern and northern margin of the complex and is composed of numerous gabbroic to ultramafic intrusions of the Layered and Marathon Series that cut a 1 km thick pile of metabasalt (Good et al., 2015; and Good and Lightfoot, 2019). Mineralization at the Marathon PGM-Cu deposit is hosted by Two Duck Lake gabbro and ultramafic rocks of the Marathon Series. The Marathon PGM-Cu deposit is an ideal site to test the passive seismic technique because of the extensive geological database and the distinct petrophysical property contrast exhibited by the various syenites and gabbros of the complex, and the underlying Archean metavolcanic rocks of intermediate composition. A preliminary noise survey was completed in 2017 to test ambient source signal-to-noise ratio. It was determined that wave action from Lake Superior generates sufficient ambient noise to proceed to a production scale survey. In 2018, a production scale survey was completed with 90 sensors deployed at 300 m spacing in an array that is elongated parallel to wave propagation in order to maximize signal pairs. 62 �The geophones used were GSX-1 single channel units, which collected data in the vertical direction. They recorded data every 4 ms for a total of 26 days (Hollis, 2018). The density and P-wave velocities for representative samples of each lithologic unit at the deposit were measured at Western University. These measurements were used to constrain interpretations of lithological boundaries determined from the 3D velocity inversion model for the survey data. Augite syenite (Vp of 5500 m/s and Rho 2650 km/m3) overlies the Two Duck Lake gabbro (Vp 6200 m/s and Rho 3100 kg/m3) while the Archean metavolcanic footwall (Vp 5000 m/s and Rho 2800 kg/m3) lies below the gabbro. The ultramafic (Vp 6800 m/s and Rho 3500 kg/m3) units that host the mineralization occur as lenses and pods that are distinguishable from the gabbro units. The geological boundary between the Two Duck Lake gabbro and the Archean metavolcanic footwall was successfully resolved by the survey. The survey also identified a high-velocity anomaly down dip from the Marathon PGM-Cu deposit at a depth of 600 m. The anomaly has a velocity value that is representative of an ultramafic unit. To validate the velocity anomaly, a 6 km gravity line was completed over the area which confirmed a high-density body at depth. By combining both passive seismic and gravity methods along with the structural association of the anomaly along feeder conduits, the anomaly is interpreted to be an accumulation of dense minerals such as magnetite, apatite, olivine and sulfide in a conduit setting. The passive seismic technique therefore identified an exploration target at depth where previous electromagnetic and magnetic surveys had not. Passive seismic geophysics is an excellent technique for the mineral exploration industry as it brings considerable depth penetration with the advantages of 3D seismic imaging, at low cost while being sensitive to the environment. References Good D.J., Epstein R., McLean, K., Linnen R., and Samson, I., 2015. Evolution of the Main Zone at the Marathon Cu-PGE Sulphide Deposit, Midcontinent Rift, Canada: Spatial Relationships in a Magma Conduit Setting. Economic Geology, v. 110, pp. 983-1008. Good D.J. and Lightfoot P.C., in press, Significance of Metasomatized Lithospheric Mantle in the Formation of Early Basalts and Cu-PGE Sulfide Mineralization in the Coldwell Complex, Midcontinent Rift, Canada, Canadian Journal of Earth Sciences, 2019. Heaman, L., Easton, M., Hart, T., Hollings, P., McDonald, C., and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario: Canadian Journal of Earth Sciences v. 44, pp. 1055-1086 Hollis D., 2018, Marathon Passive Seismic Project, internal report, Sisprobe, 24 Allee des Vulpains, 38240 Meylan, France. Mitchell, R., and Platt, R., 1977. Field guide to the aspects of the geology of the Coldwell alkaline complex: Institute on Lake Superior Geology, Technical Report 63 �The Wolf River Orogeny: Geon 14 Magmatism, Sedimentation, and Deformation in the Southern Lake Superior Region MEDARIS, L. G. Jr.1, MALONE, D. H.2, HILL, G. C.2, SINGER, B. S.1, JICHA, B. R.1, VAN LANKVELT, A.3, WILLIAMS, M. L.3, and REINERS, P. W.4 1 Department of Geoscience, University of Wisconsin–Madison, Madison, WI 53706 Department of Geography, Geology, and the Environment, Illinois State University, Normal, IL 61790 3 Department of Geosciences, University of Massachusetts–Amherst, Amherst, MA 01003 4 Department of Geosciences, University of Arizona, Tucson, AZ 85721 2 The Proterozoic Wolf River Batholith (WRB), which is the most prominent Precambrian geological feature in northeastern Wisconsin, was first described in 1975 by Van Schmus et al. and initially interpreted to represent an episode of anorogenic igneous activity by analogy with the classic Proterozoic rapakivi granites in Finland (Anderson & Cullers, 1978). Subsequently, it was recognized that the WRB is the local expression of a transcontinental belt of Geon 14 granites that were again interpreted to be anorogenic (Anderson, 1983). More recent investigations reveal that emplacement of these transcontinental Geon 14 granites along the eastern and southern margins of Laurentia was associated with an orogenic event involving continental arc magmatism, sedimentation, and deformation (Whitmeyer & Karlstrom, 2007; Daniel et al., 2013), certain aspects of which are now recognized as being related to the Wolf River event in Wisconsin. Magmatism The WRB underlies a minimum area of 1.45 x 104 km2 and consists predominantly of alkaline biotite granite and biotite–hornblende adamellite and subordinate quartz syenite, monzonite, and anorthosite (Anderson & Cullers, 1978). U–Pb zircon ages for the different plutons range from 1468 ± 4 to 1484 ± 2 Ma, with the main part of the batholith yielding an average crystallization age of 1476 ± 2 Ma (DeWayne & Van Schmus, 2007). To the west of the WRB in Marathon county, the Wausau syenite and Nine Mile granite plutons yield older crystallization ages of 1522 and 1506 Ma, respectively. Oxygen, Sm–Nd, and Lu–Hf isotopic data indicate that the WRB was derived from partial melting of the late Paleoproterozoic crust in the region (Anderson & Morrison, 2005; DeWayne & Van Schmus, 2007; Goodge & Vervoort, 2006). Sedimentation The Baldwin conglomerate occurs at the northeastern margin of the WRB, where it lies unconformably on the Geon 18 Macauley gneiss and Waupee metavolcanic and metasedimentary rocks and is intruded by the 1470 Ma Hager porphyry. The Baldwin conglomerate is polymict and chemically immature, containing clasts of the underlying lithologies set in a medium–grained arkosic matrix. A relative probability plot for detrital zircons in the Baldwin conglomerate displays a prominent Geon 14 (Wolf River) peak, subordinate Geon 16 (Mazatzal), Geon 17 (Yavapai), and Geon 18 (Penokean) peaks, and a single detrital zircon at 2690 Ma (Algoman) (Fig. 1). The maximum age of deposition (MAD) calculated from the youngest statistically homogenous population (MSWD ≤ 1.0) is 1458 ± 10 Ma. These results demonstrate that deposition of the Baldwin conglomerate was synchronous with crystallization of the WRB. Figure 1. Relative probability plot for detrital Deformation and recrystallization Evidence zircons in the Baldwin conglomerate for Geon 14 deformation associated with the WRB is best revealed by metasedimentary rocks of the post– Mazatzal Baraboo Interval. In the Baraboo Range, muscovite parallel to slatey cleavage in four samples 64 �of Seeley Slate yields 40Ar/39Ar cooling ages of 1473, 1483, 1493, and 1496 Ma (all with ± 3Ma), and muscovite decorating crenulation cleavage in Waterloo metapelite yields 1465 ± 7 Ma. In addition, cooling ages of 1472 ± 3, 1480 ± 11, and 1469 ± 11 Ma have been obtained for muscovite in breccia in the Baraboo Quartzite, in hydrothermal veins at the base of the quartzite, and in metamorphosed paleosol beneath the quartzite. Figure 3. U/Th–He ages for hematite in Baraboo metapelite Figure 2. Th map and U-Pb ages for monazite in Seeley Slate Monazite occurs as a detrital mineral in the Seeley Slate, and some grains exhibit new monazite rims that extend parallel to cleavage (Fig. 2). Electron probe microanalysis and dating of monazite were done using the UMass Ultrachron probe. Detrital monazite cores yield Penokean and Archean ages; rims yield a date of 1502 ± 30 Ma, comparable to the age of the WRB. In the Baraboo Quartzite, folded metapelite layers consisting largely of pyrophyllite contain tiny grains (50–100 m in diameter) of recrystallized hematite. Such hematite yields a mean U/Th–He age of 1507 ± 153 Ma (Fig. 3), which is consistent with the ages obtained for muscovite and monazite by other geochronologic methods. Note that the Baraboo Interval sedimentary rocks containing evidence for Geon 14 folding and recrystallization, e.g. the Baraboo and Waterloo quartzites, are located within the trans-continental belt of Geon 14 granites, whereas those located outside the transcontinental belt, e.g. the Sioux and Barron quartzites, are neither folded nor recrystallized. Despite the massive character of different Wolf River plutons and “anorogenic” appearance of the batholith itself, it is now clear that emplacement of the WRB was accompanied by Geon 14 sedimentation and deformation and can be viewed as an orogenic event. The Wolf River orogeny provides a link between the Pinwarian orogeny to the northeast and the Picuris orogeny to the southwest, thus completing the transcontinental extent of Geon 14 orogenesis in North America. References Anderson, 1983, GSA Memoir 161, 133–154; Anderson & Cullers, 1978, Precam. Res. 7, 287–324. Anderson & Morrison, 2005, Lithos 80, 45–60; Daniel et al., 2013, GSA Bull. 125, 1423–1441. DeWayne & Van Schmus, 2007, Precam. Res. 157, 215–234. Goodge & Vervoort, 2006, Earth Planet. Sci. Lett. 243, 711–731. Whitmeyer & Karlstrom, 2007, Geosphere 3, 220-259; Van Schmus et al., 1975, GSA Bull. 86, 907–914. 65 �The Importance of “Tablesetting” Intrusions in Creating Economic Ni-Cu-PGE Deposits in the Midcontinent Rift MILLER, Jim University of Minnesota Duluth (emeritus) and JDM GeoConsulting, Shuniah, ON (mille066@umn.edu) Some of the most promising targets for economic Ni- Cu-PGE sulfide deposits in the Lake Superior region are associated with small-scale ultramafic-mafic intrusions emplaced during early stages of the 1.1Ga Midcontinent Rift. While many of these intrusions share well documented attributes – small size, sub-horizontal conduit geometries, high grades and tenors of Ni-Cu-PGE sulfide ore, ultramafic host rock – one common attribute that is not so well known is the association of these mineralized intrusions with precursor intrusions. I refer to these earlier intrusions as “tablesetting” intrusions (TSI) as their emplacement appears to have played a major role in producing the well mineralized intrusions that followed. Before discussing what role TSI plays, the basic structural, lithologic and geochemical attributes of the TSI associated with four well-studied MCR ultramafic intrusions will be described. I am familiar with these intrusions through the MS thesis research of my UMD graduate students – Eagle (Mulcahy, 2018), Tamarack (Goldner, 2011), BIC (Foley, 2011), and Current Lake (Chaffee, 2015) - and through many years of discussions with exploration geologists such as Dean Rossell (Rio Tinto), Bob Mahin (Eagle/Lundin), Al MacTavish (MagmaMetal/Panoramic), and Geoff Heggie (Magma Metals/Panoramic). The discovery of the Eagle deposit in 2002 by Dean Rossell and his Rio Tinto/Kennecott crew in the Baraga Basin area north of Marquette, Michigan set off a flurry of exploration activity in the Lake Superior region that continues to this day. Eagle is the only MCR-related Ni-Cu-PGE deposit that has progress to active mining, which began in 2014, soon after the property was acquired by Lundin Mining. In 2015, continued exploration in the area revealed additional economic mineralization in the subhorizontal conduit of the nearby Eagle East intrusion. With total minelife of the Eagle and Eagle East deposits projected to end in 2023, the company is aggressively exploring for additional deposits in the area. One of the main vectoring tools being employed is to seek out pyroxenite dikes. As observed at both Eagle and Eagle East, weakly mineralized pyroxenite to melagabbro (PYX unit) occurs at the margins of the main peridotite body that hosts the bulk of the Ni-Cu-PGE mineralization. Weakly mineralized pyroxenite also occurs as xenoliths in well-mineralized peridotite - the HTBX and IBRX units (Mulcahy, 2018). It is hoped that tracing the occurrences of the tablesetting PYX rock type will lead to discovery of another mineralized peridotite body in the vicinity of Eagle/Eagle East. Concurrent with their exploration in Upper Michigan, RioTinto/Kennecott was seeking an Eagle-like occurrence in the Paleoproterozoic Animikie Basin in east-central Minnesota. This led to the discovery of significant Ni-Cu-PGE mineralization in 2008 in the Tamarack intrusion/deposit. The main zone of massive to semi-massive sulfide mineralization occurs in the “tail” section of the tadpole-shaped subhorizontal intrusion where two distinct peridodite bodies come into contact – 1) a deeper CGO unit, which is characterized by coarse cumulus olivine and significant intercumulus clinopyroxene and plagioclase, and 2) an overlying FGO unit, which is characterized by finer grained cumulus olivine and only a minor intercumulus component. Whereas Goldner (2011) concluded from petrographic and geochemical attributes that the CGO unit is the precursor intrusion, the local Rio Tinto/Kennecott crew has interpreted the FGO is the earlier intrusion. In either case, the mineralization is clearly focused where the two intrusive components come into contact. The Thunder Bay North PGE-Cu-Ni deposit associated with the Current Lake Intrusion was discovered by Magma Metals in 2006 near the occurrence of glacial boulders of well mineralized peridotite on the shoreline of Current Lake. Like Tamarack and Eagle, the mineralization is hosted by peridotite, and like Tamarack, it has a subhorizontal tadpole shape (chonolithic). However, it is intrusive 66 �into Archean granitoids and metasedimentary rocks rather than Paleoproterzoic black shales. Another significant difference is that the precursor rock is a strongly contaminated, commonly xenolith-rich lithology that Magma Metals termed the Hybrid Unit (with red and gray varieties). Petrographic studies by Chaffee (2015) determined this rock is a weakly mineralized quartz gabbro that is variably discolored by hematitic staining. Geochemical modelling also showed that the parental magma to the hybrid unit is a contaminated equivalent to the peridotitic magma that followed emplacement of the hybrid. The hybrid intrusions were emplaced in an orthogonal pattern of subhorizontal and subvertical dikes. The main mineralized peridotite tended to be emplaced at the intersections of the vertical and horizontal hybrid dikes to form chonolith-shaped bodies. The Bovine Igneous Complex (BIC) is a funnel-shaped intrusion that, like Eagle, was emplaced into the Paleoproterozoic Baraga Basin of Upper Michigan. Because it is one of the few ultramafic intrusions with surface exposures, exploration activity on BIC by Dean Rossell and his Rio Tinto/Kennecott crew began in the mid-90’s, though significant mineralization was not discovered until 2006. Detailed petrographic and geochemical studies by Foley (2011) on two drill core profiling the igneous stratigraphy of BIC showed it to be composed of two well differentiated intrusive cycles. The lower (early) sequence grades from an Ol cumulate upward to a Cpx+Ol cumulate. This is overlain by a cumulus reversal back to an Ol cumulate that grades upward to a Cpx+Ol cumulate, then an Ox+Cpx±Ol cumulate, and is capped by a Pl+Cpx+Ox cumulate. Both differentiated sequences show smooth cryptic layering of Mg/Fe ratio in olivine and augite, and Ca/Na ratio in plagioclase. Ni-Cu-PGE enriched sulfide mineralization occurs intermittently through the lower ultramafic sequence and at the basal contact between of the upper differentiated sequence. Although economic grades of mineralization appear to be lacking in these sequences, the possibility of finding more concentrated and higher tenor sulfide in the as yet undiscovered conduit to BIC seems high. The main take-aways from the observation made of these four mineralized ultramafic intrusions in regard to the role of the precursor “tablesetting” intrusions (TSI) are: 1) TSI are important for establishing the plumbing system for subsequent intrusions. 2) TSI serve to pre-heat and begin devolatilization of sulfide-bearing country rock, but because of rapid heat loss to cold country rock, they tend to generate little sulfide of low tenor. 3) Subsequent intrusions of hot ultramafic magmas into a pre-heated and structurally compromised country rock created by the TSI are able to cool slowly and create cumulate lithologies. Larger, more closed intrusions may become well differentiated (BIC), whereas in narrow chonolithic intrusions that are perhaps open to surface, large volumes of ultramafic magma can pass through resulting in the crystallization of uniformly primitive cumulates (Eagle, Tamarack, Current Lake). 4) In open chonolithic systems, the dynamic passage of large volumes of metal-rich ultramafic magma that bore through and inflate the precursor TSI, can upgrade the tenor of any early-formed sulfide in the TSI and any additional sulfide devolatilized by the new heat pulse. 5) Wherever the intrusive plumbing network creates subhorizontal sheets, channels, or tube-shaped (chonolith) conduits, this allows for gravitational concentration of enriched sulfide liquid. UMD MS Theses Chaffee, M., 2015, Petrographic and Geochemical Study of the Hybrid Rock Unit Associated with the Current Lake Intrusive Complex. Foley, D., 2011, Petrology and Cu-Ni-PGE Mineralization of the Bovine Igneous Complex, Baraga County, Northern Michigan. Goldner, B., 2011, Petrology and Cu-Ni-PGE Mineralization of the Tamarack Intrusion, Aitkin and Carlton Counties, Minnesota. Mulcahy, C., 2018, Emplacement and Crystallization Histories of Cu-Ni-PGE Sulfide-mineralized Peridotites in the Eagle and Eagle East Intrusions. 67 �Geochemical Vectoring Towards a Serpentinized Peridotite Chonolith, Eagle East Ni-Cu-Co-PGE Deposit, Upper Peninsula, Michigan NOWAK, Robert1, ESSIG, Espree1, MAHIN, Robert1 1 Eagle Mine Exploration, 200 Echelon Drive, Negaunee, MI 49866 Serpentinization is a low temperature (≤500°C), surficial to hypabyssal metasomatic process in which pyroxene [(Ca,Mg,Fe)2Si2O6] and olivine [(Mg,Fe)2SiO4] react with H2O +/CO2 to form hydrous silicates (serpentine) +/- hydroxides (brucite) +/- carbonates (dolomite, magnesite, calcite), and +/- Fe-oxides (magnetite) (Huang et al., 2017; Kelemen and Matter, 2008). These chemical reactions can result in the transfer of Mg2+, Ca2+, and Si4+, the oxidation of Fe2+ (Kelemen and Matter, 2008), and, under very reducing conditions, the formation of nickel alloys (awaruite (Ni3Fe; Preiner et al., 2018; Lawley, 2018). Pervasive serpentinization of peridotite can incorporate up to 13-15% H2O by weight and result in an estimated volume increase up to 40% (Schroeder et al., 2002; Shervais et al., 2005). This process can significantly alter the properties of ultramafic to mafic rocks, resulting in decreased density, seismic velocity (Miller and Christensen, 1997), and rheological strength (Escartin et al., 2001), in addition to an increase in magnetic susceptibility (Toft et al., 1990). The Eagle and Eagle East magmatic Ni-Cu-Co-PGE deposits, which formed during the Midcontinent Rift (MCR), are hosted within intensely serpentinized peridotite chonoliths. The aim of this study was to investigate whether a geochemical signature could be detected from drill core analyses outside the main serpentinized peridotite chonoliths and potentially utilized as an exploration vector toward mineralized peridotite. A pyroxenite sheet dike, which extends along strike and is crosscut by the Eagle East ore-hosting chonolith, was the focus of this study. Over fifty samples, collected from drill core intercepts of the pyroxenite sheet dike, were analyzed for major, minor, and trace elements (using ICP-MS and XRF methods) and utilized to generate 3-D models (using Leapfrog software). Pyroxenite intercepts ~400 meters away from any secondary intrusion (i.e. peridotite chonolith or gabbroic stock) were used as a baseline comparison to pyroxenite intercepts above and below the Eagle East peridotite chonolith. Pyroxenite samples above the Eagle East peridotite conduit have relatively enriched (on the order of 1 to 5 wt %) SiO2 and MgO values, and relatively depleted CaO contents (on the order of 1.5 to 2 wt%) relative to baseline pyroxenites. The enrichment and depletion trends become most pronounced ~50 meters above the flat-lying Eagle East conduit. Pyroxenite samples below the Eagle East chonolith contain relatively enriched SiO2, MgO, and CaO values, except for pronounced depletions within ~50 meters of the chonolith contact. Enrichment of nickel (on the order of 0.2 to 1 wt%) in the lower pyroxenites can extend up to 500 meters away from the lower chonolith keel contact (Fig. 1). The overall pattern of depletion proximal to the Eagle East chonolith is interpreted as resulting from near contact related serpentinization of the pyroxenites. The overall pattern of enrichment distal to intense serpentinization is interpreted as redistribution of these elements outside the zone of intense serpentization into less-altered pyroxenites. In summary, serpentinization is an important component to consider when modelling and interpreting the major and base metal element content of ultramafic and mafic rocks which can potentially host Ni-Cu-Co-PGE mineralization. The occurrence of proximal depletion, coupled with distal enrichment of MgO, CaO, and SiO2 may provide exploration criteria that could be 68 �used to vector towards serpentinized peridotite. The redistribution of Ni from serpentinized olivine, presumably into the mineral awaruite, displayed the most widespread detection halo (up to 500 meters outside the Eagle East system). The implications of this study could aid in determining the nickel prospectivity of a magmatic system and improve estimations on the potential size of a serpentinized system based on the scale of the geochemical halo observed. Figure 1: Long-section, looking southeast, showing nickel content (ppm) of pyroxenite samples projected onto the modelled pyroxenite plane. The modelled serpentinized Eagle East chonolith surface with massive- (red) and semi-massive sulfide ore-bodies (yellow) is also shown. References Escartin, J., Hirth, G. and Evans, B., 2001. Strength of slightly serpentinized peridotites: Implications for the tectonics of oceanic lithosphere. Geology, 29, 1023-1026. Huang, R., Lin, C., Sun, W., Ding, X., Zhan, W., Zhu, J., 2017. The production of iron oxide during peridotite serpentinization: Influence of pyroxene. Geoscience Frontiers, 8, 1311-1321. Kelemen, P.B., and Matter, J., 2008. In situ carbonation of peridotite for CO 2 storage. PNAS, 105, 17295-17300. Lawley, C., 2019. Gold and PGE mobility during serpentinization. PDAC technical session in: advances in mineral systems modelling of Ni-Cu-PGE and gold, v. 2019 Preiner, M., Xavier, J.C., Sousa, F.L., Zimorski, V., Neubeck, A., Lang, S.Q., Greenwell, H.C., Kleinermanns, K., Harun,T., McCollom,T.M., Holm, N.G., and Martin, W.F., 2018. Serpentinization: Connecting Geochemistry, Ancient Metabolism and Industrial Hydrogenation. Life, 41, 1-22. Schroeder, T., John, B. and Frost, B.R., 2002. Geologic implications of seawater circulation through peridotite exposed at slow-spreading mid-ocean ridges. Geology, 30, 367-370. Shervais, J.W., Kolesar, P. and Andreasen, K., 2005. A field and chemical study of serpentinization-Stonyford, California: Chemical flux and mass balance. International Geology Review, 47, 1-23. Toft, P.B., Arkani-Hamed, J. and Haggerty, S.E., 1990. The effects of serpentinization on density and magnetic susceptibility: a petrophysical model. Physics of the Earth and Planetary Interiors, 65, 137-157. 69 �Catchment Geology Correlation with Fish Otolith Microchemistry Across Disparate Glacial Till Depths in the Lake Michigan Basin PRICHARD, Carson G1., STUDENT, James J2., JONAS, Jory L3., WATSON, Nicole M1., and PANGLE Kevin L1. 1 Central Michigan University, Department of Biology, Mount Pleasant, Michigan, 48859 USA Central Michigan University, College of Science and Engineering, Center for Elemental and Isotopic Analysis, Mount Pleasant, Michigan, 48858 USA 3 Michigan Department of Natural Resources, Charlevoix Fisheries Research Station, Charlevoix, Michigan, 49720 USA 2 Fish otoliths are calcium carbonate boney-like structures found in fish ears that grow concentrically, and as such they preserve a chemical record of select environmental changes during life. This study used laser ablation inductively coupled plasma mass spectrometry (LAICP-MS) to record variations in signal intensities of magnesium (25Mg), calcium (43Ca), manganese (55Mn), copper (65Cu), zinc (66Zn), strontium (88Sr), barium (137Ba), and lead (208Pb) isotopes in steelhead (Oncorhynchus mykiss) otoliths. These signals were then converted to trace element concentrations (in ppm) along transects that represent a timespan when each fish resided in a particular catchment. A portion of the otolith data used in this study was previously used to build models that discriminate Wild-and Hatchery-Origin steelhead across the Lake Michigan Basin (Watson et al., 2018). The current study incorporated results from 538 WildOrigin steelhead otoliths that were collected in 2014 and 2015 (Prichard et al., 2019). A general introduction to otolith microchemistry applications, trace element uptake in otoliths, Michigan steelhead, and Michigan Geology and otolith microchemistry will be presented. Glacial deposits obfuscate much of the Michigan bedrock influence on stream chemistry, and this in turn influences the utility of Sr isotope systematics in lower peninsula Michigan streams as compared to bedrock dominated fluvial systems. A primary application of otolith microchemistry is distinguishing natal origins of individual fish within a mixed-stock fishery. Stocks must be distinguishable according to stockspecific microchemistry patterns, with accurate stock assignment contingent upon microchemistry assessment of all sources contributing to the mixed-stock fishery. However, otolith microchemistry signatures of individual fish, upon which classification models are built, likely represent only a portion of the variability that exists for the stocks corresponding to each natal source. To statistically infer expected otolith microchemistry patterns among unsampled catchment areas proximal to sampled areas, we tested the hypothesis that variation in catchment geology among 35 stream sites across the Lake Michigan basin is correlated with the variation in otolith microchemistry signatures of age-0 steelhead collected at those sites. Matrices of Mahalanobis distances between all pairs of individual fish were calculated for each of the following: (1) assignment scores from discriminant function analysis of the variation among sites based on otolith microchemistry, and (2) the geology (bedrock age, bedrock lithology, and 70 �surficial geology) underlying the catchments upstream of each of the sites where fish were sampled. Based on Mantel tests, these matrices were found to be significantly correlated, indicating that age-0 steelhead that exhibit greater differences in otolith microchemistry signatures tended to come from sites exhibiting greater differences in catchment geology. Surficial geology alone was more correlated with otolith microchemistry than bedrock age, bedrock lithology, or any combinations of the three geological datasets. The significant relationship between geology and otolith microchemistry, although weak, supports tenuous hydrologic and geologic bases for delineating natal source geographic boundaries. References Prichard, C.G., Student, J. J., Jonas, J. L., Watson, N. M., and Pangle, K. L., (2019) Geologic variability underlying stream catchment areas correlates with fish otolith microchemistry across disparate glacial till depths, Fisheries Research, submitted Dec., 2018 and is currently under revision. Watson, N. M., Prichard, C.G., Jonas, J. L., Student, J. J., and Pangle, K. L., (2018) Otolith ChemistryBased Discrimination of Wild- and Hatchery-Origin Steelhead across the Lake Michigan Basin, North American Journal of Fisheries Management, ISSN: 0275-5947 DOI: 10.1002nafm.10178. 71 �Using graphitic sedimentary rock geochemistry as an indicator of gold potential in the Shebandowan greenstone belt, northwestern Ontario PUUMALA, Mark Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, Resident Geologist Program, Suite B002, 435 James Street South, Thunder Bay, Ontario, P7E 6S7 Graphitic sedimentary rocks are a common feature of Archean greenstone belts. Due to their high carbon content, these rocks tend to be more metalliferous than non-carbonaceous sedimentary rocks. They also act as strong reducing agents to hydrothermal fluids and can sequester metals and other elements from those fluids (Barrie 2004). Springer (1985) noted that graphitic argillites in the Abitibi greenstone belt often contain anomalous concentrations of gold (up to 0.5 ppm Au), and that much higher concentrations (up to 15 ppm Au) can be found in graphitic argillites that show evidence of hydrothermal alteration (e.g., quartz veining and carbonate alteration). Given their relative abundance in Archean greenstone belts and their response to gold-bearing hydrothermal fluids, the geochemistry of graphitic sedimentary rocks should provide information to assist in the search for mesothermal gold deposits. Detailed geochemical studies completed in the Abitibi greenstone belt by Barrie (2004) demonstrated that graphitic argillite proximal to the Owl Creek, Hoyle Pond, Holloway and HoltMcDermott mines typically contains elevated concentrations of gold (Au), arsenic (As), antimony (Sb) and mercury (Hg). Based on the results of this work, Barrie (2004) developed a method of calculating a hydrothermal alteration index that is based on concentrations of these elements and is normalized to graphitic and carbonaceous (non-carbonate) carbon (C*) and sulphur (S). Normalization of the data accounts for the likelihood that the degree of metal sequestration from hydrothermal fluids will be proportional to the graphitic/carbonaceous carbon and sulphur contents of the rock. The alteration index equation is as follows: log (Au x Hg x As x Sb)/(C* x S); where concentrations of Au and Hg are in parts per billion, concentrations of As and Sb are in parts per million and concentrations of C* and S are in weight %. Alteration index (AI) values of >6.5 were deemed to be very significant and indicative of sample collection within 1 km of ore, while AI values <5.5 were considered insignificant. During the 2017 and 2018 field seasons, staff of the Thunder Bay Resident Geologist Office collected 72 samples of graphitic sedimentary rock from various locations in the Shebandowan greenstone belt west of the City of Thunder Bay. The program included the collection of outcrop samples and drill core samples. Drill core was obtained from the Ontario Geological Survey’s Thunder Bay and Conmee Township core repositories. The purpose of this sampling was to test the applicability of the graphitic argillite gold alteration index method of Barrie (2004) as an exploration targeting tool in the Shebandowan greenstone belt, and to establish a geochemical database for graphitic sedimentary rocks. Samples were analysed by the Ontario Geological Survey Geoscience Labs in Sudbury for the same comprehensive suite of major, minor and trace elements that were included in the Abitibi greenstone belt studies of Barrie (2004). This paper will focus on the work that was completed in 2017 (48 samples) near known gold and base metal occurrences, as results are still pending from the 2018 sampling program. As shown on Figure 1, Alteration index (AI) values exceeding 6.5 were obtained from samples collected near three known gold prospects located in the Shabaqua area (West Zone, Bylund and South Zone). No highly significant AI values were obtained from samples collected proximal to volcanogenic massive sulphide (VMS) or ultramafic rock-hosted Ni-Cu occurrences located further to the south in Conmee, Adrian, Sackville and Aldina townships. 72 �The highest AI value was obtained from an outcrop grab sample collected at the West Zone. Two more West Zone samples (1 outcrop and 1 drill core) also displayed elevated AI values. Gold at the West Zone is hosted in 2 brecciated, silicified and sulphide mineralized chert horizons. These horizons are both approximately 3 m wide and have assayed up to 6.87 g/t Au over 3.05 m. Anomalous AI values of 7.31 and 6.43 were obtained from two drill core samples collected proximal to the Bylund gold prospect. Gold mineralization on the Bylund property occurs in a 125 m wide zone of carbonate-altered rocks and stockwork quartz-carbonate veins. The anomalous AI value near the South Zone gold occurrence was obtained from a surface grab sample collected from a historic exploration trench. There are no known surface gold showings in proximity to this sample location. However, it is located approximately 20 m from the collar of the diamond drill hole that intersected the South Zone gold mineralization. The results of this study are consistent with the findings of Barrie (2004) and demonstrate that graphitic mudstone geochemistry can be used as a gold exploration targeting tool in the Shebandowan greenstone belt. The Bylund-West Zone-South Zone corridor in the Dawson Road Lots area has been identified as a high priority gold exploration target. Figure 1. Map illustrating gold alteration index (AI) values for graphitic sedimentary rock samples collected in the vicinity of the South Zone, West Zone and Bylund gold showings near Shabaqua, Ontario. AI values of >6.5 suggest that the sample site may be located within 1 km of a significant gold mineralized structure. Map grid is provided in UTM NAD83, Zone 16 co-ordinates. References Barrie, C.T. 2004. Geochemistry of exhalates and graphitic argillites near VMS and gold deposits, an Ontario Mineral Exploration Technologies (OMET) project; C.T. Barrie and Associates Ltd., Ottawa, ON, 126p. Springer, J. 1985. Carbon in Archean rocks of the Abitibi belt (Ontario-Quebec) and its relation to gold distribution; Canadian Journal of Earth Sciences, v.22, p.1945-1951. 73 �Wawa, undercover: Bedrock geologic and bedrock topographic mapping in north-central Minnesota RADAKOVICH, Amy1, CHANDLER, Val1, and JIRSA, Mark1 1 Minnesota Geological Survey, 2609 Territorial Road West, St. Paul, MN 55114 Recently published bedrock geology and bedrock topography maps for four counties in north-central Minnesota (Chandler and Radakovich, 2018; Jirsa and Chandler, 2016; Radakovich and Chandler, 2016a, b, c; Radakovich and Chandler 2018a, b, c) serve as a case study for mapping Precambrian geology in areas of almost complete cover by glaciogenic sediment. Production of the maps therefore relied heavily on data from geophysical investigation methods; successes and challenges are discussed herein. Some exploration drill core and minimal outcrop data locally guided bedrock geology mapping; however, aeromagnetic and gravity data proved to be the most useful tools for deciphering bedrock composition in most of the four county area. Basement geology (Fig. 1A) consists chiefly of Archean metavolcanic, metasedimentary, and metaplutonic rocks of the Wawa subprovince, intruded by a suite of northwest-trending Paleoproterozoic mafic dikes. Geophysical modeling refined the characterization of the Archean Leech Lake Structural Discontinuity and several buried Archean iron formations across the study area. Younger sedimentary rocks of the Paleoproterozoic Animikie Group overlie the basement bedrock in several separate, formerly-continuous basins in eastern parts of the study area. Sparse drilling data indicate poorly consolidated sedimentary strata overlying bedrock of all ages, particularly in and along topographic lows in the Precambrian surface. Pollen analysis verified a Cretaceous age for these poorly consolidated sedimentary rocks. Bedrock topographic surfaces (Fig. 1B) were hand-contoured based on limited bedrock elevation data from rare bedrock outcrops, exploration drill hole records, and drilling records of the small percentage of wells that reached the bedrock surface. As a result of the paucity of direct bedrock elevation information, a significant amount of data from passive seismic and conventional seismic soundings allowed bedrock elevation to be inferred over a large portion of the four-county area. Depth to bedrock calculations indicate that several hundred feet to as much as over 1000 feet of Quaternary glacial sediment covers the bedrock surface across most of the region. In some locations, bedrock composition and structure appear to have played a role in the development of paleo drainages on the bedrock surface; in others, drainage seems to have been less affected by apparent bedrock composition. One of the most difficult obstacles to depicting the bedrock topography was recognition of Cretaceous bedrock between the Precambrian weathering surface and the bottom of the Quaternary sediment. This poorly consolidated material is generally transparent to geophysical methods and rarely recognized by well drillers. 74 �A B Figure 1. A) Bedrock geologic map of Wadena, Becker, Hubbard, and Cass (WaBeHuCa) Counties in Minnesota. Includes Archean, Paleoproterozoic, and Cretaceous strata. A full legend for bedrock units can be obtained in the referenced publications and will be discussed during the talk. B) Bedrock topographic map of WaBeHuCa counties. Sun illumination angle 315°, Sun elevation 45°. 5x vertical exaggeration. References Chandler, V.W., and Radakovich, A.L., 2018, Bedrock Geology, pl. 2 of Lusardi, B.A., project manager, Geologic atlas of Hubbard County, Minnesota: Minnesota Geological Survey County Atlas C-41, pt. A, 6 pls., scale 1:100,000. Jirsa, M.A., and Chandler, V.W., 2017, Bedrock Geology, pl. 2 of Bauer, E.J., project manager, Geologic atlas of Becker County, Minnesota: Minnesota Geological Survey County Atlas C-42, pt. A, 6 pls., scale 1:100,000. Radakovich, A.L., and Chandler, V.W., 2016a, Bedrock topography and depth to bedrock, pl. 5 of Lusardi, B.A., project manager, Geologic atlas of Wadena County, Minnesota: Minnesota Geological Survey County Atlas C-40, pt. A, 5 pls., scale 1:200,000. ------ 2016b. Bedrock Geology, pl. 2 of Lusardi, B.A., project manager, Geologic atlas of Wadena County, Minnesota: Minnesota Geological Survey County Atlas C-40, pt. A, 5 pls., scale 1:100,000. ------ 2016c, Bedrock topography and depth to bedrock, pl. 6 of Bauer, E.J., project manager, Geologic atlas of Becker County, Minnesota: Minnesota Geological Survey County Atlas C-42, pt. A, 6 pls., scale 1:200,000. ------ 2018a, Bedrock Topography and Depth to Bedrock, pl. 6 of Lusardi, B.A., project manager, Geologic atlas of Hubbard County, Minnesota: Minnesota Geological Survey County Atlas C-41, pt. A, 6 pls., scale 1:200,000. ------ 2018b, Bedrock Topography and Depth to Bedrock, pl. 6 of Lusardi, B.A., project manager, Geologic atlas of Cass County, Minnesota: Minnesota Geological Survey County Atlas C-43, pt. A, 6 pls., scale 1:200,000. ------ 2018c, Bedrock Geology, pl. 2 of Lusardi, B.A., project manager, Geologic atlas of Cass County, Minnesota: Minnesota Geological Survey County Atlas C-43, pt. A, 6 pls., scale 1:200,000. 75 �Mesoarchean Chemical Sedimentary Rocks of Northwestern Ontario: Implications for Hydrosphere Composition in Deep Time 1RAMSAY, Brittany, 1FRALICK, Philip, 1BIELSKI, Paul, 2HOMANN, Martin, SANSJOFRE, Pierre2, and LALONDE, Stefan2 1 Department of Geology, Lakehead University, Thunder Bay, Canada, bjramsay@lakeheadu.ca European Institute for Marine Studies, CNRS-UMR6538 Laboratoire Géosciences Océan, Brest, France 2 The 2.88 Ga carbonate sediments on Woman Lake, within the Uchi Subprovince of the Superior Province, preserve a chemical record of the Archean hydrosphere. Chemical sediments act as proxies for ancient waters by incorporating rare earth elements (REE) and isotopic signatures into their crystal lattice as they precipitate, thereby documenting the chemical composition of the waters from which they formed (Webb et al., 2009). Comparing the sedimentologic characteristics of Archean units to modern analogues enable us to determine the depositional environments of the past. Detailed stratigraphic columns linked with geochemical and isotopic data permit a more complete understanding of the depositional environment and evolutionary processes occurring at this early time in Earth’s history. At the base of the carbonate platform, lying atop rhyolitic Archean basement, is a massive carbonate grainstone unit, interbedded with minor crinkly-stratiform silicified microbial mats. A sharp contact separates them from a thrombolite unit (TB), composed of discontinuous and clotted laminations of dark, organic rich carbonate and white carbonate cement. Up section within the TB unit colloform stromatolites develop. This unit transitions into a stromatolitic unit that is comprised of 12-15cm thick carbonate grainstone (CG) alternating with 5-8cm thick pustular stromatolites (PS) for ~5m. Geochemically interesting trends reminiscent of both Archean and modern oxygenbearing signatures are evident in shale-normalized REE spectra (Fig. 1). REE concentrations were determined from weak acetic acid leaches by ICP MS and laser ablation-ICP-MS at the European Institute for Marine Studies (IUEM). The CG’s display distinct negative Ce anomalies and slightly positive Eu anomalies while the PS show weaker negative Ce anomalies and no Eu anomaly (Fig. 1b). The TB unit displays consistent spectra with negligible Ce anomalies and positive Eu anomalies (Fig. 1a). Laser ablation-ICP-MS was used to obtain REE spectra exclusively from the TB’s white calcite cements (Fig. 1c), which show a more pronounced Ce anomaly. Stable C and O isotopes were also analyzed at IUEM. The PS have slightly lower δ 13C values compared to the CG and TB (Fig. 2), however all samples fall within a relatively restricted range (-1.19 to 1.22‰). Positive Eu and negative Ce anomalies are generally accepted as robust indicators of the influence of hydrothermal fluids and the presence of free oxygen, respectively (Derry and Jacobson, 1992). They are unique among REE in that they have two possible valence states (Eu2+,3+, and Ce3+,4+). In high temperature hydrothermal fluids, Eu3+ is reduced to Eu2+, which renders it more soluble than its trivalent neighbors, a process that enriched Archean seawater with Eu and imparted a positive Eu anomaly on precipitating carbonates. Negative Ce anomalies result from the oxidation of Ce3+ to Ce4+ and the subsequent removal of the less soluble Ce4+ from solution. Once oxidized, Ce readily adsorbs onto particulate matter, removing it from solution and permitting Ce depleted chemical precipitation. Stable carbon isotopes in marine carbonate rocks are widely used as proxies for carbon cycling through time and are commonly employed to track organic carbon burial, which 76 �preferentially sequesters 12C and leads to 13C enrichment in residual dissolved inorganic carbon (Schidlowski, 2001). Throughout most of earth’s history δ13C varies only slightly from 0‰ (Veiser, 2001), and Woman Lake carbonates are no exception. The PS are slightly more negative compared to the CG and TB, likely due to the PS being more abundant in organic matter. The geochemical anomalies and isotopic signatures present at Woman Lake seem to indicate that the carbonates precipitated from two different fluid sources. The positive Eu anomaly suggests a hydrothermal source which is characteristic of Archean seawater, while the Ce anomaly suggests fluids interacted with free oxygen. Stratigraphically, TB’s precipitated first and contain a positive Eu anomaly (Fig. 1a), suggesting they precipitated within seawater. The overlying Stromatolitic Unit (PS and CG) contain negative Ce anomalies (Fig. 1b), which imply that they were deposited nearshore, where the PS could grow and interact with oxygen-bearing freshwater. The TB cements contain a more pronounced Ce anomaly compared to the whole rock composition (Fig.1c). It is possible that the cements inherited the Ce anomaly from the overlying pustular stromatolites and grainstones. If they were subaerially exposed to rain, they may have partially dissolved and percolated through CG, reprecipitating with negative Ce anomalies in spaces within the TB producing the clotted fenestral appearance. It is also possible that the pustular stromatolites produced locally oxic conditions while the thrombolites were not as capable. References Derry, L.A., Jacobsen, S.B., (1990) The chemical evolution of Precambrian seawater: Evidence from REEs in banded iron formations. Geochim. Cosmochim. Acta 54, 29652977. Schidlowski, M. 2001. Carbon isotopes as biogeochemical recorders of life over 3.8 Ga of Earth history: Evolution of a concept. Precambrian Res., v. 106, p. 117–134. Veizer, J., 2003. Isotopic evolutions of seawater on geological time scales: sedimentological perspective, in Lentz, D.R. ed., Geochemistry of Sediments and Sedimentary Rocks: Evolutionary Considerations to Mineral DepositForming Environments: Geological Association of Canada, GeoText 4, p. 53-68. Webb, G., Nothdurft, L., Kamber, B., Kloprogge, T., and Zhao, J. 2009. Rare earth element geochemistry of scleractinian coral skeleton during meteoric diagenesis: a sequence through neomorphism of aragonite to calcite. Sedimentology, vol. 56, p. 14331463. 77 �Precambrian Geology of the Eastern Shebandowan Greenstone Belt - Insights into Stratigraphy and Structural History RATCLIFFE, Laura M.1 1 Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, Sudbury, Ontario P3E 6B5 The Shebandowan greenstone belt (SGB), located in the Wawa-Abitibi terrane of the Superior Province, extends 150 km west from Thunder Bay to Quetico Provincial Park and has an arcuate shape. It is bordered by the Quetico Subprovince to the north and wraps around the Northern Light–Perching Gull Lakes batholithic complex to the south. Paleoproterozoic sedimentary rocks overlie the SGB’s southeastern extent. This presentation reports on a multiyear project by the Ontario Geological Survey focused on updating the bedrock geology of the eastern part of the SGB, and new insights into the stratigraphy and structural history of the eastern SGB are explored. This work builds on previous work by Shegelski (1980), Williams et al. (1991), Berger (1993) and Corfu and Stott (1998). The Shebandowan greenstone belt contains a succession of supracrustal rocks and their syn-eruptive intrusive equivalents. Geochronological data and previous geological studies have defined 3 main supracrustal assemblages: the Greenwater assemblage (circa 2720 Ma), the Kashabowie assemblage (circa 2695 Ma), which is not currently recognized in the eastern SGB, and the Shebandowan assemblage (circa 2690 to 2680 Ma) (Corfu and Stott 1998). Based on the interpretation of Corfu and Stott (1998), the SGB has undergone 2 main stages of deformation (D1 and D2). D1 deformation occurred at approximately 2695 Ma, and is thought to have tectonically imbricated rocks of the Greenwater and Kashabowie assemblages across the SGB, this event is poorly understood in the eastern SGB. D2 deformation is constrained between 2685 and 2680 Ma and thought to record oblique northwest-directed compression. Regionally, the emplacement of sanukitoid plutons between 2685 and 2680 Ma is thought to have occurred during the waning of D2. In the eastern part of the SGB the Greenwater assemblage (circa 2720 Ma) comprises predominately massive, aphyric mafic volcanic flows with minor aphyric, pillowed and locally variolitic mafic flows. Among the mafic volcanic flows are thin layers of felsic and ultramafic volcanic rocks 100 to 750 m thick, as well and thin 100 m layers of terrigenous-clastic sedimentary rocks. Syn-eruptive intrusive mafic and ultramafic rocks occur throughout the Greenwater assemblage as sills and dikes. In the eastern part of the SGB the Shebandowan assemblage (circa 2690 Ma) is comprised of predominately intermediate volcaniclastic to epiclastic, heterolithic, amphibole- and plagioclase-phyric tuff, lapilli tuff, tuff breccia and course tuff breccia and/or terrigenous-clastic wacke, siltstone and conglomerate. The conglomerate commonly contains sedimentary fragments. The contact between the Greenwater and Shebandowan assemblages regionally has been interpreted to be an unconformity (Corfu and Stott 1998), however it is not been clearly demonstrated in outcrop. Mapping as part of this project has identified a distinct lithostratigraphic unit separating rocks from the Greenwater and Shebandowan assemblages. The 1 to 1.5 km thick (in plan view) “boundary zone” comprises intermediate tuffs and flows and/or wacke to siltstone, interlayered with chemical sedimentary rocks and lenses of conglomerate (containing chemical sedimentary rocks ± mafic, ± ultramafic, ± felsic volcanic fragments). The “boundary zone” may be interpreted as the inferred lower stratigraphic unit of the Shebandowan assemblage or as the rocks deposited during a transitional period between the Greenwater and Shebandowan assemblages. Work is ongoing to evaluate the geologic context of this distinctive unit and its significance with respect to the stratigraphy of the SGB. Some new constraints on the timing of deformation in the eastern SGB are provided by local outcrop observations and targeted geochronological analyses. An outcrop where structural relationships are well 78 �exposed consists of a wacke deposited after 2694±3 Ma (Davis, Ménard, and Sutcliffe 2018), that is assigned to the Shebandowan assemblage intruded by a set of tonalite dikes emplaced at 2682 ± 2 Ma (Davis, Ménard, and Sutcliffe 2018). Bedding and a layer parallel foliation in the wacke are folded and the folding is cross-cut by the tonalite dikes (Photo 1A). The tonalite dikes are also folded and boudinaged (Photo 1B), and finally a weakly penetrative cleavage overprints the previously described features. These observations indicate there were multiple phases of deformation affecting the Shebandowan assemblage rocks and that deformation continued past the emplacement of the dikes circa 2680 Ma. Corfu and Stott (1998) considered the final phase of deformation in the SGB to be between 2685 Ma and 2680 Ma and that tectonic activity was quiescent after 2680 Ma in contrast to the adjacent Quetico Subprovince, and other greenstone belts farther east in the Wawa Subprovince, where deformation is recorded after 2680 Ma. However, the previously described outcrop observations show that there were multiple phases of deformation in the SGB after circa 2680 Ma. Thus, tectonic activity in the eastern SGB continued for longer than previously interpreted. Photo 1: Photographs from a well exposed outcrop showing structural relationships. Two geochronology samples were analyzed from this location and ages are displayed (Davis, Ménard, and Sutcliffe 2018) (290782E 5363887N). Photo A) Folded primary layering and layer parallel schistosity (S0 and S1) cross cut by a tonalite dike. The dike is outlined by the black dashed line. The primary layering and layer parallel schistosity is indicated by the white dashed line. The folding event preceding dike emplacement is annotated in in white and black (F1). Photo B) Tonalite dike emplaced in a thinly bedded wacke and siltstone is folded and boudinaged. The dike is outlined by the black dashed line. The primary layering and layer parallel schistosity (S 0 and S1) is indicated by the white dashed line. The folding event following dike emplacement is annotated in white and black (F2). Compass is 22 cm long including sighting arm. The UTM co-ordinates are provided using NAD83 in Zone 16. References Berger, B.R. 1993. Geology of Adrian and Marks townships; Ontario Geological Survey, Open File Report 5862, 90p. Corfu, F. and Stott, G.M. 1998. Shebandowan greenstone belt, western Superior Province: U/Pb ages, tectonic implications, and correlations; Geological Society of America Bulletin, v.110, p.1467-1484. Davis, D.W., Ménard, J. and Sutcliffe, C.N. 2018. U-Pb geochronology of samples from northern Ontario, Part B: LA-ICP-MS; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 94p. Shegelski, R.J. 1980. Archean cratonization, emergence and red bed development, Lake Shebandowan area, Canada; Precambrian Research, v.12, p.331-347. Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L. and Sage, R.P. 1991. Wawa Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.485-541. 79 �High-technology metals in ore-forming environments and their signature in volcanichosted sulfide mineralization in northern Minnesota and Wisconsin. SCHARDT, Christian and DAVID, Mady Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 1114 Kirby Dr. Duluth, MN 55812 While the use of high-technology metals (HTMs), such as In, Ge, Ga, and Tl, is increasing in essential industrial applications and renewable energy technologies, our understanding of the sourcing and accumulation of these elements is insufficient. This applies to both their general distribution in various geological environments, their sourcing by ore-forming processes, and their deposition in selected ore deposits from which they are being mined. Typical concentrations for these metals are very low in most rock types (0.1 - 2 ppm; e.g., Terashima, 2001) and may reach concentrations > 1 % in certain ore deposit types (Murao et al., 2008; Kampunzu et al., 2009) by substituting for common metals (Zn, Cu, Sn) in familiar ore mineral such as sphalerite, chalcopyrite, or stannite (Johan, 1988, Pavlova et al, 2015). The formation of ore deposits showing elevated values of In, Ge, Ga, and Tl (volcanichosted massive sulfides, granitic tin deposits, MVT deposits) are relatively well understood but it is unclear why these metals do not accumulate in other ore-forming environments, i.e. SEDEX, SSC, or porphyry deposits, known to contain elevated concentrations of other HMTs. Little research has been conducted into the general thermodynamic behavior or potential enrichment mechanisms and there is no organized database of the concentration of these metals in various geological settings or their sourcing in ore-forming systems. Previous work (Schardt and David, 2018) initiated data collection and analysis of these metals in common rock types and an interpretation of potential sources. In this study, the database has been expanded to include most common ore-forming environments to better understand commonalities and differences with regards to selected HMTs and evaluate volcanic-hosted massive sulfide signatures from northern Minnesota (Vermilion district) and Wisconsin (Penokian Volcanic Belt). Figure 1 plots all available data (whole rock, mineral, ore material, alteration) for the most common ore deposit types and compares them to available data from the Vermilion district as well as known volcanic-hosted massive sulfide deposits in the Penokean Volcanic Belt of Wisconsin. Except for Tl, both environments show low average concentrations compared to similar formation environments (V). Data would suggest that either a) hydrothermal processes unrelated to volcanogenic massive sulfide formation (e.g., lower- temperature SEDEX, epithermal, gen. hydrothermal, SSC, MVT) are more efficient at sourcing and concentration HTMs, or b) the difference in host/source rock has a significant influence on the ability of the ore-forming system to source and accumulate HTMs. Most rock types have very similar HTM concentrations, except for Ga, which is significantly more abundant in volcanic rocks (~ 1 ppm vs. ~ 20 ppm). Higher fluid temperatures, such as those found in granitic ore systems (G in figure 1) do not exhibit any specific HTM enrichment but their whole-rock data (stippled line in figure 1) indicate that In may be more mobile under these conditions. This interpretation is speculative as available data are scattered and these elements are not routinely analyzed. While no systematic work has been conducted to assess the behavior of these metals a robust database is now available to study their distribution in hydrothermal systems and apply results to exploration efforts in Minnesota and Wisconsin. 80 �Figure 1. Plot of HTM concentrations as a function of ore-forming environment. Vertical bars represent minimum, average, and maximum values for each ore deposit type (see text). Solid lines denote trend in average concentrations (all data) while stippled line shows whole-rock concentrations only. L – low temperature (SSC, MVT); V – volcanogenic (all types of VHMS); H – hydrothermal (SEDEX, epithermal, hydrothermal), G – graniterelated (skarn, tin, porphyry). References Johan, Z, 1988, Indium and Germanium in the Structure of Sphalerite: an Example of Coupled Substitution with Copper. Mineralogy and Petrology, v. 39, p.211 – 229 Kampunzu, A.B., Cailteux J.L.H., Kamona, A.F., Intiomale, M.M., and Melcher, F., 2009, Sediment-hosted Zn–Pb– Cu deposits in the Central African Copperbelt, Ore Geology Reviews, v. 35, p. 263-297 Murao, S., Deb, M., and Furuno, M., 2008, Mineralogial evolution of indium in high grade tin-polymetallic hydrothermal veins - A comparative study from Tosham, Haryana state, India and Goka, Naegi district, Japan, Ore Geology Reviews, v. 33, p. 490-504 Pavlova, G.G., Palessky, S.V., Borisenko, A.S., Vladimirov, A.G., Seifert, T., and Phane, L.A. (2015) Indium in cassiterite and ores of tin deposits. Ore Geology Reviews, v. 66, p. 99–113 Schardt, C., and David, M., 2018, High-technology metal behavior in ore-forming environments and its implication for the Vermilion District, northern Minnesota, Proceedings of the Institute on Lake Superior Geology, v. 64, p. 91-92 Terashima, S. (2001) Determination of Indium and Tellurium in Fifty Nine Geological Reference Materials by Solvent Extraction and Graphite Furnace Atomic Absorption Spectrometry. Geostandards Newsletter, v. 25, p. 127 - 132 81 �Geochemistry of Archean Gneisses in Dickinson County, Northern Michigan SCHULZ, K.J.1, CANNON, W.F.1, WOODRUFF, L.G.2, AND AYUSO, R.A.1 1 U.S. Geological Survey, 954 National Center, Reston, VA 20192, 2 U.S. Geological Survey, Mounds View, MN 55112 A terrane composed largely of Meso- to Paleoarchean gneisses and granitic rocks occurs along the southern margin of the Neoarchean Superior Craton in the Lake Superior region. These rocks are best documented from exposures in the Minnesota River Valley (MRV) in southwestern Minnesota, but they also occur in basement uplifts in northern Michigan including the Watersmeet Dome in the MareniscoWatersmeet area, the Carney Lake Gneiss north of the Menominee iron range, and the Southern Complex south of the Marquette Trough. Recent studies in the MRV have shown a range in ages primarily between ~2.6 Ga to ~3.5 Ga, representing both primary intrusive events and metamorphic/tectonic overprints (Bickford et al., 2007 and references therein). Similarly, dating of the gneisses in the Watersmeet Dome (Miska et al., 2018) and Carney Lake Gneiss (Ayuso et.al, 2018) have a range of ages from ~1.8 Ga to ~3.6 Ga, but also several spot analyses of zircon cores and xenocrysts that date at ~3.8 Ga. Thus, the northern Michigan gneisses show evidence of an Eeoarchean component and effects of the Penokean orogeny neither of which are seen in the MRV. Here we report on the geochemistry of Archean gneisses from Dickinson County in northern Michigan including the Carney Lake Gneiss. The Archean rocks in Dickinson County are described in James et al., (1961) and Bayley et al., (1966). As is typical of Archean gneiss terranes, the rocks consist mostly of variably banded and deformed tonalite-trondhjemite-granodiorite (TTG) gneisses and granites. Geochemistry Major elements Major element geochemistry of the gneisses in Dickinson County span the compositional range from tonalite to granite. Using the granite classification of Frost et al., (2001), the gneisses are magnesian and mostly calcic, although some of the more felsic gneisses are alkali calcic. A notable feature of the gneisses is that they are weakly to strongly peraluminous and corundum normative (Fig, 1A). The Na 2O content ranges from 3 to 5 wt.%, and the K2O content ranges from ~1 to 5 wt.% (medium- to high-K range); Na2O/ K2O ratios are mostly <2. There is a positive correlation between Na2O and Al2O3 contents, and a negative correlation between K2O and Al2O3 contents. Trace elements Unlike many Archean TTG suites which are typically characterized by Sr contents >400 ppm, the Sr contents of the Dickinson County gneisses are variable but <400 ppm. Rubidium/Sr ratios are variable, ranging from <0.1 to ~1, and show a negative correlation with Al2O3. Samples exhibit significant variation in chondrite normalized REE patterns, both in terms of pattern steepness and size of the Eu anomaly (Fig. 1B, C). The (La/Yb) N for samples from the Carney Lake Gneiss varies from 29 to 183 with no to moderately negative Eu anomalies; the more felsic samples tend to have higher light REE abundances and larger negative Eu anomalies (Fig. 1B). In contrast, two biotite gneiss samples from near Felch have much flatter patterns ((La/Yb) N of 14 and 24) and moderate to large negative Eu anomalies (Fig. 1C). Two samples of Norway Lake Gneiss from north of Felch have intermediate sloped REE patterns ((La/Yb)N of 44 and 58), no Eu anomaly for the tonalite sample, and moderate negative anomaly for the more granitic sample (Fig. 1C). All samples have negative Nb and Ta anomalies on primitive mantle normalized trace element plots. Comparison with Archean TTG suites Archean TTG suites are commonly silica-rich (SiO2 >64 wt.%, but commonly ≥70 wt.%), have high Na2O (>3.0 wt.%) and Na2O/K2O (>2), and low ferromagnesian element contents (Moyen and Martin, 2012). They trend from metaluminous to slightly peraluminous (A/CNK ~1; normative corundum <1%), with A/CNK increasing in more granitic compositions. Two subgroups are recognized: most 82 �Archean TTG suites have high Al2O3 (>15 wt.% at 70 wt.% SiO2) with high Sr contents (>400 ppm) and strongly fractionated REE patterns ((La/Yb)N up to 150); the second subgroup has lower Al2O3 (<15 wt. %) as well as lower Sr and less fractionated REE patterns. The Archean gneisses and granitic rocks in Dickinson County have the geochemical characteristics of typical TTG suites with respect to Al2O3 and strongly fractionated REE patterns. However, most samples are more potassic and less sodic than typical TTG and have lower Sr contents (<400 ppm). In addition, the Dickinson County samples are all peraluminous. Discussion Most models for the genesis of typical TTG involve partial melting of garnet amphibolite or eclogite (Moyen and Martin, 2012). However, the geochemical characteristics of the Archean gneisses in Dickinson County suggest a more complex petrogenesis involving melting of preexisting evolved crustal sources. This is supported by the presence of ~3.8 Ga xenocrystic zircons in some samples (Ayuso et al., 2018). It has been proposed that the Archean gneiss terrane in the Lake Superior region is a remnant of the Wyoming Craton, which was rifted from the Superior Craton in the Paleoproterozoic. In this regard, it may be significant that the quartzofeldspathic gneisses and granitoids in the Wyoming Craton have similar geochemical characteristics to the Archean gneisses in Dickinson County including relatively high K2O, low Sr, variably steep REE patterns, and are also mostly peraluminous (Frost et al., 2006). References Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., and Jackson, J., 2018, New U-Pb zircon ages for rocks from the granite-gneiss terrane in northern Michigan: Evidence for events at ~3750, 2750, and 1850 Ma: Institute on Lake Superior Geology, Proceedings of 64th Annual meeting, Part 1: Program and Abstracts, p. 7-8. Bayley, R.W., Dutton, C.E., and Lamey, C.A., 1966, Geology of the Menominee iron-bearing district Dickinson County, Michigan and Florence and Marinette Counties Wisconsin: U.S. Geological Survey Professional Paper 513, 96 p. Bickford, M.E., Wooden, J.L., Bauer, R.L., and Schmitz, M.D., 2007, Paleoarchean gneisses in the Minnesota River Valley and northern Michigan, USA, in Van Kranendonk, M.J., Smithies, R.H., and Bennett, V.C., eds., Earth’s Oldest Rocks, Developments in Precambrian Geology, v. 15, p. 731–750. Frost, B.R., Collins, C.G., Arculus, R.J., Ellis, D.J., and Frost, C.D., 2001, A geochemical classification of granitic rocks: Journal of Petrology, v. 42, p. 2033–2048. Frost, C.D., Frost, B.R., Kirkwood, Robert, and Chamberlain, K.R., 2006, The tonalite-trondhjemite-granodiorite (TTG) to granodiorite-granite (GG) transition in the late Archean plutonic rocks of the central Wyoming Province: Canadian Journal of Earth Sciences, v. 43, p. 1419–1444. James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., 1961, Geology of central Dickinson County Michigan: U.S. Geological Survey Professional Paper 310, 176 p. Miska, M.A., Mueller, P.A., and Bermudez, Katherine, 2018, Paleoarchean crust of the Minnesota-Michigan corridor: Evidence from the Watersmeet Dome, northern Michigan: Geological Society of America Abstracts with Programs, v. 50, no. 6, doi: 10.1130/abs/2018AM-318140. Moyen, Jean-Francois, and Martin, Hervé, 2012, Forty years of TTG research: Lithos, v. 148, p. 312–336. 83 �Geologic Architecture and Precious Metal Mineralization in the Southern Abitibi; New Insights from the Larder Lake Area SHERLOCK, Ross, RUBINGH, Kate and the Metal Earth research team Mineral Exploration Research Center, Harquail School of Earth Sciences, Laurentian University, Sudbury Ontario Metal Earth is one of the largest mineral exploration research project ever undertaken and is a fully funded $104M / 7 year research project focused on the processes responsible for differential metal endowment and ore localization during the Archean. A major focus of Metal Earth is to use geological and geophysical data to define crust to mantle scale differences across ancestral fault systems and volcanic centres that have variable metal endowment. As part of the Metal Earth project, research has focused on a ~40 km long north south geologic transect that is centered over the Cadillac-Larder Lake break and extends northward into the Ben Nevis volcanic complex and to the south over the Lincoln Nipissing shear zone. The Cadillac-Larder Lake break is a regionally extensive crustal break and hosts a number of gold deposits including the Kerr Addison mine which historically produced over 11Moz of gold. The Ben Nevis volcanic complex (2696.6 ± 1.3 Ma), part of the Blake River group (2701 ± 3 – 2698.5 ± 2Ma), is correlative to the Noranda VMS camp but lacks significant metal endowment. The Lincoln Nipissing shear zone is similar to the Cadillac-Larder Lake break, in that it juxtaposes different geologic domains and is marked by ultramafic volcanic rocks, clastic sedimentary rocks and Timiskaming aged small volume intrusive rocks and associated gold prospects. At both the Cadillac-Larder Lake break and the Lincoln Nipissing shear zone, ultramafic rocks of the Larder Lake group (ca. 2710-2704 Ma) (Piché in Quebec) are unconformably overlain by clastic rocks of the Timiskaming (2677-2670 Ma) or Hearst assemblage (<ca. 2700 Ma). This suggests that the original geologic relationship was stratigraphic in nature and subsequently overprinted by deformation and alteration associated with the gold deposits, in contrast to the previous interpretations that only considered a structural emplacement. Recent geological and geophysical surveys from the Metal Earth research project indicate that the Cadillac-Larder Lake break is well-resolved using seismic methods to depths of over 30 km and has a corresponding MT conductivity anomaly. In contrast, the Lincoln Nipissing shear zone, although sharing similar characteristics to the Cadillac-Larder Lake break, is poorly resolved by seismic and MT methods, perhaps correlating with the relative lack of metal endowment along the shear zone. This is MERC-Metal Earth publication number MERC-ME2018-177. 84 �An investigation into the distribution of chalcophile elements and timing of mineralization within the Crystal Lake intrusion: A U-Pb geochronology and LA-ICP-MS study SMITH, Jennifer1, BLEEKER, Wouter1, HAMILTON, Mike2 and PETTS, Duane1 1 Geological Survey of Canada, 601 Booth Street, Ottawa, Canada; email:jennifer.smith6@canada.ca Jack Satterly Geochronology Laboratory, Dept. of Earth Sciences, University of Toronto, 22 Russell St., Toronto, Canada 2 A detailed geochemical and isotopic study is underway on the 1099.1 ± 1.2 Ma (Heaman et al. 2007) Crystal Lake intrusion, which has previously been compared to the proximal Duluth Complex (Thomas 2015). The aim of this study is to gain further insights into the controls on ore genesis within the ‘mainrift’ intrusions (Miller and Nicholson 2013). The Crystal Lake intrusion, located 47 km southwest of Thunder Bay, Ontario, Canada, outcrops as a prominent Y-shaped body within the Paleoproterozoic Animikie basin, intruding sulfur-bearing shale, argillite and greywacke of the Rove Formation. Although a number of dating studies have been undertaken on the MCR (see Heaman et al. 2007, for a relatively recent compilation), many of the intrusions either lack the precision that is now possible with routine chemical abrasion U-Pb geochronology or are yet to be dated. We are currently undertaking a comprehensive geochronology study throughout the MCR, this includes detailed dating of the Crystal Lake intrusion. In addition to refining Heaman’s et al. (2007) baddeleyite age of 1099.1 ± 1.2 Ma we aim to constrain the relationship of the northern and southern limbs of the intrusion. Furthermore, we plan to untangle the timing of the Crystal Lake intrusion relative to other MCR intrusive events including the NE-trending Pigeon River dykes, the NW-trending Cloud River dykes and the sulfide-bearing Mount Mollie intrusion developed to the east. Ni-Cu-PGE sulfide mineralization is developed within the northern and southern limbs of the Crystal Lake intrusion in association with vari-textured gabbros and irregular Cr-spinel-bearing horizons. The association of sulfides and metal enrichment with pegmatitic/taxitic units is also observed within other Ni-Cu-PGE deposits such as the ca. 1108 Ma (Heaman and Machado 1992) Coldwell Complex, Merensky Reef, Norilsk and Voisey’s Bay. Sulfide mineralization is largely disseminated, with massive sulfides (<50 cm in thickness) developed locally within the northern limb. The disseminated ores are variable in texture with globular (capped and uncapped), blebby and interstitial sulfides identified. Silicate-capped sulfide globules have been recognized in other Ni-Cu sulfide deposits (e.g. Norilsk, Insizwa Complex; Barnes et al. 2017; Le Vaillant et al. 2017) and are interpreted as being the remnants of former segregation vesicles that attached to an immiscible sulfide melt (Mungall et al. 2015). Within the Crystal Lake intrusion, the morphology of the caps, which are comprised of amphiboles, clays, chlorite and calcite, is variable. Convex silicate caps, identical to those modelled by Mungall et al. (2015), are present along with very irregular silicate attachments. The implications of degassing, which appears to be a common process within the Ni-Cu ore systems, for sulfide transportation and deposition is yet to be constrained. Furthermore, the cause (e.g. contamination, pressure changes) and timing of degassing relative to crystallization is not well understood. A detailed elemental deportment study is currently in progress, focused on characterizing the distribution and mineralogy of platinum-group minerals (PGMs). Element mapping of sulfides by LAICP-MS has been used to further investigate the control on the distribution of the chalcophile elements during sulfide fractionation. Preliminary observations indicate that Pd resides in solid solution within pentlandite (1 – 150 ppm) and as small As-Bi and Sb-bearing PGMs. Within the massive sulfides Pdbearing minerals show a strong association with nickel arsenides resulting in lower concentrations of Pd (~1 ppm) in the pentlandite than typical of other sulfide assemblages (10–150 ppm). Platinum is not compatible in any of the sulfide phases, instead occurring as discrete As and Sb-bearing PGMs. The PGMs are found either enclosed or attached to sulfides or within secondary silicates around the altered margins of the sulfides. It is yet to be established whether the crystallization of Cr-spinel and/or lowtemperature alteration of the sulfides has had any control on the mineralogy and distribution of PGEs. 85 �Element mapping of the sulfides by LA-ICPMS has revealed some interesting structural and or/mineralogical controls in the distribution of chalcophile elements. Although not observed throughout the primary sulfide assemblage, some unaltered sulfides are characterized by a strong microfabric (Fig. 1). This fabric is defined by several elements including As, Mo, Bi, Pb, Pd and Re which appear to be preferentially concentrated along thin, parallel linear features within the pyrrhotite-pentlanditechalcopyrite assemblage. The molybdenum map also shows thicker banding and elevated concentrations within pyrrhotite (Fig. 1). Interestingly this fabric is not confined to a particular sulfide phase. This is best shown by Figure 1. LA-ICP-MS element maps of primary sulfide As, Mo and Re which clearly cut across the assemblage grain boundaries of pyrrhotite, pentlandite and chalcopyrite, which suggests that this fabric was developed subsequent to crystallization of all three phases. For other elements such as Pd and Pb, the fabric is restricted to the pyrrhotite and pentlandite (Fig. 1). Silicate infilled fractures appear to cut the fabric as shown in the As map. Further work is in progress to determine the controls on selected element mobility (i.e. low temperature alteration or deformation) and to gain an understanding at what scale this remobilization is occurring. If various elements are remobilized over long distances, then it could have implications for vectoring of Ni-Cu ore systems. Element mapping by LA-ICP-MS is an extremely powerful tool, providing unparalleled detail at the micro scale. This technique provides insight into the behavior and mobility of chalcophile elements during sulfide fractionation, low temperature alteration and/or deformation and may provide a link to larger element haloes associated with some Ni-Cu-PGE deposits. References Barnes, S.J., Mungall, J., Le Vaillant, M., Godel, B., Lesher, M., Holwell, D., Lightfoot, P., Krivolutskaya, N., and Wei, B., (2017) Sulphide-silicate textures in magmatic Ni-Cu-PGE sulphide ore deposits: Disseminated and net-textured ores. American Mineralogist 102:473–506. Heaman, L.M., and Machado, N., (1992) Timing and origin of midcontinent rift alkaline magmatism, North America: evidence from the Coldwell Complex. Contributions to Mineralogy and Petrology 110:289–303. Heaman, L.M., Easton, R.M., Hart, T.R., Hollings, P., MacDonald, C.A., Smyk, M., (2007) Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences 44:1055–1086. Le Vaillant, M., Barnes, S.J., Mungall, J.E., Mungall, E.L., (2017) Role of degassing of the Noril’sk nickel deposits in the Permian–Triassic mass extinction event. Proceedings of the National Academy of Sciences 114:2485–2490. Miller, J.D., Nicholson, S.W., (2013) Geology and mineral deposits of the 1.1 Ga Midcontinent Rift in the Lake Superior region – An overview; in Field Guide to the Cu-Ni-PGE Deposits of the Lake Superior Region (ed) JD Miller. Precambrian Research Center Guidebook 13-1:1–50. Mungall, J.E., Brenan, J.M., Godel, B., Barnes, S.J., Gaillard, F., (2015) Transport of metals and sulphur in magmas by flotation of sulphide melt on vapour bubbles. Nature Geoscience 8:216–219. 86 �Seismic stratigraphy of the 1.1 Ga Midcontinent Rift beneath western Lake Superior shows evidence of differing subsidence histories for syn-magmatic sub-basins STEWART, Esther K.1, GRAUCH, V.J.S.2, WOODRUFF, Laurel G.3, and HELLER, Samuel4 1 Wisconsin Geological & Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 U.S. Geological Survey, MS 964, Federal Center, Denver, CO 80225 3 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 4 U.S. Geological Survey, MS 939, Federal Center, Denver, CO 80225 2 The nature of the Midcontinent Rift where it is hidden beneath Lake Superior can only be understood from geophysical interpretation and inferences from geologic concepts onshore. Many decades of collecting geophysical data and improving the geologic framework for the Lake Superior region have led to evolving paradigms on its structure and geologic history. We are taking a new look at the existing geophysical data considering recent age dating and geologic mapping as part of an ongoing effort to characterize the 3D geometry of the rift through time. In particular, we are constructing a seismic stratigraphic framework from reflection data collected in the 1980s to characterize the basalt-filled sub-basins present beneath the western lake. Integration with gravity and aeromagnetic data and correlation of seismic stratigraphy to onshore geology provides insight into the geometry of the sub-basins and the timing and rates of subsidence and concurrent magma accumulation. We interpret 3 reprocessed and 7 geolocated images of industry seismic sections and public seismic data (GLIMPCE Line C) using a 3D visualization software platform. Seismic sections were converted from time to depth using existing seismic refraction models as guides. Dips and thicknesses of onshore geology projected onto nearby seismic profiles tie seismic stratigraphy to mapped geologic units. Aeromagnetic modeling helps distinguish igneous rocks with strong, reversed- versus normal-polarity remanent magnetization, which constrains the cooling ages of syn-rift rocks to before or after about 1100 Ma, respectively. Seismic facies and reflection geometry image volcanic flows, sills, intrusions, and pre-rift crust. We have identified three seismic stratigraphic units. The lowest unit has clinoform reflection geometry and may represent pre-rift sediments intruded by sills. The middle and upper units have synform reflection geometry, but aeromagnetic modeling and correlations to onshore geology suggest the middle unit represents older (>1100 Ma) basalts with reversed-polarity and the upper unit represents younger (<1100 Ma) basalts with normal magnetic polarity. As pointed out by previous workers (e.g., Allen et al., 1997) we do not observe large normal faults bounding sub-basins. Except for a possible growth fault of limited extent at depth on GLIMPCE Line C (Fig. 1), we observe sub-basins that sag and thicken toward their centers, implying that syn-magmatic subsidence may have been the primary control on basin development in western Lake Superior. The sub-basins are flanked by two previously recognized seismic highs that are associated with gravity lows: Grand Marais Ridge (GMR) south of Grand 87 �Marais, MN, and White’s Ridge (WR), centered on the Bayfield Peninsula (Fig. 1; Allen et al. 1997). An isopach map of the seismic stratigraphic unit interpreted as normal-polarity basalts shows differences in thickness and age of basin fill within sub-basins surrounding GMR (Fig. 1). We correlate most of the basin fill in a western, bowl-shaped basin to the Beaver Bay Complex and younger North Shore Volcanics that were deposited or emplaced on the present-day northern lake shore of Minnesota between ca 1096 – 1094 Ma. In contrast, only 4-9 km of normal-polarity basalt infills an adjacent, elongate basin that wraps around the southeast and eastern sides of GMR. We correlate this material with the ca 1094 – 1091 Ma Portage Lake Volcanics on Michigan’s Keweenaw Peninsula. The difference in basin geometry and age indicates each basin developed independently. The western basin subsided and infilled at some 5mm per year. After the main subsidence and infilling of the western basin waned, the adjacent, shallower basin began to subside at some 3 to 1.3mm per year. Reflections interpreted as reversed- and normalpolarity basalts in the shallower basin truncate against the southern side of GMR. Truncation of these once laterally continuous basalt layers was likely caused in part by the basins’ different subsidence histories and relative uplift of GMR. A mechanism for the syn-magmatic basin subsidence and dramatic difference in subsidence rates is unclear. Fig 1: Isopach map of normal-polarity basalt. Onshore geology highlights the distribution of reversed- and normalpolarity basalt and overlying sedimentary units. Purple lines locate seismic profiles, with line GLIMPCE C labeled. Reference Allen, D.A., Hinze, W.J., Dickas, A.B., and Mudrey, M.G., Jr., 1997, Integrated geophysical modeling of the North American Midcontinent Rift System: new interpretations for western Lake Superior, northwestern Wisconsin, and eastern Minnesota, in Ojakangas, R.W., Dickas, A.B., and Green, J.C., eds., Middle Proterozoic to Cambrian Rifting, Central North America: Geological Society of America Special Paper 312, p. 47-72. 88 �Towards understanding geoarchaeological contexts in Northwestern Ontario: The newly formed lithic material comparative collection at Lakehead University SURETTE, Clarence, and TAYLOR-HOLLINGS, Jill Department of Anthropology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1; clsurett@lakeheadu.ca, jstaylo1@lakeheadu.ca Stone tools were created by flintknapping, which is the process of utilizing percussion and pressure flaking techniques on fine-grained siliceous raw materials that were carefully selected by ancient people. Excellent preservation of lithics in the boreal forest of Northwestern Ontario provides some of the best evidence in the archaeological record for reconstructing these past human activities. Professional archaeologists have been finding and interpreting stone tools from the earliest known Palaeo or Early Period (ca. 9,000-7,000 years before present) sites in the Thunder Bay region beginning in the 1950s (Dawson, 1984; MacNeish, 1952). However, little is still known about the variation and sources of them in Northwestern Ontario, despite lithics being the most commonly found artifact class. For site descriptions, archaeologists must try to identify lithic artifacts to the best of their abilities, both describing the material and then attempting to find the source from which people had obtained it. Although geologists are not typically as concerned about the macroscopic nuances of each flakeable material (e.g., chert rather than brown banded Hudson Bay Lowland chert), archaeologists note variability in order to understand ancient miners’ choices whether for better flaking, durability of formed edges, or sometimes even esthetic appeal. These descriptions often reflect superficial macroscopic observations and a misunderstanding of regional geological characterizations even in terms of major rock types (i.e., sedimentary, igneous, or metamorphic). During the cataloguing process, many archaeologically recovered lithic types are erroneously categorized, leaving many artifacts to fall into the category of “unknown material type”. Perhaps one of the biggest challenges in Northwestern Ontario is the limited availability of comparative collections to try and identify lithic raw materials. To counteract this issue, Surette, colleagues, and students in the Department of Anthropology at Lakehead University began collecting geoarchaeological samples of knappable materials in 2011, and at primary sources where possible. Secondary and tertiary source examples were also collected, even though their contexts are more complicated to understand from a geoarchaeological perspective. Geological maps were examined and geoarcheological contexts were considered (ancient hydrology, stratigraphy, etc.) to determine where these silica rich materials might be located and which would have been accessible at different times. In addition, we have started sharing and trading for samples with archaeologists and geologists from other institutes in Canada and the U.S.A. to build a representative library. To date, there are nearly 5,000 samples from both countries, of which 1,300 are from various locations in Ontario. One of the better understood hosts of raw materials is the Gunflint Formation near Thunder Bay, which has been the recent focus of much sampling for the comparative collection (e.g., Vickruck, 2018). 89 �Unfortunately, archaeologists in Canada rarely describe knappable rocks and minerals by attributes, in detail, or using correct geological terminology, even for quarry sites where there is typically one or limited sources found. This factor is problematic because it is then difficult for other researchers to determine if their samples are made of the same material - perhaps found at a different site due to trading or people obtaining quarry samples and utilizing them elsewhere. In Northwestern Ontario, there are also few basic descriptions of flakeable materials used by early Indigenous populations (Hamilton, 1981 for Lac Seul, Taylor-Hollings, 2017 regarding the Bloodvein River, and Vickruck, 2015 for the Thunder Bay region). Therefore, we aim to change that in our discipline through the study of samples in the Lakehead University lithic collection and make this information available to other researchers (eventually online). The Lakehead University lithic material comparative collection in the Department of Anthropology will provide archaeologists and geoscientists with opportunities to examine the minutia of knappable rocks and minerals, with emphasis on the Gunflint Formation but also from many other regions. This new database provides raw material examples that can be studied in many different ways, either at a large scale or microscopic studies of individual rocks or minerals. Due to having a large collection, we now know that different sources can produce similar flakeable rocks, which emphasizes the need to chemically test them to clarify provenance. The next steps are to catalogue and properly describe these samples in the collection. We also plan to develop methods for analyzing these materials with non-destructive techniques, which may be also applied to artifact characterization. Combined, this will help us address both the problem of not knowing about sources in Northwestern Ontario and illustrating these materials properly through basic lithological descriptions and in some cases, further nondestructive geoarchaeological analytical techniques. Ultimately, that will help us address the selection processes of ancient Indigenous people in the area. References Dawson, K.C.A. 1984. A history of archaeology in Northern Ontario to 1983 with bibliographic contributions. Ontario Archaeology 42:27-92. Hamilton, S. 1981. The archaeology of Wenesaga Rapids. Archaeology Research Report 17, Archaeology and Heritage Planning Branch, Ontario Ministry of Culture and Recreation, Toronto. MacNeish, R. 1952. A possible early site in the Thunder Bay district, Ontario. National Museum of Canada, Bulletin No. 126, pp. 23-47. Department of Northern Affairs and National Resources, Ottawa. Taylor-Hollings, J. 2017. “People lived there a long time ago”: Archaeology, ethnohistory, and traditional use of the Miskweyaabiziibee (Bloodvein River), northwestern Ontario. Unpublished PhD dissertation, Department of Anthropology, University of Alberta, Edmonton. Vickruck, C. 2018. Investigating the qualities of raw lithic material and the selection pressures of lithic materials from the Gunflint Formation, in Ontario Canada. Master of Environmental Studies: Northern Environments and Cultures, Lakehead University, Thunder Bay. 90 �Insights into Midcontinent Rift development resulting from a strengthened chronostratigraphic framework SWANSON-HYSELL, Nicholas L. Department of Earth and Planetary Science, University of California, Berkeley Correlation of volcanostratigraphic sequences across the Midcontinent Rift has been a long time focus of research efforts with major advances made on the basis of lithostratigraphy (e.g. Green, 1982), magnetostratigraphy (e.g. Books, 1972), chemostratigraphy (e.g. Nicholson et al., 1997), and U-Pb geochronology-based chronostratigraphy (e.g. Davis and Green, 1997). As the result of an effort to obtain high-resolution U-Pb geochronological constraints on paleomagnetic poles from the Midcontinent Rift and constrain rates of rapid plate motion, we have published 14 new chemical abrasion–isotope dilution– thermal ionization mass spectrometry (CA-ID-TIMS) 206Pb/238U dates (Swanson-Hysell et al., 2015; Fairchild et al., 2017; Swanson-Hysell et al., 2019). Single zircon analyses of chemically-abraded grains improves the geochronology of previously dated units in addition to resulting in high-precision dates developed for previously undated units. These dates can be used to construct a chronostratigraphic framework for Midcontinent Rift volcanic and sedimentary succession shown in Figure 1 that is further informed by magnetostratigraphic data and paleomagnetic pole position. Figure 1: Chronostratigraphic correlation of Midcontinent Rift volcanic sequences across the Lake Superior Region of North America, informed by new U-Pb dates. The numbered circles correspond to CA-ID-TIMS 206Pb/238U dates. The analytical uncertainty, which can be used when comparing these dates to one another, is less than the time represented by the height of the circles. Extrapolated eruption rates, paleomagnetic data (both polarity and pole position) and 207Pb/206Pb dates (not shown) inform the chronostratigraphic interpretation, but the chronostratigraphy is most robust in the proximity of the 206Pb/238U dates. Figure from Swanson-Hysell et al., 2019. 91 �To aid in the discussion of geomagnetic polarity zones, I propose naming them following the guidelines of the International Commission on Stratigraphy resulting in the Alona Bay reversed-polarity zone, the Flour Bay normal-polarity zone, the Flour Bay reversed-polarity zone and the Portage Lake normalpolarity zone with current evidence suggesting that the Portage Lake normal-polarity zone was particularly long-lived (Driscoll and Evans, 2016). The geomagnetic polarity timescale and the major inclination change recorded as the result of paleolatitude change throughout of the time period of active Midcontinent Rift magmatism continue to present opportunities to constrain the chronostratigraphy of volcanic, intrusive and sedimentary rocks of the rift. Future high-precision 206Pb/238U dates on intrusive units coupled with paleomagnetic data may offer further opportunities related to the polarity record in the earliest history of rift development. The mafic nature of the earliest Midcontinent Rift lavas (including picrites), and the associated lack of zircon, continues to present challenges to efforts to robustly constrain the chronostratigraphy of the early magmatic stage of rift development. Age constraints on the angular unconformities within the Osler Volcanic Group, between the North Shore Volcanic Group & Schroeder-Lutsen Basalts and at the base of the Oronto Group in northern Wisconsin are insightful as relates to the timescale of active rifting and the transition to the thermal subsidence stage of rift development. Post-rift unconformities can be particularly useful for constraining the timing of the end of rifting as they juxtapose underlying syn-rift strata with post-rift strata. The Brownstone Falls unconformity at which Oronto Group sedimentary rocks overlie progressively lower stratigraphic levels of the Porcupine Volcanics, Portage Lake Volcanics, and Kallander Creek Volcanics (Fig. 1) is well-explained as a post-rift unconformity. The volcanics underlying the unconformity can be interpreted as syn-rift strata with the overlying Oronto Group having been deposited during widespread thermal subsidence. The syn-rift strata in this interpretation include the Portage Lake Volcanics which could constrain the post-rift phase to postdate 1091.59 ± 0.27/0.52/1.3 Ma (Swanson-Hysell et al., 2019). The youngest well-dated magmatic product from the Midcontinent Rift is the Davieaux Island rhyolite of the Michipicoten Island Formation (1083.52 ± 0.23/0.35/1.2 Ma; Fairchild et al., 2017) although the Bear Lake volcanics are likely younger still (Kulakov et al., 2018). When combined with paleomagnetic data, these chronostratigraphic constraints provided evidence for very rapid motion of Laurentia leading up to the collisional orogenesis associated with the Ottawan phase of the Grenvillian orogeny. References Books, K., 1972, Paleomagnetism of some Lake Superior Keweenawan rocks: U.S. Geological Survey Professional Paper 760, 42 p. Davis, D., and Green, J., 1997, Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution: Canadian Journal of Earth Sciences, v. 34, p. 476–488, https://doi.org/10.1139/e17039 Driscoll, P.E., and Evans, D.A.D., 2016, Frequency of Proterozoic geomagnetic superchrons: Earth and Plan etary Science Letters, v. 437, p. 9–14, https://doi.org /10.1016/j.epsl.2015.12.035. Fairchild, L.M., Swanson-Hysell, N.L., Ramezani, J., Sprain, C.J., and Bowring, S.A., 2017, The end of Midcontinent Rift magmatism and the paleogeography of Laurentia: Lithosphere, v. 9, p. 117–133, https://doi.org/10.1130/L580.1. Green, J.C., 1982, Geology of Keweenawan extrusive rocks in GSA Memoir v. 156 Geology and Tectonics of the Lake Superior Basin https://doi.org/10.1130/MEM156-p47 Kulakov, E., Bornhorst, T. J., Deering, C., and Moore, J. B. 2018. The youngest magmatic activity of the Midcontinent Rift at Bear Lake, Keweenaw Peninsula, Michigan: ILSG Program and Abstracts, v. 64. Nicholson, S., Shirey, S., Schultz, K., and Green, J., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: Implications for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504–520, https://doi.org/10.1139 /e17041. Swanson-Hysell, N.L., Burgess, S.D., Maloof, A.C., and Bowring, S.A., 2014, Magmatic activity and plate motion during the latent stage of Midcontinent Rift development: Geology, v. 42, p. 475–478, https://doi.org/10.1130/G35271.1. Swanson-Hysell, N.L., Ramezani, J., Fairchild, L.M., and Rose, I.R., 2019, Failed rifting and fast drifting: Midcontinent Rift development, Laurentia’s rapid motion and the driver of Grenvillian orogenesis: GSA Bulletin, https://doi.org/10.1130/B31944.1. 92 �An oxygenated Paleolake Nonesuch and primary detrital hematite in the Freda river system SWANSON-HYSELL, Nicholas L., SLOTZNICK, Sarah P. and FAIRCHILD, Luke M. Department of Earth and Planetary Science, University of California, Berkeley In addition to preserving an extended interval of volcanism, sediment deposited within the thermal subsidence phase (Cannon and Hinze, 1992) of the Midcontinent Rift provide an exceptional record of late Mesoproterozoic terrestrial environments. The Oronto Group, deposited during this thermal subsidence phase commences with the Copper Harbor Conglomerate ca. 1086 Ma, which represents terrestrially-deposited alluvial fan and fluvial sediments (Elmore, 1984). The Nonesuch Formation overlies the Copper Harbor Conglomerate and is interpreted as a lacustrine facies association (Stewart and Mauk, 2017) that is exposed along a >250-km-long belt in northern Michigan and Wisconsin. The sediments of the Nonesuch Formation indicate that the lake in northern Wisconsin and Michigan (referred to here as Paleolake Nonesuch) was large and persistent. Lacustrine sedimentation continued until after the transition into the overlying Freda Formation — a transition typical based on color rather than lithofacies. The overlying Freda Formation is dominantly composed of channelized sandstone and overbank siltstone deposits that were deposited within a terrestrial fluvial environment (Ojakangas et al., 2001). Five drill cores from northern Wisconsin were used by Stewart and Mauk (2017) to develop a sequence stratigraphic framework for the Nonesuch Formation. In work published in Slotznick et al. (2018) we focused on two of these cores (DO-8 and WC-9) and have now extended our analyses to outcrop sections as well as cores from the Presque Isle syncline. In this research, we have used experimentally determined estimates of magnetization and coercivity on samples that span sections of the Nonesuch Formation. These rock magnetic data can be used to interpret three distinct magnetic facies (Slotznick et al., 2018). These three facies and their juxtaposition can be explained as the result of an oxycline in Paleolake Nonesuch. The detrital input to the lake is preserved in facies 2 and included magnetite and hematite due to weathering and oxidation of the source igneous material during transport. Magnetic facies 1 is associated with sediments in the deepest part of the lake with a magnetic mineral assemblage that shows that delivered iron oxides underwent reductive dissolution through microbial metabolic processes. Much of this iron and iron within sheet silicates reacted with sulfide to form pyrite, but sulfide availability was restricted to pore waters and not sufficient to sulfidize all of the available reactive iron. These data indicate that the sediments in the deepest part of the lack were anoxic, possibly with anoxia extending into the water column. In contrast, intermediate oxygen levels in waters throughout much of the lake allowed for the preservation of detrital magnetite and hematite in facies 2. In the shallow waters of the lake recorded in facies 3, oxic conditions prevailed and most of the detrital magnetite, as well as iron in other phases, was oxidized to hematite. In Slotznick et al. (2019), we interpreted this vertical sequence of facies to reflect a stacking of laterally distributed environments such that the transition from the deepest-water low-iron-oxide facies into the intermediate-water magnetite-rich facies and the shallower-water hematite-rich facies is the result of an oxycline within the ancient lake. The depth dependence of the oxycline is similar to that found in modern eutrophic lakes wherein the aerobic respiration of descending organic matter leads to a decrease in dissolved oxygen with depth. Overall, these data indicate that this ca. 1.1 billion-year-old lake was more deeply oxygenated than has previously been interpreted (e.g. Cumming et al., 2013) providing a hospitable environment for the diverse biota that was present in the lake that included early eukaryotes (Wellman and Strother, 2015). This framework is strengthened by new data from a section along Potato River Falls that was near a paleo-highlands. In this section, the magnetic mineralogy is dominated by Facies 2 and 3 through repeated fluvial-lacustrine cycles without the anoxia seen in the deeper part of the lake. 93 �The interpretation of a low-latitude paleolatitude of Laurentia at the time of Nonesuch and Freda deposition is reliant on magnetizations held by hematite (Henry et al., 1977). While detrital hematite in sediment can lead to a primary depositional remanent magnetization, alteration of minerals through interaction with oxygen can lead to the post-depositional formation of hematite. We have used the exceptionally-preserved fluvial sediments of the Freda Formation to gain insight into the timing of hematite remanence acquisition and its magnetic properties. This deposit contains siltstone intraclasts that were eroded from a coexisting lithofacies and redeposited within channel sandstone. Thermal demagnetization, petrography and rock magnetic experiments on these clasts reveal two generations of hematite. One population of hematite demagnetized at the highest unblocking temperatures and records paleomagnetic directions that rotated along with the clasts. This component is a primary detrital remanent magnetization. The other component is removed at lower unblocking temperatures and has a consistent direction throughout the intraclasts. This component is held by finer-grained hematite that grew and acquired a chemical remanent magnetization following deposition resulting in a population that includes superparamagnetic nanoparticles in addition to remanence-carrying grains. This primary magnetization can be successfully isolated from co-occurring authigenic hematite through high-resolution thermal demagnetization. These data lend credence to existing paleomagnetic data from the Freda Formation as well as future efforts to develop such data at higher resolution. References Cannon WF, Hinze WJ (1992) Speculations on the origin of the North American midcontinent rift. Tectonophysics 213:49–55 Cumming VM, Poulton SW, Rooney AD, Selby D (2013) Anoxia in the terrestrial environment during the late Mesoproterozoic. Geology 41:583–586. Elmore RD (1984) The Copper Harbor Conglomerate: A late Precambrian fining- upward alluvial fan sequence in northern Michigan. Geol Soc Am Bull 95:610–617. Fairchild, L. M., Swanson-Hysell, N. L., Ramezani, J., Sprain, C. J., & Bowring, S. A. (2017). The end of Midcontinent Rift magmatism and the paleogeography of Laurentia. Lithosphere, 9(1), 117-133. doi: 10.1130/L580.1 Henry, S., Mauk, F., & Van der Voo, R. (1977). Paleomagnetism of the upper Ke- weenawan sediments: Nonesuch Shale and Freda Sandstone. Canadian Journal of Earth Science, 14, 1128-1138. doi: 10.1139/e77-103 Ojakangas RW, Morey GB, Green JC (2001) The Mesoproterozoic midcontinent rift system, Lake Superior region, USA. Sediment Geol 141-142:421–442. Slotznick, S., Swanson-Hysell, N.L., and Sperling, E. (2018), An oxygenated Mesoproterozoic lake revealed through magnetic mineralogy, PNAS, doi:10.1073/pnas.1813493115 Stewart EK, Mauk JL (2017) Sedimentology, sequence-stratigraphy, and geochemical variations in the Mesoproterozoic Nonesuch formation, Northern Wisconsin, USA. Precambrian Res 294:111–132. Swanson-Hysell, N.L., Ramenzani, J., Fairchild, L.M. and Rose, I. (2019), Failed rifting and fast drifting: Midcontinent Rift development, Laurentia’s rapid motion and the driver of Grenvillian orogenesis, Geological Society of America Bulletin, doi:10.1130/B31944.1. Wellman CH, Strother PK (2015) The terrestrial biota prior to the origin of land plants (embryophytes): A review of the evidence. Palaeontology 58:601–627. 94 �New paleomagnetic constraints on the formation of the Slate Islands impact structure SWANSON-HYSELL, Nicholas L., TIKOO, Sonia M. and FAIRCHILD, L.M. Department of Earth and Planetary Science, University of California, Berkeley Pioneering paleomagnetic study by Halls (1975, 1979) was central to establishing an impact origin for the Slate Islands Impact structure. We have conducted further paleomagnetic study of both the injectite lithic breccia dikes throughout the structures as well as the target rocks which have an impactrelated overprint. These data enable further conclusions about the timescale of cratering processes and the origin of the impact-related overprint discovered by Halls (1975, 1979). In our work, we previously used lithic breccia dikes that were injected into the target rock during crater excavation to constrain the rate of crater modification within the central uplift of the Slate Islands Impact structure (Fairchild et al., 2019). We studied both the matrix as well as the clasts within the breccia dikes throughout the impact structure paleomagnetically. These data revealed uniform and linear paleomagnetic directions both in the matrix and in the clasts that are best interpreted as being due to frictionally heating above the magnetite Curie temperature (580 °C) during dike emplacement and subsequently cooling in situ through magnetic blocking temperatures. The work of Halls (1979) used data from the matrix of the breccia dikes to argue that the dikes acquired a thermal remanent magnetization during cratering and our data provide strong support for this hypothesis. The tight clustering of paleomagnetic directions from these breccia dikes indicates that the dikes cooled and locked in their magnetic remanence during a time interval in which the impact structure had reached a stable state with no major ongoing structural rotations. Applying a conductive cooling model to the thinnest sampled breccia dike demonstrates that magnetic remanence was being acquired six minutes after emplacement which indicates a stable crater structure at that time. This constraint of a stable crater structure only minutes after impact that results from breccia dike paleomagnetism is a rare case in which a geological process can be resolved on such a short time scale. We have also pursued paleomagnetic study paired with hydrocode modeling of the Slate Islands impact structure host rocks. The goal of this study is to determine the relative contribution of the competing nature of shock and thermal effects on the magnetization of moderately to highly shocked rocks within the crater. Target rocks within the central peaks of complex craters are often exposed to pressures >10 GPa (such as the case in the Slate Islands; Dressler et al., 1998) and thus experience shock heating to >200 °C (Stewart et al., 2007) as well as baked contact heating to higher temperatures associated with the emplacement of impact breccias (Fairchild et al., 2016). Under such conditions, magnetization acquired during the passage of a shock wave (shock remanent magnetization; SRM) would be largely overprinted by a higher intensity thermal overprint during post-impact cooling (Tikoo et al., 2015). In our thermal demagnetization experiments, the impact-related magnetization component was predominantly removed from samples by laboratory unblocking temperatures of ~250-350°C consistent with heating to such temperatures. These data, combined with results from paleointensity experiments, support an interpretation that the Slate Islands overprint is primarily thermal in origin. This result is consistent with paleomagnetic studies of several other terrestrial impact craters, which report thermallyinduced magnetizations in target rocks (Elbra et al., 2007; Jackson and Van der Voo, 1986). The potential for thermal and viscous remagnetization, as well as the acquisition of chemical remanence from hydrothermal activity (Pesonen et al., 1999; Quesnel et al., 2013), render it unlikely that SRM will be preserved in rocks as the dominant impact-related magnetization in the majority of settings. 95 �References Collins, G. S. (2014), Numerical simulations of impact crater formation with dilatancy, J. Geophys. Res. Planets, 119, 26002619, doi:10.1002/2014JE004708. Dressler, B. O., V. L. Sharpton, and B. C. Schuraytz (1998), Shock metamorphism and shock barometry at a complex impact structure: Slate Islands, Canada, Contrib. Min. Petrol., 130, 275-287, doi:10.1007/s004100050365. Elbra, T., A. Kontny, L. J. Pesonen, N. Schleifer, and C. Schell (2007), Petrophysical and paleomagnetic data of drill cores from the Bosumtwi impact structure, Ghana, Meteorit. Planet Sci., 42, 829-838. Fairchild, L. M., N. Swanson-Hysell, and S. M. Tikoo (2016), A matter of minutes: Breccia dike paleomagnetism provides evidence for rapid crater modification, Geology, 44, 723-726, doi:10.1130/G37927.1 Halls, H. C. (1975), Shock-induced remanent magnetisation in late Precambrian rocks from Lake Superior, Nature, 255, 692-695, doi:10.1038/255692a0. Halls, H. C. (1979), The Slate Islands meteorite impact site: a study of shock remanent magnetization., Geophys. J. Roy. Astr. S., 59, 553-591, doi:10.1111/j.1365-246X.1979.tb02573.x. Jackson, M., and R. Van der Voo (1986), A paleomagnetic estimate of the age and thermal history of the Kentland, Indiana cryptoexplosion structure, J. Geol., 94, 713-723. Pesonen, L. J., S. Elo, M. Lehtinen, T. Jokinen, R. Puranen, and L. Kivekas (1999), Lake Karikkoselka impact structure, central Finland: New geophysical and petrographic results., Geol. S. Am. S., 339, 131-147. Stewart, S. T., A. Seifter, G. B. Kennedy, M. R. Furlanetto, and A. W. Obst (2007), Measurements of emission temperatures from shocked basalt: hot spots in meteorites, Proc. 38th Lunar and Planetary Science Conference, 2413. Tikoo, S. M., J. Gattacceca, N. L. Swanson-Hysell, B. P. Weiss, C. Suavet, and C. Cournede (2015), Preservation and detectability of shock-induced magnetization, J. Geophys. Res., doi:10.1002/2015JE004840. Quesnel, Y., J. Gattacceca, G. R. Osinski, and P. Rochette (2013), Origin of the central magnetic anomaly at the Haughton impact structure, Canada, Earth Planet. Sci. Lett., 367, 116-122. 96 �Petrological and geochemical characteristics of the granitic rocks from the Dog Lake Granite Chain: Implications for the genesis of Quetico Basin WANG, Shiwei1,2, HOLLINGS, Pete2 and KUZMICH, Ben2 1 School of Resources and Environmental Engineering, Hefei University of Technology, Hefei 230009, China 2 Department of Geology, Lakehead University, Thunder Bay, Ontario, Canada, P7B 5E1. The Archean Quetico Basin is a metasedimentary terrane of the Superior Province between the Wawa-Abitibi Terrane to the south and the Western Wabigoon, Winnipeg River, and the Marmion terranes to the north. The majority of plutonic rocks within the Quectio Basin are granitoids (Williams, 1991), and are an important tool to investigate the evolution and genesis of the Quetico Basin and the nature of Archean tectonic processes (Sawyer et al., 2002). Percival (1989) studied the geology of granitic intrusions in the western Quetico basin and identified three main types: (i) an older, rare, white foliated hornblende-biotite tonalite, (ii) a pink, magnetic, medium- to coarse-grained, mostly massive, biotite leucogranite (e.g., the Lac La Croix Batholith), and (iii) a white-grey, medium-grained to pegmatitic, muscovite leucogranite (e.g., the Sturgeon Lake Batholith). He proposed that the muscovite leucogranites were S-type granites with a metasedimentary source, whereas the pink biotite granite had lower SiO2, and higher Na2O, K2O, and Sr than the S-type granite and could have been derived from either an S-type or I-type source (Percival, 1989). The Dog Lake Granite Chain, consisting of six ovoid magnetite-bearing intrusions (Shabaqua, Silver Falls, Trout Lake, Barnum Lake, White Lily and Penasen Lake), is characterized by a linear trend that parallels the tectonic boundary between the Wawa-Abitibi terrane to the south, and the Quetico Basin to the north. The intrusive rocks in the Dog Lake Granite Chain can be divided into three units: monzodiorite, microcline-phyric monzonite/quartz monzonite, and granite. The majority of monzodiorite and monzonites are metaluminous, whereas the granites are peraluminous. Whole rock data show that the Al2O3 and LREE (La, Ce, Pr and Nd) contents of the monzodiorite show a positive correlation with SiO2, whereas the monzonite and granite units show a negative correlation. Similarly, the K2O content of the monzodiorite unit correlates positively with SiO2 content, whereas the monzonite and granite units show no strong correlation. The HREE (Er, Tm, Yb and Lu) contents of the monzodiorite and monzonite units decrease with SiO2 content, whereas in the granites they increase. In combination this suggests that the three units were likely derived from different magmatic sources. Mantle-like whole rock ɛNd (+1.30 to +1.76) and zircon ɛHf (+0.34 to +7.27) values, and arclike geochemical signatures characterized by the enrichment of large ion lithophile elements (LILE) and low HSFE, suggesting that the monzodiorite unit was likely generated by partial melting of the mantle wedge above a subduction zone. The monzonite units show I-type granite 97 �signatures with the positive whole rock ɛNd (+0.81 to +1.23) and slightly enriched zircon ɛHf (0.21 to +3.81) values, consistent with them having formed from re-melting of metavolcanic rocks. The S-type granites exhibit positive Ce anomalies, negative Eu anomalies and a ranges of whole rock ɛNd (-1.75 to +0.43) and zircon ɛHf (-2.04 to +3.37) values, suggesting a mixed source comprised of arc-like orogenic sediments and minor metavolcanic rocks. Therefore, we conclude that the addition of voluminous melt generated by partial melting of arc-like orogenic sediments likely caused the transition from I-type to S-type magmas as the magmatic system evolved. References Percival, J.A., and Williams, H.R. 1989. The Late Archean Quetico accretionary complex, Superior Province, Canada. Geology 17(1), 23-25. Sawyer, D., 2002. "Discovering plate boundaries: A classroom exercise designed to allow students to discover the properties of tectonic plates and their boundaries." Rice University. http://plateboundary.rice.edu/intro.html Accessed 17 December. Williams, H.R., 1991. Quetico Subprovince. In: Thurston, P.C., Williams, H.R., Sutcliffe, R.H., and Stott, G.M. (Eds.), Geology of Ontario; Ontario Geological Survey Special Vol. 4. 98 �Mineral deposits of the Midcontinent Rift System - A new space/time classification WOODRUFF, Laurel G.1, NICHOLSON, Suzanne W.2, DICKEN, Connie L.2, and SCHULZ, Klaus J.2 1 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 U.S. Geological Survey, MS 954, 12201 Sunrise Valley Drive, Reston, VA 20192 2 The Midcontinent Rift System (MRS) hosts a diverse suite of magmatic and hydrothermal mineral deposits, largely known from rift rocks exposed at or near the surface in the Lake Superior region (Nicholson et al., 1992). Most of these deposits, which are significant past, present, and likely future providers of critical minerals, can be placed into a new space/time metallogenic framework (Fig. 1). This framework was developed using 552 mineral deposits compiled from the U.S. Geological Survey Mineral Resources Data System (MRDSa) and the Ontario Ministry of Energy, Northern Development and Mines Mineral Deposit Inventory (MDIb). Deposits were classified by deposit type, host rock age and type, and estimated mineralization age. The deposits were then put into a tectonic evolutionary framework for the MRS, which showed that many deposits formed in unique spatial and temporal stages of rift evolution. The distribution of 106 zircon/baddeleyite age dates, also compiled in this study, reflects three main magmatic MRS stages: 1) an early Plateau Stage from ~1113 to ~1105 Ma, characterized by widespread subaerial volcanism (e.g., magnetically reversed North Shore Volcanic Group, Osler Volcanics, Siemans Creek Volcanics) and related intrusive activity (e.g., Coldwell Complex, Early Gabbroic/ Felsic Series of the Duluth Complex); 2) a Rift Stage (~1102 to ~1091 Ma), characterized by eruption of thick sections of subaerial flood basalts largely confined to central, sagging basins (e.g., magnetically normal North Shore Volcanic Group, Portage Lake Volcanics) accompanied by voluminous intrusive events (e.g., Anorthosite/Troctolite Series of the Duluth Complex, Beaver Bay Complex, and Mellen Complex); and 3) a Late-Rift Stage (~1090 to ~1083 Ma) with diminished sporadic, mainly andesitic/felsic volcanism (e.g., Lakeshore Traps, Michipicoten Island Volcanics). A Post-Rift Stage is dominated by sediment deposition from the margins of the rift as subsidence within the central basins continued because of thermal collapse. This created thick sections of sedimentary rock (Copper Harbor Conglomerate, Nonesuch Formation, and Freda Sandstone that comprise the Oronto Group) that overlie stacked basalt flows within rift basins. The time frame for Oronto Group deposition and its relationship to clastic sediments of the Bayfield Group and Jacobsville Sandstone, the youngest rocks assigned to MRS history, are poorly constrained. The last event that put a close to the MRS was a Compressional Stage (~1060 and ~1040 Ma) that created reverse faults along some margins of the rift and carried older rocks over younger. Mineralization in the MRS evolved within this broad tectonic context, beginning with intrusive magmatic deposits that formed contemporaneously with intrusion of MgO-rich mafic/ultramafic magmas during the Plateau Stage. Deposit types in this stage include: 1) small conduit-type sulfide deposits (e.g., the Ni-Cu-PGE Eagle, Tamarack, and Thunder Bay North deposits, and the Cu-PGM Marathon deposit); 2) layered Ti-Fe-(V) deposits in the Duluth Complex Early Gabbro Series; and 3) Nb-U(±Th±REE) in alkaline intrusions in Ontario. Magmatic mineral deposits related to the MRS Rift Stage include large contact-type disseminated Cu-Ni-PGE sulfide deposits and small but potentially economic Ti-Fe(-V) oxide ultramafic deposits along the basal section of the Duluth Complex. With diminished volcanism and increased sedimentation during the Late/Post-Rift Stages, hydrothermal fluids became an 99 �increasingly important component of mineralization. Deposits that formed during this time frame are thought to include: 1) chalcocite in basalt (e.g., 543S); 2) copper sulfide veins (e.g., Coppercorp); and 3) Cu-Mo breccia pipes in the Mamainse Point area, which all may have a hybrid magmatic/hydrothermal origin. Additional hydrothermal deposits within this time period are: 1) a complex group of metal-bearing veins in Ontario (e.g., Ag-bearing veins of the Mainland and Island Groups, unconformity Pb-Zn-Ba(±U) veins); and 2) stratiform Cu deposits in the Nonesuch Formation (e.g., White Pine and Copperwood). The final, major MRS mineralizing event occurred during the Compressional Stage and created the world-famous Keweenawan native Cu(±Ag) deposits, contained within the Rift Stage Portage Lake Volcanics. This space/time classification of MRS mineral deposits is outlined in a USGS Story Mapc. Figure 1. Simplified geologic map of the Lake Superior region, showing pre-MRS and MRS-related rocks, distinguished by stage, type, and magnetic polarity. Mineral deposits compiled from USGS and OGS databases have been classified by deposit type and put into a time-space evolutionary model for the MRS. Geology and mineral deposits taken from the USGS MRS geodatabase. Nicholson, S.W., Cannon, W.F., and Schulz, K.J., 1992, Metallogeny of the Midcontinent rift system of North America: Precambrian Research, v. 58, p. 355-386. a https://mrdata.usgs.gov/mrds/ https://www.mndm.gov.on.ca/en/mines-and-minerals/applications/ogsearth/mineral-deposits-mdi c https://wim.usgs.gov/geonarrative/MRS_mineral_deposits/ b 100 �Sulfur mobility in arc magma systems: Implications for porphyry ore deposits WRAGE, Jackie1, FIEGE, Adrian1, KONECKE, Brian1, SIMON, Adam1, RUPRECHT, Philipp2, BEHRENS, Harald3 1 Department of Earth and Environmental Sciences, University of Michigan, Ann Arbor, Michigan, USA 2 Department of Geological Sciences and Engineering, University of Nevada, Reno, Nevada, USA 3 Institute of Mineralogy, Leibniz University Hannover, Callinstrasse 3, 30167 Hannover, Germany Porphyry ore deposits supply two-thirds of the world’s Cu and nearly all of its Mo, as well as significant amounts of Au, Ag, and critical elements such as Re, Se and Te. These deposits form as a result of arc-related volcanism, when partial melts generated by dehydration melting of the subducting basaltic ocean crust percolate upwards through the mantle wedge and accumulate at the base of the crust in a process called underplating. As these mafic magmas fractionate, felsic melts segregate and ascend to the middle and upper crust where they form magma chambers that are thought to be the source of ore fluids in porphyry systems. However, a major problem with sourcing porphyry fluids from intermediate to felsic magmas is that mass balance calculations indicate that such silicic magmas cannot supply all of the S in porphyry ore deposits. The most plausible explanation for the excess S in porphyry deposits is underplating of middle to upper crustal silicic magma chambers by decompressing, volatile-saturated mafic magma that delivers volatiles such as S, H2O, and Cl, and possibly metals into the overlying felsic magma. This study explores the effects of underplating on volatile exchange between a mafic recharge magma and felsic host magma by simulating an underplating scenario. Diffusion-couple experiments were performed wherein a cylinder of mafic magma (basaltic andesite) was juxtaposed beneath a cylinder of felsic magma (dacite) and run under a range of pressuretemperature-composition-redox conditions relevant for upper crustal arc magma (porphyry) systems. The most intriguing finding is the development of a redox gradient of ~1.8 log units fO2 at the mafic-felsic interface of the most oxidizing (FMQ+4) experiments, where the mafic melt is oxidized, and the felsic melt is reduced. Sulfur x-ray absorption near-edge structure (S-XANES) analyses also indicate complex S-speciation near the mafic-felsic interface in the most reducing (FMQ+1) experiment. Such a gradient affects the speciation of redox-sensitive elements such as S and moderates mass transfer from mafic to felsic melt, as well as affecting the metalscavenging potential of an exsolved magmatic-hydrothermal volatile phase. Studying the effects of underplating in arc systems is paramount to understanding the source(s) and mechanisms responsible for the titration of volatiles and metals into ore forming environments and could help reconcile the excess sulfur problem in volcanic systems. 101 �Multiscale Layering in the Black Sturgeon Sill, Nipigon, Ontario ZIEG, Michael J. Department of Geography, Geology, and the Environment, Slippery Rock University, Slippery Rock, PA 16057 The basic petrology of the Nipigon sills was established by Richard Sutcliffe thirty years ago: “The diabase occurs primarily as … 150 to 200 m thick sills with a textural stratigraphy indicating that the sills represent single cooling units. Compositional variation in the sills indicates that they crystallized from several magma pulses.” (Sutcliffe, 1987) “Fractionation in the sills is attributed to flowage differentiation and movement of residual liquids … toward the top of the sections.” (Sutcliffe, 1989). The current study has confirmed and amplified these interpretations of sill-scale processes. A continuous drill core through the Black Sturgeon sill (Zieg & Wallrich, 2018), a 250 m Nipigon group (Hollings et al., 2007; 2010) mafic intrusion, has yielded far more detailed stratigraphic variations than were previously available. With this newly-available data, Sutcliffe’s original model can be shown to apply at multiple scales. Specifically, although textural evidence for “non-unit” a cooling history remains elusive, smaller-scale magma batches can be recognized within the larger-scale magma pulses. Additionally, there is compositional evidence for upward movement of evolved liquids within any sub-section of the sill. These patterns have been traced down to a scale on the order of tens of centimeters, with individual textural and mineralogical discontinuities visible on the thin section scale. The relationship between classical igneous layering and the observed compositional and textural variations is unclear. Much of the data from the Black Sturgeon sill suggests that olivine (and plagioclase) concentrations in this system reflect the injection of antecryst-laden magma. A similar type of analysis on a more traditionally layered intrusion might shed light on the extent to which the processes that controlled layer formation in the BSS also contributed to layer formation in other systems. References Hollings, P., Hart, T., Richardson, A. & MacDonald, C. A. (2007). Geochemistry of the Mesoproterozoic intrusive rocks of the Nipigon Embayment, northwestern Ontario: Evaluating the earliest phases of rift development. Canadian Journal of Earth Sciences, v. 44, p. 1087–1110. Hollings, P., Smyk, M., Heaman, L. M. & Halls, H. (2010). The geochemistry, geochronology and paleomagnetism of dikes and sills associated with the Mesoproterozoic Midcontinent Rift near Thunder Bay, Ontario, Canada. Precambrian Research, v. 183, p. 553–571. Sutcliffe, R. H. (1987). Petrology of Middle Proterozoic diabases and picrites from Lake Nipigon, Canada. Contributions to Mineralogy and Petrology, v. 96, p. 201–211. Sutcliffe, R. H. (1989). Mineral variation in Proterozoic diabase sills and dykes at Lake Nipigon, Ontario. The Canadian Mineralogist, v. 27, p. 67–79. Zieg, M.J, & Wallrich, B.M. (2018). Emplacement and differentiation of the Black Sturgeon Sill, Nipigon, Ontario: A principal component analysis. Journal of Petrology, v. 59, p. 2385–2412. 102 �Figure 1. Compositional profiles. Higher Ni concentrations coincide with olivine-dominated portions of the system, regardless of scale. Distinct compositional layering can be recognized at all three scales. Figure 2. Textural profiles. Mean plagioclase length is typically lower in parts of the system that have accumulated olivine. This pattern can be traced down to olivine accumulations less than a meter thick. Although textural variations can be recognized at all three scales, there is little clear evidence for discrete cooling units. Rather, textural variations apparently reflect different magma batches (crystal cargo). 103 � https://digitalcollections.lakeheadu.ca/files/original/86325f55313fe9ffc025e80c4ac4582c.pdf b1fa39b1d9f97f9eb95b50db71f9d1b8 PDF Text Text 65th Annual Meeting Terrace Bay, Ontario - May 8-9, 2019 Institute on Lake Superior Geology Part 2 – Field Trip Guidebook �Thank you to our sponsors! Individual contributors to student travel scholarship: Al MacTavish, Mary Kay Arthur, L. Gordon Medaris, Jr., Nick Swanson-Hysell �65th Annual Meeting Institute on Lake Superior Geology May 8-9, 2019 Terrace Bay, Ontario HOSTED BY: Mark Smyk and Pete Hollings Co-Chairs Ontario Geological Survey and Lakehead University Proceedings - Volume 65 Part 2 – Field Trip Guidebook Compiled and edited by Al MacTavish and Pete Hollings Cover Photos: Left - Pillowed Archean metabasalt, Schreiber Beach, Middle - Layered Eastern Border Gabbro, Coldwell Complex. Right - Foliated Archean metavolcanic rocks, Slate Islands. ��65th Institute on Lake Superior Geology Volume 65 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trip 1: The Slate Islands Trip 2: Midcontinent Rift-Related Carbonatites and Diatremes Trip 3: Geology of the Western Schreiber-Hemlo Greenstone Belt Trip 4: Geology of the Nipigon Area Trip 5: A stratigraphic transect across the Northern flank of the Midcontinent Rift near Rossport Trip 6: Geology of the Coldwell alkaline complex Trip 7: Building and ornamental stone sites of the Marathon Area, Ontario Trip 8: Geology of the past-producing Winston Lake Cu-Zn Mine Reference to material in Part 2 should follow the example below: Magnus, S., 2019. Geology of the Western Schreiber-Hemlo Greenstone Belt. In; MacTavish, A. and Hollings, P. (Eds.), Institute on Lake Superior Geology Proceedings, 65th Annual Meeting, Terrace Bay, Ontario, Part 2 - Field trip guidebook, v.65, part 2, 3-31. Published by the 65th Institute on Lake Superior Geology and distributed by the ILSG Secretary: Pete Hollings - ILSG Secretary Department of Geology Lakehead University 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Email: peter.hollings@lakeheadu.ca ILSG website: www.lakesuperiorgeology.org ISSN 1042-9964 ��Proceedings of the 65th ILSG Annual Meeting - Part 2 Table of Contents 65th Annual Meeting..........................................................................................................A Introduction, safety considerations and acknowledgements................................................1 Field trip 1 - The Slate Islands.............................................................................................2 Field trip 2 - Midcontinent Rift-Related Carbonatites and Diatremes.................................3 Field trip 3 - Geology of the Western Schreiber-Hemlo Greenstone Belt.........................14 Field trip 4 - Geology of the Nipigon Area........................................................................43 Field trip 5 - A stratigraphic transect across the Northern flank of the Midcontinent Rift near Rossport.............................................................................................................60 Field trip 6 - Geology of the Coldwell alkaline complex..................................................75 Field trip 7 - Building and ornamental stone sites of the Marathon Area, Ontario..........105 Field trip 8 - Geology of the past-producing Winston Lake Cu-Zn Mine.......................113 -i- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Introduction, safety considerations and acknowledgements Pete Hollings Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada and Mark Smyk Resident Geologist Program, Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, Thunder Bay, Ontario, P7E 6S7, Canada This volume is intended to serve not only as a guide for 65th ILSG field trip participants but also as a reference for those planning to revisit these areas at a later date. Consequently we have included UTM coordinates in the NAD 83 datum for stops, as well as instructions on how to reach them. As some of the stops are on private and staked land, please be sure to obtain the land owners’ permission before entering their land. Contact the staff of the Resident Geologist Program in Thunder Bay for current ownership information. We are once again offering field trips onto Lake Superior. This creates a number of unique safety issues. Please exercise caution when getting in and out of the boats as the outcrops are often extremely slippery. Personal flotation devices must be worn in the boats at all times. If you are planning to revisit these sites please be very careful, as Lake Superior is dangerous; waves can often be many metres high and even in midsummer fog can appear very quickly. either major highways or busy logging roads. Please take care when crossing or parking along these roads. We would like to thank all the other authors who contributed to this field guide, all those who provided comments and/or assisted with the running of the field trips themselves (Shannon Zurevinski, Rob Cundari, Phil Fralick, Dorothy Campbell, Mark Puumala, Robert Lodge, Al MacTavish, Dave Good, John McBride, Peter Hinz, Seamus Magnus, Bill Skrepichuk). We appreciate the assistance and cooperation of the exploration and mining companies in providing us access and information concerning their properties, particularly Superior Lake Resources Ltd. (Winston Lake Mine), Plato Gold Corp. (Good Hope Niobium), Rudy Wahl (Madonna), Jerry Blakely (Shack Lake), Red Rock Indian Band (Ruby Lake), Alex Pleson (Stenlund), John Ternowesky (Dead Horse North) , Stillwater-Sibanye (Marathon) and Ontario Parks (Slate Islands and Neys). The other field trips will be visiting stops along Figure 1. Map showing the location of the eight field trips offered in 2019. -1- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 1 - The Slate Islands Pete Hollings Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada Mark Smyk Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, Thunder Bay, Ontario, P7E 6S7, Canada Bill Addison and Phil Fralick Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada The field guide for the Slate Islands has previously been published as ILSG Special Publication #1 and is available on the ILSG website - www.lakesuperiorgeology.org InstItute on Lake superIor GeoLoGy Special publication #1 FIeLd trIp GuIdebook For the sLate IsLands, ontarIo pete hoLLInGs, Mark sMyk, bILL addIson & phIL FraLIck -2- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 2 - Midcontinent Rift-Related Carbonatites and Diatremes Shannon Zurevinski Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada Dorothy Campbell and Mark Puumala Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, Thunder Bay, Ontario, P7E 6S7, Canada Introduction The Midcontinent Rift System (MCR) of North America is one of the largest known aborted rifts, and extends ~2200km from Kansas, north to the Lake Superior area. There are several MCR-related fault systems in Lake Superior and the surrounding terrain, as evidenced by seismic reflections, gravity and various magnetic anomalies. Some of the most fascinating geology of the Terrace Bay-Marathon area of Northwestern Ontario is where Archean rocks of the Superior Province are intruded by diatremes, alkalic rocks, carbonatites and ultramafic lamprophyres, which are mostly related to the MCR and the Trans-Superior Tectonic Zone (TSTZ). Many of these intrusions are located along the strike of the Big Bay-Ashburton Fault (which is the proposed northern extension of the Thiel Fault), representing the most northerly component of the TSTZ (Sage, 1982). The Coldwell Alkalic Complex, Killala Lake Alkalic Complex, Chipman, Prairie Lake and Good Hope Carbonatite occurrences, the Dead Horse Creek Diatreme, and various ultramafic lamprophyres lie along the extrapolated arm of the fault system (Fig. 1). We would like to acknowledge the earlier work of Ron Sage and David H. Watkinson, who guided an extensive field trip into the Alkalic rocks of the MCR during the 41st Annual Meeting of the Institute in 1995. This field trip will serve as an update to the prospecting and exploration completed over the last 25 years and will visit: 1) the North Dead Horse property; 2) the diamondiferous Madonna dyke; 3) the Prairie Lake Carbonatite; and 4) the Good Hope Carbonatite (Fig. 2). We also acknowledge Rudy Wahl for his hard work and boots-on-the-ground prospecting that has led to the discoveries of both the Madonna dyke and Good Hope Carbonatite. A special thanks to Seamus Magnus for providing assistance with the descriptions of the North Dead Horse subcomplex trenches. Figure 1. A regional map of the area north of Lake Superior between Terrace Bay and Marathon, showing the alkaline complexes, diatremes, carbonatites and ultramafic lamprophyres of the area. (modified from Smyk et al., 1993). Acknowledgements We are grateful to Rudy Wahl (Madonna dyke and Good Hope Carbonatite), John Ternowsky (North Dead Horse Property) and Nuinsco Resources (Prairie Lake Carbonatite) for permissions to access the properties for this field trip. Stop descriptions Stop 1: The North Dead Horse Creek Diatreme -3- UTM Coordinates 523959E 5409978N (parking lot) �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 2. General geology with field trip stops (geology from the Ontario Geological Survey). Introduction The Dead Horse Creek Diatreme is hosted by the metasedimentary and metavolcanic rocks of the Schreiber-White River greenstone belt, within the Wawa Subprovince (Sage, 1982). It is located approximately 1km from the western margin of the Coldwell Alkalic Complex (Fig. 2). Sage (1982) describes the complex as a broad spectrum of heterolithic breccias that have undergone varying degrees of alteration and are variably radioactive, and subsequently divided the complex into five subcomplexes (North, South, East, West, and Central; Fig. 3). Past exploration programs by Gulf Minerals Canada (1977) focused on the uranium mineralization of the West Dead Horse and the North Dead Horse subcomplexes (Fig. 3). Subsequent exploration on the property by Highwood Resources Ltd. (1985) focused on exploration for beryllium, yttrium, and cerium; Unocal Canada’s (1987) main interest was exploring the potential for yttrium; while Canadian International Mining (2011) focused on exploration for the Rare Earth Elements (REEs). The mineralized zone of the West Dead Horse subcomplex has been described as “diverse, exotic, hydrothermally altered, and rare metal mineralized” (Sage, 1982). Unpublished U-Pb geochronology of zircons from the Figure 3. The Dead Horse Creek Diatreme geology map. Shown here are the 5 subcomplexes. Also note the Gulf minerals drilling program (investigated for U). After Sage (1982) -4- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Dead Horse Creek West subcomplex, were found to be 1128.7 ± 6 Ma (Sage, 1995). The West Dead Horse subcomplex (400m x 1600m elongate in a N-NE direction) has been the focus of extensive geochemical and mineralogical studies, and is described as an occurrence of heterolithic and carbonate-rich breccias and veins (Smyk et al., 1993; Potter and Mitchell, 2005). Smyk et al. (1993) reported Heavy Rare Earth Element (HREE) enrichment from the mineralized zone of the West Dead Horse occurrence to be up to 1004ppm ΣHREE. Potter and Mitchell (2005) summarized the REE mineralogy of the occurrence, including a phenakite-bearing quartz vein, Ca-zirconosilicate, zircon, thorite, uraninite, apatite, xenotime-Y, monazite-Ce and rutile. Smyk et al. (1993) proposed that the volcaniclastic breccia formed during the early stages of the MCR event and incorporated clasts of Archean metasedimentary and granitic rocks. They proposed that after the emplacement of the volcaniclastic breccia, U-Be-Zr mineralization occurred along fault structures after being introduced via A-type granitic fluids, and was followed by alkaline metasomatism. The mineralized zone is not related to the igneous activity that produced the volcaniclastic breccia, rather the porous breccia allowed for the deposition of the U-Be-Zr mineralization. Potter and Mitchell (2005) provide a genetic model for the complex based on the exotic mineral suite present (including the accessory REE mineralogy), and concluded that upon emplacement of the nearby Coldwell Alkalic complex (1108 ±1 Ma; Heaman & Machado, 1992), the Dead Horse volcaniclastic breccia was subjected to thermal metamorphism and post-Pleistocene supergene alteration. Nb-Ti-V-Cr-bearing alkaline fluids were introduced into the same fault system and reacted with the initial mineralization, creating the suite of exotic minerals which was further diversified by supergene alteration (Potter and Mitchell, 2005). The remaining question: What is the source of these Nb-Ti-V-Cr-bearing fluids? While the Coldwell Alkalic Complex does contain the A-type granitic rocks that would produce these fluids, the timing of the emplacement of both the Coldwell Alkalic Complex and the Dead Horse Creek diatreme would need further geochronological studies to constrain this relationship. Canadian International Minerals Inc. (CIM) recently conducted an extensive exploration program on the North Dead Horse property, including trenching, assaying, and geophysical surveys. The showing is arranged in a cross, and the E-W trending stripping is ~250m, while the N-S stripping is ~320m in length (Fig. 4). This recently exposed trenching provides an Figure 4. The North Dead Horse Creek trench map (modified from Magnus 2019). -5- �Proceedings of the 65th ILSG Annual Meeting - Part 2 excellent opportunity to observe the breccia and its various components. Stop 1a: North Dead Horse: Trachytic Diabase Dyke UTM Coordinates 524072E 5410174N Occurring along the trail from the parking area is a trachytic diabase dyke. This is described as a mafic dyke containing feldspar phenocrysts with a magnetic matrix (Fig. 5). These dykes are widespread throughout the area and also occur within the Coldwell Alkalic Complex. Figure 6. Photograph of the grey breccia from the North Dead Horse trench, cut by an alkalic dyke. Figure 5. (a - top) Photograph of the Trachytic diabase dyke from Stop #1. (b - bottom) Close-up photograph of the feldspar phenocrysts of the Trachytic diabase dyke. Stop 1b: North Dead Horse: East-West Trench Figure 7. Photograph of the dark grey breccia from the North Dead Horse trench, showing large clasts of banded metasediments, and hematized and zoned fragments, as well as chert fragments. Grey and red heterolithic breccias are found at the North Dead Horse trenches, cut by various mafic and intermediate dykes (Fig. 4). (Figs. 6 and 7). CIM describes the matrix of the grey breccias as carbonate-rich. The clasts are generally composed of amphibolite, wacke, and granitoid rocks (Fig. 6). Some wacke clasts are several meters in length. Brecciation of the various clasts appears to be found along bedding and schistosity planes. The light grey breccia is host to relatively unaltered clasts that show their primary and structural fabrics The red breccia is dominant in the western part of the trench and is highly altered (Fig. 8). The clasts are UTM Coordinates 524210E 5410389N -6- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 8. Photographs of the red breccia from the North Dead Horse trench, showing altered hematized metasediments, and both granitic and mafic metavolcanic clasts. composed of granitic and mafic metavolcanics, altered metasediment, and lesser chert. The metasedimentary clasts are hematized and heavily altered at the rims, with the degree of alteration reduced towards the core of the clasts. Some clasts are sulphide-rich. The matrix is magnetic and there is carbonate and chlorite alteration present. Rare earth mineralization on the property appears to be associated with areas containing the red breccia unit (Fig. 4). Minerals identified by Potter and Mitchell (2005) on the property include albite, potassium feldspar, quartz, calcite, apatite, phenakite, aegirinejervisite, aegirine-natalyite, allanite, barite, barylite, coffinite, Ca-Mn-silicate, magnetite, monazite-(Ce), niobian vanadium rutile, pyrite, thorite, thoro-gummite, thortveitite, uraninite, vanadium crichtonite, xenotime(Y), ankerite-dolomite and zircon (Quist, 2011). (S. Magnus, personal communication). The east end of the E-W trench is host to a recessively weathered dyke with an unknown, possibly carbonatitic composition. Stop 2: The Madonna Dyke UTM Coordinates 530377E 5427200N Introduction Located approximately 30kms northwest of Marathon, the Madonna Dyke was discovered by Rudy Wahl in 2007 (Fig. 9). The Madonna Dyke is North-South trench The north-south trench contains mostly grey breccias. The grey breccia has large, relatively unaltered metasediment, amphibolite, and granitic clasts (as described above). Intermediate alkalic dykes, mafic dykes, and late fractures crosscut all breccias. The mafic dykes are similar to the dyke from Stop 1a (and occur throughout the area and within the Coldwell Alkalic Complex). The intermediate alkalic dykes are fine-grained and grey to pink in colour. The intermediate dykes are less common than the mafic dykes. Preliminary studies have shown the intermediate alkalic dykes to have up to 2000ppm ΣREEs, and it is reported that the altered and mineralized zones of North Dead Horse are coincident with these intermediate REE-enriched alkalic dykes Figure 9. Photo of the Madonna dyke (photo courtesy of R. Wahl). -7- �Proceedings of the 65th ILSG Annual Meeting - Part 2 approximately 1 to 2m in width, outcrops for ~50m, strikes 009°, and dips 65° towards the west (Wahl website). Sampling recovered 66 diamonds from a 1205.80 kg sample, which includes white, green, yellow, brown and grey diamonds (http://users.renegadeisp. com/~rwahl/Kimberlite%20Targets%20Available%20 for%20Option.htm). In 2018, R. Wahl completed drilling 82m to the southwest of the Madonna Dyke and intersected the same diamondiferous dyke over a width of 2.78m. The dyke is underlain by Neoarchean biotite granite gneiss and hornblende granite gneiss, as well as post tectonic quartz monzonites (Coates, 1970). The dyke intrudes near lineament intersections that could represent zones of crustal weakness that may have acted as a pathway for the ascent of mantlederived magmas (Kozlowski, 2016). Classification The Madonna dyke been classified as a diamondiferous alnöite ultramafic lamprophyre (Kozlowski, 2016). It is described as a mafic hypabyssal rock with medium- to fine-grained rounded phenocrysts and a fine-grained dark-green to black groundmass (Fig. 10). The dyke has an orange-brown rind on its weathered surface. The phenocrysts (1 to 10mm in diameter) are estimated to make up to 50% of the modal abundance and are identified as pseudomorphs after olivine, pyroxene, and oxides rimmed by carbonate and lesser late-stage calcic amphiboles (Kozlowski, 2016). The groundmass consists of calcite (after melilite), Figure 10. Photo of the Madonna dyke showing rounded phenocrysts set in a dark green groundmass. Also note the orange-brown rind on the weathered surface (photo courtesy of R. Wahl). phlogopite, magnetite, apatite and some alteration products (Kozlowski, 2016). Mineralogy of the Madonna Dyke Kozlowski (2016) summarizes the mineral chemistry of the Madonna dyke used to provide proper classification of the dyke. Pseudomorphed olivine occurs as microphenocrysts, phenocrysts, and rare macrocrysts replaced by serpentine, magnetite, and calcite. A few fresh olivine macrocrysts show mantle compositions ranging from Fo91 to Fo92 (Kozlowski, 2016). Clinopyroxenes are aluminous diopside with Al2O3 ranging from 3.11 to 14.47 wt.%. Groundmass micas have kinoshitalite–phlogopite compositions, with up to 4 wt.% BaO and 20.9 wt.% Al2O3. Spinelgroup mineral compositions follow Magnetic Trend #2 – the Titanomagnetite Trend, where spinels range in composition from aluminous magnesian chromite to titanian magnesian chromite to titanian chromite to members of the ulvöspinel-magnetite series (Fig. 11; Kozlowski, 2016). Spinel-group minerals occur as red chromium spinel phenocrysts to macrocrysts with magnesium-rich cores and iron-rich rims, often associated with olivine phenocrysts and macrocrysts. They also occur as fine-grained opaque groundmass titanomagnetites with altered cores, and as reaction products forming a necklace texture around olivine. Atoll spinel is present. Although the Madonna dyke shows some textural and petrogenetic features of kimberlites, the mineralogy, including the presence of calcite after melilite and amphibole, are analogous with an ultramafic lamprophyre of alnöitic affinity Figure 11. Reduced spinel prism with compositions of Madonna dyke spinels (blue = core; red = rim) showing a magmatic trend 2 – the titanomagnetite trend (green arrow) with a hiatus in the middle (classification after Mitchell, 1986). -8- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Table 1 Summary of petrographic features of the Madonna Dyke compared to kimberlite and UML (ultramafic lamprophyre). After Mitchell (1995b) and Birkett et al. (2004). Olivine (macrocrysts) Kimberlites common rare (phenocrysts) common common common Mica common phlogopite common phlogopite not observed (groundmass) common, phlogopite kinoshitalite common Al-biotite common, phlogopite kinoshitalite Spinels abundant, Mg-chromite to Mg ulvöspinel common, Mg-chromite to Ti-magnetite common, Mg-chromite to Ti-magnetite (atoll) very common present present (necklace) present present present Perovskite common, Sr- and REEpoor common, Sr- and REEpoor not observed Diopside absent common, Al- and Tirich common, Al- rich Apatite common, Sr- and REEpoor common, Sr- and REEpoor common, Sr- and REEpoor (skeletal) rare rare common Calcite abundant common common Melilite absent common common Amphibole absent present present (phenocrysts) (Kozlowski, 2016). Table 1 provides a summary of the features of the Madonna dyke compared to kimberlite rocks and ultramafic lamprophyres. Stop 3: The Prairie Lake Carbonatite Complex UTM Coordinates 520150E 5431100N Introduction The Prairie Lake Carbonatite Complex is located ~26 km from the shores of Lake Superior (Figs. 1 & 2). It covers a surface area of ~8.8 km2 and generally consists of foidolitic and carbonatitic rocks (Fig. 12; Sage, 1987). It is generally a small arcuate intrusion emplaced in Archean gneisses along the TSTZ (Sage, 1987). The complex has been actively explored since 1968 for U and Nb, and is considered ‘multicommodity’ with its potential residual apatite deposits. There is significant modal heterogeneity of the rocks at the Prairie Lake complex. Wu et al. (2017) summarize the variation of the rock types present: 1) calcite carbonatite; 2) biotite pyroxenite; 3) the ijolitic series rocks; 4) potassic syenites; 5) heterogeneous UML rare Madonna Dyke carbonatites; and 6) rare dolomitic carbonatite (Fig. 12). The niobium mineralization in the Prairie Lake Carbonatite complex includes Na- and Ca- pyrochlore, latrappite, loparite, U-pyrochlore, Ce-pyrochlore, Pbpyrochlore, marianoite, and wohlerite (Mitchell, 2015). The pyrochlore has a wide range of compositions and are complexly zoned and resorbed. The Nb mineralization is distributed mainly between perovskite, pyrochlore, and Nb-Zirconolite and tends to be REE-poor with in situ alteration (Mitchell, 2015). Zurevinski and Mitchell (2015) describe the only known worldwide occurrence of orbicular ijolite from within the Prairie Lake Complex (Fig. 13). This occurrence had been previously noted and described by Sage (1987, 1995). The orbicules occur in an ijolite matrix, and the mineralogy of the orbicules is similar to that of their host ijolite (nephelene, diopside, calcite, apatite, andradite-melanite garnet, titanite, etc.; Fig. 13; Zurevinski and Mitchell, 2015). Detailed mineralogy and petrology have shown that the orbicular ijolite represents an interaction of a partially crystallized quenched ijolitic melt, in contact with a second pulse -9- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 12. A map of the Prairie Lake Carbonatite complex. From Sage (1987). of consanguineous ijolite magma. Immersion in the latter resulted in a sub-solidus diffusion and annealing recrystallization (i.e., magma mixing; Zurevinski and Mitchell, 2015). Figure 13. Orbicular ijolite from the Prairie Lake Carbonatite Complex (from Zurevinski and Mitchell, 2015). Recent detailed geochronology has been completed on the Prairie Lake complex by Wu et al. (2016). In summary, U-Pb with baddeleyite from the carbonatite gave emplacement ages of 1157 ± 2.3 Ma and 1158 ± 3.8 Ma; baddeleyite from the ijolite-series rocks gave 1163 ± 3.6 Ma and U-Pb apatite from the carbonatite gave an emplacement age at ~1160 Ma. Therefore, the findings of Wu et al. (2016) reveal that all units were synchronously emplaced at ~1160 Ma. Furthermore, Sr-Nd-Hf tracer isotopic studies showed that the Prairie Lake carbonatites, ijolites, and syenite rocks had identical isotopic composition, therefore, the silicate and carbonatite rocks are co-genetic and thereby related by fractional crystallization processes (Wu et al., 2016). This data has reinforced the previous conclusions that Prairie Lake is the earliest manifestation of midcontinent rift magmatism, and is not genetically related to the nearby Coldwell or Killala Alkalic complexes (Rukhlov and Bell, 2010). - 10 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 3: Prairie Lake Calciocarbonatite (Sövite) (West) The area is noted by the prominent hill surrounded by low-lying marshy land. Generally, there is very sparse outcropping of extensively weathered calciocarbonatite. As you walk up the hill, you will view cobbles and boulders of calciocarbonatite and siliciocarbonatite rocks. (Figs. 14 and 15). The deep cuts on the right side of the trail will show deeply weathered carbonatite-rich soils, and in some areas, relict primary banding may be observed. The calcite carbonatite mineralogy varies throughout the complex, but generally contains calcite, apatite, olivine, phlogopite, pyrite, and magnetite. The modal layers and bands that are sometimes observed at Prairie Lake are represented by various oxides (commonly magnetite), interlayered between calcitedominant layers. Stop 4: The Good Hope Carbonatite UTM Coordinates 519363 E 5431721 N (Parking area) Introduction The Good Hope Carbonatite is located approximately 28km North of Hwy 17 and was discovered in 2015 by Rudy Wahl (Fig. 2). It occurs within the magnetic “low” on the Northwestern flank of the Prairie Lake Carbonatite, in low-lying marshy land. The Nb property has undergone mapping, geophysical surveys, trenching, and drilling since its discovery. Plato Gold Corp. has an option agreement with Rudy Wahl and other claim holders on the property. Figure 14. Banded sövite (calciocarbonatite) from the Prairie Lake Carbonatite Complex (photo courtesy of M. Smyk). The Good Hope property is host to carbonatite, ijolite, and alkali granite. Alkali granite is fine- to medium-grained and characterized by the abundance of orange and red potassium feldspar and quartz (Selway, 2017). At surface, the medium-grained carbonatite has a reddish-brown rind and appears in some cases to be brecciated (Fig. 16). The brecciated material shows angular to subround fragments with buff-white carbonate stringers present (Puumala et al., 2015; Fig. 16). Investigations from the drill core samples have Figure 15. Weathered carbonatite from the Prairie Lake Carbonatite Complex (photo courtesy of M. Smyk). Figure 16. Brecciated carbonatite from the Good Hope Carbonatite. - 11 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 led to the classification of the Good Hope carbonatitic rocks. Three types of carbonatite have been identified on the property: calciocarbonatite, ferrocarbonatite, and siliciocarbonatite (Cleaver, 2017). Alkali granitic breccia with carbonatite veins has been identified in the drill core (Selway, 2017). The veins correlate with the higher-grade mineralization (>1.0 wt. % Nb2O5) (Selway, 2017). Niobium mineralization is primarily concentrated in pyrochlore, which are characterized by low UO3 and are ThO2-free (https://www.platogold. com/projects/good-hope-niobium-project/).A pyrochlore-group mineral that has no ThO2 and low UO3 content is important as these radionuclides (Th, U) end up in the slag during processing and can be quite problematic. Other minerals found occurring within the carbonatite rocks include calcite, ferrodolomite, siderite, apatite, ferrocolumbite, mica, and pyrochloregroup minerals (Cleaver, 2017). Cleaver (2017) divided the carbonatites into two paragenetic varieties, pyrochlore-rich and pyrochlorepoor. Mineralogical and petrological evidence from the pyrochlore-rich carbonatites show early crystallized cumulates of apatite and Na-Ca pyrochlore minerals, while the pyrochlore-poor carbonatites appear to represent a later stage of crystallization (Cleaver, 2017). Cleaver (2017) concluded that the mineralogy of the Good Hope occurrence is different to that of the carbonatites occurring in the western and southern margins of the Prairie Lake carbonatite. The differences in mineralogy, coupled with a different magnetic signature, carbonatite texture, weathering profile, and distinct topography, indicate that the Good Hope occurrence is perhaps not directly related to the Prairie Lake carbonatite, however, the genetic relationship remains unknown (Cleaver, 2017). At this stop we will visit the areas that have been the focus of the exploration program over the last few years. Recent trenching has uncovered various outcrops of the ijolites, carbonatites and alkali granites present on the property. References Birkett, T.C., McCandless, T.E., and Hood, C.T. 2004. Petrology of the Renard igneous bodies; host rocks for diamond in the northern Otish Mountains region, Quebec. Lithos 76 (1): 475-490. Cleaver, A. 2017. Mineralogy and petrology of the Good Hope carbonatite occurrence, Marathon, Ontario. Unpublished HBSc. thesis, Lakehead University. Figure 17. Abundant pyrochlore in carbonatite at 102m, sample #1219065 (from Selway, 2017). Figure 18. Apatite mineralization in a carbonatite vein, as shown with a UV light system. Photo courtesy of R. Wahl. Coates, M.E. 1970. Geology of the Killala-Vein Lakes area, District of Thunder Bay, Ontario Department of Mines, Geology Report 81, 35p. Heaman, L.M. and Machado, N. 1992. Timing and origin of midcontinent rift alkaline magmatism, North America: evidence from the Coldwell Alkaline Complex. Contributions to Mineralogy and Petrology 110:289-303. Kozlowski, A. 2016. The mineralogy and petrology of the diamondiferous Madonna Dyke, Marathon, ON; unpublished HBSc. thesis, Lakehead University, Thunder Bay, ON, 72p. Mitchell, R.H. 1986. Kimberlites: Mineralogy, Geochemistry, and Petrology: New York, Plenum Press, 442 p. Mitchell, R.H. 1995. The role of petrography and lithogeochemistry in exploration for diamondiferous rocks. Journal of Geochemical Exploration, 53: 339350. Mitchell, R.H. 2015. Primary and secondary niobium mineral deposits associated with carbonatites. Ore Geology Reviews 64:626-641. Puumala, M.A., Campbell, D.A., Tims, A., Debicki, R.L., Pettigrew, T.K., and Brunelle, M.R. 2015. Report of Activities 2014. Resident Geologist Program, Thunder Bay South Regional Resident Geologist Report: Thunder Bay South District; Ontario Geological Survey, Open File Report 6303, 75p. Puumala, M.A., Campbell, D.A., Tuomi, R.D., Pettigrew, T.K. and Hinz, S.L.K. 2018. Report of Activities - 12 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 2017, Resident Geologist Program, Thunder Bay South Regional Resident Geologist Report: Thunder Bay South District; Ontario Geological Survey, Open File Report 6338, 101p. Sage, R.P. and Watkinson, R.H. 1995. Alkalic rocks of the Midcontinent rift, Institute of Lake Superior Geology, 41st Annual Meeting, Proceedings Vol. 41, Part 2a, Marathon, Ontario. Potter, E.G. and Mitchell, R.H. 2005. Mineralogy of the Dead Horse creek volcaniclastic breccia complex, Northwestern Ontario, Canada. Contributions to Mineralogy and Petrology, 150: 212-229. Selway, J. 2017. Assessment report for geological mapping program, Good Hope Niobium Property, Marathon, ON, Canada. Plato Gold Corp. 116p. Quist, B. 2011. Dead Horse Creek Rare Earth property, Walsh and Grain Townships, Thunder Bay Mining Division; Thunder Bay District, Assessment Report, AFRO 2.52396, 174p. Rukhlov, A.S. and Bell, K. 2010. Geochronology of carbonatites from the Canadian and Baltic Shields, and the Canadian Cordillera: clues to mantle evolution. Mineralogy and Petrology 98: 11-54. Sage, R.P. 1982. Mineralization in diatreme structures north of Lake Superior, Ontario Geological Survey Study, vol. 27, Ontario Ministry of Natural Resources, Toronto, p79. Sage, R.P. 1987. Geology of Carbonatite - Alkalic Rock Complexes in Ontario: Prairie Lake Carbonatite Complex, District of Thunder Bay; Ontario Geological Survey, Study 46, 9Ip Smyk, M.C., Taylor, R.P., Jones, P.C., and Kingston, D.M. 1993. Geology and geochemistry of the West Dead Horse Creek rare metal occurrence, Northwestern Ontario. Exploration and Mining Geology, 2:245251. Wahl, R. Wahl’s prospecting renegadeisp.com/~rwahl/ website. http://users. Wu, F.Y., Mitchell, R.H., Li, Q-L., Zhang, C., and Yang, Y-H. 2017. Emplacement age and isotopic composition of the Prairie Lake Carbonatite complex, Northwestern Ontario, Canada. Geological Magazine 154(2): 217236. Zurevinski, S.E. and Mitchell, R.H. 2015. Petrogenesis of orbicular ijolites from the Prairie Lake complex, Marathon, Ontario: Textural evidence from rare processes of carbonatitic magmatism. Lithos 239: 234-244. - 13 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 3 - Geology of the Western Schreiber-Hemlo Greenstone Belt Seamus Magnus Ontario Geological Survey, 933 Ramsey Lake Road Sudbury, ON, P3E 6B5 Canada Preface This field trip guidebook was prepared for a 1-day pre-meeting field trip held in conjunction with the Institute on Lake Superior Geology (ILSG) Annual Meeting hosted in Terrace Bay, Ontario from May 7 to 10, 2019. This geological guidebook was written to showcase the preliminary results of four years of bedrock mapping conducted by the author for the Ontario Geological Survey from 2015 to 2018 in the Schreiber–Hemlo greenstone belt (Magnus and Walker, 2015; Magnus and Arnold, 2016; Arnold et al., 2017; Magnus, 2017a,b; Magnus and Hastie, 2018). The coincidence of this meeting being held in Terrace Bay (within the mapping area) as the project was wrapping up provided the perfect opportunity to showcase these preliminary results to a broad audience. The ILSG meeting has never been hosted in Terrace Bay, however the meeting was hosted in the nearby towns of Nipigon in 2005 and Marathon in 1995. A field guide, “Geology of the Schreiber Greenstone Assemblage and its Gold and Base Metal Mineralization” was prepared for the 1995 meeting in Marathon (Smyk and Schnieders 1995); two of the five stops from that field guide are revisited in this field guide, but with updated information. The tectonically diverse geological history of the Lake Superior region has made it a playground for geologists of every discipline. The north shore of Lake Superior has over a century of mining and exploration history, including precious metals, base metals rare earth metals, and unique industrial minerals such as the colourful marbles at Ruby Lake. The location of Paleoindian sites along the north shore of Lake Superior, which was likely settled during glacial retreat at about 10,000 years before present, tend to be located in close proximity to the cherty rocks of the Gunflint formation, a source for tooling material (Norris, 2012). In fact there is evidence to show that ancient peoples south of the lake were mining, using and trading native copper as early as 4,000 years before present (Pleger, 2000). Furthermore, the Ojibway story of Nanabush and Waub-Ameek (the Giant Beaver) describes the glacial history of the Great Lakes (Snake et al., 1991), albeit it in a mythological way. Indeed, the geology of the Lake Superior area has been of interest to humans for a long time, and the author is thankful for the opportunity to learn a little more about the geological history of the area, and even more thankful to be able to share this knowledge. Safety Some of the field trip stops are located on the Trans Canada Highway 17 which is busy year-round and especially during the summer months. This highway is the major transportation route between western and eastern Canada, and as such much of the traffic along this highway includes transport trucks and logging trucks which have great momentum, especially when fully loaded. The terrane along the north shore of Lake Superior is rugged, thus the highway in this area has many hills and blind curves and the road is mostly restricted to two lanes with narrow shoulders. To maximize the safety of the field trip participants and that of the drivers on the highway, and to minimize the effect that our presence has on the flow of traffic, the author has selected field trip stops that provide ample parking space away from the shoulders of the highway and have suitable sight-lines with the traffic. The area along the north shore of Lake Superior is prone to inclement weather conditions, with dense fog possible at any time of year, causing additional risk for drivers and pedestrians; please use extreme caution during foggy periods. Care should always be exercised when parking, exiting vehicles and crossing the roads. Use of safety vests and/or bright clothing is recommended to improve your visibility to motorists. Stop 3 involves some driving along a railway This field trip guide is also available as Ontario Geological Survey, Open File Report 6357, and can be downloaded from: http://www.geologyontario.mndm.gov.on.ca/mndmaccess/mndm_dir.asp?type=pub&id=OFR6357 - 14 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 maintenance path underlain by sand, gravel and rough ground. Two-wheel drive vehicles are capable of driving on this path, but vehicles with higher clearance and all-wheel or four-wheel-drive are preferred. This stop also involves a short hike away from the parking spot, across a railroad track and up and down a steeplygraded slope. Participants should be aware of the potential for railroad traffic and “slips, trips and falls” hazards. It is recommended that anyone following this field guide individually should bring first aid supplies, food and water. Cell phone service coverage at Stop 3 is not 100% for all providers, especially in areas with more rugged terrain; participants should ensure that their cell phones have adequate connection to their networks before driving down the railroad access path, and again before hiking to the outcrop. Most of the trip routes and sites are on Crown land or public roadways, but access is on or near private property for some routes. As in all such situations, please respect the property rights of others to maintain good relationships with the landowners so that future access for geologists is not adversely affected. Terminology A number of terms used in this report are outlined below. named according to Jensen (1976). Regional Geology The bedrock along the north shore of Lake Superior hosts rocks spanning roughly 1.9 billion years of Earth’s history, from the beginning of the Mesoarchean era to the end of the Mesoproterozoic era, and include a diverse range of rocks formed in a variety of tectonic settings. Neoarchean Geological Setting The Superior Province is an Archean Craton that forms part of the North American continental shield. Rocks of the Superior Province, which range in age from circa 3.4 Ga to 2.6 Ga, are arranged in greenstone belts and plutonic domains. The Superior Province has been subdivided into terranes in which the rocks share similar lithological, geochemical, age and isotopic characteristics and structural and metamorphic histories (Stott et al., 2010). The relationship between these terranes during the early stages of their formation is unclear, however the histories of their evolution converge at circa 2700 Ma, when the terranes were amalgamated to form the Superior Craton (Stott et al., 2010). All whole rock chemical analyses that appear in this report were done at the Geoscience Laboratories, Ontario Geological Survey, Ministry of Northern Development and Mines, Sudbury. All chondrite- and primitive mantle-normalized data or diagrams referred to or shown in this report use the normalizing values of Sun and McDonough (1989). Three major terranes are present near the north shore of Lake Superior; the Wawa–Abitibi Terrane to the south, the Wabigoon Terrane to the north, and the Quetico Terrane between them (Fig. 1). The Wawa– Abitibi granite-greenstone terrane contains Neoarchean volcanic rocks erupted through juvenile oceanic crust and is interpreted to represent an oceanic arc depositional environment (Williams, 1989). The Wabigoon granitegreenstone terrane contains Neoarchean volcanic rocks erupted through and deposited upon Mesoarchean crust, is interpreted to represent a continental arc depositional environment, and is considered to have been a “protocontinent” (Williams 1989). The Quetico terrane is composed mainly of turbiditic siliciclastic rocks with sparse slivers of oceanic crust and is interpreted to represent an accretionary wedge deposited offshore of the Wabigoon “proto-continent” (Williams, 1989; Fralick et al., 2006). A preliminary compilation of geochronological data for these terranes (Fig. 2) helps to visualize the timing of events. Rock type names based on major element analyses are based on the Total Alkalis versus Silica diagram (TAS; LeMaitre, 1989), except for more ultramafic rocks such as basaltic komatiites, which have been Sedimentary rocks in the Manitouwadge greenstone belt and in the western Schreiber–Hemlo greenstone belt, both along the northern margin of the WawaAbitibi Terrane (see Fig. 1), contain detrital zircon For the sake of simplicity, the name “Wabigoon Terrane” is used in figures and in the text to refer to the collective granite and greenstone domains between the Quetico and English River metasedimentary terranes. As used in this report, “Wabigoon Terrane” includes several subdivisions included in Stott et al. (2010). Terminology for clastic sedimentary rocks, such as wacke and mudstone, follows Pettijohn (1975). Terminology for volcaniclastic rocks, such as tuffaceous conglomerates, follows Schmid (1981). - 15 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 1. Regional map of the north shore of Lake Superior, displaying Archean and Proterozoic geology. White stars indicate local past-producing and currently producing mines. Abbreviations: O-TGB = Onaman–Tashota Greenstone Belt, MGB = Manitouwadge Greenstone Belt, ESHGB = Eastern Schreiber–Hemlo Greenstone Belt, WSHGB = Western Schreiber–Hemlo Greenstone Belt, HWY 17 = Trans-Canada Highway 17. Geology from Ontario Geological Survey 2011; Terrane and domain boundaries from Stott et al. (2010). populations that are correlative with those in the Quetico Terrane and in the Beardmore–Geraldton greenstone belt (Zaleski et al., 1999; Fralick et al., 2006; Tóth, 2018; Tóth et al., 2015; Fig. 2). This suggests that during deposition of the sedimentary sequences, the Wabigoon, Quetico and Wawa–Abitibi terranes were a contiguous depositional environment. In this interpreted environment, detrital material from both the ongoing Wabigoon continental arc volcanism and from erosion of Mesoarchean crust of the Wabigoon “proto-continent” was deposited into a fore-arc accretionary wedge. As the Wawa–Abitibi oceanic arc approached the proto-continent, sediments from the continent began to fill the basin between them, eventually spilling over onto the still-active WawaAbitibi volcanic arc (Fralick et al., 2006). The end of supracrustal rock formation in the northern Lake Superior region is marked at circa 2690 Ma by crosscutting felsic plutons (Figs. 1, 2); plutonism in the region was accompanied by regional deformation and metamorphism from circa 2690 to circa 2670 Ma (Fig. 2). The three terranes were deformed synchronously during three main events; 1) early thrusting during collision of the terranes (D1), 2) upright folding during continued compression (D2), and 3) late transpressional shearing (D3; Williams, 1989). These deformational events likely represent a succession of different styles of deformation during a single protracted event; not three distinct events (Williams, 1989). The structural histories for the Shebandowan (Corfu and Stott, 1998) and Manitouwadge greenstone belts (Zaleski et al., 1999) and the eastern part of the Schreiber–Hemlo greenstone belt (Muir, 2003) are similar to Williams’ (1989) broad interpretation for the region, however the timing and development of deformation is slightly different for each greenstone belt and within each terrane. These differences are likely caused by uncertainties in the geochronological data, inconsistencies in interpretations of all of the geological and related data, and the diachronous nature of regional deformation itself (Corfu and Stott, 1998). - 16 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 2. A preliminary compilation of geochronological data for the Wabigoon, Quetico and Wawa–Abitibi Terranes. Error bars have not been illustrated, for the sake of visual simplicity. Abbreviations: Beard.–Gerald. = Beardmore–Geraldton, Win. L. = Winston Lake, Sch.–Hem. = Schreiber–Hemlo. Numbers correspond to sources for geochronology data: 1 = Anglin et al. (1988), 2 = Blackburn et al. (1985), 3 = Corfu (2000), 4 = Corfu and Muir (1989), 5 = Corfu and Stott (1986), 6 = Corfu and Stott (1998), 7 = Davis (1996), 8 = Davis and Sutcliffe (2017), 9 = Davis, Beakhouse and Jackson (1998), 10 = Davis et al. 1985, 11 = Davis, Pezzutto and Ojakangas (1990), 12 = Davis et al. (1994), 13 = Fage (2011), 14 = Fralick and Davis (1999), 15 = Fralick et al. (2006), 16 = Hart et al. (2002), 17 = Kamo (2015), 18 = Kamo (2016), 19 = Kamo and Hamilton (2017), 20 = Tóth and Lafrance (2018) and references therein, and 21 = Zaleski et al. (1999). - 17 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Paleoproterozoic Geological Setting Several Paleoproterozoic mafic dike swarms are present in the area north of Lake Superior, including dikes from the Matachewan (2480-2445 Ma; Heaman, 1997; Bleeker et al., 2012), Biscotasing (2175-2166 Ma; Buchan et al., 1993; Davis and Stott, 2003; Halls and Davis, 2004; Hamilton and Stott, 2008) and Marathon (2122-2100 Ma; Halls et al., 2008) dike swarms (Fig. 3). Sedimentary rocks of the Animikie Group unconformably overly the Superior Craton and the Paleoproterozoic dike swarms (Fig. 1). The base of the Animikie Group is defined by a thin, locally developed Kakabeka Conglomerate, which hosts carbonaceous microfossils (Gunflintia and Huronosporia) preserved in cherty stromatolites, interpreted to have formed in near-shore and shallow water environments (e.g., Wacey et al., 2013). These rocks are overlain by iron formation, carbonate rocks and siliciclastic rocks of the Gunflint Formation, which is interpreted to have been deposited during multiple marine transgressions in an extensional basin between the Penokean volcanic arc and the Superior Craton prior to their collision at circa 1.86 Ga (Fralick et al., 2002) or, alternatively, in a foreland basin north of the Penokean fold-thrust belt (Ojakangas et al., 2001). The Gunflint Formation is overlain by fine-grained argillites and slates of the Rove Formation, which contain a mixture of Archean zircons and circa 1.83-1.77 Ga zircons (Heaman and Easton, 2006) and are interpreted to have been deposited in a deep marine setting between the assembled Laurentian Craton and the circa 1.8-1.7 Ga Yavapai volcanic arc (Whitmeyer and Karlstrom, 2007). The boundary between the Gunflint Formation and the Rove Formation, and their lithostratigraphic equivalents in the USA, is marked by an unusual rock unit thought to represent distal ejecta from the circa 1.85 Ga Sudbury impact (Addison et al., 2005; Cannon et al., 2010). Mesoproterozoic Geological Setting The circa 1.4 Ga Sibley Group, a sequence of sediments deposited in alluvial-fluvial, lacustrine and eolian settings unconformably overlies the Paleoproterozoic Animikie Group (Rogala et al., 2005). The circa 1.1 Ga Keweenawan Midcontinent Rift event caused widespread magmatic activity in the Lake Figure 3. Simplified geological map of the Schreiber–Marathon area highlighting the Proterozoic formations in the area. All UTM coordinates provided using NAD83 in Zone 16. - 18 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Superior area. Pre-rift intrusive rocks are preserved north of Lake Superior, including the circa 1157-1164 Ma Prairie Lake carbonatite-ijolite complex (Rukhlov and Bell, 2010; Wu et al., 2017). Early-rift rocks north of Lake Superior include 1120-1110 Ma mafic to ultramafic intrusions such as the Thunder Bay North intrusive complex, Kitto and Seagull intrusions and the Logan diabase sills (Bleeker et al., 2018 and references therein). Most of the preserved rift-related rocks were emplaced between 1109 and 1093 Ma and include both intrusive and supracrustal rocks, including the Osler Volcanic Group, the Nipigon and Inspiration diabase sills and the circa 1108 Ma Coldwell alkalic intrusive complex (Bleeker et al., 2018 and Liikane et al., 2018, and references therein). Younger dike rocks include the circa 1099-1095 Ma Pigeon River and Cloud river dike swarms (Liikane et al., 2018). The supracrustal rocks include several packages of mafic and felsic volcanic rocks and sedimentary rocks, which are overlain by late-rift volcanic and sedimentary rocks as young as circa 1083 Ma (Miller and Nicholson, 2013 and references therein). These supracrustal rocks crop out primarily south of Lake Superior in Minnesota, Wisconsin and Michigan, and occur sporadically along the northern and eastern shores of Lake Superior. In the Terrace Bay area, mafic volcanic rocks of the <1108 Ma Osler Group unconformably overlie the circa 1400 Ma Sibley Group (Davis and Sutcliffe, 1985; Heaman and Easton, 2006), and two groups of volcanic rocks of unknown age unconformably overlie the circa 1108 Ma Coldwell Alkalic Intrusive Complex (Heaman and Machado, 1987, 1992) the Coubran Lake and Wolfcamp Lake volcanic rocks (Fig. 3; Cundari, 2012; Davis, 2016; Davis et al., 2017). Archean Geology of the Western Schreiber-Hemlo Greenstone Belt The western Schreiber–Hemlo greenstone belt is a roughly 50km long belt of supracrustal and intrusive rocks bounded on its north and west sides by Archean granitoid plutonic rocks. It extends southward under Lake Superior and is separated from the eastern Schreiber–Hemlo greenstone belt by the Mesoproterozoic Coldwell Alkalic Intrusive Complex (Fig. 1). The greenstone belt is apparently connected to a greenstone belt in the Winston Lake–Big Duck Lake area to the north by a north-trending sliver of greenstone, but the relationship between the belt and the Archean volcanic rocks on the Slate Islands, 10km to the south, is unknown (Fig. 1). Stratigraphy of the Schreiber–Hemlo Greenstone Belt Based on stratigraphic way-up indicators such as flow contacts, pillow cusps and graded bedding, the supracrustal rocks of the western Schreiber–Hemlo greenstone belt are arranged in upright, generally open folds that are locally intensified to tight and isoclinal folds proximal to pluton boundaries and in shear zones (Figs. 4 and 5). The supracrustal rocks have been subdivided into several stratigraphic packages based on common volcanic and sedimentary facies as well as geochemical characteristics and geochronological constraints (Fig. 4, inset). At the time of the ILSG field trip, geochemical and geochronological data for the rocks north and west of the Terrace Bay pluton is pending, thus the stratigraphic arrangement presented herein is tentative. East of the Terrace Bay pluton, three packages of supracrustal rocks are present: Package A, dominated by felsic metavolcaniclastic rocks; Package B, dominated by mafic metavolcanic rocks and Package D, a sequence of turbiditic wackes equivalent to the McKellar Harbour formation of Fralick et al. (2006); Package C is not present (Magnus and Walker, 2015; Magnus and Arnold, 2016; Magnus, 2017a,b). North and west of the Terrace Bay pluton, Packages A and B are present, however Package B is disconformably overlain by Package C, another sequence of distinct mafic metavolcanic rocks, and package D is not present (Magnus and Hastie, 2018). The highly strained and structurally complex area between the Terrace Bay and Santoy Lake plutons appears to mark the boundary between these two stratigraphic sections. Package A Package A is composed mainly of felsic volcaniclastic rocks, including tuffs, crystal tuffs, tuffaceous conglomerates and minor coherent flows. The crystal tuffs contain plagioclase phenocrysts and, in many cases, contain blue quartz phenocrysts. The tuffaceous conglomerates are generally clastsupported and contain pebble to cobble-sized clasts of coherent felsic rocks with similar plagioclase and blue quartz phenocrysts in a felsic tuffaceous matrix (Magnus, 2017b). This package contains minor mafic to intermediate massive to pillowed flows, including some massive flows with a high concentration of quartz and/or calcite-filled amygdules. Chert and sulphide- - 19 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 4. Simplified geological map of the western Schreiber–Hemlo greenstone belt, highlighting the major Archean rock types, some of the stratigraphic younging indicators observed during this study, all of the U-Pb zircon geochronological data in the area, and the inferred fold axial traces. An inset figure outlines the inferred depositional packages A, B, C and D. All UTM coordinates provided using NAD83 in Zone 16. Figure 5. Map of the Western Schreiber–Hemlo greenstone belt outlining domains with distinct structural and metamorphic characteristics. Abbreviations: JMMHSZ = Jackfish-Middleton-McKellar Harbour Shear Zone, HWY 17 = Trans-Canada Highway 17. All UTM coordinates provided using NAD83 in Zone 16. - 20 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 bearing chemical metasedimentary rocks are present in this package, most commonly near and along the contact between this package and Package B. An incredibly well-preserved package of felsic, intermediate and mafic metavolcanic rocks southwest of the town of Schreiber is stratigraphically correlative with Package A (i.e. below Package B, from younging directions), however the volcanic facies are different than those present in the majority of Package A as described in the previous paragraph (Magnus and Hastie, 2018). The most notable difference is the absence of the blue quartz phenocrysts present throughout the remainder of Package A. There are also fewer tuffs and crystal tuffs; the felsic rocks present are predominantly massive, plagioclase porphyritic flows and breccias with angular clasts of similar plagioclase porphyritic material. Several intermediate, plagioclase and amphibole porphyritic massive to brecciated flows are also present, as well as massive to pillowed mafic flows up to 200m thick. Interflow formations in this sequence include chert and magnetite-bearing chemical metasedimentary rocks and tuffaceous wackes and conglomerates. Package A contains the oldest known rocks in the western Schreiber–Hemlo greenstone belt; three samples from the top of the package have ages of circa 2720 Ma, which is correlative with similar felsic volcanism in the Winston Lake area, in the Manitouwadge greenstone belt and with the Greenwater assemblage in the Shebandowan greenstone belt (Davis et al., 1994; Davis and Sutcliffe, 2017). Volcanic rocks of this age have not yet been identified in the eastern Schreiber–Hemlo Greenstone Belt, however there are numerous felsic volcanic formations on that portion of the Schreiber–Hemlo belt that do not have any age information. Older phases of both the Pukaskwa and Black Pic batholiths from the eastern Schreiber–Hemlo belt, however, have yielded similar circa 2720 Ma ages (Corfu and Muir, 1989; Beakhouse and Davis, 2005). The mafic to intermediate rocks in this package contain trace element concentrations consistent with both arc volcanic and oceanic plateau volcanic settings. A-B Disconformity In the Schreiber area, a substantial sequence of interbedded graphitic argillite, chert, sulphide facies iron formation and felsic tuffaceous breccias represents a disconformity between packages A and B. The Elwood and Morley base metal sulphide occurrences occur along this horizon. Above this disconformity, there are several massive to pillowed flows which are overlain by a chemical metasedimentary sequence that includes graphitic argillite, sulphide facies iron formation and marble. So far two exposures of this horizon have been observed. In one, the marble is composed mainly of calcite with minor silicate and sulphide components; the other is a breccia, with mafic volcanic, argillite and sulphide clasts supported by a matrix of calcite. The significance of these marbles is unknown and requires further study. Is the carbonate derived from a primary sedimentary source? Could the breccia have been formed in a karstlike environment? Or was the carbonate introduced during later hydrothermal alteration? Elsewhere in the belt, similar chert and sulphide facies iron formation are concentrated along the A-B disconformity, including one occurrence of marble south of the Foxtrap Lake pluton. However, the occurrence of these rocks is more sporadic, which may be a consequence of their location in areas that are more highly strained and metamorphosed than the well-preserved occurrence in the Schreiber area. A unique feature of the A-B disconformity, which crops out along Highway 17 in Schreiber, is a sequence of interbedded turbiditic wacke and siltstone with basaltic andesitic composition that was deposited either synchronously or directly on top of the chemical metasedimentary rocks. The mafic composition of these rocks indicates they were derived primarily from a mafic volcanic source. No zircons have been found in these rocks, which precludes detrital geochronology, but a whole rock Sm-Nd isotopic study may help identify possible sources for this mafic sediment. Package B Package B is dominated by massive to pillowed mafic flows with two distinct geochemical populations based on trace elements that are consistent with both oceanic plateau volcanism and back-arc basin volcanism, respectively. Flows with trace elements consistent with an oceanic plateau volcanic setting are typical green to grey-green massive to pillowed flows, with lath-shaped plagioclase microphenocrysts visible in thin section and small vesicles concentrated around the edges of the pillows. Some massive flows with this chemistry also contain abundant calcite-filled amygdules. Flows with trace elements consistent with a back-arc volcanic setting have distinct variolitic - 21 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 textures as well as irregular-shaped cavities that likely served as conduits for volatile fluids and gases rather than typical vesicles, which are not present in these rocks (Magnus, 2017b). marked by a sequence of chert, graphitic argillite and sulphide facies iron formation that is continuous along the entire contact. One occurrence of a coherent felsic flow, with a perlitic texture and possible flow banding, has been observed near the top of this package near the Steel River (Magnus, 2017b). Package D, also known as the McKellar Harbour Formation (Fralick et al., 2006), represents the youngest known supracrustal rocks in the western Schreiber– Hemlo greenstone belt. This package is composed of a sequence of interbedded turbiditic wacke, sandstone and mudstone with normal graded bedding and sharp bedding contacts. B-C Disconformity North and west of the Terrace Bay pluton, the top of Package B is marked by a horizon of felsic volcaniclastic rocks including tuffs, tuffaceous wackes and tuffaceous conglomerates, locally interbedded with chert and sulphide facies iron formation. This horizon is located between the Terrace Bay pluton and the Lunch Lake pluton (an area known locally as the Empress Structure) and wraps around the southeastern edge of the Terrace Bay pluton (Fig. 4, inset). Package C Package C is dominated by massive to pillowed mafic to intermediate flows with trace element concentrations that are predominantly consistent with oceanic plateau volcanism, including a more thorium-enriched variety of that chemical signature and some calc-alkalic arc volcanism. The rocks in this package commonly contain medium grained equant plagioclase phenocrysts, which are uncommon in the other metavolcanic packages. This package lacks the variolitic back-arc volcanic rocks that are a distinctive feature of Package B. Apart from the felsic volcaniclastic rocks that mark the lower contact between this package and Package B, several other isolated lenses of felsic volcaniclastic material have been observed in this package. A sequence of tuffaceous metasedimentary rocks along the western edge of the Santoy Lake pluton may represent the top of Package C. This sequence includes tuffaceous wackes with abundant plagioclase phenocrysts which locally display graded bedding interbedded with tuffaceous conglomerates. The tuffaceous conglomerates are clast-supported and polymictic, with a variety of felsic and mafic metavolcanic rocks, including clasts of mafic rocks with equant plagioclase phenocrysts like those in the mafic rocks of Package C (Magnus, 2017b). B-D Disconformity East of the Terrace Bay pluton, packages B and D are in disconformable contact. This disconformity is Package D – McKellar Harbour Formation The youngest detrital zircon at the base of the package is 2696±3 Ma, which marks the maximum age of deposition for the package, and the youngest detrital zircon from the top of the package is 2693±4 Ma, which suggests that the basin had a source for young zircons during deposition (Fralick et al., 2006) 2006). The sedimentary rocks are crosscut by the 2689.6±2 Ma Steel River pluton (Kamo and Hamilton, 2017), which places a minimum age of deposition for the package, suggesting deposition occurred between 2690 and 2696 Ma. There is volcanism recorded during the 26962689 Ma interval in both the eastern Schreiber–Hemlo greenstone belt (Corfu and Muir, 1989; Davis and Lin, 2003) and in the nearby Shebandowan greenstone belt which could have provided young detrital material during deposition of the package. Older detrital zircons are present in this package, including several Mesoarchean zircons at circa 2900 Ma (Fralick et al., 2006) and a single concordant Paleoarchean grain at 3423±10 Ma (Davis and Sutcliffe, 2017). This implicates a continental component for the source of the sediments, which is interpreted to have been the Wabigoon proto-continent (Fralick et al., 2006) and/or the Minnesota River Valley terrane. Mafic to Intermediate Intrusive Rocks A series of sill-like gabbroic rocks that intrude Package B are parallel to stratigraphy and have chemistry similar to the nearby variolitic mafic metavolcanic rocks. These have been interpreted as syn-volcanic intrusions and may in some cases represent medium to coarse grained massive flows. In several of these bodies, rocks with basaltic chemistry display spinifex-like textures composed of abundant elongate amphibole crystals; invariably these rocks are associated with massive rocks of basaltic komatiitic chemistry, which may represent cumulate phases within an ultramafic flow or sill. - 22 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Mafic intrusive rocks that occur locally along the contact between packages B and D east of the Terrace Bay pluton, as well as near the contact between packages B and C northwest of the Terrace Bay pluton, typically have elongate, “plumose” amphibole crystals with interstitial plagioclase feldspar. These intrusions have chemistry similar to arc basalts, which helps distinguish them from the rocks with the spinifex-like texture described above. The Lunch Lake and Longworth Lake plutons are both composed mainly of diorite and have generally equigranular hypidiomorphic textures. The Lunch Lake pluton also includes porphyritic diorite with distinct blue quartz phenocrysts, which have not been observed in any other mafic or intermediate intrusive rock in the belt, but which are common in the felsic volcanic and volcaniclastic rocks. Felsic Intrusive Rocks Felsic intrusive rocks that surround and crosscut the greenstone belt were emplaced over a period of at least 20 million years (Figs. 1, 2). The Terrace Bay, Steel River and Syenite Lake plutons have ages between 2690-2680 Ma (Kamo, 2016; Kamo and Hamilton, 2017, 2018). These plutons are generally oblate and irregular in shape and have well-developed foliations along their contacts that penetrate up to 500m into the surrounding rocks. The Terrace Bay and Steel River plutons are both composed mainly of grey, equigranular to quartz and/or alkali feldspar porphyritic granodiorite with minor dioritic components. The Syenite Lake pluton is composed mainly of pink alkali feldspar porphyritic quartz monzonite and quartz monzodiorite. The Foxtrap Lake and Little Pic River plutons, as well as the small pluton between them, have ages between 2680-2670 Ma (Kamo and Hamilton, 2017, 2018). These plutons are round in plan view, have well-developed foliations along their margins, which continue up to 1 km into the surrounding rocks. These plutons are composed mainly of grey, equigranular to alkali-feldspar porphyritic granodiorite. The Santoy Lake pluton, with a zircon age of 2667±4 Ma (Kamo, 2016), is round to irregular in shape and has a weakly developed foliation along its contacts. This pluton is composed of pink quartz monzonite to monzonite, with local alkali feldspar porphyritic varieties, and contains a distinctly low abundance of mafic minerals. The Crossman Lake batholith (age unknown) is elongate and has a very well-developed foliation along its southern contact that penetrates up to 3km into the supracrustal rocks to the south. The intrusion is composed of equigranular white to grey trondhjemite, tonalite and granodiorite. A small pluton south of the town of Schreiber (age unknown) is irregular in shape and does not have foliations developed along its contacts. This intrusion is mostly composed of grey to pink quartz porphyritic granite with more intermediate varieties towards its southern contact. Quartz and/or feldspar porphyritic felsic dikes (age unknown) are abundant around Schreiber and northwest of the Terrace Bay pluton but are uncommon elsewhere in the greenstone belt. The “Whitesand Lake Batholith” (age unknown) is heterogeneous and appears to include more than one distinct intrusion; further mapping is required to better delineate the granitoid rocks of this batholith. Archean Structural Geology The degree of metamorphism and the nature of structural fabrics varies throughout the western Schreiber–Hemlo greenstone belt. Distinct domains containing common metamorphic mineral assemblages and structural fabrics are illustrated in Figure 5. Few crosscutting relationships have been observed between these different fabrics, thus, the structural features displayed on the map have been separated into two structural events; the early penetrative ductile fabrics, and the later more discrete brittle-ductile fabrics (Fig. 5). Observations of stratigraphic younging indicators, bedding-cleavage relationships and fold closures as well as stratigraphic correlation using geochronology have provided evidence for the upright folded Archean stratigraphy in the Schreiber–Hemlo greenstone belt (Fig. 4). The intensity of deformation and metamorphism tends to increase towards the granitoid plutons (Fig. 5). The supracrustal rocks in a 1 to 1.5 km wide zone along the northern margin of the greenstone belt have amphibolite facies mineral assemblages and display strong penetrative foliations which define tight to isoclinal fold axial planes that are parallel to the margin. South of this zone, the rocks are generally less deformed, have greenschist to amphibolite facies mineral assemblages and have east to northeast trending foliations and fold axial traces. These fold - 23 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 axial traces indicate that the rocks were deformed under northwest-directed compression during regional ductile deformation (Fig. 4). Discrete, outcrop-scale shear zones within the “Open Folding” domains (Fig. 5) are also east to northeast striking and have kinematic indicators that indicate reverse, south-side up vertical displacement with a dextral horizontal component, which is consistent with northwestdirected compression. are generally parallel with the regional ductile fabric. A thin zone of high strain up to 200 m wide is present along the margins of the circa 2690 Ma Terrace Bay pluton (Kamo and Hamilton, 2018). Kinematic indicators along the strained margins of the pluton indicate reverse, south-side up vertical displacement with a dextral horizontal component. In plan view, this pluton and the 2682.3 ± 1.1 Ma Syenite Lake pluton (Kamo and Hamilton, 2018), which are both hosted within the greenstone belt, strike northeast and resemble large-scale dextral sigma clasts. This suggests that the bulk of horizontal displacement in the greenstone belt during regional ductile deformation was dextral. The Terrace Bay pluton is the oldest known pluton in this part of the Schreiber–Hemlo greenstone belt; there has been no evidence to determine whether regional ductile deformation had commenced prior to emplacement of this pluton. Several occurrences of polymictic, clast-supported conglomerate with pebble to cobble-sized clasts and interspersed lenses of stromatolitic chert unconformably overly the Archean basement along the shore of Lake Superior southwest of Schreiber. This conglomerate represents the base of the Gunflint Formation, and hosts the famous microfossils Gunflintia and Huronosporia (e.g., Wacey et al., 2013). The 2674.1 ± 1.3 Ma Foxtrap Lake pluton (Kamo and Hamilton, 2018) truncates several east-trending fold axial traces, which are overprinted by fold axial traces that are parallel to the pluton margins. This indicates that regional ductile deformation had commenced prior to 2674 and continued after emplacement of the pluton. The 2667 ± 4 Ma Santoy Lake pluton (Kamo, 2016) truncates several east-trending fold axial traces and has very thin zones of strain along its margins. This suggests that the pluton was emplaced during the later stages of regional ductile deformation. There are two highly strained zones to note in the greenstone belt (Fig. 5). 1) The Jackfish–Middleton– McKellar Harbour shear zone is a 1-2 km wide zone of greenschist-facies rocks which are isoclinally folded and heavily sheared along lithological contacts. 2) The Empress Structure, located in a narrow band between the Terrace Bay and Lunch Lake plutons, hosts amphibolite facies rocks which are isoclinally folded and sheared with a moderate penetrative foliation throughout the zone. Throughout the greenstone belt, other thinner, unnamed ductile shear zones occur that Conjugate northwest-striking and north to northeast striking brittle-ductile shear zones and faults crosscut the greenstone belt and offset lithological contacts and the ductile fabrics. Proterozoic Geology Gunflint Formation Diabase Dikes A multitude of diabase dikes are present throughout the Schreiber-Terrace Bay area. Where possible, these dikes have been assigned to dike swarms that have been previously recognized in the area based on their orientation and chemistry. Dikes from three Paleoproterozoic dikes swarms have been recognized, including the circa 2460 Ma Matachewan Swarm, the circa 2170 Ma Biscotasing Swarm and the circa 2120 Ma Marathon Swarm (Bleeker et al., 2012; Halls and Davis, 2004; Halls et al., 2008, respectively). Three northeast-striking dikes with chemistry similar to dikes of the circa 1096 Ma Pigeon River Swarm (Liikane et al., 2018) have been observed. Previously the Pigeon River Swarm has only been recognized in the Thunder Bay area, however the dikes in the Schreiber–Hemlo greenstone belt are approximately along strike from those dikes (180 km). These dikes do not present a distinctive geophysical signature in the aeromagnetic dataset, thus tracing their extent is difficult. A series of east to northeast striking subalkalic dikes is present east of Terrace Bay which have not yet been correlated with other regional events. The author believes these to be related to the circa 1.1 Ga Keweenawan Midcontinent Rift. The most abundant dikes in the area are a series of west to northwest striking alkalic, olivine tholeiitic diabase dikes that are similar in chemistry to, and - 24 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 appear to point directly to, the Wolfcamp Lake volcanic rocks that unconformably overly the Coldwell Alkalic Intrusive Complex. If this interpretation is correct, then it is likely that the alkalic dikes are related to an episode of Keweenawan volcanism younger than 1108 Ma (the age of the Coldwell Complex which they crosscut). as massive sulphide in irregularly-shaped veins, and in sulphide-bearing quartz veins. Where these structures crosscut felsic metavolcanic rocks, the rocks are typically altered from grey to beige, and feldspar phenocrysts are altered to a distinct bright green colour (Magnus and Hastie, 2018). Mineral Potential Proterozoic rocks near the north shore of Lake Superior host a variety of commodities, including nickel, copper and platinum group elements associated with mafic to ultramafic intrusions; other transitional metals, such as niobium, tantalum and titanium and rare earth elements associated with carbonatitic (e.g., the Prairie Lake carbonatite-ijolite complex) and alkalic rocks (e.g., the Coldwell Alkalic Intrusive Complex); and diamond, associated with lamprophyric and kimberlitic dikes. The Western Schreiber–Hemlo greenstone belt has a long history of mineral exploration and has potential for a variety of styles of base and precious metal mineralization. Base metal sulphide and precious metal occurrences are associated with supracrustal rocks throughout the greenstone belt. Sulphide mineralized rocks occur near and along the upper contact of the circa 2720 Ma metavolcanic rocks of Package A. The host rocks are sulphide facies iron formation and chert, interbedded with felsic volcaniclastic rocks and garnetiferous mafic metavolcanic rocks. These rocks are lithologically similar and chronologically correlative with circa 2720 Ma metavolcanic rocks in the Winston Lake and Manitouwadge areas, which both host past-producing Zn-Cu-Ag base metal mines (see Fig. 1; Davis et al., 1994; Zaleski et al., 1999). Gold and base metal sulphide occurrences are associated with highly strained supracrustal rocks, including the JMMHSZ, the Empress Structure, strained rocks surrounding plutons and other discrete shear zones throughout the greenstone belt. Mineralization in these shear zones typically occurs in quartz and carbonate veins in the shear zones and in silicate and carbonate altered haloes adjacent to the veins. The Hemlo gold deposit, located in the eastern part of the Schreiber–Hemlo greenstone belt, is hosted in highly strained and altered supracrustal rocks (Fig. 1; e.g. Muir, 2003). Gold occurrences, with minor silver, molybdenum and copper mineralization, are associated with the Terrace Bay pluton and other granitoid rocks in the map area. Mineralization occurs in sulphide-bearing quartz veins and in altered granitoid rock adjacent to the veins. The veins are typically straight, with sharp contacts, and occur in parallel sets and in “stockwork” arrangements (Arnold et al., 2017; Marmont, 1984). Zinc, silver and lead mineralization (with minor copper and gold) is associated with north to northeast striking faults near Schreiber. Mineralization occurs Road Log Note: Caution should be taken when parking vehicles on the shoulder of the highway and when examining outcrops located along Highway 17. All UTM coordinates are provided in NAD83, Zone 16. Figures 3 and 4 show the location of the field trip stops. The primary focus of this trip is on the Archean rocks of the Schreiber–Hemlo Greenstone Belt, however because of the abundance of Proterozoic diabase dikes in the area, some of the stops will also feature Proterozoic rocks. Note the mileages in the road log are not cumulative, rather each number is the distance from one stop to another. 41.7 km - Starting in Terrace Bay, drive east along Highway 17 for roughly 42km (25 minutes). About 1.5km (1 minute) before the parking spot, you will pass under two power transmission lines and Ripple Lake on the southeast side of the road; begin to slow down at this point, to make sure vehicles behind you have time to react, as you will be pulling off the road shortly. As you approach the parking spot, you will drive downhill across the McKellar Creek bridge. The parking spot is on the right (south) side of the road at the east end of this bridge just past the guardrail. Reset odometer to zero as subsequent mileages to stops are based on starting here. If you pass the stop, turn at the junction between Highway 17 and Dead Horse Road and find a suitable location to turn around, and retrace your route back to the stop. - 25 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 40.3 km - If starting in Marathon, drive west along Highway 17 for 40.3 km (29 minutes). After crossing the Little Pic River Bridge, then passing Dead Horse Road, Stop 1 will be on the south (left) side of the highway. Beware of oncoming traffic prior to turning into the parking area; east-bound vehicles will be driving speedily downhill at this location. Stop 1. Sheared rocks, diabase and lamprophyre UTM coordinates 521355E 5407099N There are outcrops on the north and south sides of the highway at this location. The highway marks a contact between metasedimentary rocks to the north and mafic metavolcanic and intrusive rocks to the south. The metasedimentary rocks north of the highway are wacke and siltstone with local sulphide mineralization. Box folds are observed throughout these rocks, which indicate that the rocks have been subjected to beddingparallel shearing. Several biotite porphyritic ultramafic lamprophyre dikes, up to 8cm wide, crosscut the metasedimentary rocks, which exhibit iron carbonate alteration haloes adjacent to the dikes. A single 240m long, near vertically-faced outcrop is present south of the highway. In this large outcrop, the mafic rocks display varied degrees of strain. In more highly strained areas, fractures and quartz veins tend to be parallel to the strong foliation, and where the outcrop surface is parallel to the foliation plane, geometric shapes appear in the outcrop. In less strained areas, fractures are more irregular, and parallel sets of quartz veins dip shallowly to the south, nearly perpendicular to the foliation in the outcrop. These areas are most easily observed by viewing the entire outcrop from the north side of the highway. All primary textures have been obliterated in the highly strained areas, where a strong steeply dipping foliation hosts folded and boudinaged quartz veins and box folds and local sulphide mineralization. Quartz vein boudins with both sigma and delta asymmetries indicate the latest displacement was vertical; reverse motion with the northern side moving up towards the south. Two lineations are present in the outcrop; one shallowly plunging lineation which is interpreted to be a crenulation lineation, and a steeply-plunging lineation which is interpreted to be a stretching lineation, both formed during the reverse shearing event. The trend and plunge of these lineations vary throughout the outcrop as the foliations that host them waver. The two areas of more competent rock are composed of massive medium-grained equigranular aphyric mafic rock, which may represent either massive metavolcanic flows or intrusive rocks. In the western zone, a dike of granodioritic rock crosscuts the mafic rock. Near the west end of the outcrop, two adjacent alkalic diabase dikes crosscut the Archean rocks, striking west and dipping to the north. Together, these dikes are about 10m wide. These dikes are generally equigranular with fine-grained chilled margins and fractures orthogonal to the contacts. A 5cm wide feldspar porphyritic diabase dike is present at the very west end of the outcrop. Several north-striking biotite porphyritic ultramafic lamprophyre dikes are present near the middle and at the east end of the outcrop, similar to those on the north side of the highway. The mineralogy and texture of these dikes, including their associated iron carbonate alteration, are typical of lamprophyre dikes in this area. Return to vehicles and turn left to drive west on Highway 17. 12.9 km - Drive west on Highway 17 for 12.9 km (8 minutes). Two minutes before the stop, you will pass Black Fox Lake on the right (north) side. The parking location for Stop 2 will be on the left (south) side of the highway in a turn-around location, on the east side of the Steel River Bridge (Fig. 6). If you miss the stop, there is a suitable turn-around location on the west side of the bridge. Stop 2. Turbiditic Wacke UTM coordinates 508700E 5402577N There are outcrops on both the north and south sides of the highway at this location. The outcrop on the north side of the highway is shorter in length, and has cleaner outcrop surfaces, so it will be the focus of this stop. Stop 2 is the same locality as Stop 1 in Smyk and Schnieders (1995). This outcrop is composed of southward-younging normally graded wacke interbedded with mudstone (Photo 1). On the south side of the road, and in outcrops along Santoy Bay and on Lawson Island, the younging direction of graded beds switches from south to north repetitively within tens of metres. These tight younging reversals are the main evidence for isoclinal folding in the area (Fig. 6). - 26 - A pervasive schistosity throughout the outcrop, �Proceedings of the 65th ILSG Annual Meeting - Part 2 sheared during folding. Return to the vehicles and turn left to drive west on Highway 17. Photo 1. This photograph displays normally graded wacke interbedded with mudstone. Arrows point to enigmatic features which may be related to primary loading structures, folding or both. The outcrop surface is horizontal and a scale card is pointing north. axial planar to local isoclinal folding, is at a low angle to bedding. Along this schistosity, the bedding planes display a strange pattern in which the more fissile mudstone layers form tabular projections into the sandstone, whereas the sandstone layers form more lobate projections into the mudstone. This may be interpreted as a primary loading structure that has been 4.6 km - Drive west on Highway 17 for 4.6 km (3 minutes). You will be driving uphill on a moderate to steep grade, and any westbound transport trucks on this stretch of road will be driving below the speed limit with their 4-way flashers on. Do not pass these trucks; stay well behind them to ensure they do not decide to pull over and stop. After cresting the hill, the next turn will be about 1 minute away. Look for a “road intersection” sign and prepare to turn left (south). 2.2 km - Drive down this gravel road south towards Lake Superior, always keeping to the left at any junction. Eventually you will turn east and pass a gravel pit and a railway crossing. Slow down considerably and continue onward. ~2.0 km - The road turns from gravelly to sandy and narrows to a single lane. Continue driving eastward with caution. There are several branches of this trail that quickly lead to dead ends; if you reach a dead end, turn around and try another path. If at any time you feel unsafe or are unsure whether your vehicle is Figure 6. Simplified geological map of the Santoy Bay area, which includes stops 2 and 3. Several rocks of interest near Stop 3 are also indicated but will not be visited during this field trip. All UTM coordinates provided using NAD83 in Zone 16. - 27 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 capable of driving in this terrain, turn back and skip ahead to Stop 4. The parking spot for Stop 3 is located at UTM 505400E, 5402555N, in a location where the path widens and there is a clearing through the trees on the south side of the path. 250-300 m - Unpack your hiking gear, including food, water and first aid equipment. Walk south along a footpath towards the railway; there is good line-ofsight along the rails at this location to see any oncoming trains. For the 2019 ILSG trip, there will be a flagged trail through the bush towards the lakeshore outcrop (Stop 3). For future users of this guidebook, you will have to navigate directly southward to the lake. Safety Note: the hill between the railway and Lake Superior is steeply sloping and built out of cobbles. This is a health and safety risk to those with reduced mobility. If at any time you feel unsafe to continue, turn back and skip ahead to Stop 4. Stop 3. Variolitic Mafic Flows UTM coordinates 505366E 5402266N This lakeshore outcrop, which has been kept lichen free by winter ice in Lake Superior is the best exposure of variolitic mafic flows in the western Schreiber– Hemlo Greenstone Belt (Fig. 6). Exposure here is good enough to trace flow contacts and to observe various macrotextures present in at least four consecutive flows (Fig. 7). Pillows are generally 3m across, with single rinds up to 3 cm thick, and selvages filled with glassy material and hydrothermal minerals like quartz, calcite and epidote. These pillows lack vesicles or amygdules, but some pillows situated at or near the top of flow sequences contain elongate, discontinuous, quartz- and carbonate-filled cavities, which the author interprets to represent large, formerly gas-filled cavities. The most conspicuous feature of these pillows is the variolitic texture (Fig. 7). Inward from the chilled margins, the pillows are dark green and very finegrained, with only a few small varioles (up to 2 mm). Varioles become larger (up to 8 mm) and more abundant toward the core of the pillow, where they appear to have amalgamated to produce a more massive, leucocratic pillow core (Fig. 7). The interiors of the varioles are concentrically zoned with bands of calc-silicate minerals. The distribution of varioles in the pillows is not always perfectly concentric; their distribution seems to be more erratic at the tops of the pillows. Very few pillows display multiple concentric variolitic and non-variolitic bands. In pillows that contain both the gas cavities and varioles, the gas cavities are always located in the dark green, non-variolitic upper portions of the pillows. The massive flows, which may be traced in this Figure 7. A) Simplified bedrock geology map of the shoreline outcrop at Stop 3, and B) illustration of the macrotextures observed in outcrop along A–A’. All UTM coordinates provided using NAD83 in Zone 16. - 28 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 outcrop for up to 50 m in length, have widths that vary in proportion to their lengths (i.e., thinner flows are less laterally extensive). The internal structure of the flows is similar to that observed in the pillows, with non-variolitic, dark green material inside the rinds that grades into massive variolitic cores. The transition from rind to variolitic core at the base of the flows occurs over approximately 5 cm, whereas, at the top of the flows, the varioles have coalesced into lobes and pods that appear pillow-like, without the necessary rinds (i.e., pseudopillows). Randomly oriented, elongate crystals of amphibole are present in the massive, medium-grained core of the thickest flow in this outcrop. These flows and all other variolitic mafic flows in the greenstone belt have trace element contents consistent with mafic rocks erupted in “back-arc basin” volcanic environments. This is the defining characteristic of Depositional Package B, which is dominated by rocks of this chemistry intercalated with mafic rocks of “oceanic plateau” volcanic affinity. Nearby outcrops of a perlitic felsic flow (accessible along a footpath) and outcrops of spinifex-textured mafic to ultramafic rocks (accessible along Highway 17 or along the railway) are indicated in Figure 6. These rocks will not be visited during this field trip. Please use caution, especially along the highway and the railway, if you decide to visit these rocks. Return to vehicles by the same path you took to the outcrop. Turn the vehicles around and drive west along the path. ~2.0 km - Drive back along the sandy path the way you entered, until you reach the gravel pit. 2.2 km - Keep to the right and drive north until the gravel road intersects with Highway 17. 7.3 km - Turn left and drive west on Highway 17 for 7.3km (4 minutes). Along the way, you will get an excellent view over Jackfish Lake and the terraced hills to the north. Those hills are the location of the historic Empress gold mine, which produced 112 oz Au at 0.10 oz/ton from 1896-1897. After passing the lake and beginning to drive uphill, prepare to slow down for Stop 4. The parking spot for Spot 4 is a turn-around spot on the left (south side) of the highway, on the west end of a large granite and diabase outcrop. Stop 4. Granite, sheared mafic rocks and diabase dikes UTM coodinates 503233E 5411222N There are outcrops on the north and south sides of the highway at this location. On the south side of the highway, grey to pink granodiorite of the Terrace Bay pluton is crosscut by a 50 m wide plagioclase porphyritic diabase dike. This dike trends generally northward and is aligned with a north-northeast trending geophysical anomaly consistent with dikes of the Biscotasing dike swarm. There are smaller plagioclase porphyritic dikes with chill margins that crosscut the large dike. On the north side of the road, there are two outcrops composed of a series of southward dipping panels of granite, mica schist, mafic intrusive rocks and massive felsic rocks. The bottom of the western outcrop is massive grey to pink granodiorite of the Terrace Bay pluton, the top of the outcrop is a dike or sill of weakly foliated massive, fine grained aphyric felsic rock and a panel of mica schist lies between them. The mica schist is composed of biotite and chlorite with abundant quartz and calcite veins. The dominant foliation in the mica schist dips more steeply to the south than the contacts between the schist and the felsic rocks between which it is sandwiched. This looks like a C-S structural fabric; the contacts between units represent the “C” plane, and the strong foliation in the schistose rock represents the “S” plane. Quartz veins in the schist are boudinaged; asymmetric boudins (mostly sigma clasts) and the orientation of the C-S fabric both indicate south-side up reverse displacement northward. Box folds in this schist post-date the sigmoidal quartz boudins. A biotite porphyritic ultramafic lamprophyre dike crosscuts the rocks at the west end of the outcrop. The eastern outcrop is a massive sheared mafic rock composed mainly of amphibole and biotite with minor quartz, feldspar and carbonate minerals, cut by a small granitoid dike at the west end of the outcrop. Southward dipping shear zones crosscut this rock with C-S fabrics similar to those observed in the western outcrop, indicating the same south-side up reverse displacement. The trace element composition of this mafic rock is similar to that of the granitoid rocks in the Terrace Bay pluton, with higher concentrations of transitional elements such as iron, magnesium, chromium, vanadium, nickel and copper. This rock is interpreted to represent mafic country rock that was - 29 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 altered during emplacement of the Terrace Bay pluton. Return to the vehicles and turn left to drive west on Highway 17. 17.4 km - Drive west on Highway 17 for 17.4km (12 minutes). You will pass through the town of Terrace Bay. On the west side of town, just west of the Aguasabon River, turn left at the intersection between the highway and Aguasabon Gorge Road. Note the outcrops of grey granodiorite, locally altered to pink granodiorite, along the highway towards Terrace Bay. Most of the rocks in the Terrace Bay pluton look like this. 750 m - Drive to the end of the road. There will be a large parking area with outhouses and picnic tables. There is a boardwalk with railings at the south end of this parking lot that leads towards a vista with a view of the Aguasabon Falls and Lake Superior. Safety Note: Although there are small foot-paths off of the main boardwalk, do not hop over the railing to walk on these paths. Falling into the gorge would lead to death. Stop 5. Aguasabon Falls Gorge; structures and alteration UTM coordinates 490833E 5403233N This vista overlooks granitoid rocks of the Terrace Bay pluton (Photo 2). Granodiorite in the pluton is normally grey; locally, the granodiorite is altered pink, caused by hydrothermal alteration around regionalscale shear-zones and faults. North, northwest and northeast-striking faults, which correlate with similar shear zones that crosscut the supracrustal rocks of the greenstone belt, control the vertical cliff faces in the Aguasabon River Gorge. Return to vehicles, drive back along Aguasabon Gorge Road and turn left to drive west on Highway 17. 8.1 km - Drive west on Highway 17 for 8.1 km (5 minutes). Along the way you will get an excellent view of Lake Superior (Terrace Bay). One minute before the next turn, you will pass an intersection between the highway and Worthington Bay Road. You will then cross a train bridge; turn right onto Hays Lake Road 200m north of the bridge (Fig. 8). Note the large outcrop of sulphide-bearing chert and graphitic argillite at the beginning of Hays Lake Road. Drive for 800 metres along Hays Lake Road; there Photo 2. View of the Aguasabon Falls Gorge, with Lake Superior and the Slate Islands in the background. Photo taken from a vista at the end of Aguasabon Gorge Road. will be a clearing in the trees on the north side of the road. Pull your vehicle safely to the shoulder of the road and park. Stop 6. Harkness Hays and Gold Range UTM coordinates 483711E 5404933N North of the road, there is a northeast-trending ridge of outcrops that have been the subject of gold exploration for more than a century, with the earliest staking recorded in 1917. Over the following several decades, numerous adits and shafts were used to sample the bedrock in this ridge, which hosts the HarknessHays property to the west and the Gold Range property to the east. The Harkness-Hays property produced 200 oz Au at 2.58 oz/t during intermittent mining activities between 1920 and 1936; the Gold Range property produced 36.35 oz Au at 0.91 oz/t during intermittent mining activities from 1921 to 1941 (Schnieders et al., - 30 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 8. Simplified geological map of the Schreiber area, which includes field trip stops 6, 7 and 8. The mafic metasedimentary rocks and the Morley occurrences, which are correlative with the rocks at Stop 7, are indicated, as well as nearby occurrences of marble associated with other chemical metasedimentary rocks. A box outlines the area covered in Figure 9. Abbreviations: Ag = silver, Au = gold, Cu = copper, Fe = iron, g = graphite, grt = garnet, Mo = molybdenum, Pb = lead, py = pyrite, S = sulphur, Zn = zinc. All UTM coordinates provided using NAD83 in Zone 16. 1996). The focus for this stop will be on the HarknessHays property, which is the most easily accessible (and has the better historic gold grade). These outcrops are composed mainly of massive to pillowed mafic metavolcanic rocks, crosscut by quartz feldspar porphyritic felsic dikes and biotite porphyritic lamprophyre dikes. These rocks are located in a moderately strained zone along the northwest edge of the Terrace Bay pluton and contain amphibolite facies mineral assemblages (Figs. 5 and 8). Quartz veins in the northeast-striking foliation, parallel to the contact with the pluton, host gold-bearing sulphides and occurrences of native gold. Native gold at this site is found most commonly in white, vuggy quartz veins. Much of the bedrock has been blasted, and quartz vein bearing rocks are dispersed throughout the resultant pile of rocks. It is recommended that visitors search through this pile of rock, rather than scale the pile to access the steep, cliffy outcrops. Return to vehicles, turn the vehicles around and drive back towards Highway 17. 5.3 km - Turn right and drive westward on Highway 17 for 5.3 km (4 minutes). Along the way you will pass through the town of Schreiber. After passing the Villa Blanca Inn, on the west side of town you will begin driving uphill. The parking spot for Stop 7 will be on the right (north) side of the road at the top of this hill; prepare to stop. Note as you drive through Schreiber, on the right (northeast) side of the road are several outcrops of turbiditic mafic-derived metasedimentary rocks. - 31 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 7. Elwood Occurrence UTM coordinates 479511E 5407266N There are outcrops on the north and south sides of the highway at this location (Fig. 8). On both sides of the highway, the outcrops are an upright north-dipping sequence of interbedded chert and graphitic argillite with minor beds of felsic tuffaceous conglomerate. Sulphide mineralization is disseminated throughout the outcrop and occurs in calcite-sulphide veins and as conformable lenses of massive sulphide. The sulphide minerals are dominantly pyrrhotite and pyrite with minor chalcopyrite. On the north side of the highway, a massive mafic flow marks the top of this sequence and forms an erosion-resistant cap on the outcrop. At the base of this flow, the rock contains abundant siliceous xenoliths of chert ripped up from the underlying cherty units, as well as abundant quartz and calcite amygdules, likely caused by the release of volatile fluids from the underlying sediments as the mafic flow was deposited. One 15cm wide dike of similar composition crosscuts the underlying rocks and is interpreted to be a feeder dike to the flow. The sequence of sedimentary rocks is roughly 80 m thick and represents a significant disconformity between the arc volcanic rocks of Package A and the back-arc basin volcanic rocks of Package B (Fig. 4). These rocks have a strong electromagnetic signature which is traceable along strike; eastward, the anomaly coincides with the mafic metasedimentary rocks along the highway in Schreiber and with chemical metasedimentary rocks at the Morley occurrence southeast of town (Fig. 8). Several other chemical sedimentary rocks occur east of town, including two occurrences of marble interbedded with argillite and sulphide facies iron formation and the outcrop of sulphide-bearing chert and argillite observed earlier at the intersection of Highway 17 and Hays Lake Road (Fig. 8). Whether these nearby rocks represent separate sequences or are part of the same depositional sequence but separated by cryptic folding requires further, more detailed mapping. The strata in this area are arranged in open, upright folds, which is apparent in this outcrop. However, the more fissile argillite-rich zones have developed C-S fabrics, kink folds and kinematic indicators like sigma and delta clasts that all indicate a significant amount of dextral shearing has affected these rocks. Because the only place locally that this deformation has been observed is in these argillitic rocks, the timing and cause of this shearing is unknown. Return to vehicles and turn left to drive east on Highway 17. 1.2 km - Drive 1.2 km into the town of Schreiber and take the third right onto Winnipeg Street. This is the street immediately east of the Golden Rail chip truck. 600 m - Drive to the end of Winnipeg Street, where you will see a railroad museum, and turn right onto Scotia Street. 70 m - Drive 1 block west on Scotia Street, then turn left on Subway Street. 210 m - Driving south on Subway Street, you will pass beneath the railway. Take the first right onto Isbester Drive. 2.3 km - Drive south to the end of Isbester Drive. There is a parking lot at the end of the road, and an outhouse and a gazebo down near the beach. Walk down the footpath to the beach, then turn right (west) and walk towards the first outcrop. This is Stop 8. Stop 8. Schreiber Beach Outcrops UTM coordinates 478600E 5404600N The outcrops at Schreiber Beach (Photo 3) are the best exposure of a conformable sequence of Archean metavolcanic and metasedimentary rocks in the Schreiber–Hemlo greenstone belt and perhaps throughout Ontario (Fig. 9). From the easternmost outcrop to the Schreiber Channel Provincial Nature Reserve, where exposures of Proterozoic Photo 3. View of Schreiber Beach and the rocky shoreline westward, taken from a vista along the Casque-Isles Trail. - 32 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 9. Simplified geological map of the shoreline west of Schreiber Beach on Lake Superior. Metavolcanic flows and metasedimentary sequences are labelled I-XII and i-iii respectively. Proterozoic rocks of interest are labelled. Note that the Casque-Isles trail is well constrained for two km west of the Schreiber Beach, but the author was not able to get better a resolution trace of the trail further west. All UTM co-ordinates provided using NAD83 in Zone 16. rocks interrupt the Archean exposure, there are three kilometers (as the crow flies) of near-continuous rocky shoreline, with less than 5% of the shoreline covered by cobble beaches. Winter ice scraping against the rocky shoreline keeps the rocks up to several metres inland lichen-free, displaying beautiful mineral, volcanic and sedimentary textures. The rocks are not pervasively deformed; only crosscut by discrete fractures with minimal displacement and up to metrewide shear zones. There are only a few altered areas in which the primary textures of the rocks are preserved. The upright stratigraphy strikes northwest and dips shallowly to the northeast, such that the straight, eastwest trending shoreline provides for us a perfect crosssection; walking west along the shoreline guides us down through stratigraphy. The lead author only had the opportunity to map the shoreline between Schreiber Beach and Twin Harbours, and so that stretch will be the focus of this description. From compiled data, the Archean rock between Twin Harbours and the granite-greenstone contact to the west is composed mainly of mafic metavolcanic rock. The world-famous Schreiber Channel Stromatolite outcrop and associated Gunflint conglomeratic rocks occurs along this stretch of shoreline, and a Keweenawan diabase sill occupies the Archean-Proterozoic unconformity. Eastward from Twin Harbours, higher in the stratigraphy, are five consecutive mafic metavolcanic flows up to 200 m in apparent thickness (roughly 175 m in true thickness, with an estimated 30 degree dip; Fig. 9). These flows are massive at the base, with medium to coarse-grained equigranular textures that could easily be mistaken for mafic intrusive rocks. Thin beds of sulphide-bearing chert are present along the flow contacts. Flow II is crosscut by an alkalic diabase dike, flow III is crosscut by two north-trending, carbonate altered mafic dikes up to 25 m wide, and flow V is crosscut by a series of north-trending dikes that display unique Liesegang textures. A 50 m thick sequence (i) of tuffaceous conglomerates with pebble to cobble-sized mafic and felsic volcanic clasts lies atop flow V (Fig. 9). Minor graded beds of sandy to gravelly material are present within these conglomerates. These rocks are crosscut by alkalic diabase dikes, and unconformably overlain by an outlier of the Gunflint Formation basal conglomerate, including cherty stromatolite domes similar to those observed at the Schreiber Channel Provincial Nature Reserve. The conglomerate is massive, polymictic and clast-supported, with dominantly pebble to cobblesized clasts. The conglomerates are interpreted to have been deposited in a shallow water environment akin to the cobble beaches present along the shores of Lake Superior today. - 33 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Eastward from metasedimentary sequence i are two consecutive massive to pillowed mafic flows similar to flows I-V. Another outlier of the Gunflint basal conglomerates and stromatolites, crosscut by an alkalic diabase dike, unconformably overlies pillows near the top of flow VI (Fig. 9). A thin sequence of tuffaceous wacke (ii) marks the top of flow VII. This is overlain by an intermediate volcanic flow (VIII) with plagioclase and amphibole phenocrysts. This flow is generally massive, with an agglomerate of similar composition at its top. This agglomerate grades into the polymictic tuffaceous conglomerate (iii) similar to the conglomerates in sequence i. Another massive intermediate flow (IX) with plagioclase and amphibole phenocrysts overlies the conglomerates of sequence iii roughly 100 m wide, with agglomerate at its top similar to that of flow VIII. This agglomerate is overlain by a 3 m sequence of normally graded tuffaceous wacke, which is then overlain by a massive to brecciated felsic flow with feldspar phenocrysts (Flow X). Two consecutive massive to pillowed mafic flows (XI and XII) overlie Flow X with a 1 m wide sequence of banded chert and magnetite between them. Outcrop exposure ends just above the base of Flow XII, which is then covered by the sands of Schreiber Beach. This eastern-most outcrop will be the subject of our last stop on the field trip (Fig. 9). To access the shoreline to the west, one could either walk along the rocky shoreline, which is quite rugged and slippery when wet, or follow the CasqueIsles Trail, which intermittently jogs down towards the shoreline (Fig. 9). There is a stream with steep-sided banks roughly halfway between Schreiber Beach and Twin Harbours, at the contact between flows IV and V. At the mouth of this stream, there are cobbles and boulders along the shoreline that may be crossed if there are no waves on Lake Superior and the outflow from the stream is minimal. A small wooden footbridge, wide enough for one person, crosses this stream about 500 metres to the north along the Casque-Isles Trail. Neither of these choices are particularly safe, so exercise extreme caution whichever path you choose to follow. Note that aside from the Schreiber Channel stromatolites, most of the rocks observed between here and Twin Harbours are massive to pillowed mafic flows (Fig. 9). To reach the Schreiber Channel stromatolites, it is recommended to start at the west end of the trail in Rossport or to approach the location by boat. End of road log. Acknowledgments The author would like to thank the field crews from the summers of 2015 (Joseph Walker, Andrea Nywening, Matthew Hanewich and Lauren Madronich), 2016 (Kira Arnold, Mallory Metcalf, Lucas Wolfe and Haley Aldred), 2017(Kira Arnold, Joshua Nguyen, Maddison Hodder and Gabrielle Klemt) and 2018 (Evelyn Moorhouse, Cassandra Powell, Mateo DoradoTroughton, Jessica Verschoor, Shadman Islam and Rachel Bourassa) for their hard work and perseverance through the particularly rough terrain. The author would like to thank Evan Hastie for co-leading the 2018 field season. The author would also like to thank the Richards family of Terrace Bay, who hosted the crew at their Jackfish Lake cottages on Highway 17 during the 2015-2018 field seasons, with special thanks to local prospector Wayne Richards, for all of his logistical aid and for sharing his abundance of local mineral exploration knowledge. Thanks to the people of Pic River and Pic Mobert First Nations communities for their gracious blessing and for allowing us to work on their traditional lands. The author would also like to thank local prospector Rudy Wahl, Mike Koziol of Alto Ventures Ltd. And Troy Gill of Sanatana Resources for tours of their properties and allowing us access to their properties over the last several field seasons. Thanks also to Mark Smyk, Dorothy Campbell and Mark Puumala of the Resident Geologist Program Thunder Bay office for their help during this project. Thanks to Michael Easton and Riku Metsaranta for their careful edits, and to Laura Ratcliffe for help with the figures. References Addison W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S.A., Fralick, P.W. and Hammond, A.L. 2005. Discovery of distal ejecta from the 1850 Ma Sudbury impact event; Geology, v.33, p.193-196. Anglin, C.D., Franklin, J.M., Loveridge, W.D., Hunt, P.A. and Osterberg, S.A. 1988. Use of zircon U-Pb ages of felsic intrusive and extrusive rocks in eastern Wabigoon Subprovince, Ontario, to place constraints on base metal and gold mineralization; in Radiogenic age and isotopic studies: report 2, Geological Survey of Canada, Paper 88-2, p.109-115. - 34 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Arnold, K.A., Hollings, P. and Magnus, S.J. 2017. Geology and mineral potential of the Terrace Bay pluton, western Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2017, Ontario Geological Survey, Open File Report 6333, p.12-1 to 12-6. Beakhouse, G.P. and Davis, D.W. 2005. Evolution and tectonic significance of intermediate to felsic plutonism associated with the Hemlo greenstone belt, Superior Province, Canada; Precambrian Research, v.137, p.61-92. Blackburn, C.E., Bond, W.D., Breaks, F.W., Davis, D.W., Edwards, G.R., Poulsen, K.H., Trowell, N.F. and Wood, J. 1985. Evolution of Archean volcanicsedimentary sequences of the western Wabigoon Subprovince and its margins: a review; in Evolution of Archean Supracrustal Sequences, Geological Association of Canada, Special Paper 28, p.89-116. Bleeker, W., Hamilton, M.A., Ernst, R.E. and Soderlund, U. 2012. Resolving the age structure of the Matachewan event: magmatic pulses at c. 2445-2452 Ma, 245802461 Ma and 2475-2480 Ma; unpublished CAMIRO Reports A96, A97, and A98, 17p. available from www.supercontinent.org Bleeker, W., Liikane, D.A., Smith, J., Hamilton, M., Kamo, S.L., Cundari, R., Easton, M., and Hollings, P. 2018. Controls on the localization and timing of mineralized intrusions in intra-continental rift systems, with a specific focus on the ca. 1.1 Ga Mid-continent Rift system; in Targeted Geoscience Initiative: 2017 report of activities, volume 2; Geological Survey of Canada, Open File 8373, p. 15–27. https://doi. org/10.4095/306594 Buchan, K.L., Mortensen, J.K. and Card, K.D. 1993. Northeast-trending Early Proterozoic dykes of southern Superior Province: multiple episodes of emplacement recognized from integrated paleomagnetism and U-Pb geochronology; Canadian Journal of Earth Sciences, v.30, p.1286-1296. Cabanis, B. and Lecolle, M. 1989. The La/10-Y/15-Nb/8 diagram: A tool for distinguishing volcanic series and discovering crustal mixing and/or contamination; Comptes Rendus de Academie des Sciences, Série IIA, v.309, p.2023-2029. Cannon, W.F., Schulz, K.J., Horton, J.W. and Kring, D.A. 2010. The Sudbury impact layer in the Paleoproterozoic iron ranges of northern Michigan, USA; Geological Society of America Bulletin, v.122, p.50-75. Corfu, F. 2000. Extraction of Pb with artificially too-old ages during stepwise dissolution experiments on Archean zircon, Lithos, v.53, p.279-291. Corfu, F. and Muir, T.L. 1989. The Hemlo–Heron Bay greenstone belt and Hemlo Au–Mo deposit, Superior Province, Ontario, Canada, Chemical Geology (Isotope Geoscience Section), v.79, p.183-200. Corfu, F. and Stott, G.M. 1986. U-Pb ages for late magmatism and regional deformation in the Shebandowan Belt, Superior Province, Canada; Canadian Journal of Earth Sciences, v.23, p.1075-1082. Corfu, F. and Stott, G.M. 1998. Shebandowan greenstone belt, western Superior Province: U-Pb ages, tectonic implications, and correlations; Geological Society of America Bulletin, v.110, p.1467-1484. Cundari, R. 2012. Geology and geochemistry of Midcontinent Rift-related igneous rocks; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 142p. Davis, D.W. 1996. Provenance and depositional age constraints on sedimentation in the western Superior transect area from U-Pb ages of zircons; in Western Superior Transect Second Annual Workshop, Lithoprobe Secretariat, University of British Columbia, Vancouver, British Columbia, Lithoprobe Report No. 53, p.18-23. Davis, D.W., Beakhouse, G.P. and Jackson, S.L. 1998. U-Pb zircon and titanite geochronology; Part 3, p.18, in S.L. Jackson, G.P. Beakhouse and D.W. Davis, Geological Setting of the Hemlo GoldDeposit; an Interim Progress Report, Ontario Geological Survey, Open File Report 5977, 121p. Davis, D.W., Krogh, T.E., Hinzer, J. and Nakamura, E. 1985. Zircon dating of polycyclic volcanism at Sturgeon Lake and implications for base metal mineralization; Economic Geology, v.80, p.1942-1952. Davis, D.W. and Lin, S. 2003. Unravelling the geologic history of the Hemlo Archean gold deposit, Superior Province, Canada: A U-Pb geochronological study; Economic Geology, v.98, p.51-67. Davis, D.W., Pezzutto, F. and Ojakangas, R.W. 1990. The age and provenance of metasedimentary rocks in the Quetico Subprovince, Ontario, from single zircon analyses: implications for Archean sedimentation and tectonics in the Superior Province; Earth and Planetary Science Letters, v.99, p.195-205. Davis, D.W., Schandl, E.S. and Wasteneys, H.A. 1994. U-Pb dating of minerals in alteration halos of Superior Province massive sulphide deposits: Syngenesis versus metamorphism; Contributions to Mineralogy and Petrology, v.115, p.427-437. Davis, D.W. and Stott, G.M. 2003. Geochronology of two Proterozoic mafic dike swarms in northwestern Ontario; in Summary of Field Work and Other Activities 2003, Ontario Geological Survey, Open File Report 6120, p.12-1 to 12-7. Davis, D.W. and Sutcliffe, C.N. 2017. U-Pb geochronology by LA-ICPMS in samples from northern Ontario, internal report for the Ontario Geological Survey; - 35 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 131p. Davis, D.W. and Sutcliffe, R.H. 1985. U-Pb ages from the Nipigon Plate and Northern Lake Superior; Geological Society of America, Bulletin, v.96, p.1572-1579. Davis, S. 2016. Petrology and geochemistry of the Wolfcamp Lake basalts; unpublished HBSc thesis, Lakehead University, Thunder Bay, Ontario, 68p. Davis, S., Hollings, P. and Cundari, R.M. 2017. Geochemistry of the Mesoproterozoic Wolfcamp Lake basalts, northwestern Ontario; Ontario Geological Survey, Miscellaneous Release—Data 345. Fage, A. 2011. Geology, geochemistry and geochronology of the Hemlo East property, Schreiber–Hemlo greenstone belt, Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 155p. Fralick, P. and Davis, D. 1999. The Seine-Couchiching problem revisited: Sedimentology, geochronology, and geochemistry of sedimentary units in the Rainy Lake and Sioux Lookout areas; in 1999 Western Superior Transect 5th Annual Workshop, Lithoprobe Secretariat, The University of British Columbia, Vancouver, British Columbia, Lithoprobe Report No. 70, p.66-75. Fralick, P., Davis, D.W. and Kissin, S.A. 2002. The age of the Gunflint Formation, Ontario, Canada: single zircon U-Pb age determinations from reworked volcanic ash; Canadian Journal of Earth Sciences, v. 39, p.1085-1091. Fralick, P., Purdon, R.H. and Davis, D.W. 2006. Neo-Archean trans-subprovince sediment transport in southwestern Superior Province: sedimentological, geochemical and geochronological evidence; Canadian Journal of Earth Sciences, v.43, p.1055-1070. Halls, H.C. and Davis, D.W. 2004. Paleomagnetism and U-Pb geochronology of the 2.17 Ga Biscotasing dyke swarm, Ontario, Canada: evidence for verticalaxis crustal rotation across the Kapuskasing Zone; Canadian Journal of Earth Sciences, v.41, p.255-269. Halls, H.C., Davis, D.W., Stott, G.M., Ernst, R.E. and Hamilton, M.A. 2008. The Paleoproterozoic Marathon Large Igneous Province: New evidence for a 2.1 Ga long-lived mantle plume event along the southern margin of the North American Superior Province; Precambrian Research, v.162, p.327-353. Hamilton, M.A. and Stott, G.M. 2008. The significance of new U/Pb baddeleyite ages from two Paleoproterozoic diabase dikes in northern Ontario; in Summary of Field Work and Other Activities 2008, Ontario Geological Survey, Open File Report 6226, p.17-1 to 17-10. Hart, T.R., Termer, M. and Jolette, C. 2002. Precambrian geology of Kitto, Eva, Summers, Dorothea and Sandra townships, northwestern Ontario: Phoenix bedrock mapping project; Ontario Geological Survey, Open File Report 6095, 206p. Heaman, L.M. 1997. Global mafic magmatism at 2.45 Ga: Remnants of an ancient large igneous province?; Geology, v.25, p.299-302. Heaman, L.M. and Easton, R.M. 2006. Preliminary U-Pb geochronology results from the Lake Nipigon Region Geoscience Initiative; Ontario Geological Survey, Miscellaneous Release—Data 191, 1 CD-ROM. Heaman, L.M., Easton, R.M., Hart, T.R., Hollings, P., MacDonald, C.A. and Smyk, M. 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario; Canadian Journal of Earth Science, v.44, p.1055-1086. Heaman, L.M. and Machado, H. 1987. Isotope geochemistry of the Coldwell alkaline complex I—U-Pb studies on accessory minerals; in Program with Abstracts, Geological Association of Canada-Mineralogical Association of Canada, v.12, p.54. Heaman, L.M. and Machado, H. 1992. Timing and origin of Midcontinent Rift alkaline magmatism, North America: evidence from the Coldwell complex; Contributions to Mineralogy and Petrology, v.110, p.289-303. Jensen, L.S. 1976. A new cation plot for classifying subalkalic volcanic rocks; Ontario Department of Mines, Miscellaneous Paper 66, 22p. Kamo, S.L. 2015. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 1: 20152016; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. Kamo, S.L. 2016. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: Bedrock Mapping Projects, Ontario, Year 1: 20152016, internal report prepared for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 48p. Kamo, S.L. and Hamilton, M.A. 2017. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: bedrock mapping projects, Ontario, Year 2: 2016-2017; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 72p. Kamo, S.L. and Hamilton, M.A. 2018. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: Bedrock Mapping Projects, Ontario, Year 3: 2017-2018, internal report for the Ontario Geological - 36 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 44p. Le Maitre, R.W. 1989. A classification of igneous rocks and glossary of terms; Blackwell Scientific Publications, Oxford, United Kingdom, 193p. Liikane, D.A., Bleeker, W., Hamilton, M., Kamo, S., Smith, J., Hollings, P., Cundari, R. and Easton, M. 2018. Controls on the localization and timing of mineralized intrusions within the ca. 1.1 Ga Midcontinent Rift system; 64th Institute on Lake Superior Geology, Proceedings, v.64, pt.1, p.65-66. Magnus, S.J. 2017a. Precambrian geology of Tuuri and Walsh townships, northwestern Ontario; Ontario Geological Survey, Preliminary Map P.3812, scale 1:20 000. Magnus, S.J. 2017b. Geology and mineral potential of Syine Township, western Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2017, Ontario Geological Survey, Open File Report 6333, p.11-1 to 11-8. Magnus, S.J. 2019. Precambrian geology, Syine township; Ontario Geological Survey, Preliminary Map P.3826, scale 1:20 000. Magnus, S.J. and Arnold, K.A. 2016. Geology and mineral potential of the western Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2016, Ontario Geological Survey, Open File Report 6323, p.11-1 to 11-7. Magnus, S.J. and Walker, J. 2015. Geology and mineral potential of Walsh, Tuuri and Syine townships, Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2015, Ontario Geological Survey, Open File Report 6313, p.14-1 to 14-12. Magnus, S.J. and Hastie, E.C.G. 2018. Geology and Mineral Potential of Priske and Strey Townships, western Schreiber–Hemlo greenstone belt; in Summary of Field Work and Other Activities, 2018, Ontario Geological Survey, Open File Report 6350, p.10-1 to 10-10. Marmont, S. 1984. The Terrace Bay Batholith and associated mineralization; Ontario Geological Survey, Open File Report 5514, 95p. Miller, J.D., and Nicholson, S.W., 2013, Geology and mineral deposits of the Midcontinent Rift—An overview, in Field Guide to the Cu-Ni-PGE Deposits of the Lake Superior Region; Precambrian Research Center Guidebook Series 13-1, University of Minnesota Press. Muir, T.L. 2003. Structural evolution of the Hemlo greenstone belt in the vicinity of the world-class Hemlo gold deposit; Canadian Journal of Earth Sciences, v.40, p.395-430. Norris, D. 2012. Current archaeological investigations in Ontario: The discovery of and preliminary information regarding several Paleoindian sites east of Thunder Bay; The Minnesota Archaeologist, v.71, p.45-59. Ojakangas, R.W., Morey, G.B. and Southwick, D.L. 2001. Paleoproterozoic basin development and sedimentation in the Lake Superior region, North America; Sedimentary Geology, v.141-142, p.319341. Ontario Geological Survey 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release---Data 126-Revision 1. ISBN 978-1-4435-5704-7 (CD) ISBN 978- 1-4435-5705-4 [zip file] Osmani, I.A. and Stott, G.M. 1988. Regional-scale shear zones in Sachigo Subprovince and their economic significance; in Summary of Field Work and Other Activities 1988, Ontario Geological Survey, Miscellaneous Paper 141, p.53-67. Pettijohn, F.J. 1975. Sedimentary Rocks, 2nd ed.; Harper and Row Publishers, New York, 628 p. Pleger, T.C. 2000. Old copper and red ocher social complexity; Midcontinental Journal of Archaeology, v.25, p.169-190. Rogala, B., Fralick, P.W. and Metsaranta, R. 2005. Stratigraphy and sedimentology of the Mesoproterozoic Sibley Group and related igneous intrusions, northwestern Ontario: Lake Nipigon Region Geoscience Initiative; Ontario Geological Survey, Open File Report 6174, 128p. Rukhlov, A.S. and Bell, L. 2010. Geochronology of carbonatites from the Canadian and Baltic shields, and the Canadian Cordillera: clues to mantle evolution; Mineralogy and Petrology, v.98, p.11-54. Schmid, R. 1981. Descriptive nomenclature and classification of pyroclastic deposits and fragments: recommendations of the IUGS Subcommission on the systematics of igneous rocks; Geology, v.9, p.4143. Schnieders, B.R., Smyk, M.C., Speed, A.A. and McKay, D.B. 1996. Mineral occurrences in the NipigonMarathon Area, Volumes 1 and 2; Ontario Geological Survey, Open File Report 5951, 912p. Smyk, M.C. and Schnieders, B.R. 1995. Geology of the Schreiber greenstone assemblage and its gold and base metal mineralization; 41st Institute on Lake Superior Geology, Proceedings, Part 2c, field trip guidebook, 77p. Snake, S., Coatsworth, E.S., Coatsworth, D. and Kaggie, F. 1991, The adventures of Nanabush, Ojibway Indian stories; 2nd edition, Doubleday Canada, Toronto, Ontario. - 37 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stott, G.M., Corkery, M.T., Percival, J.A., Simard, M. and Goutier, J. 2010. A revised terrane subdivision of the Superior Province; in Summary of Field Work and Other Activities, 2010, Ontario Geological Survey, Open File Report 6260, p.20-1 to 20-10. Sun, S-S. and McDonough, W F. 1989. Chemical and isotopic systematics of oceanic basalts: Implications for mantle compositions and processes; in Magmatism in ocean basins: Geological Society of London, Special Publication 42, p.313-345. Tóth, Z. 2018. The geology of the Beardmore-Geraldton belt, Ontario, Canada: geochronology, tectonic evolution and gold mineralization; unpublished PhD thesis, Laurentian University, Sudbury, Ontario, 276p. Tóth, Z. and Lafrance, B. 2018. Preliminary results from the assessment of the structural evolution of the southern Geraldton–Onaman transect; in Summary of Field Work and Other Activities, 2018, Ontario Geological Survey, Open File Report 6350, p.32-1 to 32-8. Toth, Z., Lafrance, B., Dube, B., McNicoll, V.J., MercierLangevin, P., and Creaser, R.A. 2015. Banded iron formation-hosted gold mineralization in the Geraldton area, northwestern Ontario: Structural setting, mineralogical characteristics, and geochronology, In: Targeted Geoscience Initiative 4: Contributions to the Understanding of Precambrian Lode Gold Deposits and Implications for Exploration; Geological Survey of Canada, Open File 7852, p. 85–97. Wacey, D., McLoughlin, N., Kilburn, M.R., Saunders, M., Cliff, J.B., Kong, C., Barley, M.E. and Brasier, M.D. 2013. Nanoscale analysis of pyritized microfossils reveals differential heterotrophic consumption in the ~1.9-Ga Gunflint chert; Proceedings of the National Academy of Sciences of the United States of America, v.110, no.20, p.8020-8024. Whitmeyer, S.J. and Karlstrom, K.E. 2007. Tectonic model for the Proterozoic growth of North America; Geosphere, v.3, no.4, p.220-259. Williams, H.R. 1989. Geological studies of the Wabigoon, Quetico and Abitibi–Wawa subprovinces, Superior Province of Ontario, with emphasis on the structural development of the Beardmore–Geraldton belt; Ontario Geological Survey, Open File Report 5724, 189p. Wu, F-Y, Mitchell, R.H., Li, Q-L, Zhang, C. and Yang, Y-H. 2017. Emplacement age and isotopic composition of the Prairie Lake carbonatite complex, northwestern Ontario, Canada; Geological Magazine, v.154, no.2, p.217-236. Zaleski, E., van Breemen, O. and Peterson, V.L. 1999. Geological evolution of the Manitouwadge greenstone belt and Wawa-Quetico subprovince boundary, Superior Province, Ontario, constrained by U-Pb zircon dates of supracrustal and plutonic rocks; Canadian Journal of Earth Sciences, v.36, p.945-966. - 38 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Appendix 1. Geochemical data for typical samples of the major rock types in the western Schreiber–Hemlo greenstone belt. All analyses were performed at the OGS Geoscience Laboratories, Sudbury. Complete data are available in Magnus (2017) (Tuuri and Walsh townships), Magnus (2019) (Syine township) and in upcoming Miscellaneous Release—Data reports for Priske and Strey townships. Sample Number Related Stop 17SJM013C 17SJM013C 17SJM013C 15SJM068A 15SJM068A Stop 2 Stop 2 Stop 2 Rock Name mafic volcanic rock Package B back arc volcanic 505364 5402267 base of flow mafic volcanic rock Package B back arc volcanic 505364 5402267 top of flow S&S Stop 2A mafic volcanic rock Package B back arc volcanic 506861 5402734 base of flow basalt mafic volcanic rock Package B back arc volcanic 505364 5402267 middle of flow basalt back arc basin 48.67 back arc basin 50.07 basaltic andesite back arc basin 52.7 S&S Stop 2A mafic volcanic rock Package B back arc volcanic 506861 5402734 "spinifex" texture basalt back arc basin 49.56 0.9 1.02 1.05 0.66 Al2O3 12.32 14.05 14.21 Cr2O3 0.078 0.044 0.054 Formation Volcanic Setting Easting (m) Northing (m) Notes TAS rock name Tectonic Setting SiO2 (wt %) TiO2 Fe2O3 total 16SJM157 A near Stop 2 17SJM082C 15SJM204B none melanogab bro back arc basin 46.17 picrobasalt back arc basin 36.97 mafic volcanic rock Package C plateau volcanic 501620 5412183 thoriumenriched basalt S&S Stop 2B mafic volcanic rock Package B plateau volcanic 506190 5403315 trachytic texture basalt continental arc 47.89 continental arc 45.45 0.34 0.22 2.18 1.93 12.08 6.2 4.43 15.36 14.76 0.08 0.31 0.58 0.008 0.02 basalt Package B back arc volcanic 506123 5402406 n/a 12.99 12.44 11.28 12.07 11.66 10.62 14.57 13.13 MnO 0.194 0.199 0.213 0.173 0.137 0.175 0.242 0.339 MgO 11.08 8.11 6.01 11.06 24.23 26.37 2.54 3.08 CaO 10.036 8.439 9.481 9.384 5.055 6.66 12.959 8.382 0.93 1.02 2.37 1.93 0.02 <0.02 0.42 2.9 Na2O K2 O 0.23 2.59 0.27 0.08 0.01 0.03 1.19 0.73 P2O5 0.099 0.109 0.118 0.047 0.027 0.018 0.356 0.372 LOI Total Mg Number 3.41 2.99 2.98 2.99 6.02 14.12 1.73 9.1 100.95 101.11 100.75 100.12 100.18 100.19 99.47 100.2 0.82 0.78 0.74 0.83 0.92 0.93 0.49 0.56 Th (ppm) 0.382 0.424 0.436 0.212 0.097 0.062 1.151 0.765 Nb 3.098 3.47 3.706 1.29 0.597 0.478 8.719 9.274 Ta 0.205 0.23 0.236 0.071 0.033 0.024 0.562 0.528 Ti 5267 5972 6139 4028 1952 1326 12625 11341 Zr 73 80 85 42 22 14 172 142 1.08 1.07 1.07 1.05 1.08 0.98 2.53 3.65 La/LuCI Total REE 44.06 49.17 52.90 25.42 10.89 8.57 112.25 108.66 Notes: “Formation” refers to the depositional package, pluton or dike swarm that the sample is related to; “Volcanic Setting” refers to the volcanic environment inferred using different geochemical and geological parameters; “Tectonic Setting” refers to the inferred tectonic setting of mafic rocks, based on Cabanis and Lecolle (1989); “TAS Rock Name” refers to the rock name based on the Total Alkalis versus Silica diagram from Le Maitre (1989); Major element oxides are in weight %; trace element data are in parts per million; Mg number = atomic Mg/Mg + Fe, where Fe = total Fe expressed as ferrous iron; this value is dimensionless La/LuCI includes elements normalized using values from Sun & McDonough (1995); this value is dimensionless Abbreviations: LOI = loss-on-ignition; n/a = not applicable; REE = rare earth elements; S&S = Smyk and Schnieders (1995); TAS = Total Alkalis versus Silica; - 39 24- �Proceedings of the 65th ILSG Annual Meeting - Part 2 Appendix 1, continued Sample Number Related Stop Rock Name Formation Volcanic Setting Easting (m) Northing (m) Notes TAS Rock Name Tectonic Setting SiO2 (wt %) 17SJM016B 17SJM016A 16SJM213A 16WM043A Stop 8 mafic volcanic rock Package A Stop 8 mafic volcanic rock Package A none lapilli tuff none crystal tuff none tuffaceous conglomerate none tuff like Stop 2 wacke Package A Package A arc volcanic arc volcanic arc volcanic B-C Disconformity arc volcanic Package D arc volcanic B-C Disconformity arc volcanic 478600 5404609 pillow core, Flow XI 478600 5404609 base of flow, Flow XII 512241 5409187 498541 5402882 clasts up to 5cm andesite dacite dacite 499867 5412118 with interbedded chert dacite 513955 5411351 n/a basaltic andesite transitional arc 514731 5408709 quartz and feldspar phenocrysts rhyolite calc-alkaline calc-alkaline calc-alkaline calc-alkaline calc-alkaline calc-alkaline 58.89 70.26 76.1 69.22 64.6 63.96 54.7 17SJM129C 17SJM168B 16WM027A n/a dacite TiO2 1.32 0.74 0.47 0.1 0.56 0.49 0.57 Al2O3 15.92 15.16 14.66 12.05 13.14 12.35 15.92 Cr2O3 0.027 0.022 <0.002 0.01 0.009 0.024 0.03 Fe2O3total 11.85 6.44 3.75 1.35 3.61 5.66 5.99 MnO 0.11 0.084 0.052 0.047 0.034 0.17 0.094 MgO 3.31 4.53 1 0.11 4.04 3.55 3.14 CaO 2.651 5.437 3.518 0.502 2.99 6.886 3.511 Na2O 2.38 3.63 4.27 0.45 2.82 2.19 4.03 K2 O 1.84 0.51 1.35 8.99 0.61 1.09 1.79 P2O5 0.12 0.174 0.113 0.017 0.145 0.108 0.179 LOI 5.62 5.33 1.28 0.85 2.23 2.35 1.04 Total 99.86 100.98 100.76 100.64 99.42 99.5 100.28 0.60 0.79 0.59 0.31 0.86 0.77 0.74 Th (ppm) 0.559 2.941 3.728 5.249 1.816 1.764 9.035 Nb 3.969 6.999 6.471 9.436 4.74 3.572 7.056 Ta 0.251 0.468 0.603 0.918 0.335 0.275 0.471 Ti 6481 5823 2781 563 3209 2842 3462 Mg Number Zr La/LuCI 86 175 187 167 133 86 158 3.29 13.81 6.25 5.12 6.33 9.55 19.71 Total REE 61.83 163.67 101.45 124.07 80.13 58.96 181.12 Notes: “Formation” refers to the depositional package, pluton or dike swarm that the sample is related to; “Volcanic Setting” refers to the volcanic environment inferred using different geochemical and geological parameters; “Tectonic Setting” refers to the inferred tectonic setting of mafic rocks, based on Cabanis and Lecolle (1989); “TAS Rock Name” refers to the rock name based on the Total Alkalis versus Silica diagram from Le Maitre (1989); Major element oxides are in weight %; trace element data are in parts per million; Mg number = atomic Mg/Mg + Fe, where Fe = total Fe expressed as ferrous iron; this value is dimensionless La/LuCI includes elements normalized using values from Sun & McDonough (1995); this value is dimensionless Abbreviations: LOI = loss-on-ignition; n/a = not applicable; REE = rare earth elements; S&S = Smyk and Schnieders (1995); TAS = Total Alkalis versus Silica; - 4025 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Appendix 1, continued Sample Number Related Stop Rock Name 17SJM147A 17SJM038A 15JW094A 17KA040A 15JW006A 16SJM171A 17SJM200A 17SJM033B Stops 4 and 5 granodiorite none granodiorite none granite none monzogranit e none tonalite none quartz syenite none porphyritic dike Terrace Bay pluton Foxtrap Lake pluton Santoy Lake pluton Steel River pluton Syenite Lake pluton porphyritic dikes 506412 5410351 dacite Crossman Lake batholith 488795 5423075 dacite 522826 5412786 rhyolite 507404 5412210 rhyolite 511259 5405636 dacite 487180 5415617 trachy-dacite 482667 5405457 dacite 64.15 69.74 71.6 70.18 none quartz monzodiori te Little Pic River pluton 524180 5413445 trachyandesite 61.77 64.34 67.92 67.67 TiO2 0.44 0.36 0.18 0.28 0.58 0.53 0.25 0.25 Al2O3 15.19 15.19 16.04 14.99 17.36 14.05 15.82 14.87 Cr2O3 0.017 0.006 <0.002 0.004 <0.002 0.02 0.006 0.022 Formation Easting (m) Northing (m) TAS rocktype SiO2 (wt %) Fe2O3 total 4.04 3.56 1.37 2 5.1 6.1 2.21 2.73 MnO 0.078 0.062 0.024 0.025 0.086 0.082 0.04 0.042 MgO 3.03 0.98 0.5 0.55 2.17 2.91 1.03 2.49 CaO 3.89 3.282 1.122 1.129 4.113 2.713 2.067 2.937 Na2O 4.57 4.09 6.31 5.23 4.67 2.84 4.95 5.58 K2 O 2.45 1.52 2.95 4.66 2.86 2.42 4.17 0.89 P2O5 0.256 0.13 0.085 0.127 0.292 0.138 0.136 0.099 LOI 1.46 1.19 0.81 0.81 0.72 3.81 Total 99.7 100.14 101.14 100.08 99.84 100.04 99.3 101.02 Mg Number 0.80 0.60 0.67 0.60 0.70 0.72 0.72 0.83 Th (ppm) 9.668 3.212 5.215 32.841 6.723 7.388 13.741 1.795 Nb 5.497 7.076 3.251 11.59 6.843 5.554 7.084 2.326 Ta 0.378 0.726 0.194 0.683 0.385 0.42 0.612 0.193 Ti 2551 2046 1106 1661 3396 3278 1477 1438 Zr 190 160 94 287 191 144 174 87 24.83 5.45 19.80 52.14 23.09 17.57 31.85 17.62 La/LuCI 3.39 Total REE 244.97 77.55 84.17 270.77 230.87 137.44 185.46 66.37 Notes: “Formation” refers to the depositional package, pluton or dike swarm that the sample is related to; “TAS Rock Name” refers to the rock name based on the Total Alkalis versus Silica diagram from Le Maitre (1989); Major element oxides are in weight %; trace element data are in parts per million; Mg number = atomic Mg/Mg + Fe, where Fe = total Fe expressed as ferrous iron; this value is dimensionless La/LuCI includes elements normalized using values from Sun & McDonough (1995); this value is dimensionless Abbreviations: LOI = loss-on-ignition; n/a = not applicable; REE = rare earth elements; S&S = Smyk and Schnieders (1995); TAS = Total Alkalis versus Silica; - 26 41 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Appendix 1, continued Sample Number Related Stop Rock Name Formation Easting (m) Northing (m) Notes TAS rocktype Tectonic Setting SiO2 (wt %) 16SJM178D 15SJM019B 15JW072E 16SJM119C 15SJM068B 17SJM006B 16SJM139A Stop 1 none Stops 1 and 4 Stop 4 S&S Stop 2A none none alkalic alkalic lamprophyre subalkalic subalkalic subalkalic subalkalic diabase diabase diabase diabase diabase diabase rift-parallel rift-parallel lamprophyre Biscotasing Marathon Matachewan Pigeon River alkalic dikes alkalic dikes dikes 515772 523556 512213 517970 506861 498906 506570 5408307 5407664 5405453 5406078 5402734 5412471 5407803 n/a trachytic n/a n/a n/a n/a n/a texture basalt trachy-dacite foidite basalt basalt basalt basalt continental continental intercontinental continental calc-alkaline calc-alkaline continental arc arc rift arc arc 46.11 60.78 29.55 49.74 47.65 50.08 48.02 TiO2 1.19 0.49 4.35 1.22 0.74 1.45 1.93 Al2O3 14.54 15.66 3.97 14.77 13.35 13.68 16.04 Cr2O3 0.02 <0.002 0.1 0.02 0.11 0.02 0.02 Fe2O3total 12.77 8.63 15.76 14.04 10.82 14.92 13.86 MnO 0.217 0.234 0.269 0.206 0.171 0.19 0.193 MgO 5.51 0.34 15.94 6.4 10.47 6.9 6.04 CaO 10.656 1.814 13.069 10.428 11.079 8.774 9.742 Na2O 3.22 6.41 0.1 2.27 1.71 2.81 2.75 K2 O 1.53 4.82 2.35 0.4 0.31 0.39 0.48 P2O5 1.099 0.077 0.876 0.118 0.117 0.098 0.214 LOI 2.89 0.82 12.61 0.7 3.33 1.5 0.54 Total 99.92 100.08 99.04 100.32 99.87 100.83 99.86 0.70 0.18 0.85 0.71 0.84 0.72 0.70 Mg Number Th (ppm) 9.282 56.416 8.508 1.138 0.496 3.17 1.669 Nb 63.339 >277 124.706 4.426 2.866 6.18 11.382 Ta 2.739 14.976 7.901 0.28 0.129 0.435 0.774 Ti 7279 2967 >25000 6976 4623 3630 11126 Zr 180 1041 375 80 62 148 160 26.63 19.55 53.01 2.16 4.20 11.19 3.50 La/LuCI Total REE 468.09 923.53 418.37 56.87 62.82 144.04 103.50 Notes: “Formation” refers to the depositional package, pluton or dike swarm that the sample is related to; “Tectonic Setting” refers to the inferred tectonic setting of mafic rocks, based on Cabanis and Lecolle (1989); “TAS Rock Name” refers to the rock name based on the Total Alkalis versus Silica diagram from Le Maitre (1989); Major element oxides are in weight %; trace element data are in parts per million; Mg number = atomic Mg/Mg + Fe, where Fe = total Fe expressed as ferrous iron; this value is dimensionless La/LuCI includes elements normalized using values from Sun & McDonough (1995); this value is dimensionless Abbreviations: LOI = loss-on-ignition; n/a = not applicable; REE = rare earth elements; S&S = Smyk and Schnieders (1995); TAS = Total Alkalis versus Silica; - 4227 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 4 - Geology of the Nipigon Area Philip Fralick Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada and Robert Cundari Resident Geologist Program, Ontario Geological Survey, Ministry of Northern Development and Mines, Thunder Bay, Ontario, P7E 6S7, Canada Introduction The Nipigon area, west of Terrace Bay, hosts Neoarchean, amphibolite-facies metamorphic and intrusive rocks of the Quetico Subprovince, as well as overlying Mesoproterozoic Sibley Group sedimentary rocks and Midcontinent Rift-related mafic intrusive rocks. This trip buids upon previous ILSG trips, namely Fralick et al. (2000) and Smyk and Kissin (2005). A lot of this guide is taken from them. The field trip begins on Highway 17 just north of Lake Helen focusing on metamorphosed clastic sedimentary rocks, their high-grade metamorphic equivalents and derived granitic rocks of the Quetico Subprovince (Stops 1-6; Fig. 1). The trip continues east of Nipigon highlighting Proterozoic sedimentary rocks of the Sibley Group and intrusive mafic rocks related to the Midcontinent Rift (Stops 7-10; Fig. 1). Many stops, especially Stop 1 through 6, are on road cuts along a narrow section of highway. Please use extreme caution when viewing road side outcrops. Regional Geology - Quetico Subprovince The Quetico Subprovince of the Superior Province is situated between the Wabigoon and Wawa volcanoplutonic subprovinces that bound the Quetico on its northern and southern margins, respectively. This east- Figure 1: Geology of the Nipigon area field trip stop locations. - 43 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 trending subprovince has a fairly consistent width of 70 km and is composed predominantly of metasedimentary rocks and their migmatitic and anatectic derivatives (Fig. 2; Williams, 1991). In general, the boundaries of the Quetico, whether or not they may be primary and/or tectonic, have been mapped as steeply dipping surfaces across which there is commonly a distinct contrast in lithology. However, LITHOPROBE deep seismic has shown that the Quetico and Beardmore-Geraldton volcanic and sedimentary belts of the Wabigoon Subprovince are thrust over the volcanic rocks of the Onaman-Tashota terrane of the Wabigoon Subprovince at an angle of 12 degrees (Geller, 2012). This agrees well with prior interpretations of the Quetico as an accretionary prism (Devaney and Williams, 1989) and the Beardmore-Geraldton area as its associated forearc basin (Barrett and Fralick, 1989; Fralick et al., 1992) that were thrust northward onto the Wabigoon arc during Wawa arc collision. An overview of the lithologic, metamorphic, structural and tectonic characteristics of the Quetico Subprovince has most recently been provided by Easton (2000): “The intensity of metamorphism varies within the subprovince, such that rocks marginal to the subprovince tend to be at lower grade than in the interior. The lowest metamorphic grade is found along the northern boundary with the Wabigoon subprovince (Pirie and Mackasey, 1978). Locally, subgreenschist- to greenschistfacies rocks occur along the southern boundary (Borradaile, 1982), but typically, there is a rapid rise in metamorphic grade north of the Wawa subprovince, especially north of Manitouwadge, where a belt of metasedimentary granulites occurs within the Quetico subprovince close to, and parallel with, the northern margin of the Wawa Subprovince (Coates, 1968; Williams and Breaks, 1989, 1990; Pan et al., 1994). As a result, grade distribution is asymmetrical, with the maximum in temperature and pressure occurring south of the central Quetico, locally coincident with the southern margin.” In contrast, Seemayer (1992) also described an asymmetric metamorphic grade distribution that had metamorphic grade increasing from south to north across the Quetico, southwest of Lake Nipigon. The southern margin was characterized by greenschist-facies rocks, the central portions were at amphibolite facies and the northern margin displayed an abrupt decrease in grade Figure 2. Generalized regional geology of the Quetico Subprovince (after Williams, 1991). - 44 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 adjacent to the Wabigoon Subprovince to the north. The Quetico Fault, which is normally situated at the northern margin of the Quetico in this western area, lies well within the Quetico southwest of Lake Nipigon (Seemayer, 1992). Seemayer (1992) determined temperatures based on garnet-biotite thermometry ranging from 526°C at the southern margin of the Quetico Subprovince, increasing asymmetrically to a maximum of 714°C, then falling sharply at the northern margin to 517°C. Pressure calculated at near peak temperature was 5 ± 1.5 kbar. Easton (2000) described the regional metamorphic conditions and metamorphic history: “P–T conditions increase from west to east, for example, 500ºC and 2.5 kbar at the Minnesota border west of Thunder Bay (Percival, 1989), to 700–780ºC and 5.4–6.1 kbar adjacent to the Kapuskasing structural zone (Percival and McGrath, 1986; Percival, 1989). Typical conditions in the central region are on the order of 620ºC and 3.3 kbar (Percival, 1989). Granulites north of Manitouwadge yield 680–770ºC and 4.4–6.4 kbar (Pan et al., 1994; Percival, 1989). The regional variation in P–T can be ascribed to a relatively shallow level of erosion in the west (<10 km) and a deeper level in the east (>12 km) (Percival et al., 1985). Rocks located east of the Kapuskasing structural zone are believed to be generally at upper-amphibolite-facies conditions (Williams, 1991).” facies being structurally controlled within thrustbounded panels (Williams, 1991).” “In contrast, the main phase of regional metamorphism (M2), which produced the observed map-pattern (Fig. 2), occurred late syntectonically (Sawyer, 1983; Williams, 1991). The general sequence of isograds, based on the appearance of diagnostic assemblages in pelites, is Chl–Ms–Bt, Grt+And+Sil, Grt+Crd+Sil, in situ granitic leucosome, and Opx (Pirie and Mackasey,1978; Percival and Stern, 1984). The common occurrence of Grt–And in metapelites in the western Quetico subprovince is diagnostic of bathozone 2 (<3.4 kbar; Carmichael, 1978), whereas the presence of Sil–St in the eastern Quetico is diagnostic of bathozones 3 and 4 (3.4– 5.5 kbar).” “As noted by Williams (1991), tectonic thickening of the sedimentary pile and intrusion of minor I-type granitic rocks occurred prior to the thermal acme. Most of the large pre- to syntectonic granitic bodies are peraluminous and have sedimentary sources but display little evidence of thermal contact metamorphism; one exception is the South Beatty Lake pluton in the northern Quetico subprovince (Pirie & Mackasey, 1978). Steeply dipping thermal gradients, local increases in temperature around large plutons, and the general association of the highest-grade rocks with abundant generation of leucosome, indicate that the source of heat was a combination of burial, upward magmatic transport, and tectonism.” “Evidence for an earlier, medium-pressure, low-temperature, pre-tectonic or early syntectonic metamorphism comes from four areas within the subprovince. In the Atikokan region, and in northern Minnesota, both at the northern margin of the Quetico subprovince, an early M1 metamorphic peak between D1 and D2 produced Ky–St–Bt assemblages (Ayres, 1978; Tabor et al., 1989). Kyanite inclusions in plagioclase within Grt–Sil–Bt–Pl–Qtz schist near Raith, north of Thunder Bay, have been reported by Percival et al. (1985). Kehlenbeck (1976) also presented textural evidence for a polymetamorphic history along the northern margin of the Quetico subprovince north of Thunder Bay. Again, along the northern margin of the subprovince, in the Beardmore–Geraldton area (Williams, 1989), amphibolite-facies conditions were attained prior to D2 deformation, with the distribution of “In the northern Quetico, M1 metamorphism is estimated to have occurred between 2698 Ma, the maximum age of sedimentation (Davis et al., 1990) and 2688+4 Ma, the age of emplacement of the late syntectonic Blalock pluton (ibid). In the southern Quetico, M1 occurred after 2690 Ma, the maximum age of sedimentation (Zaleski et al., 1999). The timing of M2 metamorphism is less well constrained and may have been protracted. In the Manitouwadge area, (ibid) constrained regional D2 deformation to 2680– 2677 Ma and suggested that migmatization in both the northern Wawa and southern Quetico subprovinces occurred after 2679 Ma, broadly coincident with D3 deformation. This inference is consistent with observations elsewhere in - 45 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 the Quetico that contact aureoles around late plutons, dated at 2671±2 to 2665±2 Ma, as well as late granitic pegmatites dated at 2653±4 Ma, overprint regional metamorphic fabrics (Percival and Sullivan, 1988; Percival, 1989). North of Manitouwadge, Pan et al. (1998) reported a U–Pb zircon age of 2666±1 Ma from a granitic pegmatite concordant with respect to the D3 fabric, and suggested that the regional amphibolite-facies metamorphism occurred between 2671 and 2665 Ma, consistent with the ages cited above. The timing of peak granulite-facies metamorphism north of Manitouwadge appears to be some 15 Ma younger, on the basis of U–Pb zircon ages of 2650± and 2651±3 Ma from a mafic granulite and a tonalitic leucosome, respectively (ibid). Zaleski & van Breemen (1997) reported that titanite ages young with increasing metamorphic grade, ranging from ~2686 Ma in the southern Manitouwadge greenstone belt to ~2640 Ma in the southern Quetico, suggesting that the thermal effects of regional metamorphism may have lasted over ~30 million years, from 2677 to 2640 Ma, in the higher-grade parts of the Wawa and Quetico subprovinces. On the basis of their regional geological and geochronological studies, Zaleski et al. (1999) concluded that “M2 metamorphism occurred after the tectonic juxtaposition of the Quetico and Wawa subprovinces.” (14 to 15 km deep in the crust; Easton, 2000; Percival, 1989). The nomenclature of Mehnert (1968) has been used in describing migmatites in this field guide. Regional Geology - Sibley Group The Mesoproterozoic, 1.4 Ga, Sibley Group crops out in a ~175 km wide by 400 km long ovoid under Lake Superior extending north to the south and west of Lake Nipigon with a minimum thickness of 950 m inferred from drill core (Rogala et al., 2007). The group is a predominantly flat-lying red-bed sequence which can be broken up into five lithological units; Pass Lake Formation, Rossport Formation, Kama Hill Formation, Outan Island Formation and Nipigon Bay Formation (Fig. 3). The following is mainly based on research conducted and published on by Becky Rogala and Riku Metsaranta: This field trip will cover the southern half of the Quetico Subprovince, from south of Beardmore to Nipigon (Fig. 1). As mentioned above, there is an asymmetric distribution in metamorphic grade, with a gradual progression from greenschist-facies, clastic metasedimentary rocks near the northern contact with the Wabigoon Subprovince; to lower amphibolite-facies schists and gneisses; through to upper amphibolitefacies migmatites and derived granitic rocks near the southern contact with the Wawa Subprovince near Nipigon. Thermal and pressure maximum occurs south of the center of the Quetico. The metamorphic character is of high-temperature/low-pressure (Abukuma-type) metamorphism, associated with the abundant intrusion of granitoid rocks and the regional distribution of migmatites derived from the metasedimentary rocks (Kamineni et al., 1988; Percival and McGrath, 1986; Percival, 1989; Williams, 1989). Peak metamorphic assemblages in the field trip area suggest conditions >650ºC and 5 kbar, corresponding to bathozones 4 to 5 The Pass Lake Formation consists of two members; the basal Loon Lake Member conglomerates and the overlying Fork Bay Member sandstones. The basal conglomerates are typically only a few metres thick with a maximum 15 m thickness in topographic lows within the basement rock. Conglomerates of the Loon Lake Member begin the large-scale, thinning-upward succession of the lower portion of the group and are dominantly composed of associated basement material (Rogala et al., 2007). They were deposited in channels eroded into the basement below braided streams in a semi-arid environment as highlighted by dolocrete horizons in floodplain sediment. Lower energy streams deposited trough cross-stratified sandstone. In places these materials were reworked into cobble-pebble beach deposits as a lacustrine system to the south expanded northward. With transgression fining- and thinning-upwards successions of sheet sandstone were deposited offshore from river mouths, while in areas with higher sediment supply deltaic forced regression occurred. With time the location of the lacustrine system in an area of internal drainage resulted in saline conditions developing. Alternating dolomitic red and light grey thin layers of the Channel Island member, Rossport Formation, attest to cyclic changes in organic productivity leading to more organic material in the bottom sediment reducing the oxidized iron. Higher hematitic clay contents in the red layers may be the - 46 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 3. Lithostratigraphy, depositional environments and climate during Sibley Group Sedimentation (from Rogala et al., 2007). result of sediment water influx leading to more rapid deposition and therefore less organic sediment. Tectonic instability brought this phase of sedimentation to an end by flooding the lacustrine system with sheet sandstone probably derived from the south because of north down-tilting of the terrain (Rogala et al., 2007). Strandline stromatolitic dolomite of the Middlebrun Bay Member was laid down as the lake shrunk. Subaerial exposure caused the development of karst topography, terra rosa and other types of soil horizons on the dolomite. Intrabasin mass flows composed of clasts from underlying Sibley lithologies commonly developed during this time period. With the end of widespread lacustrine conditions the water table remained close to the surface resulting in the precipitation of gypsum and carbonate in extensive mudflats of the Fire Hill Member, Rossport Formation. Saline ponds with gypsiferous stromatolites and teepee structures developed in lower areas. Higher mudflats were too dry for evaporates to form. This is the end of continuous sedimentation in the lower Sibley Group. A time gap of unknown extent separates the underlying playa system from overlying deltaic - 47 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 deposits of the Kama Hill and Outen Island Formations (Rogala et al., 2007). The Kama Hill Formation represents a ~50 m thick wave-rippled and hummocky cross-stratified sandstone and mudstone package which abruptly overlies the upper Fire Hill Member of the Rossport Formation (Rogala et al., 2007). Paleocurrent indicators suggest material comprising the Kama Hill Formation was derived from the south to southeast with the unit being thickest in the central portion and thinning toward the east (Cheadle, 1986). The Kama Hill Formation is dominated by horizontally laminated siltstone and fine-grained sandstones with interbeds of mudstone and ripple-laminated, fine-grained sandstone (Rogala et al., 2007). It represents prodelta and distal bar deposits. The Outan Island Formation consists of two members; the lower Lyon Member including coarsening- and thickening-upward sandstone successions overlain by siltstone and ripple laminated sandstone and the Hele Member consisting of a fining-and thinningupward sandstone succession overlain by mud-cracked siltstone (Rogala et al., 2007). The Lyon Member represents coarsening- and thickening-upwards, sandy distributary mouth bars overlain by the Hele Member fluvial system with extensive floodplains deposited in a non-arid climate (Fralick and Zaniewski, 2012; Ielpi et al., 2018). The coarsening upwards delta lobes are on the same scale as the modern Mississippi giving the impression that this was a substantive drainage system. A slight angular unconformity exists at the top of the Outan Island Formation with the overlying Nipigon Bay Formation. Four hundred and fifty meters of Nipigon Bay sandstones make up half of the Sibley Group. They were deposited sometime between 1.4Ga and 1.1 Ga. This Formation represents large Aeolian dunes developed in an arid setting. Franklin et al. (1980) originally inferred the Sibley Basin to be a result of subsidence caused from the ~1.1 Ga Midcontinent Rift system. This was called into question with age constraints on the Sibley Group, based on U-Pb geochronology of detrital zircons and stratigraphy, less than 1440 Ma for the entire Sibley Group (Rogala et al., 2007). The basal conglomerate of the Osler Group, erosively overlying the Nipigon Bay Formation, gives a lower age constraint of 1109 Ma (Davis and Sutcliffe, 1985) for the Nipigon Bay, whereas the other formations have a tighter lower limit of 1339±33 Ma derived from diagenetic Sr geochronology (Franklin et al., 1980). Cheadle (1986) noted intercalation of English Bay Complex rhyolites with Sibley sandstones, which together with geochronology, debunks the Sibley Group as being of MCR affinity and lends the notion that the succession was in fact 250-350 m.y. older than the MCR. The current model holds that the Sibley Group was deposited in a half graben-controlled basin (Rogala et al., 2007) inferred to be a product of large-scale thermal subsidence following the ~1550 Ma thermal plume event which produced the English Bay Complex (Hollings et al., 2004). Regional Geology - Midcontinent Rift Mesoproterozoic intrusive, volcanic and minor sedimentary rocks associated with the MCR collectively constitute the Keweenawan Supergroup. On the northern margin of the MCR, Keweenawan rocks include a variety of intrusive rocks and Osler Group volcanic rocks, which represent some of the earliest magmatism in the MCR. Ages range from ca. 1140 Ma (Heaman et al., 2007) to ages younger than the magnetic polarity reversal that occurred between 1105 and 1102 Ma (Davis and Green, 1997). The majority of mafic and ultramafic rocks in the Lake Nipigon and northern Lake Superior areas, including the Nipigon and Logan sills, appear to have been emplaced in a short, magnetically reversed, interval between ca. 1115 and 1100 Ma (Heaman et al., 2007). Emplacement of alkalic intrusions, such as the 1108 Ma Coldwell Complex (Heaman and Machado, 1992), and filling of much of the submerged part of the rift in Lake Superior, also occurred in this period. This was followed by a period of magnetically normal, waning mafic and felsic magmatism, between 1096 and 1085 Ma, that is preserved mainly along the Lake Superior shore by units such as the Crystal Lake (1099±1 Ma), Moss Lake (1095±2 Ma) and Blake (1095±2 Ma) gabbros, and a Pigeon River dyke near Arrow River (1093±3 Ma; Heaman et al., 2007). Hypabyssal Mafic Rocks Diabase sills, extending from the vicinity of Thunder Bay to east of Lake Nipigon, represent the northern remnants of the Midcontinent Rift, and have previously been referred to as the Logan sills (Stockwell et al., 1972). However, a geochemical difference has been noted between the sills to the north and south of the City of Thunder Bay (Hart, 2003; Hart et al., 2005). Hollings et al. (2007a) proposed that the - 48 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 term Logan Igneous Suite, which would fall within the Midcontinent Rift Intrusive Supersuite (Miller et al., 2002), should be applied to all the diabase sills in the area north of Lake Superior, with subdivision into the informal terms, Nipigon sills for the sills north of Thunder Bay, and Logan sills to the south. Nipigon Sills Nipigon sills are commonly massive, medium- to coarse-grained, olivine-tholeiitic diabase/gabbros (Sutcliffe, 1989; Hart and MacDonald, 2007). Nipigon sills are dominantly present throughout the Lake Nipigon area but have also been recognized in the Thunder Bay area (Hollings et al., 2007b). Nipigon sills are characterized by a massive, subophitic to ophitic, plagioclase and clinopyroxene texture with trace to 3% olivine and 1-2% modal magnetite (Hart et al., 2005). Nipigon sills display a reverse magnetic polarity and generally form thick, columnar jointed sheets. Sills commonly intrude Sibley group sedimentary rocks but also can be found in contact with Archean rocks of the Quetico subprovince and the Marmion and Winnipeg River terranes. Sills often intrude earlier emplaced ultramafic units of the Nipigon Embayment as well as the 1129.0 ± 2.3 Ma Pillar lake Volcanic rocks and the 1546.5 ± 3.9 Ma English Bay Complex (Heaman et al., 2007) providing evidence for their emplacement during the second main phase of magmatism (Hart and MacDonald, 2007). The shallow dipping Nipigon diabase sills are estimated to cover an area in excess of 20,000 km2 (Sutcliffe 1991) ranging in thickness from <5 m to >180 m (Hart and Macdonald, 2007). Stops Stop 1 – Glacier Lake Batholith Leucogranite UTM coordinates 0409911E 5445897N Large road cuts on both sides of Highway 11 provide excellent exposures of the Glacier Lake Batholith (GLB). At this location, it consists of white, massive to locally foliated, muscovite < biotite, medium- to coarsegrained granite. Localized pods of tourmaline-biotitemuscovite- potassium feldspar pegmatite occur within the granite. These pods may reach 1 m in diameter and locally contain tourmaline-quartz intergrowths. Numerous, curvate, fibrolite-muscovite+tourmaline veins up to 1 cm thick also occur in the host granite. Purple fluorite is exposed on a fractured outcrop face on the west side of the highway (Fig. 4). Field Trip Road Log Stop Locality Terrace Bay to Nipigon Lake Helen stops Intersection of Hwy.’s 11 &17 in the Town of Nipigon Take Hwy 11 north 1 Glacier Lake Batholith 2 Biotite leucogranite 3 Migmatite - roadside rest area Pull-off area 4 Pegmatitic granite 5 Pegmatitic granite 6 Pegmatites in migmatite Nipigon Stops Intersection of Hwy.’s 11 &17 in the Town of Nipigon Take Hwy 17 east 7 Stendlund Barite-Amethyst 8 Ruby Lake O1 Polygonal Diabase O2 Migmatites 9 Kama Hill 10 Unconformity at Gurney O3 Sunrise-Sunset Fluorite - 49 - km 105 km 0 17.1 10.8 7.4 5.9 5.2 4.85 4.6 0 6.8 (4.1 km into site) 18.4 21.4 39.9 �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 4. Fluorite on fracture surface in sericite- and fibrolitebearing Leucogranite (Stop 1). Stop 2 – Glacier Lake Batholith Leucogranite/ Migmatite UTM coordinates 0407163E 5440408N This is another example of the southern margin of the Glacier Lake Batholith where it is in contact with metasedimentary migmatite (Fig. 5). White, biotite-muscovite leucogranite contains rare, bluegreen apatite. Foliation is developed in feldspathic segregations and biotitic seams. Leucogranite dykes and migmatite locally exhibit a lit-par-lit structure. Tight to isoclinal folds have developed in the quartzbiotite-feldspar schist (Fig. 6). All rocks display boudinage and folding. Folded leucosome suggests an early and protracted deformation history (Fig. 7). Metasedimentary migmatite enclaves are common in the white, S-type pegmatitic granites along the highway. Note that less evolved S-type granite may contain biotite as the only mica. Narrow diabase Figure 5. Contact between folded metasedimentary migmatite and Leucogranite (Stop 2). Figure 6. Folded metasedimentary migmatite (Stop 2). dykes cut the country rocks on the western side of the highway. Figure 7. Folded metasedimentary schist, leucosome development (Stop 2). - 50 - incipient �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 3 – Migmatite UTM coordinates 0406986E 5437331N (Rest Area; Coffee) Migmatites are exposed on the shore and islands of Lake Helen and on a large, glacially streamlined, whaleback outcrop behind the highway rest stop/ pulloff area. There is quite a bit of variation in the relative amounts of leucosome and restite (Fig. 8). Boudinaged and pytgmatically folded leucosome pods and veins occur in a quartzo-feldspathic matrix with narrow, biotitic restite septa. Figure 9. Pegmatitic K-spar - quartz – biotite granite (Stop 4). 10). The megacrystic, foliated granite may represent an earlier, xenolith-bearing phase that was intruded by the massive granite while still relatively warm and plastic. Figure 8. Lit-par-lit migmatites with large metasedimentary inclusions (schollen) (Stop 3 area). Stop 4 – Pegmatitic Granite UTM coordinates - 0407272E 5435208N This outcrop shows a transition in magmatism from S-type (i.e. Glacier Lake Batholith) to I-type granitoids. Pink, coarse-grained, biotite granite here consists of pink, coarse-grained, perthitic K-feldspar, quartz, coarse-grained biotite, and brown, altered plagioclase (up to 4 cm long). These weakly peraluminous, pegmatitic, biotite granites are relatively primitive and are younger than the white, two mica leucogranites. Such rocks are probably of I-type origin and typically are metasedimentary enclave free. Bulk rock levels of rare-elements are very low: 128 ppm Rb, 2.63 ppm Cs, 1.8 ppm Nb, 0.39 ppm Ta (F. Breaks, OGS, personal communication, 2004). A K-feldspar-megacrystic, pink biotite granite (Fig. 9) with a shallowly eastdipping foliation and metasedimentary xenoliths is intruded by a more massive granite at this locality. The intrusive contact is embayed and scalloped, suggestive of co-mingling magmatic textures (Fig. Figure 10. Scalloped contact (arrow) between foliated biotite granite (top) and pegmatitic granite (bottom) (Stop 4). Stop 5 – Pegmatitic Granite UTM coordinates - 0407458E 5434906N A large, bare, whaleback outcrop on the east side of the highway consists of coarse-grained to pegmatitic, massive, homogeneous pink granite (i.e. the younger granitic unit at Stop 4; Fig. 11). Crystals or interstitial patches of quartz, biotite and locally sericitized K-feldspar average 2 to 3 cm in size (Fig. 12). Individual feldspar megacrysts may attain lengths of over 60 cm. Stop 6 – Pegmatite Dykes in Migmatite UTM coordinates 0407587E 5434690N Approximately 200 m south of the massive pink granite, white pegmatite dykes intrude fine-grained - 51 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 11. “Whaleback” outcrop of coarse-grained to pegmatitic granite; note large (>60 cm) feldspar megacryst (circled) (Stop 5). metasedimentary schist. This dark, fine-grained, feldspathic, biotite schist displays a well-developed foliation and minor folds, indicative of a long, protracted ductile deformation history, perhaps coeval with dyke emplacement. Note the sub-horizontal mineral lineation. Flattening of feldspars, producing small-scale, augen structures, is also indicative of hightemperature (>500º C) deformation. The boudinaged and annealed, sericite- and biotite-bearing dykes range from a few centimetres (lit-par-lit structured) to several metres in thickness. They are mineralogically and texturally interesting, containing both cordierite and quartz-tourmaline intergrowths with alkali (locally sericitized) feldspar. Garnet is conspicuous by its absence. Breaks et al. (2003) described potassic, biotite pegmatite that grades into a medium-grained, biotite granite near this stop (UTM: Easting 408339; Northing 5433016). The pegmatite contains coarse, euhedral Figure 13. Quartz-tourmaline intergrowth in pegmatite dyke (Stop 6). Figure 12. Coarse-grained to pegmatitic granite (Stop 5). K-spar; quartz; prismatic, medium- to coarse-grained, black tourmaline (schorl-dravite; Fig. 13); and fine grained, green and blue fluor-apatite (0.3 to 0.9 weight % MnO). A biotite-rich, metasomatic contact occurs between the biotite granite and diabase. Graphic, coarse-grained cordierite (<2 cm) and tourmaline occur in the granite near a diabase contact. Lunch Stop – Nipigon Marina UTM coordinates 0408076E 5429202N (Rest Area) Stop 7 – Stenlund Amethyst- Barite Occurrence UTM coordinates 0413918E 5429779N The property is underlain by flat-lying Sibley Group sedimentary rocks of the Lower Rossport Formation. Brick-red to orange muddy dolomite are interbedded with buff-coloured units. Small beige reduction spots occur in the red, hematite-rich dolomites. In contrast to this early reduction of hematite by flecks of organic matter late alteration (reduction of the ferric iron) occurs along the joints. Veins are exposed in two locations, 60 m and 170 m north of Highway 17. At the occurrence nearer the highway, 0.5 to 6.0 cm wide quartzbarite+/-amethyst veins occupy a parallel fracture set apparently controlled by a dominant set at 070°-085° SE. Vein breccias up to 20 cm wide and larger cavities and vugs appear to have developed preferentially in a sandier unit which locally overlies the red dolostone. Brecciated rock fragments are grey and green in some cases, possibly due to the vein alteration. Small, parallel quartz-barite filled fractures and veinlets, also striking 070°, occur 110 m north of the occurrence. Pieces of barite, baritic vein and breccia float are abundant in the vicinity. However, their source was not found. The - 52 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 flat relief and lack of outcrop necessitates extensive overburden stripping. Colorless, smoky and amethystine quartz and rare citrine occur with white and pink, bladed to massive barite (Fig. 14). Drusy quartz is commonly associated with barite seams. Lavender to deep purple amethyst crystals with pyramidal terminations are 0.2 to 15 mm across at their bases. Hematite has been reduced to an unspecified mineral. Figure 15. Extracted blocks of multicoloured and banded Ruby Lake marble (Stop 8). quarry (Hinz et al., 1994). Calcite, dolomite, epidote and opaque minerals were noted in thin section by Hinz et al. (1994) from the Nipigon River quarry. A number of chemical analyses for Nipigon River marble (wt.%) were also provided (Table 1). Figure 14. Drusy white quartz and amethyst with barite (Stop 7). Stop 8 – Ruby Lake Marble Quarry UTM coordinates 0414973E 5426071N Variegated, multicoloured, banded marble has been quarried here for landscaping stone (Fig. 15). Approximately 175 tonnes of marble were quarried and shipped in 1998 (D. MacAlpine, personal communication, 1999). The dimensions of the largest, transported block was 1.8 by 0.75 by 0.50 m (1.8 tonnes; ibid). Approximately 398 tonnes of marble were quarried and shipped in 1999 (D. MacAlpine, pers. comm., 2000). This marble consists of contact metamorphosed, Mesoproterozoic, Rossport Formation dolostone and other, calcareous sedimentary rocks in the contact metamorphic aureole of Midcontinent Rift-related Nipigon diabase sills. It has previously been termed Nipigon River marble and was quarried from 1883 to ca. 1910 at a site on the eastern side of the Nipigon River, approximately 6 km west of the Ruby Lake Shallow exploration trenches on the side of the road leading to the top of Ruby Mountain (i.e. top of the upper sill) have exposed copper-mineralized marble. Fine-grained, disseminated blebs of native copper (0.1 to 1.0 mm) occur along calcite-coated, hairline fractures parallel to bedding planes. Mineralized fractures are most easily recognized where secondary (supergene) malachite has formed. The top of the adjacent (middle?) sill is exposed farther along the road (optional stop 2). Similar, copper-mineralized, calcareous units have been noted near a diabase sill contact at Hughes Point, 2.5 km to the south-southwest by Schnieders et al. (1996). At this location, 2 to 5 cm wide calcite veins contain disseminated covellite (after chalcocite?) and malachite. The orientation of the veins are roughly parallel to joints developed in the adjacent sill. Grab samples have returned up to 3.065% Cu, nil Au and nil Ag (ibid). Franklin (1970) described copper-mineralized stromatolitic units near Disraeli Lake. A variety of copper minerals, including digenite, cuprite, covellite, Table 1. Chemical analyses from the Nipigon River Marble. From Hinz et al. (1994). Sample 89MCK-09 89MCK-10 SiO2 35.21 29.19 TiO2 0.20 0.10 Al2O3 8.20 5.20 Fe2O3 2.04 2.04 FeO 1.53 0.00 - 53 - MnO 0.05 0.03 MgO 23.56 27.63 CaO 26.74 35.32 Na2O 0.00 0.00 K2O 2.38 0.41 P2O5 0.09 0.07 �Proceedings of the 65th ILSG Annual Meeting - Part 2 chalcopyrite, native copper and malachite, were identified as open-space fillings in the vuggy host rock (ibid) suggested that the copper was introduced epigenetically from a gabbro to peridotite plug that intruded these Sibley Group rocks. The close spatial association of copper mineralized rocks with diabase sill contacts in the vicinity of Ruby Lake supports this theory. Alternatively these deposits are very similar to those in the Zambian copper belt where organic matter in the stromatolitic mats reduced copper travelling in groundwater forming large syngenetic ores. The marble is stromatolitic, though this is difficult to ascertain in most places due to the lack of pinnacles. The stromatolites are best seen on the tops of beds where they protrude, causing the contact to become wavy. Hints of the presence of stromatolites are visible, giving the impression that a large amount of the horizontal layering is stromatolitic S-mat (smooth mat that has a crinkly appearance). Non-crinkly, commonly lighter coloured layers interlaminated with the S-mat are storm layers of dolomitic silt washed in by wave activity. This sequence was deposited in a strandline proximal position in the playa, as denoted by the presence of teepee structures in this horizon at other locations and lacustrine deposits below the stromatolites and sub-aerial deposits above. This infers that lake size had stabilized during this interval, eliminating the large-scale fluctuations in shoreline positioning. Optional Stop 1 – Polygonal Jointed Diabase Sill UTM coordinates 0414746E 5426684 N This optional stop is accessed via a steep, rough trail ~500 m northwest from the turnoff to the Ruby Lake Marble quarry to the top of Ruby Mountain. Exposed here is an excellent exposure of the chilled, upper margin of a Nipigon diabase sill. The diabase is massive, homogeneous, fine- to medium-grained and locally feldspar-phyric. The large pavement outcrop displays polygonal cooling joints (seen in plan view) also referred to as “tortoise-shell” texture (Fig. 16). Figure 16. Polygonal jointed (“Tortoise-shell” texture) diabase (Optional Stop 1). disrupted (Fig. 17). There is a relatively high proportion of quartzo-feldspathic neosome matrix to the xenoliths, suggesting high(er) degrees of partial melting. Patches and dykelets of coarse-grained, biotite-quartz-feldspar neosome contain garnet, cordierite and large (>10 cm), euhedral green apatite crystals. Stop 9 – Kama Hill UTM coordinates 0425272E 5428131N Diabase-capped Kama Hill provides an excellent roadside exposure of the Rossport Formation (Fig. 18). The sequence is dominated by interlaminated shaley dolostone and dolomitic red shale with layer thicknesses ranging from millimeters to ten centimeters. There is some cyclicity in layer thickness variation up through the sequence, but it is not a strong trend. Some carbonate-dominated layers contain Optional Stop 2 –Migmatites UTM coordinates 0425006E 5429614N This optional stop displays schollen (raft)-structured migmatites in which folded and schlieric, mafic, paleosome xenoliths have been highly deformed and Figure 17. Schollen structure in migmatite (Optional Stop 2). - 54 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 18. View looking northwest at Kama Hill. Lower outcrop is Rossport Formation. Upper cliff is a Nipigon diabase sill (Stop 9). coarse- to medium-grained sand grains. Desiccation cracks are present, but not common. The central areas of some thicker more dolomitic (lighter-coloured) layers contain molds where gypsum crystals have weathered out. The alternation between red, more clay-rich layers, and gray, more dolomitic layers, represents climatic fluctuations over large time periods, as secular paleomagnetic trends measured in core of this facies indicates that the average meter of sediment took thousands of years to be deposited (Rogala, 2003). This compares well with modern playa systems and is reasonable as most of the sediment is dolomite formed from precipitation of HCO3-, Ca+2 and Mg+2 from solution in inflowing water. The lake was not totally drying up so there was always a standing body of saline water for the new fresh water to mix with. This lowering of salinity should have caused dissolution prior to evaporation over time causing more precipitation. See if you can see evidence of this. Ca ratio in the lake water and may have been the result of the precipitation of gypsum in the central lake. However, this is unlikely as higher salinities are expected in the shallower, marginal areas than the lake center, and thus, gypsum precipitation should be initiated in the shallows, producing shale-gypsumdolomite triplets. The lack of these means that gypsum precipitation was not necessary to increase the Mg/Ca ratio. The ambient ratio in the lake water itself must have been high enough for the precipitation of dolomite During rainy periods, water influx into the lake brought and deposited hematite-rich clays. In dry periods, evaporation from the internal drainage system resulted in lake contraction, hypersalinity and the precipitation of dolomite. This requires a high Mg/ Towards the top of the alternating layers (cyclic facies) thick sandstone sheets start appearing. These are the harbingers of tectonic rotation of the basin causing the basin to switch from northern drainage to southern drainage. Tension is recorded in sandstone Synsedimentary deformation manifests itself as small to large slump folds and brecciation of units in places. The sequence is intruded by diabase dykes which bake immediately adjacent sediments. Some evidence has been put forward that the diabase intruded watery sediment, but we have not seen clear indications of this. Structural controls on the emplacement of local sills were suggested by Antonellini and Cambray (1992). - 55 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 layers developing vertical joints along which they rotate as they extend. The sandstones also represent the dying phases of the lake. Overlying this succession is a chert-carbonate unit (Middlebrun Bay Member; Cheadle, 1986), which locally hosts stromatolites and hydrocarbons. It represents a shoreline microbialite similar to those present in some modern lakes. It is mostly composed of S-mat (crinkly layers) with extensive brecciation and silicification. The silicification is early, probably driven by rotting organic matter creating acidic conditions that favours carbonate replacement by silica. The brecciation is also quite early likely resulting from both dewatering of underlying muds having upward progress impeded by the impermeable microbialites and rainwater causing weakening of the layer through dissolution. A freshwater altered zone can be seen at the top of the dolostone. The siltstones and fine-grained sandstones immediately above the dolostone also contain stromatolites; one of the few places siliciclastic stromatolites can be seen in the rock record. As the lake shrank sub-aerial mudflats of the upper Rossport Formation were deposited on top of the dolostone. Pieces accessible in a pile of excavated talus display ripple marks, mudcracks and rare raindrop impressions. In many areas these siltstones and mudstones contain abundant gypsum, but it is only present in core as it readily weathers out of outcrops. Stop 10 – Unconformity at Gurney UTM coordinates 0440046 E 5419045 N The basal unconformity between the Sibley Group and underlying granitoids is exposed at this location (Fig. 19). A channel is visible, eroded into basement, and containing matrix-supported conglomerates overlain by cross-stratified sandstones. It represents episodes of subaerial debris flow activity interspersed with normal flash flood runoff. Archean granitic rocks are altered at the unconformity and likely represent a pre-Sibley weathered regolith. This weathered paleosol, noted by Gill (1926) and Moorhouse (1960), was locally described by Scott (1987). Friable, blotchy, red and green granite hosts quartz-carbonate veins between exfoliation blocks. Feldspars have been hematitized and/or destroyed; ferromagnesian minerals have been chloritized (ibid). Limited sampling of drill core from the Black Sturgeon Lake area cited by Scott (1987) suggests that this Figure 19. Unconformity between weathered Archean granite and Pass Lake Formation debris flows and sandstones, Highway 17 at Gurney (Stop 10). alteration may typically involve marked increases in Fe2O3, MgO, H2O) and decreases in Na2O, CaO and perhaps K2O. Paleomagnetic data suggest that this weathering was equatorial (G. Borradaille, unpublished data, 1999). As noted by Franklin (1978), Scott (1987) and Tanton (1948), a number of uranium occurrences are associated with altered Archean granitoids and overlying sedimentary rocks within the Sibley basin, prompting comparisons with the Athabasca basin in Saskatchewan. Favourable local parameters for supergene uranium deposits include: (i) uranium-enriched basement rocks (quartz monzonites, pegmatites); (ii) onlaps of basal Sibley sandstones on Archean paleotopographic “highs”; and (iii) Keweenawan(?) faults that extend to the basement (ibid). Optional Stop 3 – Sunrise-Midday Veins UTM coordinates 0454557E 5415513N The Sunrise vein is 18 m wide in the Highway 17 roadcut. It strikes approximately northeast and dips vertically with sharp contacts. Detailed mapping has not been successful in tracing this fluorite-bearing zone along strike. To the northeast, it is covered by alder swamp and glacial till. To the southwest, there is no outcrop along the strike of the vein. The Midday vein is about 5.8 m wide, strikes 74° and dips 68° northwest. It is located to the west of the Sunrise vein in the same roadcut. It extends into a swampy area to the west and pinches out 61 m east of the road-cut. It has a possible maximum strike length of 305 m. Both the Sunrise and Midday veins contain barite and fluorite, but very little amethyst. Brecciation in the - 56 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Midday vein is not as apparent as in the Sunset vein, although epidotization is very strong, giving the zones a greenish-yellow colour on fresh surfaces. The fluorite and barite occur as narrow veinlets along irregular fractures within these zones. Assay values are erratic due to the irregular nature of the mineralization. One section across the Sunrise vein assays up to 23.11% CaF2 over 3 m. References Antonellini, M.A. and Cambray, F.W. 1992. Relations between sill intrusions and bedding parallel extensional shear zones in the Mid-continent rift system of the Lake Superior region; Tectonophysics, v.212, p.33 1-349 Ayres, L.D. 1978. Metamorphism in the Superior Province of northwestern Ontario and its relationship to crustal development; in Metamorphism in the Canadian Shield, Geological Survey of Canada, Paper 78-10, p.25-36. Barrett, T.J. and Fralick, P.W., 1989. Turbidites and iron formation, Beardmore-Geraldton Ontario: application of a combined ramp/fan model to Archean clastic and chemical sedimentation. Sedimentology, v. 36, p. 221-234. Borradaile, G.J. 1982. Comparison of Archean structural styles in two belts of the Canadian Superior Province; Precambrian Research 19, p.179-189. Breaks, F.W., Selway, J.B. and Tindle, A.G., 2003. Fertile peraluminous granites and related rare-element mineralization in pegmatites, Superior Province, northwest and northeast Ontario: Operation Treasure Hunt; Ontario Geological Survey, Open File Report 6099, 179p. Carmichael, D.M., 1978. Metamorphic bathozones and bathograds: a measure of the depth of postmetamorphic uplift and erosion on the regional scale; American Journal of Science, 278, p.769-797. Cheadle, B.A.1986. Alluvial-playa sedimentation in the lower Keweenawan Sibley Group, Thunder Bay District, Ontario; Canadian Journal of Earth Sciences, v.23, p.527-542. Coates, M.E., 1968. Stevens–Kagiano Lake area; Ontario Department of Mines, Geological Report 68, p. Davis, D.W. and Sutcliffe, R.M., 1985. U-Pb ages from the Nipigon plate and northern Lake Superior; Geological Society of America Bulletin, v.96, p.1572-1579. Davis, D.W. and Green, J.C., 1997. Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences, 34: 476-488. Davis, D.W., Pezzutto, F., & Ojakangas, R.W., 1990. The age and provenance of metasedimentary rocks in the Quetico Subprovince, Ontario, from single zircon analyses: implications for Archean sedimentation and tectonics in the Superior Province; Earth and Planetary Science Letters, v.99, p.195-205. Devaney, J.R. and Williams, H.R., 1989. Evolution of an Archean subprovince boundary: a sedimentological and structural study of part of the Wabigoon-Quetico boundary in northern Ontario. Canadian Journal of Earth Sciences, v. 26, p. 1013-1026. Fralick, P. W., Wu, J., and Williams, H.R., 1992. Trench and slope basin deposits in an Archean metasedimentary belt, Superior Province, Canada. Canadian Journal of Earth Sciences, v. 29, p. 2551-2557. Fralick, P., Smyk, M. and Mailman, M. 2000. Geology and stratigraphy of the Mesoproterozoic Sibley Group; Institute on Lake Superior Geology, 46th Annual Meeting, Thunder Bay, Ontario, Proceedings Volume 46, part 2, Field Trip Guidebook, p.1-41. Fralick, P.W. and Zaniewski, K., 2012. Sedimentology of a wet pre-vegetation floodplain assemblage. Sedimentology, v. 59, p. 1030-1049. Hart, T.R. 2003. Keweenawan mafic and ultramafic intrusive rocks of the Lake Nipigon and Crystal Lake areas, northwestern Ontario; 49th Institute on Lake Superior Geology, Proceedings volume 49, Part 1, Programs and abstracts, p.21-22. Hart, T.R., MacDonald, C.A., Hollings, P., and Richardson, A., 2005. Proterozoic intrusive rocks of the Nipigon Embayment and Midcontinent Rift. In, T.O. Tormanen and T.T Alapieti, 10th International platinum Symposium Extended Abstracts, Geology Survey of Finland, 365-368. Hart, T.R. and MacDonald, C.A., 2007. Proterozoic and Archean Geology of the Nipigon Embayment: implications for emplacement of the Mesoproterozoic Nipigon diabase sills and mafic to ultramafic intrusions. Canadian Journal of Earth Sciences 44: 1021-1040. Heaman, L.M., Easton, M., Hart, T.R., Hollings, P., Macdonald, C.A., and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of earth Sciences 44: 1055-1086. Heaman, L.M. and Machado, N., 1992. Timing and origin of midcontinent rift alkaline magmatism, North America: evidence from the Coldwell Complex; Contributions to Mineralogy and Petrology, v. 110, p. 289-303. Hollings, P.N., Fralick, P.W., and Kissin S.A., 2004. Geochemistry and geodynamic implications of the Mesoproterozoic English Bay granite-rhyolite complex, northwestern Ontario. Canadian Journal of - 57 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 boundary in the De Courcey – Smiley Lakes area, northwestern Ontario; Canadian Journal of Earth Sciences, v.13, p.737-748. Earth Sciences, v. 44, p. 389-412. Hollings, P., Hart, T., Richardson, A., and MacDonald, C.A., 2007a. Geochemistry of the mid-Proterozoic intrusive rocks of the Nipigon Embayement, northwestern Ontario. Canadian Journal of Earth Sciences 44: 1087-1110. Hollings, P.N., Smyk, M.C., and Hart, T., 2007b. Geochemistry of Midcontinent Rift-related mafic dikes and sills near Thunder Bay: New insights into geographic distribution and the geochemical affinities of Nipigon and Logan sills and Pigeon River and other dikes. 53rd Institute on Lake Superior Geology, Annual Meeting, Proceedings volume 53, Part 1. Lutsen, Minnesota, May 2007, pp. 40–41. Ielpi, A., Fralick, P.W., Ventra, D., Ghinassi, M., Lebeau, L., Marconato, A., Meek, R., and Rainbird, R.H., 2018. Fluvial floodplains prior to greening of the continents: Stratigraphic record, geodynamic setting, and modern analogs. Sedimentary Geology, v. 372, p. 140-172. Mehnert, K.R., 1968. Migmatites and the origin of granitic rocks. Elsevier. 405p. Miller, J.D., Smyk, M.C., Severson, M.J., Lavigne, M.J., and Middleton, R.S. 2002. PGE occurrences in mafic intrusions around western Lake Superior, USA and Canada; 9th International Platinum Symposium, Field Trip Guidebook, 135p. Moorhouse, W.W. 1960. The Gunflint iron range in the vicinity of Port Arthur; Ontario Department of Mines, Annual Report, v.69, pt.7, p.1-40. Pan, Y., Fleet, M.E., and Williams, H.R., 1994. Granulite facies metamorphism in the Quetico subprovince, north of Manitouwadge, Ontario; Canadian Journal of Earth Sciences, v.31, p.1427-1439. Easton, R.M. 2000. Metamorphism of the Canadian Shield, Ontario, Canada I: The Superior Province; The Canadian Mineralogist, v.38, p.287-317. Pan, Y., Fleet, M.E., and Heaman, L.M., 1998. Thermotectonic evolution of an Archean accretionary complex: U–Pb geochronological constraints on granulites from the Quetico subprovince, Ontario, Canada; Precambrian Research, v.92, p.117-128. Franklin, J.M. 1970. Metallogeny of the Proterozoic rocks of the Thunder Bay District, Ontario; unpublished Ph.D. thesis, The University of Western Ontario, London, 304p Percival, J.A and Stern, R.A., 1984. Geological synthesis in the western Superior Province, Ontario; in Current Research, Part A, Geological Survey of Canada, Paper 84-1A, p.397-407. Franklin, J.M. 1978. Uranium mineralization in the Nipigon area, Thunder Bay District, Ontario; in Current Research, Part A, Geological Survey of Canada, Paper 78-lA, p.275-282. Percival, J.A and Sullivan, R.W., 1988. Age constraints on the evolution of the Quetico belt, Superior Province, Ontario; in Radiogenic Age and Isotopic Studies: Report 2, Geological Survey of Canada, Paper 88-2, p.97-108. Franklin, J.M., McIlwaine, W., Poulsen, K., and Wanless, R. 1980 Stratigraphy and depositional setting of the Sibley Group, Thunder Bay District, Ontario, Canada. Canadian Journal of Earth Sciences, v. 17, p. 633-650. Geller, P., 2012. A review of the western Superior Lithoprobe Line 3. Unpublished Honours Thesis, Lakehead University, 45 pp. Gill, J.E. 1926. Gunflint iron-bearing formation, Ontario; Geological Survey of Canada, Summary Report, 27C, p.28c-88c. Hinz, P., Landry, R.M., and Gerow, M.C. 1994. Dimension stone occurrences and deposits in northwestern Ontario; Ontario Geological Survey, Open File Report 5890, 191p. Kamineni, D.C., Stone, D., and Johnston, P.J. 1988. Metamorphism of Quetico sedimentary rocks near Atikokan, Ontario; in Program with abstracts, Geological Association of Canada-Mineralogical Association of Canada-Canadian Society of Petroleum Geologists Annual Meeting, v.13, p.A63. Kehlenbeck, M.M., 1976. Nature of the Quetico–Wabigoon Percival, J.A., 1989. A regional perspective of the Quetico metasedimentary belt, Superior Province, Canada; Canadian Journal of Earth Sciences, v.26, p.677693. Percival, J.A. and McGrath, P.H., 1986. Deep crustal structure and tectonic history of the northern Kapuskasing uplift of Ontario: an integrated petrological-geophysical study; Tectonics, v.5, p.553-572. Percival, J.A. and McGrath, P.H., 1986. Deep crustal structure and tectonic history of the northern Kapuskasing uplift of Ontario: an integrated petrological-geophysical study; Tectonics, v.5, p.553-572. Percival, J.A., Stern, R.A., and Digel, M.R., 1985. Regional geological synthesis of western Superior Province, Ontario; in Current Research, Part A, Geological Survey of Canada, Paper 85-1A, p.385-397. Pirie, J. and Mackasey, W.O., 1978. Preliminary examination of regional metamorphism in parts of the Quetico metasedimentary belt, Superior Province, Ontario; Geological Survey of Canada, Paper 78-10, p.37-48. Rogala, B., 2003. The Sibley Group: A lithostratigraphic, - 58 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 geochemical, and paleomagnetic study. Unpublished MSc thesis, Lakehead University, Thunder Bay, 254 pp. Rogala, B., Fralick, P.W., Heaman, L.M., and Metsaranta, R., 2007. Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada. Canadian Journal of Earth Science 44, 1131-1149. Sawyer, E.W., 1983. The structural history of a part of the Archean Quetico metasedimentary belt, Superior Province, Canada; Precambrian Research, v.22, p.271-294. Schnieders, B.R., Smyk, M.C., Speed, A.A., and McKay, D.B. 1996. Mineral occurrences in the NipigonMarathon area, Volumes 1 and 2, Ontario Geological Survey, Open File Reports 5951, 912 p. Scott, J.F. 1987. Uranium occurrences of the Thunder BayNipigon-Marathon area; Ontario Geological Survey, Open File Report 5634, l58p. Seemayer, B.E., 1992. Variations in metamorphic grade in metapelites in transects across the Quetico Subprovince north of Thunder Bay, Ontario; unpublished M.Sc. thesis, Lakehead University, Thunder Bay, 163p. Smyk, M.C. and Kissin, S.A. 2005. Geology and rare element pegmatites of the Quetico Subprovince near Nipigon, 51st Institute on Lake Superior Geology, Proceedings volume 51, pt.2d,14p. Stockwell, C.H., McGlynn, J.C., Emslie, R F., Sanford, B.V., Norris, A.W., Donaldson, J.A., Fahrig, W.F., and Currie K L. 1972. Geology of the Canadian Shield, in Geology and Economic Minerals of Canada, edited by R.J.W. Douglas, Geological Survey of Canada, Economic Geology Report 1, 838 p. Sutcliffe, R.H. 1989. Mineral variation in Proterozoic diabase sills and dykes at Lake Nipigon, Ontario; Canadian Mineralogist, v.27, p.67-79 Sutcliffe, R.H., 1991. Proterozoic geology of the Lake Superior area. In: Geology of Ontario. Edited by P.C. Thurston, H.R. Williams, R.H. Sutcliffe, and G.M. Stott. Ontario Geological Survey, Special Vol. 4, Part 1, pp. 405–484. Tabor, J.R., Hudleston, P.J., and Mcloughlin, J., 1989. Metamorphism of the Quetico supracrustals north of the Vermillion granitic complex, northern Minnesota; Geological Association of Canada – Mineralogical Association of Canada, Program with Abstracts, v.14, p.A38. Tanton, T.L. 1948. Radioactive nodules in sediments of the Sibley series, Nipigon, Ontario; Transactions of the Royal Society of Canada, v. XLII, series III, section IV, p.69-75. Williams, H.R., 1989. Geological studies in the Wabigoon, Quetico, and Abitibi–Wawa subprovinces, Superior Province of Ontario, with emphasis on the structural development of the Beardmore–Geraldton belt; Ontario Geological Survey, Open File Report 5724p. Williams, H.R., 1991. Quetico subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, v.1, p.383-403. Williams, H.R. and Breaks, F.W., 1989. Geological studies in the Manitouwadge–Hornepayne area; Summary of Field Work and Other Activities 1989, Ontario Geological Survey, Miscellaneous Paper 146, p.7991. Williams, H.R. and Breaks, F.W., 1990. Geological studies in the Manitouwadge–Hornepayne area; Summary of Field Work and Other Activities 1990, Ontario Geological Survey, Miscellaneous Paper 151, p.4751. Zaleski, E. and van Breemen, O., 1997. Age constraints on plutonism, metamorphism and deformation across the Wawa–Quetico subprovince boundary near the Manitouwadge greenstone belt, northeastern Ontario; Institute on Lake Superior Geology, Program with Abstracts, v.43, p.67-68. Zaleski, E., van Breemen, O., and Peterson, V.L., 1999. Geological evolution of the Manitouwadge greenstone belt and Wawa–Quetico subprovince boundary, Superior Province, Ontario, constrained by U–Pb zircon dates of supracrustal and plutonic rocks; Canadian Journal of Earth Sciences, v.36, p.945-966. - 59 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 5 - A stratigraphic transect across the Northern flank of the Midcontinent Rift near Rossport Pete Hollings and Philip Fralick Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada Safety As this trip will be taking place on Lake Superior it will be weather dependent and could be cancelled or curtailed at very short notice. Please exercise caution when getting in and out of the boats as the outcrops are often extremely slippery. Life jackets must be worn in the boats at all times. It will probably be very cold out on the lake so please dress warmly. Introduction This field guide has been updated from Hollings and Fralick (2005) that was published as part of the 51st ILSG meeting in Nipigon, Ontario. We have updated the regional geology to include some more recent work and modified some stop descriptions to reflect changing lake levels. Regional geology Archean granites, outcropping along the shoreline near Rossport, are unconformably overlain by strata of the Gunflint Formation (Fig. 1). These sediments were deposited on a south facing shelf at approximately 1878 Ma (Fralick et al., 2002). The Formation consists of a Lower Member composed of basal stromatolitic bioherms overlain by ankeritic, interclastic grainstones. A regressive, karstified surface caps the northern portion of this assemblage (Fralick and Barrett, 1995) and is succeeded by the Upper Member. It begins with a repetition of the underlying lithologies to which, higher in the succession, are added carbonaceous shales, tuffs and rarely mafic volcanic rocks. These chemical and fine-grained siliciclastic sediments record parasequence development on a storm-dominated shelf (Pufahl and Fralick, 2004) forming the relatively stable portion of a back-arc basin (Kissin and Fralick, 1994; Hemming et al., 1995) prior to compressive northward thrusting of the arc at approximately 1860 to 1835 Ma. As the compression of the Penokean Orogeny waned the sea again transgressed over the area depositing Rove Formation black, carbonaceous shales gradually transitioning upward into turbidites (Morey, 1967). The turbitite fan fed off of a delta prograding to the SSE, with sediment sourced from the rising TransHudson Mountains (Maric and Fralick, 2005). This depositional cycle occurred at 1832 Ma (Kissin et al., 2003; Addison et al., 2005). The lower portion of the Gunflint Formation in the Rossport area is poorly exposed. Lithologies present in the limited outcrop of the Lower Gunflint are similar to those in the succession comprising the thin, basal Kakabeka Conglomerate and overlying interclastic grainstones present in exposures to the west near Thunder Bay. Good exposure of the Upper Gunflint exists on Quarry Island and consists of possible basaltic flow rocks with associated stromatolites, overlain by a succession of medium- to coarse-grained, graded, sandstone beds. The geochemistry of the sandstones is similar to Archean rocks to the north indicating probable derivation from this source. Black shales, lithically correlative with the Rove Formation, outcrop on an island approximately 5 km to the west. The shales do not overlie the turbidite succession on Quarry Island where arenites of the Sibley Group disconformably rest on an erosion surface at the top of the Gunflint sandstones. The basal unit of the Sibley Group is the Pass Lake Formation. It is composed of the conglomeratic Loon Lake Member and the overlying sandstones of the Fork Bay Member (Cheadle, 1986). Where the basal conglomerates are present they either represent: 1) large channel fills cutting down into sandstones to shales with abundant caliche zones, or; 2) more laterally continuous conglomerates interbedded and overlain by parallel laminated, medium-grained sandstones. The former represent channel gravels in braided fluvial systems and the latter coarse-grained strandline deposits. The overlying Fork Bay sandstones likewise - 60 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 A B Figure 1b Osler Group 1108 Ma Nipigon diabase Stop 8 Sibley Group Lake Superior Animikie Group Figure 1c 1105 Ma Lake Superior 200 km C 15 km Rossport N 48°30’ Archean rocks 48°30’ Stop 7 Keweenawan intrusive rocks Quarry Island Stop 6 Simpson Island Stop 5 Stop 4 Stop 3 Vein Island Osler Group volcanics Channel Island Keweenawan sediments Sibley Group sediments Gunflint Formation Archean basement Stop 2 Stop 1 Wilson Island Copper Island 87°30’ 87°45’ 48°45’ 1 km 48°45’ Figure. 1. Map showing the location of the field trip area. B) Regional geology map showing the extent of the Osler Volcanic Group. Age data from Davis and Sutcliffe (1985) and Davis and Green (1997). Modified after Sutcliffe (1986). C) Geological map of the Osler Volcanic Group showing sample locations. Modified after Giguerre (1975). record both braided fluvial deposition and subaqueous, strand proximal sand-sheet development. In addition to upward thinning and fining successions developed during transgressive systems tract formation, other sandstone assemblages thicken and coarsen upwards representing progradational, delta lobe outbuilding. The delta prograded into a lacustrine setting that isotopes (C, O, S and Sr) indicate became more saline with time (Metsaranta, 2006). This is consistent with Cheadle’s (1986) findings and those of Rogala et al. (2007) based on regional paleogeography. The increasing salinity of the water resulted in dolomite, minor gypsum and rarer barite and celestite precipitation mixed with mud deposition. A red and green banding developed in this assemblage due to periodic anoxia of the bottom sediments. The final desiccation of the lacustrine basin is recorded by the development of strandline microbial bioherms (stromatolites) which are overlain by either a terra rosa (red, wind-delivered soil) or subareal, intraformational, mass-flow conglomerates. This is succeeded upwards by mudstones with abundant gypsum nodules representing mudflats formed in an arid climatic setting where hypersaline groundwaters precipitated gypsum. Together all these fine-grained sediments comprise the Rossport Formation. It is overlain by the Kama Hill Formation; a coarsening upwards deltaic succession recording flooding of the basin and development of a more humid climate. The age of the Pass Lake and Rossport Formations can be bracketed between laser ablation MS youngest ages on detrital zircons of 1600 Ma and a Rb-Sr isochron of 1339 Ma (Franklin, 1978). The youngest detrital zircons in the deltaic succession of the Outan Island Formation is 1443±31 Ma. The Sibley Group is very well exposed along the shorelines of the islands off Rossport. The basal disconformity can be seen about two thirds of the way up the cliff face on the western side of Quarry Island where it overlies graded sandstone beds of the Gunflint Formation. Blocks of medium-grained sandstone were extracted from the cliff face on the island for use in the construction of buildings in Thunder Bay. These - 61 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 V N Portage Lake IV Volcanics Upper Suite 1100 1105 R R 1110 Portage lake Volcanics IV Kallander Creek Volcanics Siemens Creek Volcanics Mamainse Point Michipicoten Island Michipicoten Island Formation Group 7 V (Group 8) III Group 6 IV IV Osler Group Age (Ma) 1095 Copper Harbour Conglmerate Isle Royale Copper Harbour Conglomerate 1090 The oldest rift-related rocks on which U-Pb age determinations have been performed lie along the northwestern portion of the rift. These include, from NE to SW, the alkaline intrusive rocks of the Coldwell Complex (1108±1 Ma, Heaman and Machado, 1992), the lower Osler Group volcanic rocks (1108+4/-2 to 1105±2 Ma, Davis and Sutcliffe, 1985; Davis and Green, 1997), the Logan Sills (1109+4/-2 Ma, Davis and Sutcliffe, 1985), the lower portion of the North Shore Volcanics (1108±2 Ma, Davis and Green, 1997), the Isle Royale Black Bay Peninsula Lake Nipigon Upper Michigan NW Wisconsin 1085 Formation, underlies the same disconformity. This highlights the fact that approximately 600 m of erosive downcutting occurred in the Rossport area before the basal Osler was deposited. II I Bessemer Quartzite Lower Suite I Simpson Isl Cgl Nipigon Sills Schroeder Basalts V Beaver Bay Complex Mostly basalt units Duluth Complex Great Conglomerate and Group 5 Central Suite III NE Minnesota SW limb Groups 3,4 Group 2 Group 1 III II I IV IV North Shore Volcanic Group sandstones are medium- and large-scale planar crossstratified and may represent a sandflat composed of transverse bars in a braided stream or small barchan, eolian dunes. Rare pebbles indicate the former may be the case but this evidence is not conclusive. Channel and Copper Islands contain excellent exposures of the lacustrine rocks with outbuilding of channel systems along the paleolake margins. One of the best outcrops of the strandline stromatolitic carbonates occurs on Channel Island and will be visited during this field trip. On Copper Island the Rossport Formation is overlain disconformably by pebbly, fluvial conglomerates of the basal Osler. Thirty kilometers to the west the uppermost unit of the Sibley, the Nipigon Bay Ely’s Peak Basalts I, II, III Nopeming sandstone Archean Basement Figure 2. Schematic correlations of volcanic rocks of the Midcontinent Rift based on the stratigraphic position of distinctive basalt sequences, magnetic polarity and absolute age where possible. Modified after Nicholson et al. (1997). Dashed lines in Upper Michigan section separate lower and upper members of Kallander Creek and Siemens Creek volcanics. Left hand column shows magnetic polarity. Roman numerals I-IV refer to five distinctive laterally extensive basalt compositions identified on the south shore of western Lake Superior. Where equivalent basalt compositions occur in other stratigraphic successions, the appropriate Roman numeral is noted (see Nicholson et al., 1997 for data sources). Shaded regions represent intervals in which contacts are covered or obscured by plutonic rocks. - 62 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Swamper Lake Gabbro and Nathan’s Series intrusive rocks (1107 Ma, Paces and Miller, 1993) and the lower portion of the Powder Mill Group (1107±2 Ma, Davis and Green, 1997). These older units outcrop on the rift flanks where erosion has removed the younger rift sequence or, in the case of the Coldwell and Logan Sills, are intruded into older rocks immediately north of the rift. More recently a number of older sill complexes have also been identified in the vicinity of the Nipigon Embayment (Heaman et al., 2007). The Osler Volcanic Group comprise a ~3km thick sequence (Cannon et al., 1989) lying unconformably above Sibley Group metasedimentary rocks (Fig. 2). The volcanic sequence is overlain and intruded by the St. Ignace Island Volcanic-Plutonic Complex an intercalated sequence of basaltic rocks and rhyolitic flows (Sutcliffe and Smith, 1988). Detailed descriptions of the Osler Group have been provided by McIlwaine and Wallace (1976), Lightfoot et al. (1991) and Keays and Lightfoot (2015). Generally the mafic flows of the Osler Group consist of massive to amygdaloidal flows, with locally developed ropey tops and pahoehoe textures (Sutcliffe and Smith, 1988). The flows range in thickness from 5cm to 30m (Lightfoot et al., 1991) with a regional dip of ~6-15° S (Giguerre, 1975; Lightfoot et al. 1991; Hollings et al., 2007). The majority of the exposed section is magnetically reversed with only the upper 100m displaying a normal polarity (Halls, 1974). Recent work by Swanson-Hysell et al., (2014) has shown that there is a progressive change in the paleomagnetic sequence of the Osler volcanic rocks that is consistent with a ca. 25° of latitudinal motion of Laurentia. The contact between the two units is marked by the presence of the Puff Island conglomerate and a discordance between the basalt flows above and below the contact. This has been interpreted as representing a significant break in the eruption history. A felsic porphyry near the base of the Osler Group has yielded an age of 1107.5+4/-2 Ma (Davis and Sutcliffe, 1985) whereas zircons from the Agate Point rhyolite towards the top of the reversely magnetized sequence have yielded an age of 1105±2 Ma (Fig. 1b; Davis and Green, 1997). Within the Osler Group interflow sediments are typically thin and of limited extent. Field descriptions of the sedimentary successions appear in Giguere (1975) and McIlwaine and Wallace (1976). They show that there are two main zones of sedimentary rocks within the Osler Group. One occurs near the base of the volcanic pile. The other is present approximately 2700 meters higher in the succession marking the paleomagnetic reversal. Lightfoot et al. (1991) in a study of the Osler Volcanic Group exposed along the shores of the Black Bay Peninsula to the west of the field trip location proposed that the major and trace element geochemical data could be used to subdivide the flows into an Upper, Central and Lower Suite although the boundaries between the suites were not clear cut. While the geochemical compositions of the Central (750900m) and Upper suites (1900-3000m) overlap their Lower Suite (0-750m) is distinguished by elevated Mg numbers (0.55-0.7 versus 0.3-0.6), lower Al2O3 (8-12 wt% versus 13-17wt%), lower Th/Nb ratios (0.090.70 versus 0.3-0.6) but higher Gd/Ybn ratios (3.5-4.5 versus 1.6-2.6). Nicholson et al. (1997) concluded that there were five geochemically distinct flood-basalt compositions within the Mid-continent rift that are common to most sections and appear in approximately the same stratigraphic order (Fig. 2) They recognized a lower suite in the Siemens Creek Volcanics (Basalt Type 1; Fig. 2), which they suggest is analogous to the Lower Suite of Lightfoot et al. (1991). In the United States this unit is less than 100m thick whereas the Lower Suite of Lightfoot et al. (1991) is ~750m thick. They report a narrow range of εNd(1100) values for the Siemens Creek Volcanics of -0.7 to +0.7. Nicholson et al. (1997) further suggest that there is a broadly recognizable suite of basalts above this (Basalt type II) which includes the upper Siemens Creek Volcanics and in the upper part of the Grand Portage lavas (Fig. 2). The suite is characterized by slight negative Nb anomalies and a range of εNd(1100) values of -1.4 to -6.9. They suggest that this may be analogous to the most primitive members of the Central Suite of Lightfoot et al. (1991). The volcanic flows of the Osler Group on Wilson Island are all basalts or basaltic andesites (SiO2 = 4756 wt%; MgO = 5-16 wt%; Hollings et al., 2007). The basalts are characterized by LREE enrichment (La/Smn = 1.5-3.9) in conjunction with moderately fractionated HREE (Gd/Ybn = 1.5-3.7) and slight positive to moderately negative Nb anomalies (Nb/ Nb* = 0.56-1.13; Hollings et al., 2007; Fig. 3). Major and trace element data show trends of increasing SiO2 and decreasing MgO and display strong positive correlations between La/Smn, Th/La and Th/Nb with height (Fig. 4). This correlation is most pronounced - 63 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 in conjunction with LREE enrichment and strongly fractionated HREE are comparable to modern OIB, albeit at lower absolute abundances (Fig. 3). When compared to other flood basalt sequences the more primitive basalts from this study closely resemble basalts from the Parana-Etendeka flood basalt sequence (Fig. 4; Gibson et al., 2000). The εNd data from the most primitive members of the Osler Group is consistent with an enriched mantle plume rather than a contaminated depleted mantle source, given the lack of trace element evidence for contamination in these samples. Depleted mantle at 1100 Ma would have had a positive εNd perhaps as high as +6 whereas an enriched plume source would have εNd ~0 (Nicholson and Shirey, 1990; Shirey et al., 1994). Up sequence the basalts are characterized by higher SiO2, Th and La/Smn abundances in conjunction with increasingly negative Nb and Ti anomalies and εNd(1106) values of -4 to -5. This is consistent with contamination of these basalts by an older lithospheric component characterized by pronounced LREE enrichment, high Th abundances but generally unfractionated HREE (Hollings et al., 2007). 100 10 Hawaiian OIB Deccan Traps CFB Parana-Etendeka CFB 1 Th Nb La Ce Pr Nd Zr Hf Sm Eu Ti Gd Tb Dy Y Ho Er Tm Yb Lu Al V Sc Rock/Primitive Mantle 100 10 Type 1 Lower Suite 1 100 Th Nb La Ce Pr Nd Zr Hf Sm Eu Ti Gd Tb Dy Y Ho Er Tm Yb Lu Al V Sc 10 Type 2 The sedimentary successions near the base of the Osler Group constitute the Simpson Island Formation Figure 3. Comparison of primitive mantle normalized and have recently been described in detail by Hollings plots from the Osler Group with A) Phanerozoic OIB and et al. (2007). They are composed of a Lower Member Continental Flood Basalts (CFB) and B & C) the Lower and dominated by trough cross-stratified, medium-grained Central suites of Lightfoot et al. (1991). From Hollings et al. sandstones directly overlying basement and an Upper (2005). Member with a greater variety of siliclastic units. The above 400m with samples from the base of the Lower Member sits on an irregular, erosional surface stratigraphy displaying more or less constant values of cut into the underlying quartz arenites of the Nipigon these ratios (Fig. 4). Measured 143Nd/144Nd ratios for Bay Formation, Sibley Group. A massive pebble-cobble the seven Osler basalts analysed range from 0.511857- conglomerate overlies the unconformable surface 0.512286 with εNd(t=1106Ma) of +0.3 to -5.3 (Hollings et and is in turn overlain by decameter-scale layers of al., 2007). The high incompatible element abundances, coarse-grained and pebbly sandstone (Fig. 5, Section 1 900 800 Central Suite Th Nb La Ce Pr Nd Zr Hf Sm Eu Ti Gd Tb Dy Y Ho Er Tm Yb Lu Al V Sc SiO2 MgO Fe2O3 Th La/Smn Gd/Ybn Th/La Th/Nb εNd 700 600 500 400 300 200 100 0 40 50 60 5 10 15 10 15 1 2 3 4 1 2 3 4 1 2 3 4 0.05 0.10 0.15 0.20 0.10 0.15 0.20 -6 -4 -2 0 Figure 4. Geochemical stratigraphy of the Osler Group on Vein and Wilson Islands. The stratigraphic position of the samples has been calculated assuming a dip of 10° parallel to the section. From Hollings et al. (2005). - 64 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 V V V V V V V V V 6 No Longitudinal 24 Complex V V V V V Major 12 V V V V 6 V V V V V V 18 V 3 0 Section 5. V V No O/C V V V V V Nipigon Bay V V V Ripples Pebbles Hummocky Cross-Strat. V V V V V No O/C ? m. m V V V V V Complex No O/C ? m. Sandy Sheetfloods 6 Sheetfloods and Debris flows 3 V V V V V V V V V V Conglomerate Pebbly V. Coarse Medium V. Fine Siltstone Shale Sst. V V V V V V 7 15 V V V 8 m. No O/C 12 21 18 15 12 Bar V 0 Massflow 3 m. No O/C V V V V V V V V V V V V V 5 6 V V V V V V 15 V V V V V V V 10 km Older Units Section Locations Section 7. V V V V V V V V V V Sand Channels with Small Gravel Longitudinal Bars Stacked Channels 9 Sheetflood Sands with Channels 6 Small Stacked Channels Distal 6 V 4 2 1 3 VV V Sedimentary Rocks Igneous Rocks 1 27 Sandy Channels to Distributary Mouth Bar 9 V m Gravel Channels 26 V V V V V Osler Group 24 29 23 V V Section 6. 32 Longitudinal No O/C Trough Cross-Strat. Small Irregular Lenses Paleocurrent Direction V V V V V V V V V V V V V V V V V V V V V V V V V V Bar 0 0 Complex Rhyolite V V V V =318o 15 9 Longitudinal Parallel Lamination V V V V Bar Tail Sand Sheet Longitudinal Bar Nipigon Bay V Sandy Channel 3 Bar Sandy Channel V Section 4. m Basalt INTERFLOW SEDIMENTS m V Channel O/C =268o V 9 3 0 V l e g e n d 18 No 6 V V Sst. V 9 V Major Sandy Channel 0 Bar V m V 3 O/C 15 Section 1. V V No O/C V n = 38 V V 21 o Section 3. m V = 265 27 Section 2. Conglomerate Pebbly V. Coarse Medium V. Fine Siltstone Shale N m V SIMPSON ISLAND FORMATION ( Basal Sediments ) 3 0 Massflow Sheetflood Sands with Channels Figure 5. Sections of sedimentary rocks in the Osler Group. Sections 1 through 4 are the basal sedimentary succession of the Lower Member, Simpson Island Formation, at different locations (see inset map). Sections 5 and 6 are of the Upper Member, Simpson Island Formation, interlayered with basal basalt flows. Section 7 is the sedimentary assemblage near the top of the Osler Group on Puff Island. From Hollings et al. (2005). - 65 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 1). Sandstones are parallel laminated, commonly have cross-stratified tops, more rarely contain pebbly transverse ribs and chute and pool-like structures. The central portion of the succession is composed of trough cross-stratified, medium-grained sandstone organized into a stacked assemblage of lenses. Pebble stringers and pebbly sandstones commonly occur on the deeper portions of curving set boundaries (Fig. 5, Section 2). Massive pebble-cobble conglomerates sharply overly the sandstone succession (Fig. 5, Sections 2 and 4). The conglomerates contain trough cross-stratified, mediumgrained sandstone lenses; decameter- to meter-scale wedges of planar cross-stratified sandstone and are interbedded with assemblages of trough cross-stratified sandstones up to one meter thick. Another assemblage of trough cross-stratified sandstone, similar to the one in the central portion of the succession, caps the basal sedimentary assemblage (Fig. 5, Sections 3 and 4). Clast lithologies in the pebble-cobble population are dominated by quartz, chert, various types of volcanic rocks, red siltstone, metamorphosed granite and at the east end of the outcrop belt, on Copper Island, a higher proportion of unmetamorphosed red granite. Current indicators show flow was to the west, averaging 265°. A sedimentary assemblage also occurs near the upper limit of outcrop of the Osler Formation, at the top of the magnetically reversed interval (Fig. 5, Section 7). These interflow sediments are located on Puff Island and overly a felsic porphyry with a U-Pb age of 1105 Ma (Davis and Green, 1997). They contain: sharp sided assemblages of laterally continuous, pebbly, coarsegrained sandstone beds with caliche horizons which are scoured into by small lenses of conglomerate; large scours filled with trough cross-sets over a meter thick; stacked assemblages of irregular lenses filled by trough cross-stratified, coarse-grained, pebbly sandstone; and, poorly sorted, disorganized, massive boulder-cobble conglomerate. Clasts are all volcanic, ranging from quartz-feldspar porphyries to mafic compositions. Paleocurrents on large-scale sedimentary structures consistently show flow to the southeast. The Simpson Island Formation is composed of a laterally continuous sedimentary succession up to 25 meters thick and discontinuous sedimentary units up to 30 meters thick interlayered with the basal basalt flows. The lowest sedimentary beds fill channelways cut into the underlying sandstones of the Nipigon Bay Formation. The channel fills and overlying sedimentary assemblage represent a braided stream system, similar to the South Saskatchawan model (Miall, 1978), where dunes composed of coarse-grained sand migrated down the channels and gravelly longitudinal bars with chute channels and bar edge sand wedges form the higher relief areas (Fig. 5). The fluvial interpretation is consistent with Tanton (1931) and McIlwaine and Wallace (1976). Clast lithologies indicate debris was mainly derived from erosion of local lithologies. Stops The trip will depart from the public dock at Rossport and will undertake a traverse through the stratigraphy of the Midcontinent Rift, starting with the youngest rocks of the Osler Volcanic Group, proceeding through the Sibley and Gunflint Formations and finishing with a look at the granites of the Archean basement (Fig. 1). In order to make the most of the calmer weather typical of early mornings we will first travel for approximately 30 minutes to the most southerly outcrop on Wilson Island. Stop 1 – Osler volcanics, Wilson Island UTM coordinates 0462794E 5402095N Flows of the Osler Group on Wilson Island are typically >1m thick, frequently amygdaloidal towards the top and bottom of the flow with a massive core and rarely displayed a pahoehoe texture on the flow surface. The basalts are characterized by clinopyroxene and Figure 6. Well-developed vesicle column in basaltic flows, Stop 1 on Wilson Island. - 66 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 8. Pahoehoe texture at Stop 2, Wilson Island. Figure 7. Approximately 2m thick mafic flow cut by sediment filled cooling crack. Stop 1, Wilson Island. plagioclase phenocrysts in a groundmass of plagioclase, augite and Fe-Ti oxides. Rarely, basalts from the base of the sequence contained pseudomorphed olivine phenocrysts. The basalts have all been subjected to low-grade metamorphism ranging from zeolite to prehnite-pumpellyite facies (McIlwaine and Wallace, 1976). At this stop, approximately 500m above the basal conglomerates, are exposed a sequence of thin rubbly basalt flows ~50cm thick with rare massive flows ~2-4m thick and thin interflow sediments. These thick flows host well-developed vesicle columns (Fig. 6). Geochemically basalts at this outcrop are similar to the Central Suite of Lightfoot et al. (1991). Sedimentary units are predominantly quartz sandstones with thin shale partings. These are best interpreted as sands washing into small hollows on the surface of the flow units with a mud drape settling out towards the top of the layer. In places units that appear to have been deposited on the surface of basalt flows connect into sub-vertical cooling cracks in the flows. Sediment filled cooling cracks can be up to 2m deep (Fig. 7). Stop 2 – Osler Volcanics, North end of Wilson Island UTM coordinates 0461810E 5403438N Exposed at this outcrop, are the lower flows of the Osler Volcanic Group ~300m above the conglomerates of the Upper Simpson Island Formation. The basaltic Figure 9. Toe lobe in pahoehoe basalt flow at Stop 2, Wilson Island. flows are generally massive, ranging in thickness from 1-3m. Flow tops vary from rubbly to well-developed pahoehoe textures (Fig. 8). The basalts are vesicular and amygdaloidal and in places the vesicles are elongated giving them almost a pipe-like appearance. Geochemically basalts at this outcrop are similar to the Central Suite of Lightfoot et al. (1991). In some well-developed flow lobes are preserved (Fig. 9). At the north end of the outcrop the flows are cut by a 2-3m wide mafic dyke. This dyke is geochemically distinct from the flows but comparable to the older diabase intrusions in the vicinity of Lake Nipigon. Stop 3 – Upper Simpson Island Formation, Daylight Point, Wilson Island - 67 - UTM coordinates 0461450E 5404650N �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 10. Fine-grained red sandstones with thin shale partings forming the base of the deltaic deposits at Stop 3. separates this assemblage from overlying mediumto large-scale, trough cross-stratified, coarse-grained sandstones to conglomerates (Fig. 12). Paleocurrent indicators show flow to the west, though with a higher variance than other sections. Clast lithologies are probably locally derived from both Archean and Proterozoic sources. The fine-grained sandstone near the base of the section represents a wave modified deltafront (i.e., a distributary mouth bar of a small delta). The presence of small, dish-shaped scours suggests a shallow water environment with no large channels. The upper part of the sequence represents badly organized river deposits with gravelly, longitudinal bar forms and channels filled with sand. The planar cross sets at the base of the cliff were formed by transverse bars, while the trough cross beds represent migrating dunes. Figure 11. Oscilation ripples with overlying hummocky, medium-grained sandstone showing the effects of wave reworking on the sediments forming the delta front at Stop 3. Figure 12. Cross-stratified sandstone and massive conglomerate forming the upper portion of the prograding deltaic succession present in the Upper Member of the Simpson Island Formation present at Stop 3. These sediments represent a longitudinal bar-channel complex of a braided stream. A sedimentary assemblage of the Upper Member occurs on Wilson Island, overlying approximately 50 meters of basal basalt. This coarsening upwards succession has oscillation rippled, very fine-grained sandstones at its base (Fig. 5, Section 6). These coarsen upwards by the addition of increasing amounts of medium-grained, parallel laminated to hummocky cross-stratified to oscillation rippled, decimeter-scale sandstone beds (Figs. 10, 11). A covered interval Figure 13. Interlayered red siltstones and dolostones (lower unit underlying the more massive strandline carbonate with overlying mass-flow deposits) were deposited in a saline lake away either temporally or spatially from areas of coarse sand influx. The colour banding reflects the position of the redox boundary as the sediments accumulated. The grey layers commonly have slightly higher dolomite contents possibly reflecting higher organic productivity leading to more photosynthetically mitigated carbonate precipitation (higher dolomite content) and heavier organic loading to the sediment (redox boundary moving upward to at or above sediment water interface). Stop 4, Mary Ann Bay. - 68 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 14. Stromatolitic layering (smooth mat with small pinnacles) with interbedded coarse silt to very fine-grained sand storm layers (white). This is typical of strandline to sabkha environments and especially open water ponds on the sabkha. Some of the storm layers were remobilized in the form of clastic dykes and sills. Stop 4 – Sibley Group, Mary Ann Bay, Channel Island UTM coordinates 0462769E 5405999N A succession of grey dolomites interbedded with red siltstones outcrop on the shoreline here (Fig. 13). These are interpreted to be part of the cyclic facies (Channel Island Member) of the Rossport Formation as exposed at Kama Hill. The dolomite and minor gypsum indicate a hypersaline environment interpreted to be lacustrine because of the multiple deltas and sand sheets building in from a variety of directions. The cyclic facies is overlain by stromatolites and interbedded thin, sandy, carbonate storm layers (Fig. 14) of the Middlebrun Bay Member. This meter thick assemblage is similar to recent sabkha deposits on the south shore of the Persian Gulf, and in particular the open water ponds on this sabkha where sandy storm layers are well preserved. The upper few centimeters of the stromatolitic unit is altered to a grey-green layer that represents a weathered horizon interpreted to have formed as the sequence became subaerial and the stromatolites weathered in situ. This weathered zone is traceable throughout the basin with the strandline deposits below it commonly containing well developed tepee structures. The carbonates are overlain by a massflow unit with intraformational clasts of red siltstone, sandstone and dolostone up to boulder size. Although the contact between the carbonates and the mass flow Figure 15. Odd shaped structures of probable stromatolitic origin. within the Gunflint Formation. Stop 5, Quarry Island units is locally obscured by the intrusion of a sill, the transition is interpreted to represent a minor time gap based on the weathered zone which expands to thick terra rosa (soil) deposits at other locations. Stop 5 – Gunflint Formation, Quarry Island UTM coordinates 0462371E 5406786N A succession of sandstones and mafic volcanic rocks outcrop on the south shore of Quarry Island. On the northeastern end of the outcrop area a gabbro, probably related to the Midcontinental Rift, is exposed. Next to this is a small outcrop of stromatolites with a box-like appearance (Fig. 15). The rectangular to square outline of the mounds contrasts with the round to oval appearance of classic stromatolites, though their organic origin is exemplified by the high angle layering, which, when projected into the area now eroded, can be seen to form mounded structures. Areas between the stromatolites are infilled with coarser siliciclastic sandstones and cherty clasts. The next outcrop of Gunflint volcanic rocks is problematic. Mafic volcanic flow rocks occur interbedded with Upper Gunflint lithologies southwest of Thunder Bay. These are also associated with stromatolites that developed on the firm substrate of the flow tops. Thus, the igneous rocks in the Gunflint assemblage on Quarry Island could be correlative to the other flow rocks, but - 69 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 16. Well graded sandstone layers. Though these are probably turbidites they were not deposited in excessively deep water and may, in fact, represent distal tempestites. Stratigraphic position is problematic as these beds may belong to either the Upper Gunflint or Rove Formations. Stop 5, Quarry Island. it is difficult to conclusively show that these rocks are extrusive. Possible flow banding is present, as are areas of finer and coarser material in individual units. The igneous rocks are overlain by medium- to coarsegrained sandstones with bed thicknesses averaging approximately 30 cm. The sandstones are dark in appearance giving the impression they were derived Figure 17. Unusual markings on the bedding planes of the graded sandstones. Stop 5, Quarry Island. Figure 18. A second example of unusual marking observed at Stop 6. The origin of these markings is unclear. Stop 5, Quarry Island from mafic detritus, but their geochemistry indicates an intermediate source similar in composition to the Archean crust to the north. The layers are excellently graded (Fig. 16) with the only sedimentary structure being sporadically developed parallel lamination. Beds such as these are commonly thought of as typical turbidites and the deposits ascribed to reasonably deep water. However, it must be remembered that graded bedding simply means a decelerating flow deposited the bed, which can occur in any water depth. These beds may be tempestites, ie. beds formed by storm events, in this case in water deeper than storm wave base but certainly not anything approaching abyssal depths. Or they may have formed from inter- or overflow sediment-water plumes off river mouths, though lack of current reworking of bed tops makes this unlikely. Alternatively they may represent prodelta deposits formed by slumping of the delta front. All of these environments are relatively shallow which agrees with the presence of stromatolites not far stratigraphically below the graded beds. Another interesting point concerning these clastic units is that although such sandstones are common in the upper Rove Formation they are not present in the Gunflint at any other location. Thus, their stratigraphic position is debatable. The third unusual attribute is the presence of difficult to interpret structures on some bedding planes. Series of enechelon small crack-fill like features cut across bedding planes (Fig. 17). In addition a jellyfishlike impression was found on a bedding plane (Fig. 18). This feature had five-fold symmetry, similar to echinoderms, but the age of the rocks and its presence in sandstone leads to the distinct possibility that it was manufactured. - 70 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 19. Silicified alteration envelopes adjacent to quartz veins in basal Gunflint carbonates, near Rossport. Photo courtesy of Mark Smyk. Stop 6 – Pass Lake Formation, Quarry Island Figure 21. Black chert with gossan within the basal conglomerate of the Gunflint Formation at Gut Point. Photo courtesy of Mark Smyk. capable of creating such an organization of lithofacies. UTM coordinates 0461720E 5406798N - this stop The presence of the pebble is significant as only freak will be time dependant. wind-storms, such as tornadoes, can move material of This outcrop consists of a cliff-face in sandstones, this weight, and these do not form sand dunes. So, it is which were quarried, and the blocks produced used in more likely that these deposits are subaqueous but this the construction of buildings in Thunder Bay (Fralick rests only on a slim piece of evidence. et al., 2000). Here we see the basal sediments of the Sibley Group, the Pass Lake Formation. The Pass Stop 7 – Basal conglomerate of the Gunflint, Gut Lake forms a diverse group of basal coarse clastic Point deposits representing environments ranging from UTM coordinates 0461610E 5408587N braided fluvial through to subaqueous sand sheets. The unconformity and basal Gunflint are exposed The medium-grained sandstones present in this cliff are organized into a series of large-scale planar cross- at Gut Point as a thin, discontinuous veneer along the stratified sets with normal to low dip angles. Sorting is lakeshore on top of Archean basement (Fig. 19). The fairly good and only one pebble has been found in the basement is a medium-grained, equigranular granite, succession. Assemblages such as this pose a dilemma which has been altered (sausseritized/chloritized) in formulating an interpretation of their depositional beneath the basal Gunflint. The basal conglomerate is environment. Both aeolian sand dunes and sandflats up to 30 cm thick and occupies depressions in the paleocomposed of transverse bars in braided rivers are erosion surface in the basement. The conglomerate is matrix-supported, with subangular to rounded pebbles of white, sugary quartz, lesser cherty and lithic fragments and minor jasper in a medium-grained, sandy matrix (Fig. 20). A black, pyritiferous chert breccia is marked by a conspicuous gossan (Fig. 21). Sulphide mineralization may be related to a persistent, parallel fracture set at 115°. Fractures may host quartz-calcitebarite veins ranging from < 1 to 20 cm wide as well as vein breccia. A conjugate fracture set at right angles to the first is locally developed. Silicification adjacent to the veins has preserved a thin (1 to 5 cm) veneer of Figure 20. Matrix supported basal conglomerate of the Gunflint from being eroded (especially the carbonate Gunflint Formation containing rounded pebbles of quartz, units). A 2m wide diabase dyke strikes at 115° through chert and lithic fragments. Photo courtesy of Mark Smyk. the outcrop. - 71 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 km in size and intrudes the Schreiber greenstone belt; no radiometric date has been generated. Feldspar phenocrysts are typically 3-4 cm across (Fig. 23), subhedral to euhedral and in places appear to display localized alignment suggestive of flow banding. Acknowledgements We would like to thank Mark Smyk and John Scott for their help and advice in the preparation of this field guide. In particular Mark Smyk provided text and photographs for Stop 7. References Addison, W.D., Brumpton, G.R., Vallini, D.A., Davis, D.W., Kissin, S., Fralick, P.W., McNaughton, N.J., and Hammond, A., 2005. Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology, 33, 193-196. Cannon, W., Green, A., Hutchinson, D., Lee, M., Milkereit, B., Behrendt, J., Halls, H., Green, J., Dickas, A., Morey, G., Sutcliffe, R., and Spencer, C., 1989. The North American Midcontinent Rift beneath lake Superior from GLIMPCE seismic reflection profiling. Tectonics, 8, 305-332. Figure 22. Porphyritic Archean granite, Selim Point. Carter, M.W. 1988. Geology of the Schreiber-Terrace Bay area, District of Thunder Bay; Ontario Geological Survey, Open File Report 5692, 287p. Cheadle, B.A., 1986. alluvial-playa sedimentation in the lower Keweenawan Sibley Group, Thunder Bay District, Ontario. Canadian Journal of Earth Sciences, v. 23, p. 527-542. Figure 23. Feldspar phenocrysts in Archean porphyritic granite at Selim Point. Stop 8 – Archean basement, Selim Point UTM coordinates 0469219E 5409146N From the dock in Rossport return to Highway 17 and head east for ~5 km. Turn right on to Lakeshore Drive just west of Whitesand Provincial Park. Follow the dirt road to a parking spot opposite a small tombola (Fig. 22). The porpyritic granite exposed here is Archean in age and part of the Wawa Subprovince. The area was mapped by Carter (1988) who described the rocks as porphyritic pink, hornblende + biotite alkali feldspar granite, a phase of the Whitesand Lake Batholith. The porphyritic “facies” is surrounded by massive pink and grey phases of alkali feldspar granite that is not exposed at this locality. The batholith is about 8 x 16 Davis, D.W., and Green, J.C., 1997. Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences, 34, 476-488. Davis, D.W. and Sutcliffe, R., H., 1985. U-Pb ages from the Nipigon plate and northern Lake Superior. Geological Society of America Bulletin, 96, 1572-1579.Davis, D.W., and Green, J.C., 1997. Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution. Canadian Journal of Earth Sciences, 34, 476-488. Fralick, P.W. and Barrett, T.J., 1995. Depositional controls on iron formation associations in Canada. In, ed by A.G. Plint, Sedimentary Facies Analysis, Special Publication of the International Association of Sedimentologists, v. 22, p. 137-156. Fralick, P.W., Kissin, S.A. and Davis , D.W., 2002. The age of the Gunflint Formation, Ontario, Canada: single zircon U-Pb age determinations from reworked - 72 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 volcanic ash. Canadian Journal of Earth Sciences, v. 39, p. 1085-1091. Fralick, P.W., Smyk, M. and Mailman, M., 2000. Geology and stratigraphy of the Mesoproterozoic Sibley Group. In, ed. by P. Fralick, Field trip Guide Books, Forty-Sixth Annual Meeting, Institute of Lake Superior Geology. p. 7-42. Franklin, J.M., 1978. The Sibley Group, Ontario. In, Rubidium-strontium isochron age studies, Report 2, Geological Survey of Canada, Paper 77-14, p. 31-34. Giguere, J.F., 1975. Geology of St. Ignace Island and adjacent islands, District of Thunder Bay. Ontario Ministry of natural Resources, Geological Report 118, 35p. Halls, H.C., 1974. A paleomagnetic reversal in the Osler Volcanic Group, Northern Lake Superior. Canadian Journal of Earth Sciences, 11, 1200-1207/ Heaman, L.M., and Machado, N., 1992. Timing and origin of the Midcontinent Rift alkaline magmatism, North America: evidence from the Coldwell Complex. Contributions to Mineralogy and Petrology, 110, 289-303. Hemming, S.R., McLennan, S.M. and Hanson, G.N., 1995. Geochemical and Nd/Pb isotopic evinence for the provinance of the early Proterozoic verginia Formation, Minnisota. Implications for the tectonic setting of the Animikie Basin. Journal of Geology, v. 103, p. 147-168. G.R., 2003. New zircon ages from the Gunflint and Rove Formations, northwestern Ontario. Proceedings Institute of lake Superior Geology, Lightfoot, P., Sutcliffe, R., and Doherty, W., 1991. Crustal contamination identified in Keweenawan Osler Group tholeiites, Ontario: A trace element perspective. Journal of Geology, 99, 739-760. McIlwaine, W.H., and Wallace, H., 1976. Geology of the Black Bay Peninsula Area, District of Thunder Bay, Accompanied by Map 2304, scale 1 inch to 1 mile. Ontario Division of Mines, GR133, 54p. Maric, M. and Fralick, P.W., 2005. Sedimentology of the Rove and Virginia formations and their tectonic significance. Institute of Lake Superior Geology, v. 51, p. 41-42. Metsaranta, R.T., 2006. Sedimentology and geochemistry of the Mesoproterozoic Pass Lake and Rosport Formations. Sibley Group. Unpublished MSc. Thesis, Lakehead University, 217 pp. Morey, G.B., 1967. Stratigraphy and sedimentology of the Middle Precambrian Rove Formation in northeastern Minnesota. Journal of Sedimentary Petrology, v. 37, p. 1154-1162. Miall, A.D., 1978. Lithofacies types and vertical profile models in braided river deposits: A summary. In ed. A.D. Miall, Fluvial Sedimentology, Canadian Society of Petroleum Geologists Memoir 5, 597-604. Heaman, L.M., Easton, M., Hart, T.R., Hollings, P., MacDonald, C.A., and Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon Region, Ontario. Canadian Journal of Earth Sciences, 44, 1055-1086. Nicholson, S.W., Shirey, S., Schulz, K., Green. J., 1997. Riftwide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development. Canadian Journal of Earth Sciences, 34, 504-520. Hollings, P., and Fralick, P., 2005. A stratigraphic transect across the northern flank of the Midcontinent Rift near Rossport. In; Hollings, P. (Ed.), Institute on Lake Superior Geology Proceedings, 51st Annual Meeting, Nipigon, Ontario, Part 2 - Field trip guidebook, v.51, part 2, 57-70. Paces, J.B., and Miller, J.D, Jr., 1993. Precise U-Pb ages of Duluth Complex and related mafic intrusions, northeastern Minnesota; geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagnetic processes associated with the 1.1 Ga Midcontinent Rift System. Journal of Geophysical Research, B, Solid Earth and Planets, vol.98, no.8, pp.13,997-14,013. Hollings, P., Fralick, P. and Cousens, B., 2007. Geochemistry and sedimentology of the Osler Formation: Evaluating rifting in the Proterozoic. Canadian Journal of Earth Sciences, 44, 389-412. Keays, R. and Lightfoot, P., 2015. Geochemical Stratigraphy of the Keweenawan Midcontinent Rift Volcanic Rocks with Regional Implications for the Genesis of Associated Ni, Cu, Co, and Platinum Group Element Sulfide Mineralization. Economic Geology, 110, 1235–1267. Kissin, S.A. and Fralick, P.W., 1994. Early Proterozoic volcanics of the Animikie Group, Ontario and Michigan, and their tectonic significance. Proceedings Institute of Lake Superior Geology, v. 40, p. 18-19. Kissin, S.A., Vallina, D.A., Addison,W,D. and Brumpton, Pufahl, P.K. and Fralick, P.W., 2004. Depositional controls on paleoproterozoic shallow-water iron formation accumulation, Gogebic Range, Wisconsin, U.S.A. Sedimentology, v. 54, p. 791-808. Rogala, B., Fralick, P.W., Heaman, L.M. and Metsaranta, R., 2007. Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario, Canada. Canadian Journal of Earth Sciences, v. 44, p. 1131-1149. Shirey, S., Lewin, K., Berg, J., and Carlson, R., 1994. Temporal changes in the sources of flood basalts: Isotopic and trace element evidence from the 1100 Ma old Keweenawan Mamainse Point Formation, - 73 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Ontario, Canada. Geochimica et Cosmochimica Acta, 58, 4475-4490. Sutcliffe, R. H., 1986. The petrology, mineral chemistry and tectonics of Proterozoic rift-related igneous rocks at Lake Nipigon, Ontario. Unpublished Ph.D. thesis, University of Western Ontario, London, 325p. Sutcliffe, R.H., and Smith, A.R., 1988. Project number 87-17. Geology of the St. Ignace Island volcanicplutonic complex. Summary of Fieldwork and Other Activities 1988. Ontario Geological Survey Miscellaneous Paper 141, 368-371. Swanson-Hysell, N. L., Vaughan, A. A., Mustain, M. R. and Asp, K. E., 2014. Confirmation of progressive plate motion during the Midcontinent Rift’s early magmatic stage from the Osler Volcanic Group, Ontario, Canada. Geochem. Geophys. Geosyst., 15, 2039–2047, doi:10.1002/2013GC005180 Tanton, T.L., 1931. Fort William and Port Arthur, and Thunder Cape map areas, Thunder Bay District, Ontario. Geological survey of Canada Memoir 167, 222p. - 74 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 6 - Geology of the Coldwell alkaline complex Allan MacTavish Panoramic PGMs (Canada) Limited, Thunder Bay, ON, Canada Mark Smyk Resident Geologist Program, Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, Thunder Bay, Ontario, P7E 6S7, Canada David Good Earth Sciences Dept., Western University, London, Ontario, Canada and John McBride Stillwater Canada Inc., Marathon, Ontario, Canada The Coldwell Alkaline Complex Trip will consist of two parts comprising 1) a west to east transect through southern part of the complex and 2) a visit to the Marathon Cu-PGE Deposit. This guide is modified from the 2017 ILSG Field Guide. Part 1: Transect Through the Coldwell Alkaline Complex Allan MacTavish and Mark Smyk A variety of Mesoproterozoic, Midcontinent Riftrelated alkalic and carbonatitic rocks occur within several intrusive complexes on or near the northern shore of Lake Superior (Figs. 1 and 2). They include the Coldwell and Killala Lake alkaline complexes, the Prairie Lake and Chipman Lake carbonatites, and numerous diatremes and related dikes in the vicinity of Dead Horse Creek (Sage, 1982, 1985, 1987; Fig. 2). These complexes are spatially localized and structurally controlled by the Trans-Superior Tectonic Zone (TSTZ), a north-northeast-trending structure that extends for over 600km and includes the Thiel Fault in Lake Superior (Klasner et al., 1982). Alkaline magmatism related to Midcontinent rifting occurred along the TSTZ from approximately 1.2 to 1.0 Ga (Table 1). It has been postulated that the TSTZ may represent Figure 1. Midcontinent Rift geology and the locations of mafic/ultramafic intrusions (After Miller et al., 1995). - 75 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 alkaline complexes are both thought by some to have formed as the result of ring fracturing and caldera collapse. The abundance of observed xenolithic blocks and roof pendants suggests that these complexes are presently exposed at relatively high structural levels. Figure 2. Regional geology (Sage, 1991) in the vicinity of the Trans-Superior Tectonic Zone (TSTZ), extension of the Thiel Fault (B). Key to numbering: 30 – Chipman Lake fenites / carbonatite dikes; 31 – Killala Lake alkaline complex; 32 – Prairie Lake Carbonatite; 33 – Coldwell alkaline complex; 36 – Slate Islands; 47 – Dead Horse Creek diatremes; 48 – McKellar Creek diatreme; 49 – Gold Range Diatreme; 50 – Neys Diatreme; A – Michipicoten Fault; C – Killala Lake Deformation Zone. part of a failed arm of a Keweenawan-age triple junction (Weiblen, 1982; Mitchell and Platt, 1982b) or the intersection of a late fracture system with the rift (Mitchell et al., 1983). Local alkalic and carbonatite complexes have been emplaced at inflections in the trends of major structural zones, or at sites of crossfaulting (Sage, 1991). The Coldwell and Killala Lake Similar ages for numerous mafic intrusions in the Nipigon Embayment (cf. Heaman et al., 2007) and the alkalic rocks of the Coldwell Complex (1108 ± 1 Ma; Heaman and Machado, 1992) indicate the contemporaneous production of tholeiitic and alkalic magmas during Midcontinent rifting. The oldest magnetization, found in the gabbros and augite syenites on the eastern side of the complex, records a concordant pole position with reversed polarity at about 1109 ± 5 Ma on the Keweenawan segment of the Precambrian apparent polar wander path (Lewchuk and Symons, 1990). The localization of the alkalic magmatism offaxis, dominantly northeast of the central rift, prompted Heaman and Machado (1992) to suggest that this may have been a region of maximum lithospheric extension during rifting. U/Pb data (Heaman and Machado, 1992) demonstrate that most rock units in the Coldwell Complex were emplaced within a relatively short time span (<3 million years) ca. 1108 Ma, and support the contention that the complex experienced relatively rapid cooling from initial emplacement temperatures to at least ~500º C. Strontium-, neodymium- and lead-isotopic compositions of selected minerals from different phases of the complex (Heaman and Machado, 1992) display considerable scatter, suggesting that their magmas had different isotopic compositions. The initial strontium- and neodymium-isotopic compositions of clinopyroxene and plagioclase from one of the earliest gabbroic phases are identical to data derived from primitive olivine tholeiites from the Midcontinent Rift and indicate that the majority of magmas, both Table 1 MCR-related Alkaline Magmatism Occurring Along the TSTZ Lithologic Unit/Complex Coldwell Alkaline Complex Be-Zr Zone crosscutting Dead Horse Creek diatreme Prairie Lake Carbonatite Lamprophyre Dyke, McKellar Harbour Gabbro (biotite), Killala Lake Complex Syenite, Killala Lake Complex ( -2.49% discordant; 1.82% discordant) 1 Age(s) (Method) 1108 ± 1 Ma (U/Pb) 1112.7 ± 4 Ma (U/Pb) 1128.7 ± 6 Ma (U/Pb) 1130 ± 10Ma (Rb/Sr) 1145 + 15/10 Ma (U/Pb) ~1160 Ma (U/Pb) 1185 ± 90 Ma (K/Ar) 1 2 1050 ± 35 Ma (Rb/Sr) 2 - 76 - Reference Heaman and Michado (1987) Krogh and Wilkinson (M. Smyk pers. Comm., 1995) Pollock (1987) Queen et al. (1996) Wu et al. (2016) Coats (1970) Bell and Blenkinsop (1980) �Proceedings of the 65th ILSG Annual Meeting - Part 2 tholeiitic and alkaline, have a uniform, nearly chondritic isotopic composition (ibid). Samarium- neodymium data, supported by oxygen-isotopic and whole-rock geochemical data, indicate that crustal contamination played a small, varied role in the generation of the Coldwell magmas (Bohay, 1997). In addition to small, variable amounts of assimilation of upper and lower crust, the parental plume magmas also interacted with the lithospheric upper mantle to a small degree (ibid). Local alkalic and carbonatitic intrusive rocks host a variety of characteristic base, precious, titaniferous, phosphate, and rare metal occurrences (cf. Smyk and Sage, 1995). They include the following: 1. Magmatic Cu-Ni-PGE (± Au, Ag) in gabbros of the Killala Lake and Coldwell complexes; 2. Magmatic Ti-V±apatite deposits in the Eastern Border Gabbros of the Coldwell Complex; 3. Magmatic U, Nb (+ wollastonite, apatite) in the Prairie Lake carbonatite (Sage, 1987); 4. Late-stage magmatic Nb-Y-F-family rare earth elements in syenite pegmatites (Alexander, 2007); 5. A Be-Zr-U-Th-Y mineralized zone crosscutting the Dead Horse Creek diatreme (Smyk et al., 1993; Potter, 2004); and 6. Pb-Zn-Ag-mineralized quartz-carbonate veins (Kissin and McCuaig, 1988). The Coldwell Alkaline Complex (Fig. 3) covers an area of ~580km2, making it one of the largest alkalic complexes in the world and the largest in North America. It was emplaced during the early stages of the Midcontinent rift system, which includes: early large and small mafic to ultramafic intrusions (i.e. Seagull Lake Complex, Thunder Bay North Intrusive Complex); Keweenawan flood basalts, the Duluth Complex, the Nipigon and Logan sills, and a variety of non-diabase mafic to ultramafic dyke-rocks. The Coldwell Complex was mapped by Kerr (1910a, 1910b), Puskas (1967), and Walker et al. (1993b, 1993c), and comprises three, superimposed ring subcomplexes or magmatic centers (Mitchell and Platt, 1978) that young progressively (Centers 1 to 3) to the southwest (Fig. 4). Walker et al. (1993) and Barrie et al. (2002) dispute the series of ring dykes or sheeted cones interpretation and suggest that the complex is a composite lopolith or sill. The intrusive centres can be generally described as follows: Center 1: Generally silica-saturated rocks with oversaturated residue; chiefly consisting of the Eastern and Western border gabbros (the oldest rocks within the complex) and later iron-rich augite syenite and syenite-syenodiorite (Mitchell and Platt, 1978, 1982; Mulja, 1989); Center 2: Generally silica-undersaturated alkalic rocks with oversaturated residue; consisting of locally nepheline- and hastingsite-bearing miaskitic nepheline syenite, and numerous volumetrically minor alkaline lamprophyre and analcime tinguaite dykes (Mitchell and Platt, 1978, 1982; Laderoute, 1989; and Mulja, 1989); and Center 3: Silica-oversaturated alkalic rocks with oversaturated residue; consisting of magnesiohornblende syenites, quartz syenites, and minor granites (Mitchell and Platt, 1994; LukosiusSanders, 1988). The mineralogy of the main lithologic units is listed in Table 2. The superimposition of the three intrusive centres and a complex, protracted magmatic history has produced a myriad of hybrid rocks, igneous breccias, and ambiguous crosscutting relationships. The wide variety of lamprophyric and other dyke rocks occurring within the complex (as described by Mitchell and Platt, 1994) include (in order of emplacement): Figure 3. Generalized geology of the Coldwell Alkaline Complex (after Walker et al., 1993) with field trip stops. - 77 - 1. Mafic ocellar lamprophyre (camptonitic variety) 2. Quartz-bearing, mafic lamprophyres (camptonitic variety) �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 4. Coldwell Alkaline Complex magmatic centres: CI (Centre 1), CII (Centre 2), and CIII (Centre 3). Generalized geology after Mitchell and Platt (1994). 3. Sannaite-type lamprophyres 4. Monchiquitic-type lamprophyres 5. Feldspar glomeroporphyritic and alkali basalt dikes 6. Analcime tinguaite (heronite) Abundant large rafts and/or roof pendants of mafic volcanic rocks are mapped throughout the Coldwell Complex and in places exhibit horizontal extensional cooling cracks on a ten to hundreds of metres scale that are thought consistent by some workers with subhorizontal bedding. For the most part the roof pendants may be the lowermost portions of the Keweenawan flood basalt sequences suggesting that the complex is barely unroofed and is exposed at a very shallow crustal level (Mitchell and Platt, 1994; Sage and Watkinson, 1995; Barrie et al., 2002). It is also highly probably that some, or most of the mafic rafts observed within the complex that could not be roof pendants are detached portions of chilled complex roof or wall rocks (finegrained gabbros). The mafic intrusive rocks occurring within Centers Table 2. Mineralized Zones Associated with the Mafic Intrusive Rocks of Centres 1 and 2. Intrusion (Centre) Eastern Gabbro (1) Western Gabbro (1) Two Duck Lake (1) Malpas Lake (1) Geordie Lake (2) Alkalic gabbro (2) Lithologic Units Layered gabbro cumulates (olivine gabbro, gabbro, troctolite, anorthositic (leuco-) gabbro); Fe-Ti oxide ± apatite cumulates Massive and layered series gabbro; olivine-bearing Gabbronorite, olivine gabbronorite, olivine-bearing gabbro, leucogabbro Hornblende gabbro to monzodiorite; olivine ferrogabbro to ferrodiorite; olivine gabbro to diorite Amphibole-bearing olivine gabbro Troctolite, olivine gabbro Biotite gabbro Biotite- and olivine-gabbro - 78 - Reference(s) Shaw (1994, 1997); Lum (1973); Barrie et al. (2002) Penczak (1992); Wilkinson (1983) Shaw (1994, 1997) Dahl et al. (1987) Shaw (1994, 1997) Mulja (1989); MacTavish et al. (1987); Good, pers. com. (2019) Mitchell and Platt (1982b) Walker et al. (1993a) �Proceedings of the 65th ILSG Annual Meeting - Part 2 1 and 2 are tabulated, with their associated mineralized zones (Table 2). Magmatic, gabbro-hosted Cu-Ni-PGE deposits in the Coldwell Complex have been the focus of much exploration and research for the past 60 years. Mineralized zones occur within the border gabbro at the eastern (Marathon deposit; Skipper Lake Zone) and western (Middleton occurrences) margins of the complex, and within its interior at Geordie Lake. The Geordie Lake mineralized zones, hosted by a younger (?) gabbro, are enriched in tellurium and silver and have higher Pd:Pt ratios (~19) (Mulja and Mitchell, 1991) than the border gabbro-hosted deposits (~4; Smyk, 2001). Geochemical variations in mineralized zones in the Coldwell Complex are shown in Figure 5. A table of selected Coldwell Complex deposits and mineralized zones is shown below (Table 3). Marked similarities exist between the mineralization style, geochemistry, and host rocks of Coldwell Complex-, Duluth Complex-, and the Crystal Lake gabbro-hosted deposits near Thunder Bay. Similarities include mineral textures, abundance and compositions, crystallization paths for the host gabbros, silicatesulphide associations, trace-element trends and Figure 5. Discrimination plot for PGE-mineralized samples for Coldwell and other Midcontinent Rift-related intrusions. Data from Good (1992), MacTavish (unpublished data, Resident Geologist’s Files, Thunder Bay), Watkinson et al. (1983), Wilkinson (1983), and unpublished data, Resident Geologist’s Files, Thunder Bay. Duluth Complex composite data from Hauck et al. (1997). chalcophile element fractionation trends (Good and Crockett, 1994a). Research by Watkinson and Ohnenstetter (1992) and Good and Crockett (1994a, 1994b) produced debate between the relative importance of magmatic and hydrothermal processes in local copper-nickel-PGE Table 3. Selected Coldwell Alkaline Complex Deposits and Mineralized Zones Mineralized Zone Marathon Geordie Lake Grade / Significant Assays Ore Mineralogy Reference(s) Measured and Indicated InPit Resources: 114.8 Mt @ 0.775 g/t Pd, 0.228 g/t Pt, 0.083 g/t Au, 0.241% Cu, 1.567 g/t Ag; Proven and Probable In-Pit Reserves: 91.447 Mt @ 0.832 g/t Pd, 0.237 g/t Pt, 0.085 g/t Au, 0.247% Cu, 1.440 g/t Ag (January 2010) Measured and Indicated Resources (above $13.00/t cut-off): 32.42 Mt @ 0.61 g/t Pd, 0.04 g/t Pt, 0.05 g/t Au, 0.37% Cu, 2.93 g/t Ag Chalcopyrite ≤ pyrrhotite >> pentlandite > cubanite ≤ pyrite; sphalerite, hollingworthite, atokitezvyaginstevite, sperrylite, Bikotulskite, michenerite, merenskyite, monceite, stibiopalladinite, paolovite, merteite II, palladoarsenide, unnamed (Pd As ), nickeline, majakite, argentian gold Chalcopyrite, bornite, pyrite, millerite, siegenite, pentlandite, galena, chalcocite, melonite, hessite, unnamed (Ag Te ), altaite, kotulskite, merenskyite, michenerite, sopcheite, Pdbismuthotelluride, paolovite, Pdarsenide, guanglinite, Pdantimonide, sperrylite, electrum, Pd1.6As1.5Ni, AgSb Chalcopyrite, pyrrhotite, pentlandite, sphalerite, pyrite Chalcopyrite, bornite, pentlandite, cobaltite, galena, chalcocite; telargpalite, polarite, kotulskite, taimyrite, merteite, zvyagintsevite, plumbopalladinite, majakite, tetraferroplatinum n/a Marathon PGM Corporation Ohnenstetter et al. (1991); Watkinson and Ohnenstetter (1992); Good and Crocket (1994a, 1994b) 5 2 3 4 Middleton Skipper Lake average grade of 1.05 g/t Pd+Pt+Au over 12 m Area 41 0.48 g/t Pt+Pd+Au over 202 m, incl. 1.23 g/t Pt+Pd+Au over 61 m - 79 - 2 news release, Marathon PGM Corporation, May 04, 2010 Mulja (1989); Mulja and Mitchell (1990, 1991) Penczak (1992) MacTavish (2000) Benton Resources Corp. �Proceedings of the 65th ILSG Annual Meeting - Part 2 mineralization processes. Watkinson and Ohnenstetter (1992) presented field, petrographic and mineralchemical data that support the interaction of magmatic sulphide mineral assemblages with a chlorine-rich mixture of magmatic (deuteric) fluid and volatile species generated by the breakdown of assimilated xenoliths at low temperatures. However, Good and Crockett (1994a, 1994b) contended that element migration took place over only very short distances and that the original, bulk sulphides were not enriched in copper and PGE by later fluids. The information within this field trip guide was taken from a variety of sources, including guidebooks from previous field trips to the Coldwell Complex: Puskas (1970); Loubat (1972); Mitchell and Platt (1977, 1982a, 1994); Smyk and Sage (1995), Smyk (2001), Smyk (2010), and unpublished field observations and mapping completed by A. MacTavish (1992). All UTM co-ordinates listed are NAD83 Zone 16 with locations shown on Figure 6.   Stop descriptions Stop C1: Natrolite-Bearing Syenite and Massive FeTi-oxides UTM coordinates 525528E 5405511N 29.4 to 29.9 km west of the Highway 626 and Highway 17 junction Description: This exposure displays natrolitebearing, pegmatitic syenite (Photo 1). Reddish orange natrolite (an acicular or prismatic zeolite mineral replacing nepheline) patches up to 15cm in diameter, crystals of perthitic feldspar up to 30cm in length, and crystals or black amphibole up to 25cm in length comprise the bulk of this syenite (Photo 2). Mitchell and Platt (1994) reported accessory pleochroic clinopyroxene, zircon, titanite, and biotite. Natrolite has locally been ascribed to the hydrothermal alteration of primary nepheline and has also been referred to as “hydronepheline” by local workers. The syenite is intruded by a camptonite lamprophyre dike (Mitchell and Platt, 1994) and also hosts large, medium-grained gabbro xenoliths (Photo 3), up to 1m in thickness and sometimes up to 5m in length (west-side of the highway), that exhibit 1 to 2cm wide, dark reaction rims adjacent to the enclosing syenite. To the east, the pegmatitic syenite gives way to finer grained nepheline syenite in which chalky-weathering nepheline may be Photo 1. Pink, natrolite-bearing, pegmatitic augite syenite. Photo credit D. Campbell. Photo 2. Pegmatitic syenite containing reddish orange patches of natrolite, light pinkish perthitic feldspar, and black amphibole. Photo credit A. MacTavish. Photo 3. Large gabbro xenolith located on the west side of the highway. Please note that the xenolith has been crosscut by fine-grained syenite veins and that the syenite below the xenolith is varitextured to pegmatitic in texture, whereas the syenite above is medium- to coarse-grained. Photo credit A. MacTavish. - 80 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 6a. Northern portion of Neys Shoreline geological map starting at Prisoner’s Cove, Neys Provincial Park with Field Stop locations. - 81 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 6b. Central portion of Neys Shoreline geological map, Neys Provincial Park with Field Stop locations. - 82 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 6c. Southern portion of Neys Shoreline geological map, Neys Provincial Park with Field Stop locations. - 83 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 observed. Rare natrolite grains are also present. Farther east, a variety of equigranular and pegmatitic syenites are exposed. Near the eastern end of the outcrop (UTM 525625E, 5404825N), a large xenolith of gabbro-hosted, massive titaniferous magnetite has been exposed. Minor clinopyroxene, plagioclase and apatite occur within the massive oxide unit. Analyses completed in 1951 and reported by Hinz and Landry (1994) indicated total iron and titanium values ranging between 33 and 45%, and 4.5 and 13.5%, respectively; phosphorus contents ranged up to 0.371%. Photo 4. Syenite outcrop on the south side of the highway containing blocks of mafic xenoliths often occurring in elongated, semi-continuous rafts. Photo credit A. MacTavish. Photo 5. Elongated, jig-saw-fit xenolithic block exhibiting both angular and lobate/cuspate (amoeboid) margins. Please note the lighter-coloured, pinkish-grey elongated xenolithic block with apparently sharp margins located below the dark grey block. This lower block appears to be coarser-grained and may possibly be different in composition. Photo credit A. MacTavish. Stop C2: Little Pic River Breccia Zone UTM coordinates 527478E 5405531N 27.3 km west of the Highway 626 and Highway 17 junction These road cuts, particularly along the south side of the highway, expose spectacular intrusive breccias within the youngest rocks of the complex, along the east side of the fault zone that the Little Pic River occupies. The breccias often occur as semi-continuous, fragmented, elongate rafts (western end of southern rock cut, Photo 4) that consist of angular to rounded blocks of fine- to medium-grained, equigranular, mafic (gabbroic?) rocks within a groundmass of pink, mediumgrained, quartz syenite. In some cases blocks can exhibit both angular and lobate to cuspate (amoeboid) margins (see Photo 5). The mafic rocks comprising the blocks were interpreted as oligoclase-bearing basalt by Mitchell and Platt (1982a). Subsequent discussion and study has led to the suggestion of perhaps 2 texturally discernable types of basic xenoliths, those with: (1) sharp, angular margins, and (2) those with lobate to cuspate margins. In this model, the angular xenoliths represent synplutonic basalts which are now preserved elsewhere as megaxenoliths in younger intrusions. The cuspate-margined xenoliths may represent the effects of mixing between two contemporaneous gabbroic/ basaltic and syenite magmas (i.e., magma mixing or co-mingling). Cuspate, possible chilled margins with quench-textured clinopyroxene, plagioclase and skeletal olivine have been noted in similar xenoliths to the south on the Coldwell Peninsula by G. Shore (personal communication with M. Smyk, 1995) and suggest the quenching of the basic magma against the cooler, syenitic magma. These are reasonable hypotheses and there are definitely at least two types and textures of xenoliths; however, they do not completely explain the presence of blocks exhibiting both margin types as observed by the senior author of this guide and shown in Photo 5. Texturally there also seems to be three different types of xenoliths: the most abundant are dark grey to black, very finegrained xenoliths (Photo 4); medium-grained, greyish pink xenoliths with somewhat less distinct, but still relatively sharp margins (Photo 5), and several unusual zones where there is are subvertical zones of rounded, dark grey, amphibole-phyric xenoliths within a pinkish, mafic groundmass. Are these some sort of breccia dykes or just hybridized zones of xenoliths (Photo 6; what do you think?)? Although isolated xenoliths are - 84 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 granites and have been interpreted to be the result of fractional crystallization of mantle-derived, basaltic magma (Lukosius-Sanders, 1988; Mitchell et al., 1993). Stop C3: Prisoners Cove, Neys Provincial Park (Sample Collecting Prohibited!) UTM coordinates 527984E 5402537N The descriptions (unpublished mapping/field descriptions, MacTavish, 1992) herein are for 14 substops along the shoreline south of Prisoner’s Cove; however, due to time constraints only the northernmost stops will be visited. 23.5km west of the Hwy 626 and Hwy 17 junction; 2.8km south of Hwy 17 to the park headquarters and then south along the shoreline trail Photo 6. Rounded dark grey, amphibole-phyric xenoliths/ inclusions within a fine- to medium-grained, pinkish mafic groundmass (breccia dyke?). Photo credit M. Puumala. common, there are many areas within these outcrops where incipient or in-situ brecciation characterized by syenite dykes and “jig-saw puzzle/jig-saw fit” breccias are observed, where brecciated fragments can be fitted back together. Miarolitic cavities, up to several centimetres in width, contain euhedral quartz, feldspar, and calcite crystals. The breccia zone persists to the east, towards the scenic lookout located 800m to the east. The south side of the highway is underlain by oligoclase gabbro and quartz syenite, while various, xenolithic-bearing syenitic rocks are exposed on the north side. These pyroxene- and amphibole-(ferro-edenite) bearing syenites contain xenoliths of alkali gabbro, alkali diorite and other, equigranular to porphyritic syenites. Near the lookout turnoff, gray, nepheline-bearing syenite intrudes the mafic rocks and contains orange natrolite. Sannaite and ocellar, camptonitic lamprophyre dikes have been reported near this site by Mitchell and Platt (1994) who proposed the following order of local emplacement: Mg-hornblende syenite → contaminated Fe-edenite syenite → Fe-edenite syenite → quartz syenite (earliest → latest) Lukosius-Sanders (1988) classified the local rocks as miaskitic, metaluminous syenites enriched in U, Th, REE and Zr. These syenites have affinities to A-type General Description: The wave-washed, glacially polished outcrops along the shoreline of Lake Superior at Prisoner Cove and south for over a kilometre along the western side of the Coldwell Peninsula exhibit a variety of lithologic, textural, and crosscutting features that characterize much of the Center 2 magmatism in the Coldwell Complex. In its simplest sense, this composite stop displays the contact between alkalic biotite gabbro and amphibole-nepheline syenite, but the enigmatic effects of assimilation and hybridization have severely complicated and obscured many of the primary features. In all cases within the nepheline syenitic rocks exposed along the shoreline at this stop the nepheline has been completely altered to the zeolite mineral natrolite which weathers to orange-coloured pits. Medium- to coarse-grained, olivine- and enclavebearing, biotite gabbro comprises much of the eastern portion of the outcrops. Gabbro xenoliths occur within the syenite and within hybrid phases along their mutual contact, which trends roughly north-south, parallel to the shoreline. The outcrops often exhibit a pitted surface resulting from the preferential weathering of mafic enclaves consisting of biotite-olivine gabbro to biotite-clinopyroxene gabbro or leucogabbro (Walker et al., 1992) within a more syenitic groundmass. The syenitic groundmass consists of fine- to coarse-grained nepheline (altered to natrolite) syenite with minor acicular amphibole and poikilitic biotite. Mitchell and Platt (1994) have identified the amphibole as hastingsite. - 85 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Distinct to diffuse layering, a nebulous to locally distinct igneous foliation, and localized soft-sediment style magmatic deformation exists within the amphibole-nepheline syenite. Identifiable, undisturbed layering may be oriented parallel to the apparent syenite/gabbro contact and dips from very steeply to vertically in the north to shallow to moderate to the east (where measurable) in the south. Observed soft-sediment deformation features consist of flame structures, fluid-escape features, slump folds, and isolated well-layered syenite blocks surrounded by obvious fluid escape textures. Much past discussion has focused on whether the observed structures have resulted simply from igneous process, syn- or postintrusion shearing, or a combination of these processes. The present author’s strongly favour igneous processes since the observed fracturing is very localized, is late and brittle, and does not appear to have affected the foliation or layering within the surrounding rocks in any observable way. It is highly probable that the crystallizing magma chamber was often shocked by MCR tectonic activity. These earthquakes then caused the slumping of unstable crystallizing layers along chamber walls; allowed the isolation of broken, but relatively intact layered blocks; and allowed trapped deuteric fluids formed during the fractionation process to escape upwards through the broken layers. Upon close examination the fracturing presently observed in outcrop obviously took place after the chamber was completely crystallized and was able to deform in a brittle manner. Sub-Stop C3a (527980E, 5402571N): This area, located on the point to the north and west of the old Photo 7. Wispy, relatively mafic in appearance, hybrid amphibole-nepheline syenite exhibiting diffuse discontinuous layers. Photo credit A. MacTavish. flat-bottomed boats, mainly consists of foliated, wispy, hybridized amphibole-nepheline syenite with diffuse discontinuous “layers” (Photo 7). The core of this outcrop is flanked to the northeast by a heterogeneous zone containing large numbers of rounded to angular, variably assimilated (metasomatized?) and disaggregated inclusions/xenoliths of biotite gabbro. Reaction rims around these inclusions are readily visible. Also the inclusion-rich zone, as a whole, seems to be enclosed within a diffuse reaction zone when compared to the hybrid syenites adjacent to the west. The western margin of the exposure is a medium- to coarse-grained hybridized syenite with numerous very coarse-grained to pegmatitic inclusions of amphibolenepheline syenite. At the northwestern tip of the outcrop is an elongate, diffuse zone of apparently nonhybridized, non-foliated syenite (possibly the original parent syenite?). Sub-Stop C3b (527950E, 5402535N): This stop, located 30m west-southwest of the old boats near the shoreline, consists of a 4 to 5m wide, west-northweststriking, brittle fracture zone hosting a 70 to 100cm thick, dark greenish-grey, ocellar lamprophyre dyke at its northern margin near the water’s edge. The ocellae present within the dyke are composed of reddish, recessive-weathering carbonate (±zeolites?) which are elongated parallel to dyke margins (elongated by flow?). The lamprophyre dyke is also enveloped by a brick-red alteration halo that is not completely within the fracture zone and also extends into the unfractured hybrid syenites to the north for up to 5m. This red halo could be due to either hematization or K-alteration. Similar, subparallel fracture zones can also be observed about 10m and 23m to the south. Sub-Stop C3c (527966E, 5402475N): This stop is located ~50m east-southeast of Sub-stop C3b, and consists of a zone of large blocks (?) of coarse-grained, natrolite-bearing, biotite gabbro to biotite melagabbro that are surrounded by fine- to medium-grained amphibole-nepheline syenite containing diffuse gabbro xenolith ghosts. It is distinctly possible that this is not a zone of xenoliths/inclusions at all, but the exposed upper contact of an underlying biotite gabbro that is part of the biotite gabbro body located about 40m to the southeast (see Sub-Stop C3e, below) where the syenite is observed to overly the gabbro. These blocks (?) are cross-cut by narrow horizontal and subvertical syenite veins and dykes (Photo 8). - 86 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Photo 8. Large biotite gabbro xenolith/inclusion (?) crosscut by veins and dykes of amphibole-nepheline syenite. Photo credit A. MacTavish. Photo 9. Biotite gabbro xenoliths/inclusions separating from and original larger block and beginning to assimilate (?) into the surrounding syenite melt through a process of metasomatism and disaggregation. Photo credit A. MacTavish. Photo 10. Disaggregating biotite gabbro xenoliths exhibiting subparallel reaction haloes. Photo credit M. Puumala. Sub-Stop C3d (528004E, 5402440N): This location (45m southeast of Sub-stop C3c) consists of an irregular zone of gabbro xenoliths/inclusions, surrounded by fine- to medium-grained, weakly foliated syenite (Photos 9 and 10). The xenoliths are in the process of being broken down in stages from originally angular, cohesive blocks to diffuse groupings of amoeboid to wispy mafic remnants within a mafic mineral-rich, hybrid syenite. This process is probably not assimilation in the strictest sense, but is more likely a process of chemical (rather than thermal) invasion through metasomatism that over time breaks down the xenoliths and then eventually disaggregates the mineral constituents of the blocks to the point where they are then assimilated into the syenite melt. The hybridized (?) syenite surrounding the xenoliths exhibit aligned amphibole grains that may indicate flow (?) around and between fragments. There are also places where there are noticeable (up to 15cm thick) halos surrounding zones of xenoliths that consist of aligned amhibole grains that are somewhat separated into diffuse bands. Sub-Stop C3e (528014E, 5402433N): Located only 12m southeast of Sub-Stop 3d and consists of coarsegrained, knobby-weathering, biotite gabbro that has been cross-cut by numerous hair thin to 5cm thick, very fine- to fine-grained syenite stringers and veins and the occasional, larger, fine-grained to pegmatitic syenite vein (pegmatite is in centre of these veins; Photo 11). There are numerous leucocratic clots (oikocrysts?) of plagioclase (Photo 12) throughout. Sub-Stop C3f (527982E, 5402324N): This substop (~115m west-southwest of Sub-stop C3e) consists of an irregular, variably assimilated zone of mafic Photo 11. Biotite gabbro crosscut by fine-grained to pegmatitic syenite dyke (centre) and thinner syenite veins (centre left). Photo credit A. MacTavish. - 87 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 magmatic layering dips shallowly to the west and west-southwest at between 20 and 26° and there is a possible weak alignment of K-feldspar laths parallel to layering. The bases of the undulating modal layers are defined by a sharp increase in amphibole content. The best defined layering is near the lake with layering becoming increasingly more diffuse, disrupted, folded (slumping?), and contorted to the east until it becomes unrecognizable. (gabbroic?) xenoliths of highly variable size ranges. Many blocks are in the last stages of assimilation where the original xenoliths are now merely ghosts infilled with isolated mafic remnants and considerable numbers of hornblende grains. Sub-Stop C3h (527971E, 5402265N): At this location (~27m south of Sub-stop C3g) are two, subparallel, aphanitic to fine-grained, ocellar lamprophyre dykes (Photo 13) occupying a narrow southeaststriking fracture zone. The dykes dip to the northeast between 54° and subvertical. The ocellae (immiscible liquid droplets) are usually centralized within the dykes away from the strongly chilled dyke margins and are infilled with several minerals including applegreen and greyish minerals (zeolites?), and possibly white calcite. Sub-Stop C3g (527978E, 5402292N): This location (~30m south of Sub-stop C3f) consists of locally welldeveloped modal layering within medium- to locally coarse-grained amphibole-nepheline syenite. The Sub-Stop C3i (527987E, 5402198N): At this location (~70m south of Sub-stop C3h) is a zone of leopard mottles in moderately mafic, often grain-sizelayered (?) amphibole nepheline syenite. Photo 12. Leucocratic clots of plagioclase (oikocrysts) within biotite gabbro. Photo credit A. MacTavish. Sub-Stop C3j (527971E, 5402265N): This location (~80m south of Sub-stop C3i) is, for lack of a better name, a “Layer Breccia Zone” where there has been strong disruption, localized rotation, and folding of original magmatic syenite layers (Fig. 7). Finergrained syenite containing acicular amphibole grains has flowed around the layer blocks and alignment of those amphibole grains mirrors flow directions. The zone is surrounded by a highly disturbed hybrid mixtite with few measurable features. Thinner blocks consist of a series of thin modal layers of highly variable textures. The thicker layers are usually the coarsest, are sometimes size-graded, and contain glomerocrysts of K-feldspar (with include amphibole and natrolite after nepheline) up to 1.5cm in diameter surrounded by acicular amphibole grains and recessive-weathering altered nepheline. Photo 13. Ocellar lamprophyre dyke in narrow fracture zone. Photo credit A. MacTavish. Sub-Stop C3k (528000E, 5402039N): Located 77m south of Sub-stop C3j. This sub-stop comprises a well-layered block of amphibole-nepheline syenite (~6.5m by 3.5m in size) that is surrounded by a highly distorted zone of fine- to very-coarse-grained (varitextured) syenitic material that appears to have flowed around the block (Photo 14 and Fig. 8). This - 88 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Photo 14. A well-layered block of amphibole-nepheline syenite that is surrounded by a highly distorted zone of fineto very-coarse-grained (varitextured) syenitic material that appears to have flowed around the block. Photo credit A. MacTavish. Also refer to Figure 9. Figure 7. Hand drawn detailed map of Layer Breccia Zone ‘A’. Mapping by A. MacTavish (1992). Abbreviation key: vcgr = very coarse-grained; f-cgr = fine- to coarse-grained; m-vcgr = medium-verycoarse-grained; int = intermediate; LS = Leopard spots (mottles). isolated, Spectacular Block ‘B’, consists of a sequence of three thick layers where amphibole and K-feldspar are aligned subparallel to layer bases. The base of each layer is undulatory on the scale of a single very coarse feldspar crystal. Figure 8. Hand drawn detailed map of Layer Breccia Zone ‘A’. Mapping by A. MacTavish (1992). Abbreviation key: fgr = fine-grained; mgr = medium-grained; cgr = coarse-grained; vcgr = very coarse-grained; f-cgr = fine- to coarse-grained; m-vcgr = medium-very coarse-grained; int = intermediate. - 89 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Sub-Stop C3l (528004E, 5402016N): A further 23m south of Sub-stop C3k is a zone characterized by well-developed syenite layering (Photo 15), some possible magmatic channel scours, some localized soft-sediment-style deformation, and a few zones of intense, localized layer disruption. Most of the layers within the southern part of the zone are quite flat lying (18°W dip). Photo 16. Crosscutting, vein-like body possibly resulting from the movement of volatile-rich magmatic fluids. Photo credit A. MacTavish. Photo 15. Well-developed, relatively flat-lying, magmatic layering within amphibole-nepheline syenite. Photo credit A. MacTavish. Sub-Stop C3m (528011E, 5402006N): This substop is located 12m southeast of Sub-stop C3l and is directly adjacent to it. It consists of a zone of disturbed and distorted syenite layering similar to that observed north of the isolated block observed at Sub-stop C3k. Contorted and convoluted layering is common and folding is observed locally with the disturbance increasing in intensity to the south. Most noticeable in this area is a deformed, crosscutting, texturally variable, vein-like body (Photo 16) composed of mobile “flowbanded” material. The margins of this “Vein C” (Fig. 9) are often irregular, possibly due to volatile fluid seepage (?) and it is often cored by coarse-grained to pegmatitic veinlets and pods. It is possible that this structure has erupted from the nose of a slump fold. Sub-Stop C3n (528020E, 5401977N): This final sub-stop is located 30m east-southeast of Sub-stop C3m and consists of a large slump-fold (Fig. 10) composed of medium- to very coarse-grained, modallyand normally grain-size graded syenite layers (Photo 17). This was interpreted as slump folding due to the presence of at least three axial planar directions present within three separate folds all in close proximity to each other. Unfortunately since mapping was completed in Figure 9. Hand drawn detailed map of Vein ‘C’. Mapping by A. MacTavish (1992). Abbreviation key: fgr = fine-grained; mgr = medium-grained; cgr = coarse-grained; vcgr = very coarse-grained; f-mgr = fine to medium-grained; c-vcgr = coarse to very coarse-grained; LM = Leopard mottles; peg = pegmatitic; int = intermediate 1992 this exposure has become partially obscured by the growth of lichen. - 90 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 tholeiitic lineage, contemporaneous with the Coldwell Complex. Fresh, metasomatized and hornfelsed, andesine-oligoclase basalt flows are estimated to attain a thickness of 5 km (Mitchell and Platt, 1994; Nicol, 1990). Assimilation and brecciation of the flows by subsequent gabbroic to syenitic magmatism has resulted in the widespread development of basaltic xenoliths ranging from 1m to over 1 km in size, comprising a roof pendant in the central part of the complex (Walker et al., 1992). Walker et al. (1992) subdivided these basaltic rocks into three main units: Figure 10. Hand drawn detailed map of the Area ‘D’ slump fold. Mapping by A. MacTavish (1992). Abbreviation key: fgr = fine-grained; cgr = coarse-grained; f-mgr = fine to medium-grained; f-cgr = fine to coarse-grained; int = intermediate 1. Aphanitic to fine-grained, massive, locally amygdaloidal (?) / ocellar basalt; 2. Medium-grained, diabasic (ophitic) basalt; and 3. Aphanitic to medium-grained, feldspar-phyric, diabasic (ophitic) basalt. At Wolf Camp Lake, aphanitic basalts contain round to amoeboid, epidote- and quartz-filled structures up to 2cm in diameter that have been interpreted as amygdules (Photo 18). Well-defined, amygdule-bearing zones dip 8° to the southwest in this vicinity (Walker et al., 1992). The basaltic roof pendant is locally underlain and enveloped by feldspar-phyric amphibole syenite and Fe-rich augite syenite. Photo 17. Lichen-obscured slump fold within layered amphibole-nepheline syenite. Photo credit A. MacTavish. Stop C4: Hornfelsed Basaltic Roof Pendants, Wolf Camp Lake UTM coordinates 541775E 5404189N 8.6 km west of the Highway 626 and Highway 17 junction Description: Hornfelsed basaltic rocks overlying the complex were recognized early in its mapping by Tuominen (1967) and Puskas (1970) and likely represent a volcanic edifice that has been subsequently eroded (Sage 1986). Mitchell and Platt (1994) and Nicol (1990) have considered these basalts to have a Photo 18. Amygdules within the basaltic roof pendant located near Wolf Camp Lake. Photo credit D. Campbell. Stop C5: Layered Fe-rich Augite Syenite (Alternate stop if time allows) UTM coordinates 544782E 5398443N 680 m west along the shoreline of Lake Superior from the end of the James River industrial road along the waterfront in Marathon; OR 150m south of Carden Cove road, 0.3 km past CPR tracks (park at 544864E, 5398750N) - 91 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Description: Broad expanses of glacially polished and wave-washed, massive Fe-rich augite syenite occur all along this part of the Lake Superior shoreline near Marathon. Fresh surfaces vary from dark green-brown to black, despite a buff to white weathered surface. Small dimension stone quarries were developed and produced in this area during the 1930’s. Much of the stone was shipped to larger centres in the American mid-west and Toronto. Fe-rich augite syenite (formerly referred to as ferroaugite syenite) comprises a large portion of the exposure in the eastern half of the Coldwell Complex. It appears to be a sheet-like intrusion that dips approximately 15° toward the center of the complex, sandwiched between the underlying Eastern Border Gabbro and an overlying, recrystallized amphibolequartz syenite; it also intrudes the basaltic roof pendants (Walker et al., 1992; 1993a). Crystallization of the syenite inwards from its upper and lower contacts produced mineralogical and compositional variations across it (Walker et al. 1993a). Constituent minerals include iridescent, lathlike, cryptoperthitic feldspar (up to 30% interstitial), and variable amounts of fayalite, amphibole, aenigmatite, and rare quartz. Coarse-grained to pegmatitic portions of the syenite host a variety of REE-bearing fluoro-carbonates, quartz, chalcedony, and molybdenite. Iridescent feldspar, known locally as “spectrolite”, was recently (2010) commercially extracted on a very small-scale from pegmatite at Shack Lake near Marathon. Although this unit is typically massive, rhythmic to chaotic layering is locally developed and where observed commonly dips shallowly towards the centre of the complex. At this site, layering strikes at 070° and dips 60° north. The layering is unusual in that it is defined by an intercumulus mineral (augite) rather that by cumulus phases (feldspar). at ~45°. This thickly layered sequence is underlain by massive gabbro near the contact with the Archean country rocks. The macrorhythmic layering is laterally discontinuous, pinching out over distances of 5 to 10m and contacts are sharp and conformable (Shaw, 1994, 1997). Rhythmic layering is modal and has been related to variation in the respective proportions of plagioclase (An60-35), augite (Fo67-43), minor orthopyroxene (En55-66), and Fe-Ti-oxides by Lum (1973). Modal plagioclase varies from approximately 60 to 80% in the leucocratic layers and 20 to 35% in the meso- to melanocratic layers (Shaw, 1994). A second band of layered gabbro, separated from the first by massive gabbro, is exposed on top of the long rock cut (Photo 19). Here, the macrorhythmic layering (Photo 20) produces relatively thin (1 to 5cm) to medium thick (5 to 100cm) layers that can be traced for over 35m along Photo 19. Macrorhythmic layering within the Eastern Border Gabbro. Photo credit M. Smyk. Stop C6: Layered Eastern (Border) Gabbro UTM coordinates 549199E 5398010N 1.7 km east of the Highway 626 and Highway 17 junction Description: Layering in the Eastern Border Gabbro shows distinct variations in style, is usually parallel to the eastern contact of the gabbro, and dips 20° to 60° toward the center of the complex (Shaw, 1994, 1997). At this stop, layering strikes approximately north and dips west towards the rest of the complex Photo 20. Macrorhythmic modal layering within the Eastern Border Gabbro. Photo credit A. MacTavish. - 92 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 strike. Layer contacts are sharp, locally scalloped and conformable. Trough cross-bedding has been noted on vertical faces by Shaw (1994). This stop is also close to the contact between the Eastern Border Gabbro and the Fe-rich augite syenite to the west. Pegmatitic syenite dykes intrude the gabbro at this locality and contain miarolitic cavities. McLaughlin (1990) has reported the presence of a variety of REE-bearing fluorocarbonates (bastnaesite, parisite, synchisite), Nb-bearing phases, and zircon in pegmatitic syenite with quartz, feldspar, and sodic amphibole. Stop C7 (Alternate Stop): Eastern Contact of the Coldwell Complex UTM coordinates 549656E 5396238N 3.3 to 3.6km east of the Highway 626 and Highway 17 junction Description: A number of highway rock cuts and outcrops expose the eastern contact of the Coldwell Complex with the enclosing Archean greenstone belt country rocks. Center 1 gabbros, the oldest rocks of the complex, form a ring dyke that forms the eastern and northern margins of the complex where it is in contact with Archean supracrustal and granitoid rocks. The reverse magnetization of these gabbros (Lilley, 1964) produces prominent magnetic “lows” on aeromagnetic maps. The most recent and comprehensive study of the Eastern Border Gabbro was conducted by Shaw (1994, 1997) who noted that more than 90% of the unit consists of layered gabbro. At this location, varitextured, unlayered Eastern Border Gabbro is in contact with, and contains numerous xenoliths of, Archean metasedimentary rocks. This has produced hybrid and contaminated phases and rheomorphic breccia. Crosscutting Center 1 syenite dikes are commonly pegmatitic. Amethystine quartz, calcite, and molybdenite occur in vugs within this chaotic contact zone. Disseminated iron- and copper-sulphides occur in biotite-rich, varitextured gabbro (Dunlop Occurrence), which has experienced sporadic exploration since the discovery of copper in the early 1950’s. It was last drilled in 1992 by Noranda Inc. with the best assay intervals grading 0.35% Cu/6.0m and 0.42% Cu/4.0m, respectively (Resident Geologist’s Files, Thunder Bay). A grab sample of rusty-weathering, moderately magnetic, fine- to medium-grained gabbro with coarse biotite and blebby chalcopyrite graded 5090ppm Cu, 494ppm Ni, 241ppm Zn, 8ppb Pd, 2ppb Pt and 22ppb Au (ibid). Overgrown pits are located just inside the tree line, west of the highway (UTM 549575E, 5396290N). Shaw (1994; 1997), Walker et al. (1993a, 1993b, 1993c), Currie (1980), and Tucker (1995) have documented a number of occurrences of rheomorphic breccia associated with the Eastern Border Gabbro along its intrusive, basal contact with the Archean supracrustal country rocks. Breccia units are characterized by chaotic flow fabrics that surround flow-oriented clasts situated in a medium-grained, granitic matrix. This unit has been somewhat enigmatic, having been alternatively described by earlier workers as conglomerate and ignimbrite (Resident Geologist’s Files, Thunder Bay). Similar exposures of this map unit also occur along the western contact of the complex, north of Middleton (cf. Wilkinson, 1983). Locally, pods of breccia vary from 20 to 75m in width and are up to 250m long. The breccia exposed along Highway 17 at this site contains mainly hornfelsed Archean clastic metasedimentary and metavolcanic rocks and massive vein quartz. In the vicinity of Two Duck Lake, the breccia contains fine-grained gabbro clasts (Tucker, 1995). The breccia varies from clast- to matrix-supported; the matrix consists of equigranular quartz, feldspar, and minor biotite, clino- and orthopyroxene, and opaque minerals; and tourmaline and prehnite overgrowths have been noted (Tucker, 1995). Rounded to angular clasts range in size from 0.5 to over 100cm and locally have developed 1 to 2cm wide, chlorite-rich reaction rims that are thickest where they are matrix-supported (Shaw, 1994). Magnetite and quartz¬feldspar-tourmaline veins cut both matrix and clasts. Quartzo-feldspathic rinds and crosscutting veinlets have been interpreted to be the result of partial melting of the felsic material during assimilation. The close association between rheomorphic breccia and the Eastern Border Gabbro suggests that the intrusion of the gabbro led to the brecciation and partial melting of the country rocks (Shaw, 1994, 1997; Tucker, 1995). - 93 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 PART 2: Marathon Cu-PGM Deposit David Good and John McBride Introduction to the Marathon Deposit The Cu-PGM sulphide mineralization of the Marathon deposit is hosted by the Two Duck Lake Gabbro, the latest mafic intrusive event and consequently the most continuous gabbroic body within the Eastern Gabbro Suite at the Marathon deposit. The Eastern Gabbro Suite, located around the eastern and northern margin of the Coldwell, was composed initially of a thick sequence of tholeiitic basalt that was subsequently intruded by a much larger volume of leucocratic to ultramafic intrusions that caused contact metamorphism of the basalt to pyroxene-hornfels grade (Good et al., 2015). All of these units are represented at the Marathon deposit (Fig. 1). The topography of the Coldwell is characterized by deep valleys and steep cliffs that form strong surface lineaments. Two lineaments at the Marathon deposit correspond to north dipping normal faults (north side down) with displacement of approximately 50 metres. Two Duck Lake Intrusion The Two Duck Lake intrusion is irregular in shape and elongated north-south (Fig. 2). The dip at the east contact is variable from nearly flat (at the south end) Figure 1. Geology of the Marathon deposit (after Good et al., 2015) highlighting location of field trip stops. Stops are marked with red dot and labelled as stop 1a, etc. Note two normal faults that correspond to strong surface lineaments (dashed lines) to vertical and locally overturned where the footwall overhangs the intrusion. The intrusion is composed of coarse-grained to pegmatitic olivine gabbro and troctolite. Modal layering is rare. The TDL gabbro was interpreted to have formed by intrusion of a nearly homogeneous plagioclase crystal mush by Good and Crocket (1994). But recent work suggests the intrusion formed by accumulation of several pulses of magma in a conduit setting (Good, Figure 2. 3d isometric view of the Two Duck Lake intrusion (from Good et al., 2015). Three coloured portions indicate blocks that were offset by normal faults with north side down by up to 60 metres. Note that numerous intrusions of mineralized Mt-Ol-Cpx-Ap rock (yellow) occur in the vicinity of major feeder zones, but those above the 6300 feeder but are not shown. - 94 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 2010; Ruthart, 2012; Good et al., 2015; and Shahabi Far, 2017). Multiple feeder channels were inferred by Good et al. (2015) to occur in the vicinity of several coincident features, including: deep V- or U-shaped channels in the footwall contact; topographic lineaments; very thick mineralized intervals; and irregular-shaped intrusions of olivine-magnetite-clinopyroxene-apatite. Age relationships Evidence suggests that all units in the Coldwell were emplaced within a short time interval between about 1108 Ma and 1105 Ma (Heaman and Machado, 1992; Good et al. in preparation). Age relationships, based on cross-cutting contacts and U-Pb age dating for the various mafic units are summarized in Figure 3. The metabasalt is interpreted to correlate with Mamainse Point Volcanic Group 1 (Fig. 4) which was emplaced at approximately 1108 Ma (Keays and Lightfoot, 2015). Two Duck Lake gabbro and associated breccia (Fig. 5) and occurs within several thick and continuous shallow-dipping lenses that parallel the footwall contact. The disseminated sulphides are concentrated in troughs along the footwall contact that approximately follow topographic lineaments (Fig. 6). The lenses are referred to as the Footwall, Main, and Hangingwall zones and the W Horizon. Sulfides in the Footwall, Main, and Hanging-wall zones consist predominantly of chalcopyrite and pyrrhotite with minor amounts of cubanite, bornite, pentlandite, cobaltite, and pyrite. Sulfides occur interstitial to primary silicates and also in association with hydrous silicates such as amphibole, chlorite, and minor serpentine (Watkinson and Ohnenstetter, 1992; Samson et al., 2008). Chalcopyrite occurs as separate grains or as rims on pyrrhotite grains. Some chalcopyrite is intergrown with highly calcic plagioclase (An70–An80) in replacement zones at the margins of plagioclase crystals (Good and Crocket, 1994; Shahabifar, 2016). The metabasalt was subsequently intruded by the following units, listed in order from oldest to youngest, layered troctolite sill of the Marathon Series, gabbroic anorthosite and olivine gabbro of the Layered Series, Two Duck Lake gabbro and various ultramafic units composed of magnetite +/- olivine +/- apatite +/clinopyroxene of the Marathon Series, Malpa Lake intrusion, and syenite. The W horizon is characterised by extreme PGE enrichment relative to Cu with several 2m thick drill hole intersections having 20 to 70 ppm Pd and Cu/Pd as low as 3 (e.g., Fig. 7, top and bottom photos). The best intersection contains 34 ppm Pd and 9.6 ppm Pt over 10 m. Mass balance considerations, assuming initial magma contained 10 ppb Pd, would require a magma column on the order of 34 km to generate the 34 ppm Pd in this interval. Disseminated sulfide mineralization is hosted by the The W Horizon is commonly difficult to identify in drill core because it typically contains only trace sulfides, but if sulfides are present, they consist of Mineralization Figure 3. Relative timing of mafic metavolcanic and intrusive events (age dates after Good et. al, in preparation) in the Eastern Gabbro Suite of the Coldwell Alkaline Complex. - 95 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 4. Correlation diagram showing range of ages for the Coldwell units compared to volcanic and intrusive units in the Midcontinent Rift (after Keays and Lightfoot, 2015). Figure 5. Stratigraphic section through the Main zone and overlying troctolite sill. Note the saw tooth pattern for Cu, Pd and Cu/Pd indicating individual pulses of sulphide-bearing crystal slurry. Unit 2d, breccia of metabasalt blocks and Two Duck Lake gabbro; unit 3bd, coarse grained ophitic and pegmatitic Two Duck Lake gabbro; unit 4a, breccia of footwall blocks and Two Duck Lake gabbro (from Good et al., 2015). - 96 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 6. Three versions of top view for the Marathon deposit showing 3d topography (green surface) and contoured footwall surface models. Note the troughs and ridges (left hand image) correspond to surface lineaments. Note the higher-grade assays for Cu (>0.5%) and Pd (>3 ppm) are aligned within zones that parallel troughs within the 3d footwall surface model. Figure 7. PGE rich samples from the W Horizon contain fresh clinopyroxene, olivine and plagioclase. Top photo sample with 107 ppm Pd+Pt+Au and 203 ppm Cu. Bottom photo sample with 70 ppm Pd+Pt+Au and 0.86 % Cu. chalcopyrite and bornite with minor pyrrhotite and trace amounts of pentlandite, cobaltite, and pyrite (Ruthart, 2012). Deposit Model Exploration strategies in the Coldwell are based on the conduit model and a schematic magmatic plumbing system such as that envisioned by Barnes et al. (2016) - 97 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 (Fig. 8). Evidence for a magma conduit setting at the Marathon deposit were described by Good et al. (2015), and include: consistent with rotation of sill from west dipping in the north to sub horizontal or north dipping in the south. • association with metavolcanic rocks • fault control (mineralization parallel to topographic lineaments) • brecciation and assimilation • accumulation in trough setting • flow through PGE upgrading • tube shaped intrusion • gravity driven back flow (Mt-Ol-Cpx-Ap cumulates) • high heat flux (wide zone of pyroxenehornfels grade metavolcanic rocks) Figure 9. Map at south end of Marathon deposit showing location of stops 1 and 2. Figure 8. Step 3 of schematic illustration for magmatic plumbing system (after Barnes et al., 2016) Field Trip Stops Stop 1a: South end of Troctolite Sill (Figs. 10b) UTM coordinates 549995E 5403560N Trench exposure with coarse-grained mottled augite troctolite shows large fresh oikocrysts of olivine (brown), clinopyroxene (black) and magnetite (black) and subhedral plagioclase (white). The layered troctolite sill is an important marker horizon because it occurs just above the top of the Main Mineralized zone and is an indicator of the relative fault offset that occurred along E-W–trending normal faults at 5404500 and 5404900 North (Fig. 1). The sill dips moderately west at the north end, but flattens out in the south to sub-horizontal (Fig. 9b). Note layering is approximately east west at Stop 1a, Figure 10. 3d image (iso view) of geology at south end of Marathon deposit showing location of stops 1 and 2 on trenched outcrops (black polygons): (a) shows orientation of footwall surface troughs approximately perpendicular to contact, and the red centre line of the W horizon at surface on the splat trench; (b) top (light blue) and bottom surfaces (dark blue) of the troctolite sill. Gap in the troctolite sill surfaces represents location where W Horizon and TDL gabbro cuts the through the sill; (c) surface model of W Horizon (yellow). - 98 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 1b: Southeast corner of Marathon deposit (Figs. 10 and 11) UTM coordinates 550090E 5403395N Trench outcrop exposure of Two Duck Lake gabbro within a shallow dipping bowl-shaped depression in the footwall (Fig. 10a) includes W Horizon and Main zone type mineralization. Figure 11. Plan map of trench at the southeastcorner of the Marathon deposit including assay table for samples 44 to 56 located in channel immediately north of the historic trench. Red circle marks location of historic trench (ca. mid 1960’s) with high copper mineralization. The unit was not assayed for Pd until 2005. The channel sample located just north of the red circle returned assays of 3.37 ppm Pd+Pt+Au, and 0.35% Cu over 18.6 m. East-west layering in TDL gabbro is visible just south of trench. Outcrop shows textural evidence for cross cutting intrusions of subophitic olivine gabbro. - 99 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 2a: Splat trench and the W Horizon (Fig.12) UTM coordinates 549900E 5403804N Stops 2a and 2b are located on the east and northwest branches of the splat trench, respectively, and highlight two locations along the W horizon as it drapes over top of a north plunging ridge in the footwall. Figure 12. Plan map of the stripped outcrop on the east arm of the splat trench highlights the distinct coarse-grained to pegmatitic subophitic olivine gabbro of the Two Duck Lake intrusion. Stratigraphic top of the section is to the northeast (compare to Fig. 9a and 9c). Note the large xenolith/breccia of metavolcanic rock along the east edge of the outcrop. Codes: 3a, medium-grained (1-5 mm) Two Duck Lake gabbro; 3b, coarse-grained (5mm to 1cm); 3d, pegmatitic gabbro; 3f, magnetite and clinopyroxene rich gabbro; 4, breccia; 2a, metavolcanic rock. Mineralization consists of disseminated cpy, bn and minor po. Assay table for samples 9 to 24 within the saw-cut channel on the outcrop have an average grade of 2.64 g/t Pd+Pt+Au and 0.1% Cu over 25.1 m. - 100 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 2b: Splat trench and the W Horizon (Fig. 13) UTM coordinates 549810E 5403825N Stop 2b: Malachite zone - Splat trench Figure 13. Plan map of the northwest outcrop on the splat trench highlights the distinct coarse-grained to pegmatitic subophitic olivine gabbro of the Two Duck Lake intrusion. Stratigraphic top of the section is to the north (compare to Fig. 9a and 9c). Note the large xenolith/breccia zone of metavolcanic rock along the east edge of the outcrop. Codes: 3a, mediumgrained (1-5 mm) Two Duck Lake gabbro; 3b, coarse-grained (5mm to 1cm); 3d, pegmatitic gabbro; 3f, magnetite and clinopyroxene rich gabbro; 4, breccia; 2a, metavolcanic rock. Mineralization consists of disseminated cpy, bn and minor po. Assay table for samples 84 to 93 within the saw-cut channel on the outcrop have an average grade of 2.13 g/t Pd+Pt+Au and 0.36% Cu over 17 m. References in ferroan olivine gabbros of the Coldwell Complex, Ontario; in The Geology, Geochemistry, Mineralogy and Mineral Beneficiation of Platinum-Group Elements. Edited by L.J. Cabri; Canadian Institute of Mining and Metallurgy and Petroleum, Special Volume 54, p.321-337. Alexander, M. 2007. The mineralogy of NYF pegmatites from the Coldwell Alkaline Complex, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario. Barnes, S.J., Cruden, A.R., Arndt N., Saumur B.M. 2016. The mineral system approach applied to magmatic Ni–Cu–PGE sulphide deposits, Ore Geology Reviews 76, 296-316. Barrie, C. Tucker, MacTavish, A.D., Walford, P.C., Chataway, R., and Middaugh, R., 2002. Contacttype and magnetite reef-type Pd-Cu mineralization Bell, K. and Blenkinsop, J. 1980. Grant 42: Ages and initial 87Sr-86Sr ratios from alkaline complexes of Ontario; in Geoscience Research Grant Program, Summary of Research, 1974-1980, Ontario Geological Survey, Miscellaneous Paper 93, p.16-23. Bohay, T.J. 1997. The Coldwell alkaline complex, Ontario: Magmatic affinity as determined by an isotopic - 101 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 and geochemical study; unpublished MSc thesis, McMaster University, Hamilton, Ontario, 135p. and deposits in northwestern Ontario; Ontario Geological Survey, Open File Report 5889, 145p. Coates, M.E. 1970. Geology of the Killala–Vein lakes area, Ontario; Ontario Department of Mines, Geological Report 81, 35p. Keays, R.R. and Lightfoot, P.C. 2015. Geochemical Stratigraphy of the Keweenawan Midcontinent Rift Volcanic Rocks with Regional Implications for the Genesis of Associated Ni, Cu, Co, and Platinum Group Element Sulfide Mineralization, Econ Geol, 110, 1235-1267. Good, D.J. 1992. Genesis of copper-precious metal sulfide deposits in the Port Coldwell alkalic complex, Ontario; unpublished PhD thesis, McMaster University, Hamilton, Ontario, 203p. Good, D.J. 1993. Genesis of copper-precious metal sulfide deposits in the Port Coldwell Alkalic Complex, Ontario Geoscience Research Grant Program, Grant No. 341, Ontario Geological Survey, Open File Report 5839, 23. Good, D.J. and Crockett, J.H. 1994a. Genesis of the Marathon Cu-platinum-group element deposit, Port Coldwell alkaline complex, Ontario: A Midcontinent Rift-related magmatic sulfide deposit; Economic Geology, v.89, p.131-149. Good, D.J. and Crockett, J.H., 1994b. Origin of albite pods in the Geordie Lake gabbro, Port Coldwell alkaline complex, northwestern Ontario: Evidence for latestage hydrothermal Cu-Pd mineralization; The Canadian Mineralogist, v.32, p.681-701. Good, D.J, Epstein, R., McLean, K., Linnen, R.L., & Samson, I.M. 2015. Evolution of the Main Zone at the Marathon Cu-PGE Sulfide Deposit, Midcontinent Rift, Canada: Spatial Relationships in a Magma Conduit Setting, Econ Geol v.110, p.983-1008. Hauck, S.A., Severson, M.J, Zanko, L., Barnes, S.-J., Morton, P., Alminas, H., Foord, E.E. and Dahlberg, E.H. 1997. An overview of the geology and oxide, sulfide and platinum-group element mineralization along the western and northern contacts of the Duluth Complex; Geological Society of America, Special Paper 312, p.137-185. Heaman, L.M., Easton, R.M., Hart, T.R., Hollings, P., MacDonald, C.A. and Smyk, M. 2007. Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon Region, Ontario; Canadian Journal of Earth Sciences, v.44, no.8, p.1055-1086. Heaman, L.M. and Machado, N. 1987. Isotope geochemistry of the Coldwell alkaline complex: 1. U-Pb studies on accessory minerals; Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Saskatoon, Saskatchewan, Program with abstracts, p.54. Heaman, L.M. and Machado, N 1992. Timing and origin of the Midcontinent Rift alkaline magmatism, North America: Evidence from the Coldwell complex; Contributions to Mineralogy and Petrology, v.110, p.289-303. Hinz, P. and Landry, R. 1994. Industrial mineral occurrences Kerr, H.L. 1910a. Geological map of part of the north shore of Lake Superior, District of Thunder Bay; Ontario Bureau of Mines, Annual Report Map 19B, scale 1:63 360. Kerr, H.L. 1910b. Nepheline syenites of Port Coldwell; Ontario Bureau of Mines, Annual Report, v.19, p.194-232. Kissin, S.A. and McCuaig, T.C. 1988. The genesis of silver vein deposits in the Thunder Bay area, northwestern Ontario: Geoscience Research Grant Program, Summary of Research, 1987-1988; Ontario Geological Survey, Miscellaneous Paper 140, p.146156. Klasner, J.S., Cannon, W.F. and Van Schmus, E.R. 1982. The Pre-Keweenawan tectonic history of the southern Canadian Shield and its influence on the formation of the Midcontinent Rift; in Geology and Tectonics of the Lake Superior Basin, Geological Society of America, Memoir 156, p.27-46. Lewchuk, M.T. and Symons, D.T.A. 1990. Paleomagnetism of the late Precambrian Coldwell complex, Ontario, Canada; Tectonophysics, v.184, p.73-86. Laderoute, D.G. 1987. The petrology, geochemistry, and petrogenesis of alkaline dyke rocks from the Coldwell Alkaline complex; unpublished M.Sc. Thesis, Lakehead University, Thunder Bay, Ontario, 89p. Tectonophysics, v.184, p.73-86. Lilley, F.E.M. 1964. An analysis of the magnetic features of the Port Coldwell intrusion; unpublished BSc thesis, University of Western Ontario, London, Ontario, 89p. Lukosius-Sanders, J. 1988. Petrology of the syenites from Center III of the Coldwell alkaline complex, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 141p. Lum, H.K. 1973. Petrology of the eastern gabbro and associated sulphide mineralization of the Coldwell alkaline complex, Ontario; unpublished BSc thesis, Carleton University, Ottawa, Ontario, 68p. MacTavish, A. 2000. A new style of PGE mineralization within the Coldwell alkaline complex, northwestern Ontario; Ontario Exploration and Geoscience Symposium, Toronto, December 11-12, 2000, Speaker Abstracts, p.3. MacTavish, A., Lukosius-Sanders, J. and Jowett, R. 1987. Geological report of the Joa Option (Geordie Lake - 102 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 property), St. Joe Canada Inc.; unpublished report, Resident Geologist’s Files, Thunder Bay, 7p. MacTavish, A Smyk, M., Good, D., and McBride, J., 2017. Transect Through the Coldwell Alkaline Complex In; MacTavish, A. and Hollings, P. (Eds.), Institute on Lake Superior Geology Proceedings, 63rd Annual Meeting, Wawa, Ontario, Part 2 - Field trip guidebook, v.63, part 2, 1-44. McLaughlin, R.M. 1990. Accessory rare metal mineralization in the Coldwell alkaline complex, northwest Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 123p. Miller, J.D., Jr., Nicholson, S.W., and Cannon, W.F. 1995. The Midcontinent rift in the Lake Superior region, in Miller, J.D., Jr., ed., Field trip guidebook for the geology of ore deposits of the Midcontinent rift in the Lake Superior region; Minnesota Geological Survey Guidebook Series, no. 20, p.1-22. Mitchell, R.H. and Platt, G. R. 1978. Mafic mineralogy of ferroaugite syenite from the Coldwell alkaline complex, Ontario, Canada; Journal of Petrology, v.19, p.627-651. Mitchell, R.H. and Platt, G. R. 1982a. The Coldwell alkaline complex; in Field Trip Guidebook, Proterozoic geology of the northern Lake Superior area, Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Winnipeg, Manitoba, p.42-61. Mitchell, R.H. and Platt, G. R. 1982b. Mineralogy and petrology of nepheline syenites from the Coldwell alkaline complex, Ontario, Canada; Journal of Petrology, v.23, p.186-214. Mitchell, R.H. and Platt, G. R. 1994. Aspects of the geology of the Coldwell alkaline complex: Field trip A2, Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Waterloo, Ontario, 36p. Mitchell, R.H., Platt, G.R., Lukosius-Sanders, J., ArtistDowney, M. and Moogk-Pickard, S. 1993. Petrology of syenites from Center III of the Coldwell alkaline complex, northwestern Ontario, Canada; Canadian Journal of Earth Sciences, v.30, p.145-158. Mitchell, R.H., Platt, R.G. and Cheadle, S.P. 1983. A gravity study of the Coldwell complex, northwestern Ontario, and its petrological significance; Canadian Journal of Earth Sciences, v.20, p.1631-1638. Mulja, T. 1989. Petrology, geochemistry, sulphide- and platinum-group element mineralization of the Geordie Lake intrusion; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 234p. Mulja, T. and Mitchell, R.H. 1990. Platinum-group minerals and tellurides from the Geordie Lake intrusion, Coldwell complex, northwestern Ontario; Canadian Mineralogist, v.28, p.489-501. Mulja, T. and Mitchell, R.H. 1991. The Geordie Lake intrusion, Coldwell Complex, Ontario: Palladiumand tellurium-rich disseminated sulfide occurrence derived from an evolved tholeiitic magma; Economic Geology, v.86, p.1050-1069. Nicol, D.N. 1990. Assimilation of basic xenoliths with Center 3 syenites of the Coldwell Complex, Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 59p. Ohnenstetter, D., Watkinson, D.H. and Dahl, R. 1991. Zoned hollingworthite from the Two Duck Lake intrusion, Coldwell complex, Ontario; American Mineralogist, v.76, p.1694-1700. Penczak, R.S. 1992. Petrology and mineral chemistry of the Middleton copper occurrence of the Western gabbro, Coldwell alkaline complex, Ontario; unpublished BSc thesis, University of Waterloo, Ontario. Pollock, S.J. 1987. The isotopic geochemistry of the Prairie Lake carbonatite complex; unpublished MSc thesis, Carleton University, Ottawa, Ontario, 71p. Potter, E.G. 2004. The rare and exotic mineralogy of the western subcomplex of the Dead Horse Creek diatreme, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario. Puskas, F.W. 1967. Port Coldwell area; Ontario Department of Mines, Preliminary Map P.114, scale 1:31 680. Puskas, F.W. 1970. The Port Coldwell alkali complex; in Proceedings, 16th Institute on Lake Superior Geology, Thunder Bay, Ontario, p.87-100. Queen, M., Heaman, L.M., Hanes, J.A., Archibald, D.A. and Farrar, E. 1996. 40Ar/39Ar phlogopite and U-Pb perovskite dating of lamprophyre dykes from the eastern Lake Superior region: Evidence for a 1.14 Ga magmatic precursor to Midcontinent Rift volcanism; Canadian Journal of Earth Sciences, v.33, p.958-965. Ruthart R. 2012. Characterization of high-PGE, low-sulphur mineralization at the Marathon PGE-Cu deposit, Ontario: M.Sc. thesis, Waterloo, ON, University of Waterloo, 145 p. Sage, R.P. 1982. Mineralization in diatreme structures north of Lake Superior; Ontario Geological Survey, Study 27, 79p. Sage, R.P. 1985. Geology of carbonatite-alkaline rock complexes in Ontario: Chipman Lake area; Ontario Geological Survey, Study 44, 40p. Sage, R.P. 1986. Alkalic rock complexes – carbonatites of northern Ontario and their economic potential; unpublished PhD thesis, Carleton University, Ottawa, Ontario, 335p. Sage, R.P. 1987. Geology of carbonatite-alkaline rock complexes in Ontario: Prairie Lake carbonatite - 103 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 complex, District of Thunder Bay; Ontario Geological Survey, Study 46, 91p. Tuominen, H.V. 1967. Port Coldwell area; Ontario Department of Mines, Map P.114, scale 1:15 840. Sage, R.P. 1991. Alkaline rock, carbonatite and kimberlite complexes of Ontario, Superior Province; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p. 683-709. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T. and Penczak, R.S. 1992. Geology of the Port Coldwell alkaline complex; in Summary of Field Work, 1992, Ontario Geological Survey, Miscellaneous Paper 160, p.108-119. Sage, R.P. and Watkinson, D.H. 1995. Alkalic rocks of the Midcontinent rift; Institute on Lake Superior Geology, Marathon, ON, Proceedings Volume 41:2A, 79p. Samson, I.M., Fryer, B.J., and Gagnon, J.E. 2008. The Marathon Cu-PGE deposit, Ontario: Insights from sulphide chemistry and textures, in Goldschmidt conference, p. 820. Shaw, C.S.J. 1994. Petrogenesis of the eastern gabbro, Coldwell alkaline complex, Ontario; unpublished PhD thesis, University of Western Ontario, London, Ontario, 292p. Shahabi Far, M.. 2016. The magmatic and volatile evolution of gabbros hosting the Marathon PGE-Cu deposit: evolution of a conduit system, PhD thesis, University of Windsor, Ontario. Shaw, C.S.J.1997. The petrology of the layered gabbro intrusion, eastern gabbro, Coldwell alkaline complex, northwestern Ontario, Canada: Evidence for multiple phases of intrusion in a ring dyke; Lithos, v.40. p.243-259. Smyk, M.C., Taylor, R.P., Jones, P.C. and Kingston, D.M. 1993. Geology and geochemistry of the West Dead Horse Creek rare-metal occurrence, northwestern Ontario; Exploration and Mining Geology, v.2, no.3, p.245-251. Smyk, M.C. and Sage, R.P. 1995. Geology and mineralization of intrusive complexes of the Marathon, Ontario area; in Field Trip Guidebook for the Geology and Ore Deposits of the Midcontinent Rift in the Lake Superior region, International Geological Correlation Program, Project 336, Field Conference and Symposium, Duluth, Minnesota, August 19 to September 1, 1995, p.182-193. Tucker, C. 1995. Origin of breccia associated with the Eastern Gabbro, Coldwell alkaline complex, northwestern Ontario; unpublished BSc thesis, University of Western Ontario, London, 57p. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T. and Penczak, R.S 1993a. Precambrian geology of the Coldwell Alkaline Complex; Ontario Geological Survey, Open File Report 5868, 30p. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T. and Penczak, R.S 1993b. Precambrian geology, Port Coldwell complex, west half; Ontario Geological Survey, Preliminary Map P.3232, scale 1:20 000. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T. and Penczak, R.S 1993c. Precambrian geology, Port Coldwell complex, east half; Ontario Geological Survey, Preliminary Map P.3233, scale 1:20 000. Watkinson, D.H., Whittaker, P.J. and Jones, P.L. 1983. Platinum group elements in the eastern gabbro, Coldwell complex, northwestern Ontario; Ontario Geological Survey, Miscellaneous Paper 113, p.183191. Watkinson, D.H. and Ohnenstetter, D. 1992. Hydrothermal origin of platinum-group mineralization in the Two Duck Lake intrusion, Coldwell Complex, Northwestern Ontario: Canadian Mineralogist, v. 30, p. 121-136. Weiblen, P.W. 1982. Keweenawan intrusive rocks; Geological Society of America Memoir 156, p.57-82. Wilkinson, S.J. 1983. Geology and sulphide mineralization of the marginal phases of the Coldwell complex, northwestern Ontario; unpublished MSc thesis, Carleton University, Ottawa, Ontario, 129p. Wu, F.Y., Mitchell, R.H., Li, Q-L., Zhang, C., and Yang, Y-H. 2017. Emplacement age and isotopic composition of the Prairie Lake Carbonatite complex, Northwestern Ontario, Canada. Geological Magazine 154(2): 217236. - 104 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 7 - Building and ornamental stone sites of the Marathon Area, Ontario Peter Hinz Mineral Exploration & Development Section, Ontario Ministry of Energy, Northern Development and Mines, 435 James Street South, Suite B002 Thunder Bay, Ontario, P7E 6S7, Canada Foreward This tour will examine two past-producing “granite” quarries, two “granite” exploration sites in the Marathon area, one ornamental stone occurrence and, time permitting, the possible source of the recently popular “yooperlite” cobbles found in the Upper Peninsula of Michigan. Introduction (taken from Hinz et al., 1994) The dimension and monument stone industry in northwestern Ontario has a long history and is linked to the development and prosperity of the region. One of the earliest commercial operations was located on Vert Island in Nipigon Bay of Lake Superior. The Mesoproterozoic Sibley Group yielded an attractive red sandstone which was extracted by the Chicago Verte Island Sandstone Company. The stone was used locally for the construction of the Canadian National Railway and shipped to Chicago, Winnipeg, southern Ontario, and other U.S. cities for construction uses. Development of some of the earliest quarries in the Marathon and Nipigon areas was directly related to the construction of the Canadian Pacific Railway in the late 1880’s. Syenites of the Coldwell Alkaline Complex in the vicinity of Marathon and sandstones south of Nipigon were used in the construction of railway trestles to span the Black, Pic, Little Pic, Steel, and Nipigon rivers. Today these trestles show very little wear and are a testament to the long-standing durability of the stones. Although markets for dimension stone decreased in the early 1900’s, production continued at the Simpson Island sandstone quarry (1900 to 1910) and at the Bannerman and Horne quarry (1912 to 1915) near Ignace. The next period of quarry development took place during the late 1920’s to early 1930’s. Five small scale quarries operated northwest of Marathon along the Canadian Pacific Railway. Black and brown granites were extracted and shipped to customers in Toronto, Buffalo, Chicago, and Detroit. In 1932, the last of these quarries closed due to the loss of market. opened a quarry approximately 12 km southwest of the town of Vermilion Bay. This highly popular pink granite began production in 1954 and continued sporadically under various names until 1991 when the quarry, now named Granite Quarriers (GQI) Inc., closed. In 1981, Nelson Granite Limited of Sussex, New Brunswick began production of an identical granite from a quarry immediately south of the highway from the Granite Quarriers Inc. site. This quarry has operated year-round since that time and is still in production. Currently, 2019, Nelson Granite Limited is the only stone producer operating in northwestern Ontario. Nelson Granite produces a range of colours including pink, yellow, green, brown, and white granite from four quarries located north of Kenora and west of Vermilion Bay. Northwestern Ontario stone is shipped around the world for a range of uses including: building stone for interior and exterior uses; monumental stone; and landscape uses including pavers, benches, and accent pieces. Detailed descriptions of the historic quarries, their operations, geology and geotechnical test results are provided in Hinz et al. (1994). Descriptions of current producers are available in the Kenora portion of the Report of Activities 2018 (Paterson et al., 2019). Geologic setting Puumala (2018) provides the following general geological description of the Coldwell Alkaline Complex. “The geology of this area is dominated by rocks of the Coldwell Alkaline Complex. The Coldwell Complex was emplaced into Neoarchean supracrustal rocks of the Wawa Subprovince of the Superior Province during the Mesoproterozoic Midcontinent Rift event at ca. 1108 +/- 1 Ma (Heaman and Machado, 1992). The complex approximately bisects the Schreiber-Hemlo greenstone belt and is located at the north end of the Thiel fault, a zone of faulting which separates grabens with different subsidence history in the rift (Cannon et al., 1989). In 1948, the Vermilion Pink Granite Company - 105 - The Coldwell Complex was mapped by Kerr (1910a, �Proceedings of the 65th ILSG Annual Meeting - Part 2 1910b), Puskas (1967), and Walker et al. (1993), and comprises three, superimposed ring sub-complexes or magmatic centers (Mitchell and Platt, 1978) that young progressively (Centers 1 to 3) to the southwest. The rocks of Center 1 are silica-saturated and include the Eastern and Western border gabbros (the oldest rocks within the complex) and later iron-rich augite syenite and syenite-syenodiorite (Mitchell and Platt, 1978, 1982; Mulja, 1989). Center 2 includes silica-undersaturated alkalic rocks with oversaturated residue. Rock types include nepheline- and hastingsite-bearing miaskitic nepheline syenite, and numerous volumetrically minor alkaline lamprophyre and analcime tinguaite dykes (Mitchell and Platt, 1978, 1982; Laderoute, 1987; and Mulja, 1989). Center 3 includes silica-oversaturated alkalic rocks with oversaturated residue consisting of magnesio-hornblende syenites, quartz syenites, and minor granites (Mitchell and Platt, 1994; LukosiusSanders, 1988). The map area covers much of the Eastern border gabbro, which hosts numerous occurrences of Magmatic Cu-Ni-PGE (± Au, Ag) and Ti-V±apatite mineralization (MacTavish and Smyk, 2017). The Marathon Cu-PGE deposit is the most notable example of the deposits hosted in the Eastern border gabbro. Another gabbroic intrusion within the interior of the Coldwell Complex hosts the Geordie Lake Cu-PGE deposit (MacTavish and Smyk, 2017). Centre 1 augite and amphibole syenite has previously been quarried as dimension stone in Marathon, just to the south of the Coldwell Complex map area (Hinz et al., 1994), and these rocks continue to see periodic exploration interest. Late-stage syenite pegmatites that host occurrences of Nb-Y-F-rare earth elements also occur in the area (Alexander, 2007).” The current field trip stops will feature rocks of all intrusive centres as shown in Figure 1: 1. Center 1: Stops 1, 2, 3, and 4; 2. Center 2: Stops 5 and 6; 3. Center 3: Stop 7. MacTavish and Smyk (2017), Walker et. al. (1993) and Hinz et. al. (1994) provide descriptions of the lithologies which will be visited. From MacTavish and Smyk (2017), “Fe-rich augite syenite (formerly referred to as ferroaugite syenite) comprises a large portion of the exposure in the eastern half of the Coldwell Complex. It appears Figure 1. Coldwell Complex maps showing field trip stop locations within the three intrusive centres. to be a sheet-like intrusion that dips approximately 15° toward the center of the complex, sandwiched between the underlying Eastern Border Gabbro and an overlying, recrystallized amphibole-quartz syenite; it also intrudes the basaltic roof pendants (Walker et al., 1992, 1993a). Crystallization of the syenite inwards from its upper and lower contacts produced mineralogical and compositional variations across it (Walker et al.; 1993a). Constituent minerals include iridescent, lathlike, cryptoperthitic feldspar (up to 30% interstitial), and variable amounts of fayalite, amphibole, aenigmatite, and rare quartz. Coarsegrained to pegmatitic portions of the syenite host a variety of REE-bearing fluoro-carbonates, quartz, chalcedony, and molybdenite. Iridescent feldspar, known locally as “spectrolite”, was recently (2010) commercially extracted on a very small-scale from pegmatite at Shack Lake near Marathon. Feldsparporphyritic amphibole syenite contains two textural variants, a feldspar porphyritic amphibole syenite with an aphanitic to medium-grained groundmass and interstitial amphibole; and a later intrusion of mediumgrained amphibole syenite with columnar feldspar and interstitial amphibole.” Walker et al. (1993) stated that “the amphibolenepheline syenite (Unit 13) is white to red, mesocratic to leucocratic, medium-grained with variable proportions of feldspar, nepheline, amphibole, biotite, apatite, and zeolites. Locally the nepheline syenite is well-layered - 106 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 with melanocratic olivine-nepheline syenite grading into mesocratic syenite. Spectacular orbicular layering occurs on the south shore of Pic Island. An intergranular texture resulting from intergrown feldspar, amphibole, and nepheline is typical of the unit. Near lineaments and lithological contacts the amphibole-nepheline syenite becomes red. Texturally different varieties of amphibole-nepheline syenite occur near the contacts and include mesocratic nepheline-amphibole syenite with near-equant euhedral amphibole prisms and mesocratic amphibole-nepheline syenite with interstitial amphibole and euhedral columnar feldspar. The amphibole-natrolite-nepheline syenite (Unit 14) is an extremely variable rock unit that intrudes the roof pendant mafic volcanics, gabbro, iron-rich augite syenite, amphibole syenite, and amphibolenepheline syenite. The main rock type within this unit is a gray to pink, mesocratic, amphibole-natrolitenepheline syenite with variable amounts of natrolite, lath feldspar and acicular amphibole. The textural complexities of the amphibole-natrolite-nepheline syenite is considered to be a product of assimilation and mixing of a variety of rock compositions in a solid, semi-molten or molten state and synplutonic intrusion of the alkaline gabbro.” From Hinz et al. (1994), “Amphibole-natrolitenepheline syenite: contains primarily anhedral, “turbid” crystals of perthitic feldspar. The turbid areas are caused by the presence of numerous vacuoles. The reddish colour of the stone may be due to ironstaining of the vacuoles and fractures within the feldspar crystals. Anhedral pyroxene (augite), biotite, amphibole (hornblende), and opaque minerals occur together.” Field Trip Stops Field trip stops are illustrated in Figure 2. Stop 1: Peninsula Quarry (1927-1932) UTM coordinates 544191E 5399826N In the 1880’s, prior to commercial production, several small quarries were operated supplying stone for the construction of river abutments and railway trestles in support of the construction of the Canadian National Railway. In 1927, “commercial operations were initiated by Peninsula Granite Quarries Ltd. on 17 claims located on the east shore of Carden Cove” (Hinz et al., 1994). Figure 2. Geology of the Coldwell Alkaline Complex with field trip stop locations. - 107 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Peninsula Granite Quarries Ltd. operated four quarries at various points along Carden Cove, north of the town of Marathon. Most of the historic infrastructure related to the operations is lost however remains of the original derrick, grout shovel, and steam engine can still be seen on-site (Figs. 3 and 4). sheet below to a depth of 10ft did not reach another sheeting plane. The rift is roughly parallel to the sheeting planes.” Stop 2: Coldspring Quarry (1931-1938?) UTM coordinates 545332E 5398469N From Hinz et al. (1994): “The ground covering the black granite was purchased by the Cold Spring Granite Company from the Peninsula Granite Quarries Ltd. During the year, a new quarry was opened with a new derrick, drilling equipment and power plant. Twelve men were employed to quarry blocks that weighted up to 35 tons. In 1931, twenty car loads of black granite were sent to Cold Spring, Minnesota for fabrication (Thomson, 1932; The Northern Miner, 1931). In the late 1930s the quarry operations were abandoned due to the lack of market.” The claims are underlain by Fe-rich augite syenite as described above in Walker et al. (1992, 1993a) and MacTavish and Smyk (2017). The stone is mediumto coarse-grained, dark brown to black in fresh-cut surfaces. Two sets of jointing are documented, the most prominent is 335° with a dip of 50° south and a second is 005° with a vertical dip. Sheeting (horizontal) fractures range from 0.45m to 2.4m and dip 8-10° west (Thomson, 1932). Figure 3. Remains of the derrick at the Peninsula Quarry. Stop 3: Shack Lake Spectrolite (1963-present) UTM coordinates 546698E 5399605N Figure 4. Remains of the steam winch at the Peninsula Quarry. The Peninsula Quarry site is underlain by iron-rich augite syenite and amphibole syenite as described above in Walker et al. (1992, 1993a) and MacTavish and Smyk (2017).” Thomson (1931) described the jointing, “Two sets of joints are seen in the quarry. The most prominent strikes almost due north and varies in dip from vertical to 70°W. The cross-jointing strikes east-west and is nearly vertical. At the quarry the north-south joints are 70ft apart and run parallel for at least 500ft. Rectangular blocks of a size limited only by plant capacity can be quarried. The sheets lie horizontally and exhibit an even and well-defined floor. The first sheet quarried had a maximum thickness of 14ft. Drilling in the next From Hinz et al. (1994): “The Shack Lake occurrence was first staked in 1963 by C.S. Downey, sporadic exploration work including diamond drilling and blasting was conducted on the site since then. For a time in the early 1990’s the property owners of the time, Jon and Audrey Ferguson considered opening a “pickyour-own” operation similar to the amethyst operations in the Thunder Bay area, this never came to fruition. At the time the Town of Marathon adopted spectrolite as the “town mineral”. Currently the property is held by G. Blakely who has conducted additional diamond drilling and blasting. Spectrolite is a variation of labradorite, with a deeper and wider range of colours (full spectrum) hence its name. Spectrolite was first identified in Finland. The spectrolite occurs within the syenite as two phases: large crystals up to 10cm across in pegmatite dikes cross-cutting the syenite; and - 108 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 smaller crystals within the contact zone between the pegmatite and medium-grained host syenite. In both cases the spectrolite displays bluish to yellow-gold schillerescence. Ribbe (1983) states that‘Schiller’ may be used to refer to diffuse, often silvery reflections from mutually oriented, platy inclusions, especially common in labradorite parallel to (010) (Rayleigh, 1923). Portions of ferro-augite syenite (commercially known as “black granite”) in the eastern part of the Coldwell Alkalic Complex near Marathon are coarse grained to pegmatitic. Large feldspar crystals may display yellow-orange to blue schillerescence and have been locally termed “spectrolite”. Two properties near Shack Lake, 2 km northwest of Marathon, have undergone limited exploration in the past but are being re-examined by owners Don Wilkinson, and Jon and Audrey Ferguson, respectively. The main potential usage of the “spectrolite”-rich syenite is as decorative or ornamental stone for use in cabochons, bookends and perhaps tiles.” From Schnieders et al. (1991): “Pegmatitic zones are commonly deeply weathered and are not amenable to large block quarrying. Hand picking and sorting can be undertaken on a small scale. Prospectors should investigate any coarse-grained to pegmatitic sections of syenite or dikes for feldspars that display this characteristic schiller effect. In deeply weathered outcrops, the feldspars commonly remain intact and retain their schiller colours. Stripping and blasting may be required to obtain “fresher”, unfractured material. X-ray diffraction analysis of the “spectrolite” shows the presence of plagioclase and minor K-feldspar (antiperthite). Examination of the mineral in oils shows that it is oligoclase. The schiller effects may be brought about by diffraction that occurs at the boundary of the exsolution lamellae (H. DeSouza, Ontario Geological Survey, personal communication, 1990).” The stone is a generally coarse-grained black to olivebrown with some greenish sections. Compositionally the stone is an iron-rich augite syenite as described above in Walker et al. (1992, 1993a) and MacTavish and Smyk (2017). Hinz et al. (1994) reports testing done by Cold Spring Granite (Canada) Ltd. yielded the following physical properties: Bulk specific gravity: 2.738 Percent absorption (48 hours): 0.560 Compressive strength: 20,130 (psi) dry, 18,420 (psi) wet Modulus of Rupture: 1,420 (psi) dry, 1,530 (psi) wet Testing completed by “Twin City Testing Corp., St. Paul, Minnesota (Assessment Files, Thunder Bay). Stop 5: Yooperlite Source Location UTM coordinates 536816E 5404785N In 2018 the US media was a-buzz over the discovery of a new “mineral” with the unofficial name of “yooperlite”. In a 2018 paper, Laughlin et al. (2018) reported that yooperlite is a fluorescent sodalitebearing syenite which occurs in the Upper Peninsula of Michigan as beach pebbles and cobbles. The authors indicate that it is probable the bedrock source is likely the Coldwell Alkaline Complex in Ontario. Centre 2 of the Coldwell Complex hosts amphibolenatrolite-nepheline syenite (Unit 14) within which, the fluorescent mineral hackmanite, a sulphur-bearing variety of sodalite, has been identified. It can be postulated that the yooperlite pebbles and cobbles found along the Lake Superior shoreline of Michigan originated from this unit and were glacially transported to their current location and subsequently wave-washed and tumbled. From Sage & Watkinson (1995). “Stop 20: Biotite gabbro intruded by various types of nepheline syenite. This stop is at a broad curve in Highway 17 and one should be very CAREFUL OF VEHICLES. Stop 4: Marathon Black Occurrence (1990-1994) UTM coordinates 543576E 5402637N The Marathon Black occurrence was staked by D. Petrunka in the early 1990’s when interest in building stone was high and the Cold Spring Granite (Canada) Ltd. was active in the area. Mr. Petrunka was able to secure funding to remove test blocks from the site, however further development did not occur. Starting at approximately 18.8km outcrops on the east side of the highway of grey to buff pyroxeneamphibole syenite contain orange fluorescing hackmanite, a variety of sodalite. This mineral can only readily be seen with a UV lamp. The syenite contains numerous mafic xenoliths up to 25 to 30cm in maximum length. The xenoliths are subangular to - 109 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 subrounded and the mafic minerals within the syenite tend to occur in clots. The larger xenoliths may have feldspar phenocrysts or porphyroblasts up to 1.0cm. The phenocrysts or porphyroblasts display a seriate size distribution and comprise up to 5 % of the rock. At the centre of the curve at 19.2km an alkalic biotite gabbro is exposed on the north side of the highway. The medium- to coarse-grained gabbro is extensively cut by medium- to coarse-grained syenite and some of the nepheline has been altered to reddish orange “hydronephelinite”. There may be two ages of alkalic gabbro with the older coarser grained gabbro displaying dark selvages up to 7 or 8cm wide and an irregular shape suggesting that it behaved in a more or less ductile manner. The alkalic gabbros have a clotty mafic mineral assemblage. The surface of the coarser grained biotite gabbro is pitted from the weathering of mafic clots. Dikes of nepheline syenite pegmatite occur on the north side of the highway at the inflection in the curve. At this site, two ages of nepheline syenite pegmatite of essentially identical composition cut each other. The trends of these dikes are approximately 340° dipping 60° south and 090° dipping 50° north. The dike trending 090° cuts the dike trending 340° and both are on the order of 20 to 30 cm in width. The dikes have been sketched by Puskas (1970). Both dikes are zoned from an amphibole-rich margin to a feldspar-natrolite (hydronephelinite)-rich core. The central parts of these dikes are commonly relatively rich in reddish orange “hydronephelinite”. West from the two dikes toward the hackmanite-bearing outcrop, coarse-grained alkalic biotite gabbro is intruded by medium- to coarsegrained pyroxene, amphibole syenite with traces of nepheline. Both of these rock types are in turn cut by coarse-grained nepheline syenite dikes. Brecciation is so intensive that the outcrop is an igneous breccia. West toward the hackmanite-bearing outcrop the syenite is pink to grey, fine- to medium-grained, inequigranular seriate with clotty assemblages of mafic minerals. On the south side of the highway coarse-grained alkalic gabbro appears to be cut by medium-grained alkalic gabbro which is in turn intruded by mottled pink to grey, inequigranular seriate amphibole syenite. There is a slight coarsening in texture next to the medium-grained gabbro and a dike of similar material projects from the contact with the medium-grained gabbro through both phases of alkalic gabbro” Stop 6: Marathon Red Occurrence (1960-1994) UTM coordinates 532030E 5402401N The first attempts to quarry red syenite occurred in the late 1880’s near Port Coldwell. From the 1960’s to mid-1980’s exposures of red granite east of Neys Lunch along the Trans-Canada Highway were staked several times, however no further work was recorded. D. Petrunka staked the claims in 1985 and optioned the ground to Cold Spring Granite (Canada) Ltd. Coldspring removed blocks from the highway exposures to evaluate the colour and market suitability. Diamond drilling conducted by Coldspring determined that the red colour of the syenite was limited to top 9.1m of the unit below which it changed to pink. In 1990 Cold Spring terminated the option agreement. The stone is an amphibole-natrolite-nepheline syenite described above by Walker et al. (1993). From Hinz et al. (1994): “In thin section, the stone is composed of primarily anhedral, “turbid” crystals of perthitic feldspar. The turbid areas are caused by the presence of numerous vacuoles. The reddish colour of the stone may be due to iron-staining of the vacuoles and fractures within the feldspar crystals. Anhedral pyroxene (augite), biotite, amphibole (hornblende), and opaque minerals occur together.” The stone is red-brown on fresh surface and orangered on the weathered surface, medium- to coarsegrained with some randomly distributed mafic knots. Hinz et al. (1994) reports samples sent to the Geoscience Laboratories in Sudbury yielded the following physical properties: Bulk specific gravity: 2.75 Percent absorption (2 hours): 0.22, (48 hours): 0.32 Compressive strength: 22,049 (psi) Modulus of Rupture: 1,394 (psi) Stop 7: Little Pic River Quarry (circa. 1884) UTM coordinates 527367E 5405024N This optional stop will be dependant on timing and weather. The construction of the Canadian Pacific Railway (C.P.R.) in the mid 1880’s required good stone quality for piers and abutments at river crossings (Fig. 5). Prospecting parties advanced along proposed rights-ofway ahead of construction in order to identify locations - 110 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 5. CPR construction across the Little Pic River, circa. 1884. Photo compliments of the Thunder Bay Historical Society. of quarriable stone. This quarry, previously unknown to Ministry staff, was one such location where the stone was quarried and land transported to the bridge construction site. Fig. 6 shows a work crew at the Little Pic River Quarry during the construction of the C.P.R., circa. 1884. References Alexander, M. 2007. The mineralogy of NYF pegmatites from the Coldwell Alkaline Complex, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario. Cannon, W.F., Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dikas, A.B., Morey, G.B., Sutcliffe, R. and Spencer, C. 1989. The North Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling; Tectonics, v.8, p.305-332. Heaman, L.M. and Machado, N. 1987. Isotope geochemistry of the Coldwell alkaline complex: 1. U-Pb studies on accessory minerals; Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Saskatoon, Saskatchewan, Program with abstracts, p.54. Hinz, P., Landry, R.M. and Gerow, M.C. 1994. Dimension stone occurrences and deposits in northwestern Ontario; Ontario Geological Survey, Open File Report 5890, 191.p. Kerr, H.L. 1910a. Geological map of part of the north shore of Lake Superior, District of Thunder Bay; Ontario Bureau of Mines, Annual Report Map 19B, scale 1:63 360. Kerr, H.L. 1910b. Nepheline syenites of Port Coldwell; Ontario Bureau of Mines, Annual Report, v.19, p.194-232. Laderoute, D.G. 1987. The petrology, geochemistry, and petrogenesis of alkaline dyke rocks from the Figure 6. Little Pic River Quarry, circa. 1884. Photo compliments of the Thunder Bay Historical Society. Coldwell Alkaline complex; unpublished M.Sc. Thesis, Lakehead University, Thunder Bay, Ontario, 89p. Tectonophysics, v.184, p.73-86. Laughlin, R., Carlson, S.M., Olds, T.A. and Miller 2018. A New Find of Fluorescent Sodalite From Michigan’s Upper Peninsula. Mineral News, Vol. 34, No. 5, May, 2018. Lukosius-Sanders, J. 1988. Petrology of the syenites from Center III of the Coldwell alkaline complex, northwestern Ontario; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 141p. MacTavish. A. and Smyk, M.C. 2017. Archean and Proterozoic geology of the Marathon-Hemlo area, 63rd Institute on Lake Superior Geology Proceedings, v. 63, Part 1, Field Trip Guidebook, p. 1-31. Mitchell, R.H. and Platt, G. R. 1978. Mafic mineralogy of ferroaugite syenite from the Coldwell alkaline complex, Ontario, Canada; Journal of Petrology, v.19, p.627-651. Mitchell, R.H. and Platt, G. R. 1982a. The Coldwell alkaline complex; in Field Trip Guidebook, Proterozoic geology of the northern Lake Superior area, Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Winnipeg, Manitoba, p.42-61. Mitchell, R.H. and Platt, G. R. 1994. Aspects of the geology of the Coldwell alkaline complex: Field trip A2, Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Waterloo, Ontario, 36p. Mulja, T. 1989. Petrology, geochemistry, sulphide- and platinum-group element mineralization of the Geordie Lake intrusion; unpublished MSc thesis, Lakehead University, Thunder Bay, Ontario, 234p. Paterson, W.P.E., Lichtblau, A.F., Ravnaas, C., Lewis, S.O., Tuomi, R.D., Fudge, S.P., Pettigrew, T.K. and Wiebe, K. 2019. Report of Activities 2018, Resident Geologist Program, Red Lake Regional Resident - 111 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Geologist Report: Red Lake and Kenora Districts; Ontario Geological Survey, Open File Report 6351, 127p Puskas, F.P. 1967. Geology of the Port Coldwell Area, District of Thunder Bay; Ontario Department of Mines , Open File Report 5014, 92p. Puskas, F.W. 1970. The Port Coldwell alkali complex; in Proceedings, 16th Institute on Lake Superior Geology, Thunder Bay, Ontario, p.87-100. Puumala, M.A. 2018. Geological description of the Coldwell Alkaline Complex. Personal correspondence. Unpublished Report, p.1. Rayleigh, Lord. 1923. Studies of Iridescent Colour, and the Structure Producing It. III; The Colours of Labrador Feldspar. Proceedings of the Royal Society (Londaon) 103A, p.34-45. Ribbe, P.H. (1983) Chemistry, structure and nomenclature of feldspars. in: Feldspar Mineralogy. (P.H. Ribbe, editor). Reviews in Mineralogy 2. Mineralogical Society of America, Washington, D.C. p. 1-19. Sage, R.P. and Watkinson, D.H. A. 1995. Alkalic Rocks of the Midcontinent Rift, 41st Institute on Lake Superior Geology Proceedings, v. 41, Part 2a, Field Trip Guidebook, 79p. Schnieders, B.R., Smyk, M.C. and Hinz, P. 1991. SchreiberHemlo Resident Geologist’s District; in Report of Activities 1990, Resident Geologists, Ontario Geological Survey, Miscellaneous Paper 152, p.141171. Thomson, J.E. 1931. Geology of the Heron Bay Area, District of Thunder Bay; Ontario Department of Mines, v.XL, pt .2, p.21-39. Accompanied by map 40d. Thomson, J.E. 1933. Geology of the Heron Bay Area, Thunder Bay District, Ontario; Ontario Department of Mines, Annual Report 1932, v.XLI, pt.6,,p.34-37. Walker, E.C., Sutcliffe, R.H. , Shaw, C.S.J. , Shore, G.T. and Penczak, R.S. 1992. Geology of the Coldwell Alkaline Complex; in Summary of Field Work and Other Activitie s 1992, Ontario Geological Survey, Miscellaneous Paper 160, p. 108-119. Walker, E.C., Sutcliffe, R.H., Shaw, C.S.J., Shore, G.T., and Penczak, R.S. 1993. Precambrian geology of the Coldwell Alkalic Complex; Ontario Geological Survey, Open File Report 5868, 30p. - 112 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Field trip 8 - Geology of the past-producing Winston Lake Cu-Zn Mine Robert W.D. Lodge Department of Geology, University of Wisconsin-Eau Claire, WI 54702-4004 Mark Smyk and Mark Puumala Resident Geologist Program, Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, Thunder Bay, Ontario, P7E 6S7, Canada Introduction The Winston Lake greenstone belt is best known for hosting economic volcanogenic massive sulphide (VMS) deposits totalling ~ 6 million tons of Zn-CuPb ore (Ontario Geological Survey, 2011). The belt is located along the northern margin of the Wawa Subprovince of the Wawa-Abitibi terrane and is about 20 km north of Schreiber, Ontario (Pye, 1964; Severin et al., 1991). The Winston Lake greenstone belt is tectono-stratigraphically equivalent to ca. 2720 Ma greenstone belts along the northern margin of the Wawa Subprovince, such as the Shebandowan (Corfu and Stott, 1998), Manitouwadge (Zaleski et al., 1999), and Vermilion (Peterson et al., 2001) greenstone belts (Fig. 1). Regional metamorphic grade in the belt is lower amphibolite facies (Williams et al., 1991). et al., 2008; Lodge et al., 2015). There are also several field trips and published guidebooks (Severin et al., 1991; Smyk and Schnieders, 1995; Lodge, 2012). The previous research in these areas has been extremely valuable throughout the planning of this field trip and preparation of the guidebook. Note that, in the descriptions of some of the units, primary igneous names are used rather than their metamorphic names (e.g., gabbro versus amphibolite). The geochemical and isotopic data presented in this guidebook are available for download through the Ontario Geological Survey (Lodge and Chartrand, 2013). The Winston Lake Greenstone Belt (Fig. 2) is a small belt located directly north of, and almost connected to the Schreiber-Hemlo greenstone belt (Williams et al., 1991); however, the contact relationship of these belts is poorly constrained (Carter, 1982b, a). Unlike the many other greenstone belts in the region, the Winston Lake greenstone belt has not been mapped at a regional scale since the 1960’s (Pye, 1964). The belt is bound to the north by the Quetico Subprovince, to the west by the Winston Lake batholith, and to the south by the Crossman Lake Batholith (Severin et al., 1991). Rocks in the western part of the belt that host the past-producing Winston Lake Mine were initially interpreted as metasedimentary rocks because of the presence of aluminosilicate minerals (Pye, 1964). They were later interpreted to be hydrothermally altered felsic and mafic volcanic assemblages. The Winston Lake greenstone belt and its VMShosting strata have received a considerable amount of research at a property scale (e.g., Osterberg, 1993; Schandl et al., 1995), a belt scale (Lodge et al., 2014), and a subprovince scale (e.g., Polat et al., 1999; Kerrich Figure 1. Geochronology of northern most greenstone belts in the Western Wawa Subprovince. Most of the belts have experienced most of their formation ca. 2720 Ma. Figure from Lodge et al. (2014) and was compiled from numerous geochronological studies (Corfu and Stott, 1998; Zaleski et al., 1999; Peterson et al., 2001; Lodge et al., 2013, 2014). - 113 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 2: Geology of the Winston Lake greenstone belt. Figure compiled by Lodge et al. (2015) from various published and unpublished maps (Pye, 1964; Ritcey, 1992; Osterberg, 1993). Regional Geology The Winston Lake belt has been informally subdivided into two main lithotectonic assemblages: the Winston Lake Assemblage (Fig. 3A) and the Big Duck Lake Assemblage (a thick mafic unit composing most of the belt in Fig. 3B; Severin et al., 1991; Polat et al., 1999; Lodge et al., 2014). The Winston Lake Assemblage is host to the VMS deposits and is composed of calc-alkaline, bimodal volcanic and siliciclastic rocks (Gorton and Schandl, 1995). The Big Duck Lake Assemblage consists of Mg- to Fe-tholeiitic basalts, quartz-feldspar porphyry dykes and sills, and their brecciated equivalents (Ritcey, 1992; Polat et al., 1999). It has been assumed that the Big Duck Lake Assemblage conformably overlies the Winston Lake Assemblage and that the contact was intruded by a thick differentiated gabbro (Osterberg, 1993). This field trip will not examine the Big Duck Lake Assemblage and therefore it will not be discussed further. If interested, please consult the references cited above (in particular: Ritcey, 1992). Prior to the research of Lodge et al. (2014), only one U-Pb age of 2723 ± 3 Ma was obtained from a felsic volcanic rock associated with the Winston Lake orebody (Davis et al., 1994). More recent geochronological data indicate that the entire Winston Lake Assemblage and the Zenith gabbro are all ca. 2720 Ma (Lodge et al., 2014). The structural history of the belt is also poorly constrained and two main structural events are Figure 3: Geology of the Winston Lake greenstone belt. The belt is subdivided into the VMS-hosting Winston Lake Assemblage (A) and Big Duck Lake Assemblage (B). Figure compiled by Lodge et al. (2014). - 114 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 interpreted: D1 manifested as tilting of stratigraphy and a foliation development (north-northwest striking foliation in the Winston Lake Assemblage; weststriking foliation in the Big Duck Lake Assemblage), and D¬2 represented by minor folds and faulting that offset contacts at the map scale (Osterberg, 1993). The VMS-hosting Winston Lake Assemblage is dominated by felsic volcanic and siliciclastic rocks. Despite the high degree of metamorphism and relatively high degree of deformation, many primary volcanic features are preserved in the volcanic rocks. Reliable younging directions obtained from pillowed flows and cross-bedding in volcaniclastic rocks suggest an eastward-younging stratigraphy. The oldest supracrustal strata in this part of the belt are felsic volcaniclastic and siliciclastic rocks. These are conformably overlain by a quartz-feldspar porphyry flow that is associated with the Pick Lake VMS deposit. Altered mafic flows (named Ladder and Middle mafic units) are interlayered with the Camp and Main felsic units that host the Winston Lake VMS deposit. The feldspar-phyric felsic volcanic rocks were then presumably overlain by mafic flows of the Big Duck Assemblage, which was followed by the emplacement of a thick, synvolcanic differentiated gabbro at that contact; the gabbro sill hosts the Zn-rich Zenith orebody. The general stratigraphy of the Winston Lake Assemblage is illustrated in Figure 4. History of Exploration Mining and Mineral Much of the mining and exploration history for the Winston Lake greenstone belt is published internally within the companies that have explored and mined these deposits. Most of these are not externally available. However, the Mineral Deposit Inventory published by the Ontario Geological Survey (2019) has summarized current and historical information available for these deposits. Much of the historical information provided in this section is summarized from this database. The size and grade of the deposits in the Winston Lake area are summarized in Table 1. Massive Zn-mineralization, in what is now interpreted to be a synvolcanic gabbroic sill, was first discovered in 1879 by prospectors and became the Zenith Mine. A total of 1065 tons of ore, averaging approximately 45% Zn, was shipped to a smelter between 1891 and 1899. Between 1899 and 1901, 2700 tons of sphalerite-rich ore was mined and concentrated by Grand Calumet Mining Company Limited (Ontario Geological Survey, 2019) Very little exploration was undertaken in the area until the grounds were claimed by Zenmac Metal Mines Ltd. in 1952. In the late Figure 4. Schematic cross-section looking north-northwest through the strata hosting the VMS orebodies of the Winston Lake area. Figure modified from unpublished Inmet Mining Corp. figures based on stratigraphic terminology from Lodge et al. (2014). - 115 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Table 1. Summary of mining activity in the Winston Lake area (Resident Geologist’s Files, Thunder Bay South District, Thunder Bay). Mine Winston Lake Years of Production 1988-1999 Ore Milled (tonnes) 3 268 698 Zenith Zenmac 1891–1901 1966–1970 3416 164 200 1960’s, almost 165,000 t at a grade of 16.5% Zn were mined from the Zenmac deposit. After the Zenith Mine closed, the property was stagnant until 1978 when Corporation Falconbridge Copper (CFC) completed reconnaissance geological mapping and lithogeochemical sampling in the region. The “metasediments” (Bartley, 1940; Pye, 1964) were re-interpreted as metamorphosed felsic volcanics. This was followed by more detailed property mapping, lithogeochemistry, and geophysical surveys, which defined the alteration zone that is in the immediate footwall to the gabbro. Areas of Na2O depletion and FeO, MgO, and Zn enrichment were outlined in the calc-alkaline volcanic rocks. CFC geologists also realized that the presence of massive sphalerite in the gabbro was unusual. The presence of VMS-like lithogeochemical signatures in the footwall to the gabbro led to the interpretation that the Zenith orebody Commodities and Grade 1.04% Cu, 14.56% Zn, 32.32 g/t Ag, 1.41 g/t Au 45% Zn 16.5% Zn was likely a large xenolith from a larger VMS orebody hosted at the top of the calc-alkaline felsic volcanic strata below the gabbro. With this newly recognized VMS potential, diamond drilling began in 1981 and targeted the felsic volcanic rocks at the base of the gabbro. In 1982, after only drilling five holes, CFC intersected 2.1 metres of massive sulphides containing 1.1% Cu, 19.1% Zn, 22.2 g/t Ag and 0.73 g/t Au. Mining began in 1988 and continued until the mine was officially closed in 1998. Later research by Osterberg (1993) and Lodge et al. (2014) outlines the lateral extent of alteration in the Winston Lake camp (Fig. 5). The surface expression of the Pick Lake orebody, the Anderson occurrence, was first reported by local prospectors in 1952. There was some shallow diamond drilling completed in the area but nothing materialized. CFC picked up the claims following the discovery of the Winston Lake deposit in 1982. In 1984, diamond Figure 5. Lateral variations in the stratigraphic thicknesses in the Winston Lake assemblage. Distance between sections is not to scale. Strata are hung from the bottom of the Ladder mafic unit. Figure from Lodge et al. (2014). For location of figure and legend, refer to Figure 3. - 116 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 drilling testing the down-dip extension of the Anderson occurrence discovered the Pick Lake orebody. In 1993, Metall Mining (formally Minnova, and operators of the mine at the time) began a 2200 metre drift to mine the Pick Lake deposit through the mine workings at Winston Lake. Doiron et al. (1997) proposed that the dyke-like nature of the deposit, combined with durchbewegung ore textures and sulphide injection structures, suggested that the Pick Lake deposit was a remobilized massive sulphide ore body. The Pick Lake Mine was abandoned when the Winston Lake Mine shut down in 1998. After the shut-down of the mine limited exploration was carried out on the property. There were several mapping and geophysical projects carried out on select parts of Superior Lake Resources’ current property between 2000-2011 by various exploration companies, but no further resources were outlined. In 2017-18, the mineral properties overlying the Pick Lake and Winston Lake deposits were acquired by Superior Lake Resources Limited. Superior Lake have completed new drilling programs that updated the resources at the Pick Lake and Winston Lake deposits (Table 2) and ground geophysical surveys that defined new targets for future exploration (Fig. 6). Field Trip Stops All coordinates are reported in NAD 83, Zone 16 The purpose of this portion of the field trip is to introduce the geological setting of strata hosting the VMS ore bodies at Winston Lake. As most of the lithofacies of the greenstone belt are difficult to access, this part of the trip will focus on camp-scale features and what role they played in the discovery of the orebodies. The stops will focus on the immediate footwall strata to the Winston Lake deposit and will highlight some of the geochemical and geochronologic data obtained from recent studies (Lodge et al., 2014). The field stops are illustrated in Figure 7. Figure 6. Recent ground EM geophysical results in the vicinity of the Winston Lake and Pick Lake ore bodies. Figure obtained from Superior Lake Resources website (www.superiorlake.com.au). Table 2. Updated resources at the Pick Lake and Winston Lake deposits (www.superiorlake.com.au). Resource Category Indicated Inferred Total // Weighted Average Tonnage (Mt) 2.07 0.28 2.35 Zn (%) 18.0 16.2 17.7 (News Release, Superior Lake Resources Limited, March 7, 2019) - 117 - Cu (%) 0.9 1.0 0.9 Au (g/t) 0.38 0.31 0.38 Ag (g/t) 34 37 34 �Proceedings of the 65th ILSG Annual Meeting - Part 2 leucocratic gabbro to pyroxenite (this is more apparent at Stop 2). The Zenith orebody appears to be associated with the transition from gabbro to pyroxenite. The Zenith orebody is essentially mined out but there are a few metre-scale slivers of massive sphalerite remaining. The ores are strictly Zn-rich and contain only minor amounts of pyrite, pyrrhotite, and chalcopyrite. Like most zinc-rich ores, they commonly are coated by white chalky zinc oxide minerals. The grain size of the sphalerite is generally coarse, most likely recrystallized during regional metamorphism. Stop 2 – Differentiated Gabbro UTM Coordinates 472337E 5425082N Figure 7. Geology of the Winston Lake mine area highlighting the location of field trip stops. Figure modified from Lodge (2012) based on data presented by Lodge et al. (2014). For location of figure and legend, refer to Figure 3. Stop 1 – Zenith Mine Open Pit UTM Coordinates 473182E 5424996N This stop is inside of the main Winston Lake Mine gate. It is a short drive along mine property roads. From where we park, walk along the base of the cliff toward the lake. WARNING: The walk into the main pit area is on a narrow and steep-sided path. Walking to the exposures of sulphides in a large group is discouraged. This stop represents the remnants of the original discovery in the Winston Lake area. The gabbro (now amphibolite) in this area is massive with some low-angle shears cutting up the outcrop. The gabbro is differentiated and consist of phases ranging from This stop is on the main road about 100 metres back from the Winston Lake mine gate along the side of the road. The different phases of the gabbro intrusion are well-exposed on either side of the road. This outcrop is very large and is dissected by a stream. We will not be crossing the stream. This stop highlights the complexity and multiple phases of the gabbro. On either side of the road, near the gate to the Winston Lake mine site, are perfectly exposed road cuts and stripped exposures of the gabbro that hosts the Zenith orebody. At this stop, we are less than 100 metres from the base of the intrusion. Particularly on the south side of the road toward the creek, the gabbro has a layered appearance with layers of pyroxenite (now mostly hornblende) and more plagioclase-rich leucocratic layers. In other areas, there are pegmatitic patches with large centimetre-scale plagioclase crystals. The compositional layering in this intrusion ranges from centimetre-scale to metre-scale Photo 1. (Left) Zenith mine open pit. (Right) Sphalerite-rich xenolith in gabbro. - 118 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 At this stop, the contact between the gabbro and the underlying volcanic rocks is exposed along road cuts leading toward the mine property. Immediately below the contact with the gabbro is a layered siliceous tuff unit that is known as the Winston Lake Interval. About 450 metres down dip of this unit is the Winston Lake main orebody. Photo 2. Differentiated gabbro showing melanocratic and leucocratic layers. layers. Most of these variations are not mappable. A leucocratic pegmatitic phase of the gabbro was sampled for U-Pb geochronology to determine the age of the gabbro. Zircons separated from the gabbro were low-U, typical of gabbroic zircons, indicating that they were magmatic in origin rather than xenocrystic. These grains were analyzed using TIMS analysis at the University of Toronto, yielding a U-Pb age of 2719.2 ± 4.0 Ma. This age indicates that the gabbro that intruded and entrained the Winston Lake VMS body is synvolcanic, and the same age (or slightly younger) as the host rocks. The finely laminated layers range in thickness from a few millimetres to 1-2 centimetres and range in composition from felsic tuff, chert, and lesser mafic tuff. This unit is laterally extensive (Fig. 3) and continues almost the entire length of the Winston Lake assemblage. Trace element and REE patterns suggest that this ash layer has an FII to FIII type felsic composition (Hart et al., 2004). There is little variation in this volcaniclastic unit, both compositionally and texturally, although it is locally interlayered with mafic flows. Stop 4 – Altered Felsic Volcanic Rocks in Footwall UTM Coordinates 472123E 5424987N Continue south(west) along the road for about 150 metres to the next stop. This stop is near the trail entrance to stops W5 to W7. There is about 100 metres of roadcut outcrop of variably altered felsic volcanic rocks here to examine. This stop is along the main road about 100 metres west from Stop 2. It is a large roadcut outcrop of the contact between the gabbro and the underlying felsic volcanic rocks. This stop, and nearby outcrops typify the alteration facies (assemblages) found within the uppermost Main and Camp felsic units. These altered units are laterally extensive, but relatively thin (Figs. 3 & 4) and it is not known how many flows are represented within this unit. Unaltered equivalents of this rock are usually massive, coherent units that are quartz- and plagioclase-phyric, Photo 3. Bedded felsic volcaniclastic unit known as the Winston Lake Horizon. Photo 4. Altered felsic volcanic rock now a quartzmuscovite-sillimanite-biotite schist. Stop 3 – Winston Lake Horizon UTM Coordinates 472200E 5425057N - 119 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 and locally contain flow banding. Minor felsic tuffs and tuff breccias have also been described in this unit elsewhere in the camp (Osterberg, 1993). This unit has a U-Pb age of 2723 ± 3 Ma (Davis et al., 1994). Altered versions of these quartz- and feldspar-phyric flows contain variable amounts of biotite, muscovite, sillimanite, and cordierite. Phenocrysts may be locally preserved, but are more difficult to see. Lesser altered equivalents contain quartz-muscovite-biotite-feldspar assemblages. With increasing degree of alteration, the rocks contain cordierite, sillimanite knots, garnet and anthophyllite. Major element geochemistry shows extensive Na2O depletion and local Fe-Mg enrichment. Trace element and REE patterns indicate that this unit is a FIII-type felsic volcanic (Fig. 8). the main road and we will be returning on the same trail and we will look at the lithofacies passed over on the return trip. WARNING: This is an ATV trail and there are many wet places and irregular surfaces. Please walk with caution and try and stay dry. This outcrop of the Ladder Mafic Unit contains some of the most spectacular features that will be observed Stop W5 – “Ladder” Mafic Flow UTM Coordinates 471800E 5425200N From the last stop, enter the trail that leads westward from the main road. The next stop is approximately 500 metres in along the trail. This is the furthest stop from Photo 5. (Top) Coarse grained orthoamphibole. (Bottom) Orthoamphibole-garnet-biotite schist. Figure 8. Geochemical discrimination plots for the felsic units in the Winston Lake greenstone belt. One notable geochemical distinguisher for the different felsic units in the Winston Lake Assemblage is the differences in Zr/Ti. Figure from Lodge et al. (2014). Fields in plots A and B are from Lesher et al. (1986) and Hart et al. (2004). - 120 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 during this field trip. In addition to the coarse-grained mineral assemblages associated with metamorphosed hydrothermal alteration, the overall lack of significant deformation in this area has preserved the volcanic textures in the rock (Fig. 9). Younging directions are still determinable in this unit despite strong alteration and they indicate an eastward younging of the strata. As with most units in the Winston Lake assemblage, this unit is laterally extensive, but relatively thin. Unaltered equivalents of the Ladder Flow contain plagioclase phenocrysts. At the eastern limit of this outcrop are spectacular exposures of an orthoamphibole-cordierite assemblage within the altered pillowed mafic flows. The pillows are metre-scale and their selvages are resistive weathering. There appears to be very little compositional difference between the pillow interior and the selvages, with the exception of slightly more biotite. Stratigraphically below the pillows is a 10 metre thick massive basalt flow. This massive flow is very homogeneous and the only variation are gradational changes in the size of the orthoamphibole crystals. These crystals can be up to 10 centimetres in size and are usually randomly oriented. The matrix is mostly medium grained cordierite and minor biotite in this massive lithofacies. Near the lower contact with the altered felsic volcanic rocks, the Ladder Flow contains abundant clots, sheets, and veins of porphyroblastic garnet. In addition to garnet-orthoamphibole-cordierite, there are also local concentrations of biotite and chlorite. This alteration style is interesting, but difficult to interpret. In some places it appears as if the garnet clots represent altered clasts in a breccia. In other locations, they may be altered pillow margins or may be deformed veins. Regardless, the sharp transition to massive orthoamphibole-cordierite altered flow into a more chaotic orthoamphibole-garnet-cordierite zone may indicate the transition from a massive to breccia facies of this unit. It may also represent a different chemical gradient within the alteration zone. Geochemically, the garnet-bearing rocks are still mafic in composition. The geochemical characteristics from the Ladder Mafic Unit, as well as other mafic units throughout the greenstone belt, are summarized in Figure 10. All of the mafic flows in the footwall to the Winston Lake mine have similar geochemical characteristics. Most noteworthy is that they are all calc-alkalic to transitional in their magmatic affinities and have pronounced negative Nb anomalies. In the hangingwall strata (i.e. Big Duck Lake Assemblage), the flows are tholeiitic and have flat rare earth and trace element patterns on normalized element plots (Lodge et al., 2014). Stop 6 – Trail Showing UTM Coordinates 471867E 5425051N From the previous stop, return eastward back toward the main road for approximately 150 metres. This stop is a flat, rusty outcrop that we walked over to get to the Figure 9. Outcrop sketch of the contact between the altered mafic “Ladder flow” and the underlying altered felsic volcanics in the Winston Lake area. Sketch only incorporates areas that were cleaned and there are additional outcrops in the area. - 121 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Figure 10. Various geochemical discrimination plots for mafic rocks from the Winston Lake greenstone belt. Basalts from the Ladder and Middle units tend to be have more calc-alkalic to transitional magmatic affinities. Sills and flows from the Big Duck Lake assemblge are mainly tholeiititc. Figure from Lodge et al. (2014). Fields in A from Shervais (1982). Fields in B from Ross & Bedard (2009). Fields in C from Piercey et al. (2002). Fields in D modified from Polat (2009). Ladder Flow. This stop represents one of the many smaller mineralized intervals in the Winston Lake camp. The mineralization in the “Trail Showing” is located near the contact between the “Ladder Flow” and underlying felsic volcanic rock and it is hosted within a siliceous, bedded volcaniclastic unit up to 15 centimetres thick. It contains over 6000 ppm Cu and Zn-bearing metamorphic minerals such as gahnite are also present. The bedded volcaniclastic unit is altered to a quartz-cordierite-biotite-garnet mineral assemblage containing variable amounts of orthoamphibole and sillimanite. The layering in the rock appears to be relict primary volcanic texture. This unit is distinct and is present at the contact between the Ladder mafic flow and the underlying massive, altered quartz-feldspar-phyric Main felsic flow, which the trail passes over from there to the main road. The alteration assemblages in the massive flow are the same as those observed at Stop 4. Stop 7 – Contact of Ladder Flow and Altered Felsic Volcanic Rocks UTM Coordinates 471918E 5424992N Continuing back toward the main road along the trail, this stop is approximately 100 metres from the previous stop. This is the last stop on the trail before returning to the main road. Photo 6. Felsic volcaniclastic rock hosting disseminated sulphides. This stop shows the contact between the Ladder mafic flow and the overlying altered Main felsic volcanic rocks, which constitute the immediate footwall strata to the Winston Lake deposit. Cleaning of this otherwise black and featureless exposure resulted in a near-perfect exposure of an altered basaltic flow top breccia that is intermingled with the overlying felsic tuff. This is one of a few places where these flow features are exposed at surface. The variety of flow facies exposed within - 122 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 rocks that dip eastward underneath this unit (Fig. 4). We will not see the host rocks to the Pick Lake deposit on this trip because they are not easily accessible. The surface expression of the Pick Deposit is known as the Anderson Occurrence and is 850 m west of this location. Photo 7. Upper contact of altered mafic flow (Ladder Flow) and altered felsic volcanic. the Ladder mafic flow indicates that the contact at this stop is not a peperite caused by a sill mingling with overlying unconsolidated tuff. Rather, it is more likely that the felsic tuff settled into on top of the basaltic flow top breccia. The flow here is altered to an orthoamphibole-garnet assemblage with retrograde chlorite and biotite. The garnet porphyroblasts are evenly distributed and occur in the felsic volcanic above the contact. This suggests some chemical exchange between the lithofacies at the contact during metamorphism, as previously suggested by Gorton and Schandl (1995). The overlying felsic tuff is altered to a quartz-muscovite-sillimanite-biotite mineral assemblage. Stop 8 – Pick Lake Vent Raise Area UTM Coordinates 471579E 5424177N About 150 metres south of the trail entrance on the main road is the gate to the Pick Lake deposit. The Pick Lake vent raise, and the next stop are about 1.1 kilometres from the gate. There are plenty of outcrops around the former vent raise to examine that show a variety of alteration facies. WARNING: Although the shaft has a concrete cap and is safe to walk on, please do not walk directly on old mine workings. This stop is in the middle of the thickest part of the quartz-feldspar-phyric felsic flow that forms the Main felsic unit of the Winston Lake assemblage and the hanging wall to the Pick Lake orebody. The Pick Lake deposit is about 300 metres directly below the mine workings near this stop. The orebody is not associated with the rocks exposed here, rather is associated with felsic volcaniclastic and siliciclastic Based on this location alone, it is not clear whether this part of the Main Felsic Unit is definitively an intrusion or extrusion (e.g. Osterberg, 1993). There appears to be very little textural variation in the unit throughout the area. The unit is massive and the only variations are in the degree of alteration and mineral assemblages. The alteration is pervasive and mineralogical changes are gradational and do not appear to represent primary compositional layering. If it is an extrusive unit, it is a very massive flow with only a thin brecciated carapace (seen at Stop W9). Stratigraphically above the Ladder Mafic Unit, the Main Felsic Unit appears to exhibit phenocryst sizes in unaltered parts of this unit that are much larger (3-4 millimetres) compared to the Camp Felsic Unit in the footwall to the Winston Lake orebody. A sample of this unit was submitted for U-Pb dating by TIMS analysis at the University of Toronto. The sample yielded a homogeneous population of zircon that produced an age of 2721 ± 1 Ma. This confirms that the unit is the same age as the host rocks to the Winston Lake orebody. The Main felsic unit has a notably higher Zr/Ti ratio that the Camp felsic unit. All the felsic magmas in this camp are FIII type felsic magmas. Photo 8. Quartz-muscovite-biotite-garnet schist near the Pick Lake mine shaft. - 123 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Stop 9 – Pick Lake Felsic Breccias UTM Coordinates 471998E 5424667N This stop is a roadside outcrop on the Pick Lake road approximately 100 metres away from the gate. This is the last stop inside the Pick Lake gated road. At this stop, the monolithic breccia phase of the felsic quartz-feldspar-phyric flow from the Main Felsic Unit is well exposed on the roadside. It is not certain if this breccia is a volcanic or structural texture. Given that the main part of this body is thick, massive and homogeneous, and that it has been confirmed to be the same age of the surrounding volcanic rocks, it is possible that this may represent the synvolcanic intrusion that was the feeder for the overlying felsic flows and volcaniclastic units above the Ladder mafic unit. It is common for hypabyssal synvolcanic intrusions to be brecciated at their margins. Alternatively, it could represent the breccia margin of a flow. Opinions are encouraged! The fragments are lenticular and stretched defining a pronounced stretching lineation. They contain quartz and plagioclase phenocrysts that compose up to 15% of the fragment. The anastomosing matrix is composed of quartz-biotite-muscovite mineral assemblages and composes up to 25% of the rock. There is no obvious layering in the rock but there is some variation in the abundance and size of the clasts that may represent crude bedding. Stop 10 – Tuffaceous Metasedimentary Rocks UTM Coordinates 472699E 5423145N This stop is 1.8 kilometres southward on the main road from the Pick Lake gate and is adjacent to the power lines. There are several roadcut outcrops that are worth checking out. This is the final stop of the field trip. The tuffaceous metasedimentary rocks are along strike and south from the orebodies, and do not contain mineral assemblages indicative of significant hydrothermal alteration. They were classified as “intermediate volcaniclastics” by Osterberg (1993). There are many sedimentary structures in this unit, such as cross-bedding, that suggesting it is a reworked volcaniclastic deposit. The composition of the rock suggests that the provenance is mostly mafic with only a minor felsic component. In this stop, mineral assemblages range from biotitequartz-garnet to biotite-quartz-hornblende and even Photo 10. Low-angle cross-bedded mafic-intermediate tuffaceous metasedimentary rock. local concentrations of lapilli-sized clasts of mafic composition. These compositional variations are on the metre- to outcrop-scale. A sample of this unit within the finer-grained, cross-bedded part of the exposure was sent for detrital zircon analysis at the LA-ICP-MS lab at Laurentian University. The results show a single peak centered around 2720 Ma, suggesting that the source of detritus was local and from 2720 Ma volcanic units (Lodge et al., 2014). The composite, mafic-felsic composition of these rocks suggests that they are not primary volcanic deposits. They may represent distal, reworked facies of slump fans deposited during rifting and sourced from bimodal volcanic units within the Winston Lake assemblage. Photo 9. Matrix-supported quartz-phyric felsic volcanic breccia. - 124 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Acknowledgements The guidebook represents a revised version of a field trip guidebook published by the Ontario Geological Survey (OFR 6282). Much of the text and images has been re-used from this publication. A complete citation for that publication is below: Lodge, RWD (2012). Winston Lake and Manitouwadge revisited: Modern views of two volcanogenic massive sulphide (VMS)-endowed greenstone belts: A field trip guidebook. Ontario Geological Survey, Open File Report 6282, 37 p. In addition, the authors would like to thank management of Superior Lake Resources for allowing members of the Institute on Lake Superior Geology to access their mineral exploration property. References Bartley, M.W. 1940. Geology of the Big Duck-Aguasabon Lakes area. Ontario Geological Survey, Map 49k. Carter, M.W. 1982a. Precambrian geology of the Terrace Bay area, northeast sheet, Thunder Bay District. Ontario Geological Survey, Preliminary Map 2557. Carter, M.W. 1982b. Precambrian geology of the Terrace Bay area, northwest sheet, Thunder Bay District. Ontario Geological Survey, Preliminary Map 2556. Corfu, F. and Stott, G.M. 1998. Shebandowan greenstone belt, western Superior Province: U-Pb ages, tectonic implications, and correlations. Geological Society of America Bulletin 110, p. 1467-1484. Davis, D.W., Schandl, E.S., and Wasteneys, H.A. 1994. U-Pb dating of minerals in alteration halos of Superior Province massive sulfide deposits; syngenesis versus metamorphism. Contributions to Mineralogy and Petrology 115, p. 427-437. Doiron, D., Siddiqui, M., and Smyk, M.C. 1997. Preliminary investigations of the Pick Lake deposit, Winston Lake Mine, Ontario: A remobilized massive sulphide orebody; 43rd Institute on Lake Superior Geology, Program with Abstracts, Sudbury, Ontario, p.17-18. Gorton, M.P. and Schandl, E.S. 1995. An unusual sink for rare earth elements: the rhyolite-basalt contact of the Archean Winston Lake volcanogenic massive sulphide deposit, Superior Province, Canada. Economic Geology 90, p. 2065-2072. Hart, T.R., Gibson, H.L., and Lesher, C.M. 2004. Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic massive Cu-Zn-Pb sulfide deposits. Economic Geology 99, p. 1003-1013. Kerrich, R., Polat, A., and Xie, Q. 2008. Geochemical systematics of a 2.7 Ga Kinojevis Group (Abitibi), and Manitouwadge and Winston Lake (Wawa) Ferich basalt-rhyolite associations: Backarc rift oceanic crust? Lithos 101, p. 1-23. Lesher, C.M., Goodwin, A.M., Campbell, I.H., and Gorton, M.P. 1986. Trace-element geochemistry of oreassociated and barren, felsic metavolcanic rocks in the Superior Province, Canada. Canadian Journal of Earth Sciences 23, p. 222-237. Lodge, R.W.D. 2012. Winston Lake and Manitouwadge revisited: Modern views of two volcanogenic massive sulphide (VMS)-endowed greenstone belts. A field trip guidebook., Ontario Geological Survey Open File Report 6282, p. 37. Lodge, R.W.D. and Chartrand, J.E. 2013. Establishing regional geodynamic settings and the metallogeny of volcanogenic massive sulphide mineralization of greenstone belt assemblages (circa 2720 Ma) of the Wawa Subprovince via geochemical comparisons, Ontario Geological Survey, Miscellaneous Release Data 306. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Franklin, J.M., and Hamilton, M.A. 2014. Geodynamic reconstruction of the Winston Lake greenstone belt and VMS deposits: New trace element geochemistry and U-Pb geochronology. Economic Geology 109, p. 1291-1313. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Franklin, J.M., and Hudak, G.J. 2015. Geodyamic setting, crustal architecture, and VMS metallogeny of ca. 2720 Ma greenstone belt assemblages of the northern Wawa subprovince, Superior Province. Canadian Journal of Earth Sciences 52, p. 196-214. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Hudak, G.J., and Jirsa, M. 2013. New U-Pb geochronology from Timiskaming-type assemblages in the Shebandowan and Vermilion greenstone belts, Wawa Subprovince, Superior Craton: Implications for the Neoarchean development of the southwestern Superior Province. Precambrian Research 235, p. 264-277. Ontario Geological Survey 2019. Mineral Deposit Inventory; Ontario Geological Survey, Mineral Deposit Inventory (April 2019 update), online database.Osterberg, S.A. 1993. Stratigraphy, physical volcanology, and hydrothermal alteration of the footwall rocks to the Winston Lake massive sulfide deposits, northwestern Ontario. University of Minnesota at Minneapolis, 351 p.. Peterson, D., Gallup, C., Jirsa, M., and Davis, D.W. 2001. Correlation of the Archean assemblages across the U.S.-Canadian border: Phase I geochronology, 47th Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 47, Part 1 - Program and Abstracts, p. 77-78. - 125 - �Proceedings of the 65th ILSG Annual Meeting - Part 2 Piercey, S.J., Mortensen, J.K., Murphy, D.C., Paradis, S., and Creaser, R.A. 2002. Geochemistry and tectonic significance of alkalic mafic magmatism in the Yukon-Tanana terrane, Finlayson Lake region, Yukon. Canadian Journal of Earth Sciences 39, 17291744. Polat, A. 2009. The geochemistry of Neoarchean (ca. 2700 Ma) tholeiitic basalts, transitional to alkaline basalts, and gabbros, Wawa Subprovince, Canada: Implications for petrogenetic and geodynamic processes. Precambrian Research 168, p. 83-105. Polat, A., Kerrich, R., and Wyman, D.A. 1999. Geochemical diversity in oceanic komatiites and basalts from the late Archean Wawa greenstone belts, Superior Province, Canada: trace element and Nd isotope evidence for a heterogeneous mantle. Precambrian Research 94, p. 139-173. Pye, E.G. 1964. Mineral deposits of the Big Duck Lake area, district of Thunder Bay, Ontario Department of Mines Geological Report 27, 58 p.. Ritcey, D.J. 1992. Geology and mineralization in the vicinity of Big Duck Lake, Ontario. Unpublished MSc. thesis, University of Ottawa, 235 p.. Ross, P.-S. and Bédard, J.H. 2009. Magmatic affinity of modern and ancient subalkaline volcanic rocks determined from trace-element discriminant diagrams. Canadian Journal of Earth Sciences 46, p. 823-839. Schandl, E.S., Gorton, M.P., and Wasteneys, H.A. 1995. Rare earth element geochemistry of the metamorphosed volcanogenic massive sulfide deposits of the Manitouwadge mining camp, Superior Province, Canada; a potential exploration tool? Economic Geology and the Bulletin of the Society of Economic Geologists 90, p. 1217-1236. Severin, P.W.A., Balint, F., and Sim, R. 1991. Geological setting of the Winston Lake massive sulphide deposit, Mineral Deposits in the Western Superior Province, Ontario, Geological Survey of Canada Open File 2164, p. 58-73. Shervais, J.W. 1982. Ti-V plots and the petrogenesis of modern and ophiolitic lavas. Earth and Planetary Science Letters 59, p. 101-118. Smyk, M.C. and Schnieders, B.R. 1995. Geology of the Schreiber greenstone assemblage and its gold and base metal mineralization; 41st Institute on Lake Superior Geology, Proceedings volume 41, pt.2c, Marathon, Ontario, 77 p.Williams, H.R., Stott, G.M., Heather, K.B., Muir, T.L., Sage, R.P., 1991. Wawa Subprovince, in: Thurston, P.C., Williams, H.R., Sutcliffe, R.H., Stott, G.M. (Eds.), Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p. 485-541. Zaleski, E., van Breemen, O., and Peterson, V.L. 1999. Geological evolution of the Manitouwadge greenstone belt and Wawa-Quetico subprovince boundary, Superior Province, Ontario, constrained by U-Pb zircon dates of supracrustal and plutonic rocks. Canadian Journal of Earth Sciences 36, 945-966. - 126 - � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2019 Description An account of the resource Institute on Lake Superior Geology. Terrace Bay, Ontario. May 8-9, 2019. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2019 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/88e4ca0cb9d8c30674f915adad33010c.pdf 00a00568506a2bb2ddd39848707548a8 PDF Text Text www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Isle Royale: Keweenaw Rift Geology Figure 1: Native copper in a vein on Washington Island, Isle Royale (photo by Justin Olson). This occurrence of copper was found all over the Keweenaw and Isle Royale, but humans dug them out and made pits and small mines to extract the precious metal. It was traded across the North American continent by Native Americans. Later Europeans re-excavated the indigenous pits and eventually developed major mining activity. This mining of copper was an economic pay-off of a geologic event that brought deep-seated heavy elements to earth’s surface more than one billion years ago. The wilderness preservation of Isle Royale may explain why such occurrences happen there but not on the Keweenaw, except in underwater places like Great Sand Bay. Physical Volcanology of Large Lava Flows Middle Proterozoic Continental Tholeiitic Flood Basalts of the 1.1 Ga Keweenaw Rift (Rodinia). Field trip Institute of Lake Superior Geology, May 25-30, 2013 Bill Rose, Justin Olson Michigan Technological University 1 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Table of Contents Topic Page No. Purpose and Philosophy Introduction (Basic References) Broad Background (Some geo background with web links) Specific Background pages: Basalt (the mother liquor of the planets) Paleomagnetism (great tool to see geologic history) Geochemistry (esoteric? geochemistry) Basalt types (field hand specimen petrology) Physical features of large lava flows Columnar Joints (entablature, colonade) Mafic Volcaniclastic Deposits (pyroclastic rocks of the rift) Isle Royale Lava Stratigraphy (nomenclature of flows) Ophitic texture (understanding an unusual igneous texture) Pegmatite (in situ differentiation of thick lava flows--also pegmatoid, dolerite) Amygdaloid (lava flow tops with bubble holes filled with colored minerals) Copper (Why native copper here?) Conglomerate (alluvial fan and fluvial sediments) LIDAR (new 2 m resolution LIDAR topography data) Specific field areas we will visit: Washington Harbor (Windigo, Grace Island) NWCoast (Hugginin Cove, Wendigo) McCargoe Cove (Minong Mine) Amygdaloid Island (Amygdaloid channel, Belle Isle, MVD) Blake Point (Upper and lower ophite, pegmatite) Passage Island Snug Harbor Scoville Point (entablature jointing) Lookout Louise (Monument Rock) Red Rock Point Raspberry Island (Segregation cylinders, vesicle cylinders, pegmatites) Tookers and Davidson Mott Island (Conglomerate) Lighthouse (Amygdaloidal flow top minerals) Ojibway What to take Home (why Isle Royale is geologically unique--what it is known for) Acknowledgements Bibliography (where this information comes from) Latitude-Longitude locations 2 3 4 6 11 14 16 19 21 28 29 32 33 35 38 40 42 43 45 48 50 51 53 56 56 57 60 62 65 70 71 74 76 77 78 79 85 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Purpose and Philosophy This field guide aims to give anyone interested in geology and Isle Royale an interpretation of things that can be seen outside in this unique National Park. We try to avoid jargon, in spite of some of the words above on this page. Each part of the earth’s surface offers part of the evidence of past events for us to interpret. On Isle Royale we see rocks which reflect earth about 1.1 Billion years ago, and we can interpret what this rock record means. These interpretations are speculative and they evolve constantly, reflecting new observations. This field guide is an update of a guide from 1994. One very important source is a geologic map done by N. King Huber of the US Geological Survey. On this geologic map of Isle Royale, this geologic map (Figure 2) the western part is mostly tan, and the eastern part is mostly green. The colors reflect glacial outwash gravels and moraines that mostly bury the bedrock in the west, while those materials are absent in the east. Because of our interest in the rift lavas, this trip focuses on the Eastern part of Isle Royale, which has only minimal glacial cover, although we do pass through Washington Harbor and part of the western portion. Isle Royale has remarkably few visitors, especially considering that it is a national park and is, for many people, within a couple of days travel. This lack of tourism can be explained partially by the park's island location and by the fact that a trip to Isle Royale seems to require a deeper commitment, of both time and money, than other vacations might require. But the very people whom you would most expect to want to visit Isle Royale don't go. When you compare the popularity of various national parks, the public's avoidance of Isle Royale is obvious, perhaps even more obvious to me because of my position. As a professor at Michigan Technological University for more than 40 years, I have had direct contact with hundreds of ecologically-minded students, many of them geology majors, who are committed to the outdoors and to field experiences. However, very few of these students go to the park, even though they live for years in Houghton, MI, which is the home of the Ranger III, one of the principal transporters of visitors to and from Isle Royale. Likewise, many of the geologists I have known have visited all of the geological sites around Lake Superior and the other Great Lakes, but only a few of them have been to Isle Royale. This is a remarkable contradiction, something I'm at a loss to explain. It seems to attest to America's addiction to the automobile; maybe people just can't stomach the thought of being separated from their car for a few days! At any rate, I hope that this guide and its website (http://www.geo.mtu.edu/~raman/SilverI/ IRKeweenawRift) will encourage more geologists, as well as other people, to visit the park. Besides the fact that Isle Royale has outstanding geological sites, a trip there can be made at moderate expense, and the park offers comfortable facilities and logistics that most geologists would find agreeable. I recommend taking a week to visit and using kayak, canoe or motor boat (bring along or rent from the park concession) to allow access to the many wave-washed outcrops. ...Bill Rose April 2013 3 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Introduction Before you go to Isle Royale on the field trip, you may wish to read some geological sources. Which ones you read could depend on your interests. One source for all who are interested in the geology is Huber (1975): USGS Bulletin 1309 (http://pubs.usgs.gov/bul/1309/report.pdf). This booklet covers much of Isle Royale geology and is well illustrated. A more academic version of Huber’s geology is USGS Prof Paper 754-C-also downloadable for free (http://pubs.usgs.gov/pp/0754c/report.pdf). There is also a report on the glacial geology, more useful in Western Isle Royale (http://pubs.usgs.gov/pp/0754a/ report.pdf). The geologic map (Figure 2) is downloadable also and can be used in GIS format with Google Earth or other base maps. For the Keweenaw Peninsula, GIS data on the geology and mineral deposits is available from Cannon et al., USGS OFR 99-149, 1999 (http://pubs.usgs.gov/of/1999/ of99-149/). The age information on the Keweenawan rocks is one of the most vital pieces of data. Those interested in age should consult Davis & Paces, 1990, and Nicholson et al., 1997. The petrology and geochemistry of the Volcanic Rocks of the Portage Lake Volcanics is thoroughly explored by Paces, 1988. Figure 3: Schematic cross section of Isle Royale, showing tilted lava and conglomerate layers. The sections which follow are specifically designed to provide background information on various geologic topics. Figure 2 (next page): Geologic map of Isle Royale National Park (Huber, 1973). (http://www.nature.nps.gov/geology/inventory/publications/map_graphics/isro_map_graphic.pdf) 4 �40 15 pu 8 pu pu 16 45 : :: 26 2018 % 12 psp 15 18 12 5 37 6 : : pu pu pu 5 : Qal pm 8 :: 5 : 5 18 5 10 Qal 17 15 13 5 pu 19 pu Ì : 20 pu 55 45 pth 15 pei » »»»» 15 7 : 6 18 : pu 13 18 pu Qal 20 45 4 5 : 10 15 50 15 : Ì Qal : pg 10 »» »» 15 10 5 15 20 10 10 50 psp Qal 35 »»» 15 45 15 10 15 45 pu 40 20 10 Qal Qal 15 10 40 15 Ì 18 pu 12 15 20 17 15 18 10 5 National Park Service U.S. Department of the Interior pei 17 pu Ì 10 ¹ 10 15 16 13 pu ¹ pu : : 15 10 May 2008 ¹ pu 22 20 20 :: : :: : : : 20 ¹ Qal Ì 40 pu 30 30 ¹ 10 21 ¹ 15 17 ¹ 16 17 ¹ pei 18 18 ¹ o ¹ ¹ o oo oo ooo ¹¹ ¹ ¹ ¹¹ o oo o o o oo o o o ooo ¹ ¹ o o o o oo oo oo o o o ¹ ¹ ¹ ¹ ¹ ¹ o o oo ¹ ¹ %% % %%% o %%% ¹ % % % %%% %%%% %% 10 Miles 15 Kilometers ¹ 35 Qal 15 15 ¹ psp 15 ¹¹ php pu psp 18 ¹ Ì 20 10 pu Qal 10 :: pu ¹¹ 9 13 7 17 ¹ Qal 14 pu pu 11 » ¹ 20 Qal 23 15 ¹ pu pm 22 16 ¹ pu pu pth 9 8 : : :: : : : :: : :: : : : : : :: :: : :: : : : : : : : : : :: : : : 13 pu ¹ 18 pu 8 12 ¹ pg psp 12 12 ¹ 28 pu 14 10 ¹ 17 Qal 9 10 ¹ 5 ¹ cb 10 ¹ 8 ¹ » 6 pei Qal : 8 5 7.5 o ¹ o 10 ¹ 9 28 ¹ 18 14 9 26 o 18 psp 13 9 26 o php 14 pu 8 27 3.75 oo o¹ o pg pu 12 ¹ 15 Qts 27 ¹ pu 16 9 ¹ ¹ Qal Qal 15 9 25 o 17 11 18 ¹ o ¹ ¹¹ pth pu 20 12 ¹ ¹¹ ¹ ¹ ¹ ¹ ¹ ¹ 13 15 20 ¹ ¹o o o oo ¹ ¹ ¹ oo ¹o o o oo o oo o o o o o ¹ oooo ¹¹ oo ¹ ¹ ¹ ¹ ¹ ¹ ¹ ¹ ¹ ¹ ¹o oo¹ o ¹¹o ¹¹ oo o ¹ ¹ ¹ ¹ o o pu Qal 14 14 14 20 o oo o o 17 o 16 20 o oo oo o o ¹ oo o o o 20 o 18 Qal 17 17 17 cp o 28 Qal 17 19 18 19 o Qal Qts 17 19 15 o o oo ¹ o o ¹ ¹ ¹ oo ¹¹ ¹ ¹ o¹ ¹ ¹ ¹ 5 pu 8 15 19 18 21 o o o 17 17 18 o 22 15 pm cb 18 19 ooo 25 php Qts 16 10 cu Qal o o o Qal : 10 14 11 18 0 2.5 : pu 13 10 16 18 ¹¹¹ ¹ ¹ oo oo o Qal 11 Qal 18 18 11 12 Qal 17 18 % 18 pg 19 17 13 o o oo o 14 11 14 ¹ : : : Qg pu : 13 o strike and dip of lava flows prospect pit pu Qal Qal % 12 o downthrown side of fault strike and dip of beds Geologic Attitude and Observation Localities : mine shaft Mine Feature Localiteis » diamond drill hole Ì # Paleoshorelines beach crest, known or certain wave-cut cliff, known or certain wave-cut cliff, approximate unknown offset/displacement, concealed unknown offset/displacemen, approximate unknown offset/displacement, know or certain kettle, known or certain drumlin, known or certain direction of glacial movement, known or certain Glacial Line Features Faults 12 base of pegmatitic zone in Greenstone Flow (pg) Linear Geologic Units Qal pg psp 15 pu 13 cu 10 12 cb % pu 15 15 % ¹¹ ¹ ¹ ¹ %%%%% % 9 12 11 o ¹ % %%%% %%% % %% %%%% %% cb 0 ¹ Isle Royale National Park Michigan 11 pwi : : Geologic Map of Isle Royale National Park known or certain Geologic Contacts approximate concealed shoreline shoreline, approximate Qal - alluvium Qts - talus, slopewash, and glacial drift Qg - glacial till cu - Copper Harbor Conglomerate, undivided cc - Copper Harbor Conglomerate, chiefly conglomerate 22 # 9 10 : cb - Copper Harbor Conglomerate, boulder and cobble conglomerate cp - Copper Harbor Conglomerate, pebble conglomerate cs - Copper Harbor Conglomerate, chiefly sandstone pu - Portage Lake Volcanics, lava flows, undivided php 18 pth pu psc - Portage Lake Volcanics, sandstone and conglomerate pp - Portage Lake Volcanics, pyroclastic rocks psp - Portage Lake Volcanics, Scoville Point Flow pei - Portage Lake Volcanics, Edwards Island Flow pmp - Portage Lake Volcanics, Middle Point Flow 19 pu : # # # : 12 ¬ 5 pli - Portage Lake Volcanics, Long Island Flow # # 12 : pth - Portage Lake Volcanics, Tobin Harbor Flow 17 # # # Qal : pwi - Portage Lake Volcanics, Washington Island Flow # ## pm : pg pwi : 10 : pg - Portage Lake Volcanics, Greenstone Flow 20 #15 # #15 # 10 11 : : :: : : pgi - Portage Lake Volcanics, Grace Island Flow 18 pu 9 9 : : pm - Portage Lake Volcanics, Minong Flow 20 12 pu pu cu 10 cu 9 : : : ph - Portage Lake Volcanics, Huginnin Flow 15 Qts 8 9 : : php - Portage Lake Volcanics, Hill Point Flow 14 %%% % 15 17 Qal : pai - Portage Lake Volcanics, Amygdaloid Island Flow water 18 %%%% ¹ 13 10 : : : % % % % % % % % % % % % % % % % % % % % % % gradational o ¹ : Produced by Geologic Resources Division : Geologic Units 13 10 : : oo o oo o o www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Broad Background Any individual place on Earth exhibits only tiny windows of Earth history. In the Keweenaw and Isle Royale, we can see into events that range from about 1.2 billion years ago until perhaps about 0.9 billion (Davis and Paces, 1990), and we can also see the deposits of the glacial periods of the last few million years. To see the record of other times we must travel to where we can see rocks of those ages are at the surface. This is the very best place to see the exposed rocks of the midcontinent rift (Figure 4). This rift extended from at least Kansas to Detroit, but it is exposed only near Lake Superior. At the time of rifting there were huge differences in the configuration of the continents and a huge supercontinent, Rodinia, was assembled, a hodgepodge of pieces of what is now North America, Antarctica, Europe and South America. And it was beginning to break up. In the Keweenaw we get a remarkable opportunity to look at rocks produced during the rifting period of Rodinia, which preceded the orogens shown in green in Figure 5. The orogens mark the areas where continental blocks approached each other at about 1.1 by ago. The orogeny in eastern North America, which eventually ended the Keweenaw Rifting episode, produced an orogen known as the Grenville Front. (Cannon, 1994). Figure 4 Map of the Mesoproterozoic Midcontinent Rift System, showing insets A: the extent of the rift as currently known and B: The main copper districts. from Bornhorst and Barron, 2012. 6 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 5 Two schematic maps of the Rodinia supercontinent showing how pieces of various modern continents are thought to have been assembled more than 1 billion years ago. Sources: John Goodge (left) and KE Karlstrom et al., 1999 (right). Rodinia’s assembly acted like a great blanket for a large area of Earth’s surface, preventing heat loss and creating an opportunity for heat to build up underneath. A great hot spot formed under the blanket. The continent began to split with very hot dike swarms. When the splitting opened the rift, magma was erupted in huge amounts—a supereruption. The ancient Earth contained more radioactive heat producers so the potential for big eruptions was greater. We still think that most of Earth’s heat comes from radioactivity, and we still expect Large Igneous Provinces (LIPs) to develop when and where mantle hot spots occur. But perhaps LIPs are getting smaller as time passes and natural radioactivity declines. Figure 6: Schematic view of the mantle plume head which developed over a hotspot, and which is thought to have led to the midcontinent rift, the great ponded flood basalt lavas of the Keweenaw and Isle Royale. High heat flow focussed on the Lake Superior region led to continental splitting and spreading, forming a rift basin (shown in red) which curved around the current Keweenaw Peninsula. From K Schulz, pers comm., USGS. 7 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Heat flow on Earth is declining with time as natural radioactivity continues to be spent. Convection of Earth’s core and mantle do not produce steady heat transfer from earth’s core to the surface. Since volcanism is driven by higher than average heat flow, volcanism comes and goes as heat flow changes in time and place. Overall, heat declines, but in any time or place, it can vary markedly in both directions. Super-eruptions result from very high heat flow conditions. Figure 7: Map of Supereruptions of the past 2 million years on Earth. Note correlation with the ring of fire. From Geological Society of London. The Midcontinent rift was driven by high heat flow, and it certainly represents a type of supereruption or Large Igneous Province (LIP). To explain the distributions of LIPs in time and place, volcanologists refer to plates, hotspots and/or mantle plumes, much of which which are far from our direct access. These plates, hotspots and plumes come and go, plates move over hotspots and/or plumes, and time/space series patterns are not clearly defined or predictable. This requires volcanologists to consider the deep thermal origin of volcanism, which is fundamental geophysics of the deep Earth and especially the mantle and core. We lack explanations to explain why deep Earth heat transfer leads to massive volcanism at rare intervals and in widely scattered surficial locations. The surface manifestations may be huge volumes of volcanic rocks. The environmental consequences must be large, but are mostly uncertain. From the recent record of LIPs, a relationship of the timing of LIPs with extinctions of living species is advanced. Volcanologists agree that super-eruptions lie in Earth’s future, but the time and place is uncertain. 8 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 8: Map of some LIPs on Earth, plotted in red with yellow hotspot locations and at right, their ages and with extinction evidence, based on the number of animal families represented. Interruptions in the trend toward diversity occur at extinctions. From Bresson, 2011. Heat flow bottom line: The Keweenaw Rift record shows how the earth has highly irregular deep-seated convective events that help shape the planet. They come and go in time and space. Once the hot spot of the whole world, the Keweenaw now has heat flow that is far below average. Figure 9: Stratigraphic units of the Keweenawan from Bornhorst and Barron, 2012. On Isle Royale the upper part of the Portage Lake Volcanics and the Copper Harbor Conglomerate are found. Figure 9 shows this mid-Proterozoic Keweenawan Supergroup, which contains all the formations of the rift. These consist of lavas from the deep earth and redbed sediments, shed off of the top of Rodinia into the gaping rift. On Isle Royale we find only the Portage Lake Volcanics and the Copper Harbor Conglomerate, while on the Keweenaw we have all the formations. The Lavas of the Portage Lake Lava Series are the result of a continental rift, very much like the currently active Red Sea. Existence of a rift is a way to explain how such huge volumes of lava could have been erupted. It also helps explain the syncline shown in Figure 11. A great crack across North America formed, stretching from Kansas to the UP and then on to Detroit. Figures 4 and 6 show the western limb of a feature called the “mid-continent gravity high,” a linear feature that extends from Kansas to Lake Superior where it coincides with the Lake Superior Syncline. This feature is mostly completely invisible, but was detected by geophysicists working with gravity meters, who showed that the gravity attraction of 9 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift earth to the instrument is measurably higher, indicating dense rock underneath. Figure 6 shows a buried dense rock region colored red. The dense rock could be the dense black lava flows we have in the Keweenaw, and their gravity shows that the rift was hundreds of miles long. Drill holes have penetrated the lavas in Kansas and Iowa, so we know that lavas are there—it is not just gravity detection. A second geophysical anomaly, this one even more deeply buried, has been discovered extending from Lake Superior southward to near Toledo Ohio (Figure 4 or 6). This adds to the definition of the hypothesized Keweenaw rift, which is sometimes described as a continental scale fissure, which resembles what happened in the Atlantic to separate Europe from North America. The rift breaks through all the older rock units (Figure 10). Figure 10: Map of Minnesota, Wisconsin, Iowa and Upper Michigan, showing the rift rocks in grey, over the proposed geologic terrane map of Precambrian basement rocks in the northern U.S. continental interior. WRB: Wolf River batholith. Underlying gray-toned base map is the newly compiled regional aeromagnetic anomaly map “Craton margin domain” represents sedimentary and volcanic rocks deposited during the interval 2.3–1.77 Ga; stippled pattern represents area affected by Penokean deformation; cross-hatched pattern represents area termed ‘gneiss dome corridor’ which was affected by Yavapai-interval deformation (Schneider et al., 2004). GIPB: Green Island plutonic belt; BS: Baraboo syncline. Figure and caption from Holm et al., Pre-C Res., 2007. Figure 11: The Synclinal nature of the layers on the Keweenaw and Isle Royale, visualized by NK Huber, USGS. 10 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift The idea of a syncline comes from observed features in geology. In the Keweenaw we cannot see the whole syncline--far from it! We just see the rocks dipping toward the north at Copper Harbor and those dipping to the south on Isle Royale. In between is how geologists earn their money! Figure 12 shows a confirmation of the synclinal nature of the rift rocks, based on seismic geophysics. Implications of this hypothesis: 1. Layers of rock extended from the Keweenaw to Isle Royale, apparently filling a basin. 2. Something caused the basin to subside. 3. The basin has influenced the formation of Lake Superior. 4. The basin may continue beyond the Lake. 5. Its importance could extend much farther than explaining the tilting. Figure 12: Profile across eastern Lake Superior, confirming the geometry of the rift with seismic geophysics (Modified from Behrendt et al. (1988). Basalt Isle Royale is mainly underlain by basaltic lava, the result of hundreds of successive eruptions from the Rift. Mostly this basalt made its way to the surface rapidly, but some was held in magma chambers and evolved before erupting. Basalt is the most common composition of lava rocks that cool from magma, liquid rock that rises from the deep Earth at volcanoes. Today basalt is forming at many active rifts, including Iceland, the East African Rift Valley, the Red Sea and the Rio Grande Valley of New Mexico and Colorado. 11 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Basalt is the result of partial melting of meteoritic material, so it forms on other terrestrial planets as well as Earth, making it the “mother liquor” of volcanoes on terrestrial planets. It is found all over Earth, but especially under the oceans and in other areas where Earth’s crust is thin. It formed in the Isle Royale-Keweenaw region because of the Midcontinent Rift. Most of Earth’s surface is basalt lava, but basalt makes up only a small fraction of continents. Keweenaw lavas are mainly basaltic: continental flood basalts with isotopic signatures close to bulk composition of Earth (Paces, 1988). Within the sequence of flows there are several cycles of evolution in subcrustal magma chambers. Overall the lavas become slightly more primitive with time. The ages are well established from U-Pb dating of zircons. Most of the great outpouring of rift lavas occurred in about 2 million years. Figure 13: U-Pb dates on zircons from pegmatite zones of the Portage Lake Volcanics, Keweenaw Peninsula (Paces and Miller, 1993). Lane (1911) first recognized and described the mirror-image geological and lithological similarity of the PLV and the CHC on both sides of the Syncline (Figure 14), and further suggested that the great lava flow of the Keweenaw Peninsula (Greenstone Flow, Figure 13) and the large flow of Isle Royale are the same. Huber (1973a) strongly supports Lane's correlations. Longo (1984), after extensive field mapping and sampling at Isle Royale and the Keweenaw, gives field observations and geochemical data that also strongly confirms the correlation of the Greenstone flow. Figure 14: Sketch map of Lane, 1911, which suggest the correlations of layers between the Keweenaw and Isle Royale. 12 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 15: Summary of ages and correlations of Keweenawan age rocks around Lake Superior (from K Schulz, pers comm, USGS, modified from Nicholson et al., 1997). This correlation means that the Greenstone flow is one of the earth's largest lava flows; according to Longo (1984), it has an aggregate volume of 1650 km3 (396 mi3), comparable to the Roza flow of the Columbia River Flood basalts, which is estimated to be 1300 km3 (312 mi3) by Swanson et al. (1975). The areal extent of the Roza, 40,000 km2 (15,450 mi2), is much larger than the Greenstone flow, 5000 km2 (1930 mi2), a comparison which results from the ponding of the Greenstone within the rift basin. Thus, the solidification of the Greenstone flow is a kind of magma ocean experiment, the likes of which is rare on this planet. Table 1 at left is from Self et al., 1998. 13 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Paleomagnetism The conceptual model for Earth’s magnetic field is that of a dipole (i.e., bar magnet) positioned at Earth’s center and aligned with the rotational axis of the Earth. This allows us to predict the direction of the magnetic field at any location on Earth’s surface using the fundamental equations of a dipole field. This equation gives a direct relation between magnetic inclination and geographic latitude at the point of observation. The geomagnetic field irregularly reverses (i.e. a magnetic compass which points north will now point south and vice versa) and these reversals are symmetrical (i.e. the normal and reversed field directions are exactly anti-parallel). The above is the fundamental assumption used to reconstruct continents to their past positions using the ancient magnetic field recorded in rocks (fossil magnetism). The record of the strength and direction of Earth’s magnetic field (paleomagnetism, or fossil magnetism) is an important source of our knowledge about Earth’s evolution throughout the entire geological history. This record is preserved by many rocks from the time of their formation. The paleomagnetic data have played an instrumental role in deciphering the history of our planet including a decisive evidence for continental drift and global plate tectonics. The data have also been crucial for better understanding the problems of regional and local tectonics, geodynamics, and thermal history of our planet. The ~1.1 billion-year-old North American Midcontinent Rift paleomagnetism has been intensively studied since early 1960s (for example, see a review in Halls and Pesonen, 1982). The rifting began during an interval of reversed polarity of geomagnetic field. The reversely magnetized (“reversed”) lavas (the Siemens Creek Formation of Powder Mill Group, the lowermost part of North Shore Volcanics, Osler Volcanics, and the lower part of Mamainse Point Formation) are found in many locations around Lake Superior (see figure 15). This early stage magmatism occurred from 1108 to approximately 1105 million years ago. The period of active magmatism was followed by a quiescence period when a geomagnetic field reversal took place. Magmatism renewed by 1102 Ma (Ojakangas et al., 2001) during the normal polarity interval. During this interval, a sequence of Portage Lake lava flows erupted within a two to three million year interval around 1095 million years ago. These rocks represent the main stage of the rift-related magmatism. All younger sedimentary and igneous suites exposed on the Keweenaw peninsula (the Copper Harbor conglomerate, LST, etc) have normal polarity magnetization. However, the geomagnetic field reversal mentioned above is characterized by an asymmetry, manifested in natural magnetization recorded by Keweenawan rocks that crop out around the Lake Superior (e.g., Palmer, 1970; Halls and Pesonen, 1982; Pesonen and Halls, 1983; Schmidt and Williams, 2003). Most but not all of the reversely magnetized lava flows and dikes of this age consistently have characteristic 14 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 16: Equal area projection of the western hemisphere showing the Logan Loop on the polar wandering curve. Letters are keyed to the table below. From Robertson and Fahrig, 1971. 15 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift directions of magnetization that are about 20 to 40 degrees steeper in inclination than their normally magnetized (“normal”) equivalents, while declinations show the expected 180 degree relationship. The paleomagnetic pole positions derived from these normally and reversely magnetized rocks define a noticeable amount of apparent polar wander that forms the western arm of the so-called “Logan Loop” (Robertson & Fahrig, 1971). The two most favored hypotheses for this reversal asymmetry are either apparent polar wander during Keweenawan times (Davis and Green, 1997; Schmidt and Williams, 2003) or the presence of a persistent non-dipole field causing the geomagnetic field to depart from a geocentric axial dipole geometry (Pesonen and Nevanlinna, 1981; Halls and Pesonen, 1982; Nevanlinna and Pesonen, 1983; Pesonen and Halls, 1983). The recent study of this problem (Swanson-Hysell et al., 2009) on lavas from Mamainse Point shows that the geomagnetic reversal asymmetry observed in rocks of Keweenawan age is an artifact of the rapid motion of North America during this time. The other study by Kern et al, (2012) on rocks of the alkaline Coldwell Complex (Ontario, Canada) also suggests no asymmetry in geomagnetic reversal during Keweenawan time. Basalt Geochemistry and field types of basalt on Isle Royale Paces (1988) conducted detailed study of the composition of the lavas of the PLV, studying a complete section on the Keweenaw Peninsula. He provided a description of the texture and thickness (see Basalt Types); chemical composition (Table 2); mineral chemistry (Figure 2); and petrography (Table 3). The lavas resemble other younger examples of continental flood basalts (see also LIPS sources) with their main composition being olivine tholeiite that contains high MgO and Ni, but also have enrichment of highly incompatible elements. There are only minor amounts of more evolved (have more complicated history) magmas and overall the magmas become more primitive (less complicated history) with time. Isotopically (Nd and Sr) the lavas are very close to bulk earth values. Paces (1988) describes the rocks: PLV lava flows display a relatively limited number of textures based on the relationships between dominant mineralogical constituents. These components originally included groundmass plagioclase, olivine, clinopyroxene, iron-titanium oxide, volcanic glass or mesostasis, occasional phenocrysts or microphenocrysts of plagioclase, and sometimes olivine. Textures that developed within the coarsest portion of different lava flows range from fine-grained intergranular through subophitic and ophitic. This same range in textures can be observed in individual, thick lava flows which grade from intergranular chilled flow margins to a coarsely ophitic flow interior. True quench textures (Lofgren, 1971) including skeletal, dendritic or spherulitic olivine and pyroxene, have not been observed in PLV basalts. PLV lava flows do not preserve evidence of an extensive pre-eruptive crystallization history. Chilled margins are generally aphanite. Occasionally, lavas contain minor amounts (usually less than 1%) of small euhedral phenocrysts of plagioclase (often with melt inclusion-rich cores) and sometimes olivine. When present, both of these phases commonly exhibit glomeroporphyritic tendencies. Neither the plagioclase nor olivine phenocrysts show obvious evidence of 16 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Table 2 17 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Table 3 disequilibrium with the liquid. Except for rounded plagioclase cores, both olivine and plagioclase phenocrysts are in apparent textural equilibrium with the liquid. Slightly porphyritic lavas frequently exhibit serrate textures. The dominant textural element in all lavas is the framework of groundmass plagioclase laths. This framework is a randomly-oriented, felt-like structure of interlocking euhedral to subhedral laths. Rarely, the partial alignment of laths forms crude trachytic fabric, indicating movement of magma after at least partial crystallization. The second most prominent textural element is defined by clinopyroxene crystals and their relationships to the plagioclase lath framework. In all cases, clinopyroxene has clearly crystallized later than olivine and plagioclase. Clinopyroxene crystals exhibit 18 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift intergranular to ophitic textures depending both on the size of the clinopyroxene crystals as well as the size of the plagioclase laths. Melaphyric flows and chilled flow margins contain small, blocky clinopyroxene crystals intergranular to the plagioclase framework. In many (but not all) thicker flows, clinopyroxene grains begin to enclose subophitically, and eventually ophitically, plagioclase and olivine crystals as the massive flow interior is approached. The boundary between subophitic and ophitic textures is gradational and is exceeded when a significant number of plagioclase laths are completely enclosed by the surrounding clinopyroxene oikocrysts. Absolute size of the oikocryst is not definitive: a large clinopyroxene grain may only subophitically enclose large groundmass plagioclase laths, however the same sized grain may ophitically enclose plagioclase laths of smaller dimensions. Thus, over half, 60-70% (volume basis), of most PLV lavaflows are typically composed of a plagioclase lath framework with loosely packed clinopyroxene oikocrysts. The remaining interstitial space within the plagioclase framework and between oikocrysts is filled with variable proportions of intergranular olivine, iron-titanium oxides, and intersertal volcanic "glass. " Evidence of gas exsolution is preserved in some flow interiors as vesicular cavities of ellipsoidal to highly irregular shapes. Diktytaxitic textures, however, are not apparent. Vesicles are particularly well preserved in thinner flows which quenched rapidly; however, they are observable in some thicker flow interiors as well. --Paces 1988 We conclude that the lavas of the Portage Lake Volcanics are typical of basaltic LIPs on earth and also chemically resemble the basalts of the moon and Mars. In the field, we can see some textural variety of basalts. Basalt is mainly made of two minerals: Plagioclase feldspar and pyroxene. Basalt has several textural varieties such as glassy, massive, porphyritic, vesicular, scoriaceous. Porphyrite or porphyritic basalt (see photos above) is characterized by obvious crystals, usually of plagioclase, which is often white or tan in color. These crystals are typically 19 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift interpreted as phases that formed before eruption, where magma was being stored (in a “magma chamber”).On Isle Royale, there are five main examples of porphyritic basalt flows: The Scoville Point (psp), Hill Point (php) Tobin Harbor (pth) Grace Island (pgi) and Huginnin (ph). “Trap,” melaphyre, or massive basalt typically has no conspicuous crystals, and in its interior regions has a uniform grey or grey brown color (see photos above). On Isle Royale there are four large “Trap” flows: Edwards Island (pei) Long Island (pli), Minong (pm) and Amygdaloid Island (pai). Ophite or Ophitic basalt (see photos above) exhibits a sometimes subtle, knobby texture with equidimensional pyroxenes usually between 0.5 and about 3 cm. On Isle Royale there are 3 main ophitic flows: Washington Island (pwi), Greenstone (pg) and Hill Point (php) . 20 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift These textural types of basalt reflect environment of deposition in part. Thicker flows which cooled more slowly are more likely to be ophitic, as figure 17 shows. Figure 17: Histogram plots of numbers of flows within the Portage Lake Volcanics which had melaphryic (Trap), and Ophitic textures. Subophitic textures are intermediary between ophitic and melaphyric. From Paces, 1988. Two conclusions emerge from Paces’ work: (1) the lavas are compositionally similar throughout the section and generally are high magnesium, olivine tholeiites; and (2) the flows range from less than 10 m (33 ft.) to more than 100 m (330 ft.) thick, and the thicker ones are more likely to have ophitic textures. Physical features of lava flows A summary statement from a review paper about basalt flows (Self et al., 1998): The most common rock type at the surface of the Earth, and on the other terrestrial planets, is basalt. Basaltic lavas come in two forms: aa and pahoehoe (from the Hawaiian ‘a’ā and pāhoehoe). Pahoehoe flows have often been thought of as small, slow-moving, inconsequential lavas. It is thus not surprising that the processes involved in the emplacement of large, fastmoving, channelized aa flows have received greater attention (see Kilburn & Luongo 1993, Crisp & Baloga 1994, Pinkerton & Wilson 1994, and references therein). However, as in the fable of the tortoise and the hare, it is the slow but unrelenting pahoehoe lava flows that 21 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift ultimately grow larger and longer than the spectacular but short-lived channelized rivers of lava that produce aa flows. In terms of both areal coverage and total volume, pahoehoe flows dominate basaltic lavas in the subaerial and submarine environments on Earth. The most abundant type of lava, submarine pillows, is closely related to pahoehoe in their style of emplacement (e.g. Macdonald 1953, Williams & McBirney 1979). A compilation of the rather sparse information on intermediate length (50–100 km) and long (>100 km) lava flows on the Earth (Table 1) shows that pahoehoe is far more common in these larger flows. Several large extraterrestrial flows also seem to be pahoehoe (e.g. Theilig & Greeley 1986, Bruno et al 1992, Campbell & Campbell 1992). The emplacement of pahoehoe flows is therefore a fundamental process in crustal formation on the Earth and the other terrestrial planetary bodies. Isle Royale and Keweenaw lava flows exhibit pahoehoe features, and do not show pillows or other subaqueous physical aspects. Therefore, here we use descriptive material from volcanological literature that describe pahoehoe flood basalts (Hon et al. 1994; Goff, 1996, Self et al., 1998 and Thordarson & Self, 2012). A generalized cross section of an “inflated” pahoehoe flood basalt is shown in Figure 18. Figure 18: Idealized cartoon of the cross section through an inflated pahoehoe lobe. The lobe is divided into three sections on the basis of vesicle structures, jointing, and crystal texture. The upper crust makes up 40–60% of the lobe and the lower crust is 20–100 cm thick, irrespective of the total lobe thickness. Upper crust: Vesicular, often with discrete horizontal vesicular zones (VZs) that form during active inflation. Bubble size increases with depth. Prismatic or irregular jointing, sometimes equivalent to the entablature in thick lava flows. Petrographic texture ranges from hypohyaline to hypocrystalline (90–10% glass). Core: Very few vesicles. Porosity is dominated by diktytaxitic voids. Vesicles are mostly in the silicic residuum, which forms vesicle cylinders (VCs) and vesicle sheets (VSs). Holocrystalline (<10% glass). Lower crust: Nearly as vesicular as the upper crust, few joints, and 50– 90% glass. from Self et al., 1998. 22 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Development of the layers shown in figure 18 is likely the result of a sequence of inflation and deflation events as observed at Kilauea and Mauna Loa and depicted and explained by Hon et al., in Figure 18 and below: Inflated pahoehoe sheet flows have a distinctive horizontal upper surface, which can be several hundred meters across, and are bounded by steep monoclinal uplifts. The inflated sheet flows we studied ranged from 1 to 5 m in thickness, but initially propagated as thin sheets of fluid pahoehoe lava, generally 20-30 cm thick. Individual lobes originated at outbreaks from the inflated front of a prior sheet-flow lobe and initially moved rapidly away from their source. Velocities slowed greatly within hours due to radial spreading and to depletion of lava stored within the source flow. As the outward flow velocity decreases, cooling promotes rapid crustal growth. At first, the crust behaves plastically as pahoehoe toes form. After the crust attains a thickness of 2-5 cm, it behaves more rigidly and develops enough strength to retain incoming lava, thus increasing the hydrostatic head at the flow front. The increased hydrostatic pressure is distributed evenly through the liquid lava core of the flow, resulting in uniform uplift of the entire sheet-flow lobe. Initial uplift rates are rapid (flows thicken to 1 m in 1-2 hours), but rates decline sharply as crustal thickness increases, and as outbreaks occur from the margins of the inflating lobe. One flow reached a final thickness of nearly 4 m after 350 hr. Inflation data define powerlaw curves, whereas crustal cooling follows square root of time relationships; the combination of data can be used to construct simple models of inflated sheet flows. As the flow advances, preferred pathways develop in the older portions of the liquid-cored flow; these pathways can evolve into lava tube systems within a few weeks. Formation of lava tubes results in highly efficient delivery of lava at velocities of several kilometers per hour to a flow front that may be moving 1-2 orders of magnitude slower. If advance of the sheet flow is terminated, the tube remains filled with lava that crystallizes in situ rather than draining to form the cave-like lava tubes commonly associated with pahoehoe flows. Inflated sheet flows from Kilauea and Mauna Loa are morphologically similar to some thick Icelandic and submarine sheet flows, suggesting a similar mechanism of emplacement. The planar, sheet-like geometry of flood-basalt flows may also result from inflation of sequentially emplaced flow lobes rather than nearly instantaneous emplacement as literal floods of lava. 23 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 19: A figure schematically describing the development of Hawaiian pahoehoe lavas with inflation and deflation (from Hon et al., 1994). 24 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift An inflated view of lava flow sections is probably appropriate for Isle Royale, given the ponded constraints of the rift valley and the thickness of flows observed in the Portage Lake Volcanics. Figure 20 shows a sequential interpretative development of layers in Portage Lake lava flows, based on numerous examples of flows exposed in cross section. This view has similarities with Figure 18, and also is analogous with solidification in sills, based on work by Bruce Marsh (Figure 21; Marsh et al., 1991; Mangan & Marsh, 1992), and also shows how liquid can be squeezed out of mush below forming a cylindrical feature moving up and then trapped in a horizontal layer. Figure 20: Cross section cartoons of Keweenawan lava flows at various stages of solidification, from Paces (1988). A is an early stage when crust has formed on the top and bottom of the flow, While B and C show later stages in solidification as liquid (darkest color) is progressively restricted to the interior, away from the cooling margins where magma is becoming a crystal mush and eventually a solid, and segregations develop from mushy regions. Figure 21: Cross section of a sill, solidifying from its top and bottom and building crystal mush layers from its cooling surfaces both above and below a liquid layer near the sill’s center. Temperature and crystal size vary with height as shown and segregations form and can rise from the lower part of the flow, but get trapped in tabular zones above the center. 25 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Vesicular zones in the PLV tend to be mineralized by zeolite and prehnite-pumpellyite facies minerals, which is an overprint over strictly physical volcanological features. Figure 21 shows a typical pattern of vesicular zones within these lavas and Figure 23 shows some typical segregation cylinders. Goff (1996) has made an extensive study of vesicle cylinders which we suggest are equivalent to segregation cylinders, and develop above the lower solidification front of the lava flow. Figure 22: Cross section of an idealized lava flow within the Portage Lake Volcanics, showing four kinds of regions where gas filled vesicles typically later become mineralized by hydrothermal fluids. Pipe Vesicles (see Figure 23) develop at the base of the flow, perhaps the result of boiling of trapped meteoric water from the soil below the flow. Segregation cylinders or vesicle cylinders (Figure 24) develop above the solidifying lower contact zone, and rise to the flow center or beyond, creating vertical vesicle rich features. At the flow top, vesicles develop as the lava crust thickens and solidifies, with vesicles being more numerous and smaller at the top and less numerous and larger across the first meter or so of flow thickness. A pegmatite zone is found occasionally above the flow midpoint, marked by tabular zones, with thin flows marked by vesicular layers (called vesicle sheets) and thicker flows being termed doleritic, with larger and more conspicuous plagioclase laths. 26 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 23: Pipe vesicles, filled with Calcite and laumontite, seen at the base of a 5m thick PLV lava flow from near Eagle Harbor on the Keweenaw Peninsula. Figure 24: Two examples of mineralized vesicle cylinders or segregation cylinders, from the Keweenaw (left) and Isle Royale (right). These features are generally found below the flow midpoint, and have variable vertical extension. Some conclusions about the lavas of the Keweenaw Rift: 1. The overall physical characteristics resemble other examples from much younger flood basalts and other basaltic volcanoes. 2. The PLV are subaerial, inflated pahoehoe flows which are ponded and do not deflate after eruption. 3. The volumes of PLV flows are as large as any known in other flood basalts. 4. Because their thicknesses are in excess of hundreds of feet, PLV flows show more pronounced in situ differentiation than other examples. 27 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Columnar Joints Like mudcracks, columnar joints form from volume contraction. In the mudcracks, volume decreases with drying, while in lava flows or volcanic tuffs it is cooling that drives the contraction. Lava flows display a variety of columns (Figure 25), often with a stratigraphic pattern. Colonnade is a coarser, more regular pattern often found at the base of the flow. Entablature is more irregular, and often found near the top. Sometimes there is a sandwich colonnadeentablature-colonnade structure like Figure 25 (Long & Wood, 1986). The Portage Lake Volcanics show columnar joints in many places. They also exhibit difference scales and styles of jointing. Figure 25: Schematic diagram of columnar jointing pattern in the Columbia River flood basalt near Bend, Oregon (left), compared with an actual photograph of one good example of a lava cross section. Individual sections never match perfectly because of environmental variables. The recognition of the role of water infiltration in the formation of certain kinds of entablature jointing (see above) in the Columbia River Flood basalts by Long & Wood, 1986 was an especially important insight (see Iceland examples especially), as was the detailed work on column formation by DeGraff and Aydin (1993) and DeGraff et al. (1989). On Isle Royale, colonnade style jointing can be seen in many places, although it is less perfectly developed than many worldwide examples. Entablature jointing is also prominent at Isle Royale, especially in the Edwards Island flow (pei) and the top of the Greenstone Flow (pg). To demonstrate the variability of columnar joints in lavas and tuffs, the field trip website explores a large collection of columnar joint photographs (http://www.geo.mtu.edu/~raman/SilverI/ IRKeweenawRift/Columnar_Joints/Columnar_Joints.html). 28 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 26: Colonnade style jointing on Isle Royale. Left photo shows Monument Rock, an exhumed (sea stack) column several meters in diameter. Right photo shows rude 5m diameter columns in the Greenstone Flow (pg). For more, see also (http://www.geo.mtu.edu/~raman/ SilverI/IRKeweenawRift/IR_Column_examples/IR_Column_examples.html) Figure 27: Entablature style jointing in the Edwards Island Flow, Scoville Point, Isle Royale. Scale of these joints is 7-12 cm. Mafic Volcaniclastic Deposits Kilauea and Iceland mainly produce lava flows like those on Isle Royale, but near their vents we find compositionally similar pyroclastic rocks of a variety of types. These pyroclastic rocks, called mafic volcaniclastic deposits (MVD) are also a minor part of the rock record at Isle Royale. We note that such rocks are well known at most continental flood basalt provinces (see Ross et al. 2005). Mechanisms for generation of these deposits include magmatic and phreatomagmatic processes. On both Isle Royale and the Keweenaw such deposits are noted in a few stratigraphic horizons. 29 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift In a review paper, Ross et al., 2005, have summarized worldwide occurrences of MVD: Flood volcanic provinces are assumed generally to consist exclusively of thick lavas and shallow intrusive rocks (mostly sills), with any pyroclastic rocks limited to silicic compositions. However, mafic volcaniclastic deposits (MVDs) exist in many provinces, and the eruptions that formed such deposits are potentially meaningful in terms of potential atmospheric impacts and links with mass extinctions. The province where MVDs are the most voluminous—the Siberian Traps—is also the one temporally associated with the greatest Phanerozoic mass extinction. A lot remains to be learned about these deposits and eruptions before a convincing genetic link can be established, but as a first step, this contribution reviews in some detail the current knowledge on MVDs for the provinces in which they are better known, i.e., the North Atlantic Igneous Province (including Greenland, the Faeroe Islands, the British Isles, and tephra layers in the North Sea basin and vicinity), the Ontong Java plateau, the Ferrar, and the Karoo. We also provide a brief overview of what is known about MVDs in other provinces such as the Columbia River Basalts, the Afro-Arabian province, the Deccan Traps, the Siberian Traps, the Emeishan, and an Archean example from Australia. The thickest accumulations of MVDs occur in flood basalt provinces where they underlie the lava pile (Faeroes: >1 km, Ferrar province: >400 m, Siberian Traps: 700 m). In the Faeroes case, the great thickness of MVDs can be attributed to accumulation in a local sedimentary basin, but in the Ferrar and Siberian provinces the deposits are widespread (>3x105 km2 for the latter). On the Ontong Java plateau over 300 m of MVDs occur in one drill hole without any overlying lavas. Where the volcaniclastic deposits are sandwiched between lavas, their thickness is much less. In most of the cases reviewed, primary MVDs are predominantly of phreatomagmatic origin, as indicated by the clast assemblage generally consisting of basaltic clasts of variable vesicularity (dominantly non- to poorly-vesicular) mixed with abundant country rock debris. The accidental lithic components often include loose quartz particles derived from poorly consolidated sandstones in underlying sedimentary basins (East Greenland, Ferrar, Karoo). These underlying sediments or sedimentary rocks were not only a source for debris but also aquifers that supplied water to fuel phreatomagmatic activity. In the Parana´–Etendeka, by contrast, the climate was apparently very dry when the lavas were emplaced (aeolian sand dunes) and no MVDs are reported. Volcanic vents filled with mafic volcaniclastic material, a few tens of metres to about 5 km across, are documented in several provinces (Deccan, North Atlantic, Ferrar, Karoo); they are thought to have been excavated in relatively soft country rocks (rarely in flood lavas) by phreatomagmatic activity in a manner analogous to diatreme formation. 30 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift On Isle Royale at least three occurrences of MVDs were noted by NK Huber: 1. A breccia found above the Amygdaloid Island Flow (pai) (Figure 28), 2. a Tuff-breccia unit at the top of the Minong Flow (pm) and 3. A Tuff-breccia above the Greenstone Flow (pg). On the field trip we plan to visit the Amygdaloid Island occurrence. We will also see sedimentary units on Mott Island which resemble MVD. Figure 28: Breccia occurring above the Amygdaloid Island Flow, collected from the south shore near the E end of the island (from Huber, 1973). Lava Stratigraphy on Isle Royale. Huber (1973) named eleven distinctive lava flows (Figure 30) from the sequence of lava units on Isle Royale, using their field characteristics (see above). These units can be traced across the island generally paralleling the elongation of the whole island. These named flows are generally the thickest and most resistant to erosion so they make topographic highs and project as islands at the margins of the main island, accounting for the smaller units of the archipelago. This layered stratigraphy is quite regular (Figures 29, 30, 31). Figure 29: Cliff section of Icelandic lavas, showing a sequence of parallel layers with variable thicknesses. We do not generally have vertical sequences like this at Isle Royale, but the layers must have very similar geometry. Photo from along the south coast of Iceland near Hof, 2008. 31 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 30: Named lava flows of Isle Royale (from Huber, 1973). Map symbols in red. psp pei pmp pli pth pwi pp pg pgi pm ph php pp pai 32 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 31: Longitudinal Stratigraphic section showing variations in thickness of the Portage Lake Volcanics and Copper Harbor Conglomerate on Isle Royale. From Huber, 1973. Ophitic Texture and Ophite significance Keweenaw rift rocks include a somewhat rare textural variety of basalt called ophite or ophitic basalt. Ophitic texture is defined inconsistently, but it is an important variety of basalt texture where pyroxene (or occasionally olivine) forms larger crystals and typically contains numerous crystals of plagioclase (Figure 32). Pyroxenes may vary from < 1 to 10 cm and may include as many as hundreds of plagioclases. In the field the pyroxenes are often 1-2 cm in diameter and give the rock a distinctive aspect. There may be a brownish or orange region surrounding the pyroxenes which may represent a glassy remnant of magma melt. Overall the ophite is thought to represent a solidified remnant of a dendritic crystal mush. 33 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Crystal size and form in volcanic rocks is known to be influenced by the rates of cooling in the immediate vicinity of the growing crystal. Slow cooling in a pluton leads to large, Figure 32: Ophitic cobbles (left) and a wet surface of an ophitic lava flow (right) are common on Isle Royale and the Keweenaw, and quite rare elsewhere in the world. The ovoid features in both photos are clinopyroxenes, while the orange or reddish material surrounding the pyroxene is typically a glassy mesostasis which is now altered to chlorite or corrensite. equidimensional crystals, while very rapid cooling can lead to no crystals at all (glass or obsidian). Intermediate cooling rates can lead to unusual shapes of crystals (spherulites, “bow ties”, spinifex, and ophitic) as crystals nucleate or grow at accelerated rates as crystallization, which requires more time than allowed by the environmental cooling of the lava, cannot keep pace and exhibits disequilibrium (Lofgren, 1980). The rate of heat loss (undercooling or supercooling) during the solidification is thus thought to cause ophitic texture, where pyroxene is growing rapidly and plagioclase is forming many more nuclei. Because ophites may completely crystallize and can be coarse-grained, especially with respect to pyroxene, some are termed gabbro rather than basalt. At first geologists looking at ophitic lava flows in the Keweenaw wondered whether they were sills. There is a tendency for ophitic textures to be found in large basaltic intrusive rock bodies such as sills, suggesting that overall they reflect relatively slow solidification. Overall ophitic texture is ubiquitous and could be a hallmark of the Keweenaw Rift lavas. Paces (1988; see Figure 17) found that the average thickness of ophitic Keweenaw flows was 33 m (range 11-140m), while subophitic ones were 12 m (range 4-45 m), and traps (melaphyres) about 5 m thick (range 2-60m). We note that the overall average thickness of Keweenawan flows is about 10-11m, much greater than what we see at modern volcanoes like Kilauea (average flow about 0.5 m thick). The differences are likely the result of ponding within the rift valley, where volcanism filled the rift basin rather than running off a slope away from the vent, as happens at Kilauea. So ophitic texture is a hallmark of slow cooling that is apparently related to ponding of the lavas. 34 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Pegmatites or pegmatoids in lava flows In Keweenawan flows pegmatite (or pegmatoid) layers are conspicuous (Fig 33), especially in the thicker flows. They appear to be analogous to vesicle sheets that are found in most flood basalts, but they may result from more evolution during the longer solidification times. Figure 33: This collection of beach cobbles shows obvious texture of Keweenawan lava pegmatite—note conspicuous plagioclase laths. These layers have vesicular texture and are typically mineralized with zeolite facies minerals. The thickest lava flows in the Keweenawan Portage Lake Volcanics contain horizons called “pegmatites,” “pegmatoids,” or “dolerites.” The following description of these features is from Longo (1984): Lacroix (1928, 1929) coins the term ''pegmatitoide" to describe the coarse-grained zones considered to represent the final stages of differentiation in basaltic lavas of France. The lavas of Michigan's Copper Country show similar differentiates for which Lane (1893) applies the term "doleritic." Cornwall (1951) adopts the textural term "pegmatite" from the usage of Butler and Burbank (1929). He changed the confusing "doleritic" term to "pegmatitic facies, " and subsequently described such units in the Greenstone flow, Big Trap, and several other large flows within the PLV on the Keweenaw Peninsula. For the present study, the term "pegmatoid zone" from Lindsley et al. (1971) is adopted to encompass the portion of the Greenstone flow with numerous en echelon, lens-shaped pegmatoids, associated granophyric phases, and subophitic layers. Texturally, pegmatoids are coarse grained when compared to ophitic zones. Coarse plagioclase laths dominate with interstitial, subhedral clinopyroxene and abundant interstitial to somewhat poikilitic magnetite and ilmenite. Consequently, the pegmatoids are strongly magnetic compared to ophitic units. This suggests that a higher titaniferous magnetite/ ilmenite ratio for magmatoids than for ophites. Visual inspection generally reveals a greater overall opaque (oxide) concentration in the pegmatoids. Subophitic layers are often found hosting the en echelon pegmatoids. These layers, like pegmatoids, are strongly magnetic and very coarse grained stratiform features, but contain less abundant, smaller sized pyroxene. The 35 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift contacts between pegmatoids and subophitic units are usually sharp, although instances of gradational contacts have been observed. Subophitic layers grade into the ophites and seem to occupy the greatest volume of the pegmatoid zone. They have been observed to pinch out within pegmatoid units and may not be continuous planar features throughout the flow. Perhaps pegmatoid units are not only lens-shaped but also flattened amoeboid-like features interfingering with subophitic layers. The frequency of pegmatoids and subophitic layers increases proportionally with increasing flow thicknesses. Both vary in thickness and shape and typically occur in the upper half of a lava flow. Pegmatoids have also been observed as auto intrusions, such as in the entablature on Isle Royale and the upper ophite on the Keweenaw Peninsula. The stratiform pegmatoids are usually found armoring the tops of cliffs formed of the lower ophite. The extension of weak vertical joint patterns into the pegmatoid (forming crude large columns) suggests that pegmatoids may be part of the colonnade. In most cases pegmatoid zones separate a basal colonnade from an upper colonnade. Pegmatoids are not unique to thick flows of the PLV. Lindsley et at. (1971) assert that three of the thicker flows from the Picture Gorge Basalt contained pegmatoid lenses. Santin (1969) discusses the presence of pegmatoids in horizontal basalts of the Lanzarote and Fuerteventura Islands in the Canarian Archipelago. --Longo 1984 Pegmatites are found to be especially well developed in thicker flows such as the Greenstone (pg), which can be more than 1200 ft thick. Pegmatite layers up to 30 ft thick are found above the flow’s midpoint at a stratigraphic layer analogous to the vesicle sheets near the top of the core of idealized pahoehoe flows as described by Self et al., 1998 (see Flow Structure section, above). Cornwall (1951) shows a Greenstone flow section from the Keweenaw in Figure 34. Upper Ophite pg Lower Ophite Figure 34: Columnar section (left) and cross section (above) of the Greenstone flow (pg) as exposed in overlapping diamond drillholes from Delaware, Michigan (Keweenaw Peninsula). Pegmatite is shown as black layers and occurs in the upper part of an unusually thick (1300 ft) lava flow. Granophyre was not found in the cores but is projected based on field data (from Cornwall, 1951). 36 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift The texture of pegmatite in thick flows is coarser and the plagioclase laths may be as large as several cm (fig 35). Figure 35: Polished surface of pegmatite boulder from Passage Island, Isle Royale, showing plagioclase laths of several cm. In thinner flows pegmatite layers are thin (often a few cm) and resemble vesicle sheets (see figure 36). Figure 36: Thin pegmatite or vesicle sheet from 6 m thick lava flow of Lake Shore Traps, Silver Island, Keweenaw Peninsula. Pegmatite layers or vesicle sheets in thinner flows are texturally similar to segregation cylinders and lie stratigraphically above them (Figure 37). 37 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 37: Tabular pegmatite layer within horizontally fractured section of 20-30 m thick flow on Raspberry Island. This 6 cm thick layer is about 5 m stratigraphically above segregation cylinders. Amygdaloidal Minerals in Portage Lake Volcanics on Isle Royale To find minerals at Isle Royale or in the Keweenaw, you should walk the coastlines, especially those that are well wave-washed. The waves expose the minerals and pebbles of various minerals, can be found on adjacent beaches. Using a canoe or small boat, and watershoes and taking plenty of time, walk the shore and watch for veins and amygdaloids. Observe the interiors of basalt flows where vesicle cylinders, pegmatites, joints and veins may expose these distinctive minerals (Figure 22). Individual minerals are sometimes difficult to identity, even for experts, but certain groups of minerals can be distinguished very easily (see Table below). The colors of amygdaloidal minerals are highly variable and distinctive (Figure 38). 38 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 38: A selection of beach pebbles showing various colors of amygdaloidal minerals. See also photos of specific minerals (http://www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift/ Amygdaloid/Pages/Amygdaloid_2.html) For another test, you can use your finger nail. The phyllosilicates, chlorite, corrensite, and saponite are all of green color and very soft minerals. You can easily scratch them with your finger nail. The other green minerals as pumpellyite or prehnite are much harder and you will not be able to scratch them with your fingernail. In fact, in pebbles along the shore they stand out, since they are not as easily eroded as the surrounding rock. The pink unusual color of prehnite of Isle Royale often is the result of very tiny inclusions of native copper which makes it similar to the zeolite thomsonite.The zeolite family is in general difficult to identify, but the zeolite, laumontite, can easily be recognized. It is of white or pink color and if you touch it with your finger nail it will split up into small fibers. What’s next? After mastering the mineral identifications in the boulders, students can also look at amygdular minerals to study the order that minerals were deposited in those vesicles, what mineralogists call paragenesis. 39 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Native Copper and the mining The Midcontinent Rift is the most important and notable location on Earth for native copper. This is truly a cosmic oddity, because copper in nature is typically found as a sulfide. Indeed, Goldschmidt classified copper with a group of elements called “chalcophile”. So why does copper occur in the Midcontinent Rift as native copper (Fig 39)? This is a major puzzle. Figure 39: NATIVE COPPER VEIN ON WASHINGTON ISLAND, ISLE ROYALE NATIONAL PARK. THIS VIEW MAY BE LIKE WHAT NATIVE AMERICANS FOUND WHEN THEY FIRST VISITED THE COPPER COUNTRY. SUCH OCCURRENCES ARE NOT COMMON ANYMORE —THEY WERE DUG OUT OF THE WAVE-WASHED SHORELINES. Could sulfur have been purged from the magma source region or from its magma chambers? This idea is suggested by the early ultramafic dikes which apparently represent the beginning of Midcontinent Rift and which contain apparently immiscible sulfide bodies containing Ni, Cu and rare earth elements (Ding et al., 2012). These dikes could represent magmas derived from mantle material that was melted more completely than when the mantle produces basalt. And this magma may have exsolved sulfide liquid before it was intruded into dikes. Loss of sulfur from the source region or a magma chamber may result in a sulfur-depleted environment favoring native copper? This is a speculation! Another explanation of sulfur loss is that loss of sulfur through degassing of magma from magma oceans would be facilitated by the ponding and long solidification times. Awareness of sulfur emissions from eruptions is heightened by recent studies of eruptions and climate. Could extensive degassing during Keweenawan rifting play a role in eventually forming native copper ore deposits? Speculation! Keweenawan native copper deposits seem to be associated with widespread hydrothermallyinduced zeolite and prehnite-pumpellyite facies metamorphism (Stoiber & Davidson, 1959; Jolly, 1974) which mineralized the permeable lava flow tops and sediment layers of the Portage Lake Volcanics, apparently about 30 ma after the rift volcanism, during the period of Grenvilleinduced deformation of the rift syncline (Bornhorst & Barron, 2012; Nicholson et al., 40 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift 1997 ,Cannon, 1994; Bornhorst et al., 1988) when there was faulting of the rift which enhanced fluid flow within the syncline. There is a rich lore about indigenous ancient copper mining in the Lake Superior region. Most of it is highly speculative and is unsupported, but it is fervently believed. The abundant archeaological copper relicts (Figure 40) leave no doubt that copper was mined at Isle Royale thousands of years ago and traded across North America and beyond. These early mines found native copper in veins at the surface. They left behind pits and dumps. Figure 40: Archeological Copper relicts of midcontinent rift native copper from the Michigan Tech Archives. These materials and open pits left behind show that ancient people mined copper in the Keweenaw and on Isle Royale. Mining by Europeans started in the 1800s on both the Keweenaw and Isle Royale. The Isle Royale mines were all marginal efforts and did not last more than a few years. 41 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Copper Harbor Conglomerate The Copper Harbor Conglomerate occurs on the SW sector of Isle Royale and has been studied by N K Huber, USGS OFR 754-B. Huber gives the following introductory comments: The Copper Harbor Conglomerate, in its type area on the Keweenaw Peninsula of Michigan, was named and defined so as to include a thick sequence of sedimentary rocks, previously separated (in ascending order) into the Great, Middle, and Outer Conglomerates, with intervening lava flows, the Lake Shore Traps (Lane and Seaman, 1907, p. 690-691; Lane, 1911, p. 37-40). On the Keweenaw Peninsula, the Copper Harbor Conglomerate conformably overlies the Portage Lake Volcanics (middle Keweenawan), and locally the two formations interfinger. The Portage Lake Volcanics consists primarily of lava flows; minor sedimentary rocks, similar to those within the Copper Harbor Conglomerate, are intercalated between flows (hereafter referred to as interflow sedimentary rocks). The transition between the two formations reflects a gradual cessation of volcanic activity and the growing dominance of a sedimentary regime. The Copper Harbor Conglomerate is overlain by the Nonesuch Shale and Freda Sandstone (upper Keweenawan, Fig 41). Figure 41: Local Stratigraphic units. Approximately four-fifths of Isle Royale is underlain by volcanic flows and minor clastic rocks of the Portage Lake Volcanics, which dip 10°-20° to the southeast in the vicinity of their contact with the overlying Copper Harbor Conglomerate (Huber, 1973b, Wolff & Huber, 1973). The Copper Harbor Conglomerate underlies the remaining one-fifth of Isle Royale and is confined to the southwestern part of the archipelago; it dips 5°-28° to the southeast. The contact between the Copper Harbor Conglomerate and the Portage Lake Volcanics appears to be conformable; the top of the Copper Harbor Conglomerate, however, is not exposed. If the Nonesuch Shale and other formations that overlie the Copper Harbor Conglomerate on the Keweenaw Peninsula are present in the Isle Royale area, they lie beneath Lake Superior to the southeast. Consisting of fluvial subaerial sandstones, siltstones and conglomerates, The CHC shows transport directions that generally spill into the rift valley (see Fig 42). Huber gives many details of the CHC on Isle Royale in his OFR. 42 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 42: Plot of observed and interpreted paleocurrents seen in the Copper Harbor Conglomerate, Isle Royale (Wolff & Huber, 1973). On the Keweenaw Peninsula the Copper Harbor Conglomerate is partly made up of alluvial fans (Elmore, 1984). On Isle Royale the sandy and silty units are more abundant and cobble sizes are generally smaller. LIDAR Topographic Surveys of Isle Royale LIDAR (LIght Detection and Ranging or Laser Imaging Detection and Ranging) survey of all of Isle Royale, with a nominal resolution of about 2 m is a new resource for understanding landscapes. The data we show here came from Seth De Pasqual, at Isle Royale National Park. It reveals a striking topography which shows the dipping lava beds, and the prominent large lava flows, like the the region NE of Windigo. Differential erosion of lava flows occurs when soft material, like what is found in the amygdaloidal flow tops and along faults is preferentially removed and makes a topographic low, while the massive flow interiors resist erosion and become topographic highs. Glacial deposits mask the lava layers in part, especially southward in the image, where the flows are mostly covered, but protrude through glacial cover. The glacial materials are softer, but they also reveal wonderful geological information. Drumlins are asymmetrical glacial features (Figure 43) which reveal the direction of glacial movement. Figure 44 shows an area near Lily Lake, which depicts conspicuous drumlins south of the lake. The pattern shows the direction of movement (from east to west) clearly, and the degree of elongation is also indicative of the rate of movement. 43 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 43: Schematic diagram of a drumlin, showing how its shape may be related to the direction of ice movement (www.geography-site.co.uk). psp Figure 44: LIDAR topography image of Lily Lake region, Isle Royale National Park, showing multiple drumlins. LIDAR is advantageous over conventional DEM (digital elevation models) for glacial features, but the good resolution of LIDAR also clarifies structural information on the lava flows. The second LIDAR image (Figure 45) shows dramatic bending of the lava flow layering that is remarkably regular in most places on Isle Royale. The bending likely reflects deformation related to faulting associated with McCargoe Cove. The LIDAR offers an opportunity to do interpretation, which will reveal details of the rift formation and its subsequent deformation. Figure 45: (next page) LIDAR topographic image of Pickerel Cove area, Isle Royale. The layered lava flow sequences of Isle Royale stand out clearly as resistant flow interiors resist erosion and stand up to higher levels. Faults which offset the flow layers are also detected as eroded topographic lows. Here there is apparent bending of the lava flows. 44 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Specific Field areas we will visit: Washington Harbor and Windigo The field trip starts at Grand Portage, Minnesota, where we will take the Voyageur II east to Washington Harbor and Windigo about 35 km (22 mi) offshore. Between the Minnesota shoreline and Isle Royale, the strike of Keweenawan rocks, known as the North Shore Volcanics in Minnesota (1109-1100 Ma), changes from E-W to about N 55° E, where the PLV formation (1096-1094 Ma; Figure 13) at Isle Royale begins. This discontinuity could be partially related to the Isle Royale Fault (IRF), which the Voyageur crosses between Grand Portage and Isle Royale. This is a thrust fault which bounds the north flank of the rift, apparently associated with the inversion of the Midcontinent Rift. The IRF was detected in the GLIMPCE (Great Lakes International Multidisciplinary Program on Crustal Evolution) seismic profile (Figure 12) collected on a NS line E of the Keweenaw Peninsula, far from Isle Royale, but along the north flank of the rift zone. It is thought to extend W to at least the SW end of Isle Royale, where Isle Royale is mantled with a much thicker portion of glacial cover and the glacial features are much more prominent (see pp. 20-21 and 41-54 in Huber 1983). 45 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift The bedrock geology of the Washington and Grace Harbor areas (Figure 46) includes four large flows that continue all the way to the other end of the island. The Greenstone Flow (pg) crosses the center of Washington Island and outcrops in several places SSE of Windigo. The Tobin Harbor flow (pth) outcrops at South Rock, SW of Washington Island. The Minong Flow (pm) outcrops S of McGinty Cove, and the Scoville Point Flow (psp) outcrops near Middle Point on the S side of Grace Harbor. The thickest flows in this area are the Washington Island Flow (pwi) and the Grace Island Flow (pgi). Both of these flows occur only locally, from the end of Washington Island to a point between Windigo and Sugar Mountain, a distance of about 14.5 km (9 rni) along strike. The lava flows here dip at 15-20° SE, an attitude that is similar for younger flows on Isle Royale. Vertical N-S trending fractures, with little offset, cut across the bedrock strata near Washington Harbor (Figure 46, 47, 48). Huber (1983) interprets these as structures related to the warping of the Lake Superior Syncline. South of Grace Harbor, the bedrock of the island is buried by till. Washington Harbor ! ! Grace Harbor Figure 46: Portion of figure 2 (Geologic map of Isle Royale, Huber, 1973) showing the area along the west end of Isle Royale. Most of the map indicates glacial deposits, shown in tan, which cover much of the lavas and conglomerates. The prominent locations where bedrock penetrates the glacial deposits are shown in bright colors. Most of eastern Isle Royale has little glacial cover. 46 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 47: Oblique Google Earth view of Washington Island, looking E, showing N-S faults and tilted lava flows. Note: the blue triangle indicates direction and angle of dip (right at about 20 deg.) Figure 48: Oblique Google Earth view of Washington Harbor, looking E. The Windigo area was the site of the last serious mining on Isle Royale, from 1890 to 1892. After failure and closure of mines farther E, the Wendigo Copper Company (renamed from the Isle Royale Land Corporation) founded a mining venture on 8000 acres of land at Washington Harbor, under the leadership of Jacob Houghton, brother of Douglass Houghton. The town site was named Ghyllbank and was located near the present site known as Windigo. The mine site, about 2 km (1.25 mi) inland to the NE, was named Wendigo. People built roads all around the W end of Isle Royale, and 135 people lived at the mine site. The company did diamond drill exploration, as well as extensive trenching. In 1892, the miners gave up and left. When mining stopped, the company tried to sell land to tourists and resort owners (Rakestraw 1965). 47 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift The N Side of Isle Royale: The Hill Point Flow After Windigo, we travel along a straight section of the coast following the Hill Point Flow (php). Muted by glacial deposits, the layered strata of lava flows shows in the geomorphology. Note the cross cutting faults (dotted lines in Figure 49), which are conspicuous in this area. These faults may have formed during deformation of the rift during its subsidence and during the Grenville Orogeny. The faults may have enhanced fluid flow, zeolite facies metamorphism, and copper mineralization. The faulted Windigo area is one place where some mining occurred. Figure 49: Oblique Google Earth View, looking S from Hugginin Cove, directly across the stratigraphy, with flows dipping away from the view. With the flows dipping SE, moving toward the N side of the island takes us further into the PLV section, until we reach the horizon of the Hill Point Flow (php). This is an ophitic flow, forming imposing cliffs along the shore from Hugginin Cove all the way to Todd Harbor, a distance of about 24 km (15 mi). This flow also makes up the majority of shoreline from Pickerel Cove all the way to Hill Point itself, at the W end of Five Finger Bay, about 64 km (40 mi) from Windigo. The tilted strata along the shore make the shoreline steep, and the prevailing winds from the NNW can make conditions treacherous for small boats. The Hill Point Flow is a coarse-grained, ophitic unit with augite oikocrysts of 2 cm (0.8 in) or more. The vertical fractures superimposed across the dipping strata are noticeable throughout the entire flow. From the west area of the flow to the east area, the fractures gradually begin to change from N-S to more N-E trending. According to Longo (1984), the Hill Point Flow may correlate with a large flow on the Keweenaw Peninsula, the Scales Creek Ophite, which extends all along the Keweenaw Peninsula for more than 160 km (100 mi) of strike length, and right through Houghton, which is about 110 km (68 mi) SSE of Hugginin Cove. 48 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift LIDAR survey (Figure 50), from Seth De Pasqual, at Isle Royale National Park, reveals a striking topography which shows the dipping lava beds, and the prominent large lava flows, like the Hill Point Flow (php) and the Minong flow (pm) in this image of the region NE of Windigo. Differential erosion of lava flows occurs when soft material, like what is found in the amygdaloidal flow tops and along faults is preferentially softer and evolves to a topographic low, while the massive flow interiors resist erosion and become topographic highs. The prominent NS faulting of the lava layers is obvious, as are less extensively altered NE trending faults. Glacial deposits partially mask the lava layers, especially southward in the image, where the Grace Island (pgi) and Greenstone Flows (pg) are mostly covered, but protrude through glacial cover. Trails are plotted in yellow. This LIDAR data is advantageous for structural geology study because of its sensitivity to faults. It also reveals details of glacial (drumlins, outwash, kames, etc.) and postglacial features (shorelines, mine pits, dumps and roads). php pm php php pm pm pgi pg pgi pg Figure 50: LIDAR topographic image of comparable area to Figure 48, showing how LIDAR is advantageous for structural studies. (Seth De Pasqual, NPS). Trails in yellow. McCargoe Cove 49 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift At the midpoint of the island is McCargoe Cove, which is a linear, 3.2 km (2 mi) long inlet (Figure 51) that follows a large fracture zone, trending N 30º E to a campground site located along an ancient Native American portage route and near a mine, the Minong Mine. Native Americans left hundreds of ancient pits as relics of mining over centuries at this site, and in 1874 three companies were formed in Detroit to exploit the potential here. They built a dock and a warehouse, and started to build a railroad. Some large masses of copper were successfully mined, and the community here grew for several years in spite of difficult winter conditions. But mining did not last beyond 1885 (Rakestraw 1965). Figure 51: Oblique Google Earth View of McCargoe Cove, looking SW. Lava layers are dipping to the left with steeper dips below and shallower ones above. LIDAR survey (nominal resolution of about 2 m; Figure 52), from Seth De Pasqual, at Isle Royale National Park, reveals a striking topography which shows the dipping lava beds, and the prominent large lava flows, like the the Minong Flow (pm) in this image of the region W of the McCargoe Campground. 50 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift pm pm Figure 52: LIDAR topography of the area west of the McCargoe Campground, showing the pits and dumps of the Minong Mine. These features cannot be resolved in DEM-based topo maps with lesser resolution. As in previous examples, increased erosion of lava flows in the amygdaloidal flow tops and along faults makes topographic lows, and flow layers and prominent NE trending faulting is obvious. Here the mine pits and dumps associated with the Minong Mine are also easily resolved, which shows how LIDAR can map topographic features that are difficult to resolve through vegetative cover. Copper mineralization in the area above a thick lava flow is common, perhaps due to the effect of channeling fluid, as the flow interiors are relatively impermeable and act as a hydrologic dam. The Amygdaloid Channel From McCargoe Cove, we will continue to the NE, passing through the Amygdaloid Channel (Figure 53). Amygdaloid Island is composed of the oldest lavas of the PLV on Isle Royale and is supported by a large flow, the Amygdaloid Island flow (pai), which is a fine-grained basalt (termed "trap"). At the W end of Amygdaloid Island is the National Park Service (NPS) ranger station near Kjaringa Kjeft. Crystal Cove, 3.2 km (2 mi) E of the station, was, beginning in 1906, a private residence and fishery. 51 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 53: Oblique Google Earth View of the Amygdaloid Channel, looking NE, with Amygdaloid Island to the left and beds dipping to the right at increasingly shallow angles. As we travel through the Amygdaloid Channel, drowned ridge and valley topography of Isle Royale will become very visible, with more resistant lava flows holding up linear islands. Amygdaloid Island is the site of mafic volcaniclastic deposits (pp on Amygdaloid Island in Fig 53). It also has a sea arch (left) which is located almost directly opposite the keyhole. (Fig 54). Shipwrecks are numerous on the many "reefs" found all around the NE end of Isle Royale. Opposite Crystal Cove on the south side of Amygdaloid Island is Belle Isle, a beautiful campground accessible only by boat and canoe, located on the site of a resort that operated in the 1920s, when it served the grand lake steamers of that period. Figure 54: Sea Arch on Amygdaloid Island. 52 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Blake Point—a key locality As we round the tip of Isle Royale to Blake Point (Figure 55), we are moving up in the stratigraphic sequence. We will first cross the Hill Point Flow (php) at Hill Point, then the Minong Flow (pm) near Locke Point, and finally the Greenstone Flow (pg) at the Palisades. The Greenstone Flow is perhaps Earth's largest lava flow. Blake Pt Locke Pt Hill Pt Figure 55: Blake Point segment of Geologic Map of Isle Royale (Huber, 1973). At right, photos of columns at the Palisades on the anti-dip slope just west of Blake Point. The following are comments by Longo (1984): Similarities in the stratigraphic sequence of Isle Royale and the Keweenaw Peninsula of Michigan were recognized by numerous workers prior to 1851. The first thorough study of both areas, conducted by Lane (1893, 1911), resulted in the correlations of specific rock units. One unit in particular, due to its persistence as a prominent ridge on both Isle Royale and the Keweenaw Peninsula, became Lane's most convincing evidence for a correlation across this section of the Lake Superior syncline. 53 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Lane (1893) states, "The backbone ridge thus agrees in every way with the great corresponding ridge on the Keweenaw Point." Outcrop and drill core data by Lane (1893) reveal this unit as a single immense, differentiated lava flow. Lane (1893) refers to the flow as "the Greenstone, the 'backbone' and biggest ophite of all, with the bed at its base we correlate as the Allouez Conglomerate. " The Greenstone's great thickness and differentiated nature led some workers to consider it as an intrusive sill (Seaman and Seaman 1944; Van Hise and Leith 1911). However, convincing data have proven this unit to be a lava flow (Lane 1893,1911; Butler and Burbank 1929; Broderick 1935, 1946; Cornwall 1951), and henceforth known as the Greenstone flow. Huber (1973a) confirms the similarities of the Greenstone flow on Isle Royale and the Keweenaw Peninsula, and he supported the correlation. --Longo 1984 The shoreline around Blake Point offers the best view of the Greenstone Flow, better than any other sites at Isle Royale or the Keweenaw Peninsula (Figure 56). On the way to the campground in Merrit Lane, the starting point of our Blake Point walk, we will pass the NW side of Edwards Island, which has good exposures of entablature columnar joints in the Edwards Island flow (pei). The boat will let us off at the Merrit Lane Campground for our walk to Blake Point. We will follow the shoreline from Merrit Lane around the point, remaining close to the wave-washed rocks, yet trying to keep our feet dry. Most of the walk is on the upper ophite unit of the Greenstone flow. (The entablature part of the Greenstone and its flow top is underneath Merrit Lane, and we will see parts of this from the boat later). The upper ophite exhibits a poorly-developed columnar structure all along the walk, with the columns perpendicular to the bedding. The size of the oikocrysts increases from top to bottom. After rounding the corner, we will cut through the bushes to descend a cliff that marks the lower anti-dip face of the upper ophite. At the base of this cliff, we will see wave-washed exposures of the pegmatoid, here about 23 m (75 ft.) thick. The contact here appears to be quite sharp, although Huber (1973a) says it is frequently gradational. The pegmatoid underlies the low shoreline and also the area under the light tower. A section of the Greenstone flow is exposed on Passage Island, a 2 km (1.2 mi) long island that can be seen about 4 km (2.5 mi) offshore from Blake Point. Around the corner from the tower and vertically down about 4 m (13 ft.) is the contact with the lower ophite (which is too difficult for us to reach safely). Longo (1984) describes the contact as a gradation over about 1 m of thickness. From here, we will return by the same route to Merrit Lane. Weather permitting, we will travel around the point in the boat to examine the lower ophite cliffs along the Palisades. The columns exposed on the anti-dip slope are up to several meters across. The base of the Greenstone flow is not exposed here. Figure 56 (next page): Oblique Google Earth View of Blake Point, looking SW, with beds dipping about 25 degrees to the SE. Most of the large land mass is underlain by the Greenstone flow, and its three distinct layers can be seen and outlined here better than anywhere else. Below is a photo from just offshore at Blake Point, at the water level. 54 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 57: LIDAR survey of Blake Point area. Most of the land imaged here is the Greenstone Flow and this image is remarkable in showing indications of layering, and also the different character of layers. On the south side of the main land body, facing Merrit Lane, are eroded remnants of the Upper Ophite layer, with its columnar jointing and dipslope aspect. On the North side there is a steep slope (Palisades) where the Lower Ophite is exposed on an antidip slope. Between these two layers lies the pegmatite of the Greenstone, which is 75 ft thick and which appears to be eroding in an irregular, wavy pattern. It is remarkable that the LIDAR shows information about these three layers. This information is valuable because we typically do not have very good exposures. Seth DePasqual, IRNP. peg peg upper ophite peg upper ophite peg upper ophite 55 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Passage Island and Gull Rocks Farther to the NE, off of Blake Point, Passage Island, 3.5 miles from Blake Point, and Gull Rocks, 8.7 miles away (Figure 58), are both built of rift lava, including the Greenstone Flow, which is found at the E end of Passage I. These are the most easterly subaerial exposures of the PLV near Isle Royale. Figure 58: Oblique Google Earth View of Passage I and Gull Rocks (same scale, but offset). To the right is a piece of the Isle Royale Geologic Map (Huber, 1973). Snug Harbor At this wonderful location in Rock Harbor, the National Park Service has chosen to concentrate its Isle Royale services and concessions for visitors. The Lodge and Visitor Center is where the field trippers will sleep, catch their boat rides and have evening discussions. The location coincides with two of Huber’s named lava flows: the Scoville Point (psp) porphyrite and the Edwards Island Trap (pei) (Figure 59). This location allows boat access to both Tobin and Rock Harbor, as well as foot trails to Scoville Point, Mount Franklin, and Daisy Farm, and is a safe harbor. 56 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 59: Oblique Google Earth View of Snug Harbor, Looking NE. Scoville Point For part of this day's trip we will walk on the rocky dip slope of the Scoville Point flow (psp), facing Rock Harbor along the shore (Figure 60). Huber (1973) describes the basalt of this flow as containing "fine, equant, millimeter sized, plagioclase crystals distributed uniformly through a fine grained matrix." He says the thickness is 30-60 m (l00-200 ft.). There are not many features that can be seen in outcrop, but the flow is very resistant to erosion and buttresses the shoreline. We will take the Stoll Trail (white line in Fig. 60), which goes along the shore of Rock Harbor. Along here, we will see 5000-year-old Nipissing shorelines and glacially grooved outcrops of the Scoville Point flow. Outwash cover here is meager, but kettle lakes and morainal zones occur. On the upper map in Fig. 60, GPS markers identify the points of interest/inquiry. Also, we will be able to see the ophitic flows above and below the Scoville Point flow along the way. About 0.8 km (0.5 mi) from the Lodge lie ancient mine pits, attributed to Native Americans who occupied this area from about 5000 yrs BP during the period of the Nipissing stage. The mining was apparently informal and quite limited in any one place, but there are more than 1000 such pits all over Isle Royale according to Rakestraw (1965). 57 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 60: Oblique Google Earth Views of Stoll Trail to Scoville Point looking N and SW. Trail is a white line, and marked points are GPS marked locations. As we near Scoville Point, the Scoville Point flow (psp) dominates the shoreline and has steep smooth exposures. At the point itself, we will look at the excellent exposures of the Scoville Point flow, the ophitic flows below it, and the Edwards Island flow (pei), which underlies the companion point located just to the NW of Scoville Point. 58 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift There is a good exposure of cellular amygdaloid in one of the ophitic flows, and the Edwards Island Flow shows well developed entablature jointing (Figures 61, 62). Figure 61: Shoreline exposure of Edwards Island flow near Dashler Cabin, Scoville Point (entablature joints in pei). This view shows a vertical cross section of the joints on the antidip slope. Figure 62: Columnar joints in the Edwards Island Trap (pei) at the Dashler Cabin near Scoville Point. This view is perpendicular to the joints and shows their polygonal forms. The scale of the polygons is about 7-10 cm. While looking at the columnar joints in the Edwards Island flow (pei) at Scoville Point near the Dashler Cabin (Figs 61, 62), we should discuss whether this jointing pattern is indeed entablature jointing in the sense of Long & Wood (1986), and whether we should infer that the Edwards Island flow was indeed cooled in part by being flooded by surface water. (see also section on columnar jointing above) 59 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift We will return to the lodge via the Tobin Harbor Trail, which is easier to hike. It stays near the shore of Tobin Harbor, mostly atop the Edwards Island flow. Just NE of the Rock Harbor Lodge on the return trail is the site of the Smithwick Mine remains; this mine was discovered in 1843 and actually operated in 1847 and 1848. The work done here mostly consisted of exploratory shafts and excavations, and it is unclear whether much ore was found (Rakestraw, 1965). Lookout Louise and Monument Rock From Mirror Lake to Lookout Louise, we will hike about 1.6 km (1 mi) long and 85 m (280ft.) up (Figure 63). We will begin on the Tobin Harbor flow, but after passing the Lake we will walk Figure 63: Oblique Google Earth View, looking SW at Lookout Louise. Trails plotted in white. on the Greenstone Flow, following a dip slope up to Lookout Louise. At about the halfway point, the trail passes Monument Rock (Figure 64), an individual column from the colonnade of the upper ophite that is exposed as an erosional remnant. 60 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 64: Woodcut from Ackerman Lithographers, New York, showing Monument Rock in the 1840s. This old view is advantageous because the modern forest blocks an overall view like this one. Huber (1983, see especially pp 47-55) suggests that Monument Rock was formed by wave cut shoreline processes along a former "raised" shoreline, which he associates with glacial Lake Minong, about 10 Ka. From Lookout Louise we will look over the steep, anti-dip slope of the lower ophite and see Five Finger Bay, Duncan Narrows, and Amygdaloid Island. LIDAR topographic survey (Figure 65) came from Seth De Pasqual, at Isle Royale National Park. It reveals a striking topography which shows the dipping lava beds. Prominent large lava flows, like the Greenstone flow (pg) are obvious features in this image. Differential erosion of lava flows occurs when soft material, such as that found in the amygdaloidal flow tops and along faults, is preferentially removed and makes a topographic low, while the massive flow interiors resist erosion and become topographic highs. In this image we can also see the different layers of the Greenstone flow, including the Upper Ophite, the Pegmatite, and the Lower Ophite. 61 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift old shorelines lower ophite pegmatite upper ophite pg old shorelines Monument Rock upper ophite pth pth Figure 65: Shaded LIDAR topographic map of area near Lookout Louise, Isle Royale. This image shows what are thought to be ancient lake shorelines, which demonstrate that Monument Rock, far from the shore today, was once close to the lake shore and could have had a sea stack aspect. The image also shows layering textures in the Greenstone Flow which could reflect the Upper and Lower Ophite and the Pegmatite. Image from Seth De Pasqual, IRNP. Trails are shown in yellow. Post glacial shorelines can be seen in parts of this image also, and including in the vicinity of Monument Rock, itself far from the current shoreline. This arrangement suggests that the freestanding form of Monument Rock is consistent with its formation as a “sea stack”, and a remnant of the upper ophite of the Greenstone Flow, which is mostly eroded from this place. This interpretation was first suggested by N.K. Huber. Red Rock Point and Porter Island At Red Rock Point (Figure 66), we will pass excellent examples of entablature jointing of the upper part of the Greenstone flow. The basalt of the entablature is melaphyre (“trap”), very fine grained. The curvi-columnar nature of a few of the columns resembles some of the Columbia River basalt descriptions (Figure 25). Long and Wood (1986) suggest that entablature jointing 62 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift results when extensive floods that are created from disrupted drainages cause dramatic quenches of solidifying flood basalts. Figure 66: Oblique Google Earth image of eastern parts of Tobin Harbor, looking SW. Around the corner of Red Rock Point is a feature that Longo (1984) describes as follows: A large autointrusive dike was found intruding (N 20° W, 65° E) the columnar-jointed melaphyre at Red Rock Point. Despite an apparent lack of aplites, the dike is texturally similar to the stratiform pegmatoid. It is composed of randomly oriented, euhedral plagioclase laths with interstitial, subhedral augite and pigeonite (no poikilitic textures occur). The plagioclase laths are immense by comparison to the microlites of a typical ophitic unit. Three characteristic features of the dike are: (1) the abundant plagioclase phenocrysts (up to 1 cm (0.4 in)), (2) a blue-green hue from plagioclase altered to chlorite in the dike, and (3) alignment of plagioclase laths parallel to the dike contact, forming an igneous lamination. Amygdules are more abundant along the dike contact also. The process of autointrusion is similar to the mechanisms of pegmatoid formation, except that after the residual liquid is pressed out of the hosting crystal mesh, the differentiated magma is squeezed up into the vertical tensional fractures. --Longo 1984 63 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 67: Sag-Flowout structure, from McKee and Stradling (1970). Longo (1984) interprets the auto intrusion to be related to a sag flowout structure (Figure 67), described by McKee and Stradling (1970) as: a large structure that develops as the crust of a partly solidified flow founders and causes the upward escape of the flow's fluid interior (see figure, above). Below the water level at Red Rock Point is an occurrence of coarse grained granophyric rock, which can be found in beach cobbles and boulders and may occur within the Greenstone flow itself. The origin of granophyres in sills are not well understood, and Figure 68 from Marsh et al., 1991 shows some of his ideas, including existing silicic material which was carried in during intrusion. Figure 68: Some possible positions of granophyre within sheet-like intrusions. the left two panels show residual fluids forming lenses. The panel at right shows an accumulation of granophyre at the upper contact which may have existed upon emplacement (Marsh et al., 1991). We will also pass Porter's Island, which includes exposures of a fragmental rock that Huber (1973a) interprets as pyroclastic (pp). The same unit can be found on the Tobin Harbor shoreline opposite Newman Island. However, according to Longo (1984), these exposures may represent the fragmental top of the greenstone flow. The breccia unit, which is about 1-5 m (3.3-16.4 ft.) thick, contains rounded and semirounded fragments of the Greenstone flow set in a finer matrix that has amoeboid-shaped, agate amygdules. Longo did an extensive petrographic study but could not find any evidence of shards or pumice. He did, however, find bow-tie spherulitic plagioclases in the matrix, which suggests an undercooled texture for the basaltic material there. This unit occurs at the top of the Greenstone flow along about 15 km (9.3 mi) of strike length (approximately to Mt. Ojibway), according to Huber's map. Similar units are found at the top of the Greenstone flow on the Keweenaw Peninsula (Longo 1984). 64 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 69: Oblique Google Earth Image of Raspberry Island, looking N. Raspberry Island At Raspberry Island (Fig 69), about 0.5 km (0.3 mi) SE of Rock Harbor Lodge, we spend the day looking at a remarkable set of exposures nearby that provide an impression of some of the solidification features of an ophitic flow (approximately 20-30 m thick). At least since 2000, and in increasing amounts, low lake levels have made these exposures more numerous and accessible. One of many small islands along the S side of Rock Harbor, Raspberry Island is three ophitic flows of the undivided PLV (pu) dipping 15° SE. The uppermost of these flows is extensively exposed on a wave-washed dip slope. This shoreline receives strong storm waves and, fortunately has wave-washed exposures about 1 km (0.6 mi) long. They expose the flow interior, with the top of the flow eroded away and the base buried. A loop trail goes around the W half of the island, marked by informative signs about the unique ecosystem of this island, which features frequent fog and damp, moss-rich swamps. Among the unusual plants is the pitcher plant (Sarracenia puerperia), which is an insectivorous plant that flourishes in the swamp along the loop trail. 65 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift First, we will visit the W end of the island, where the regional attitude of the lava flows is seen in the view along strike toward Smithwick Island across the Smithwick Channel (Fig 70). Dip Slope Smithwick I 3 flows dip SE Anti-dip Slope Figure 70: Photo of Smithwick Island taken from Raspberry Island, showing a gently dipping sequence of three lava flows with obvious dip and anti-dip slopes. The dip of 20-25 degrees to the SE is typical of Isle Royale. The point on Raspberry Island facing the Channel is underlain by the oldest of the three flows on the island. We will walk on a dip slope that shows some of the jointing pattern we will also observe on the SE sides of Davidson and Smithwick Islands. Next, we will head to the SE corner of the island to observe some poorly-developed columns in the uppermost Raspberry Island flow, before looking at vesicle and segregation cylinders, and vesicle sheets or pegmatites. On the wave-washed SE shore are two zones of exposures of vesicle cylinders. Paces (1988) describes vesicle cylinders (Goff 1996) in the PLV: Vesicle pipes are elongated, tube-like structures, 10-30 cm (4-12 in) in diameter and 0.5-2 m (1.6-6.6 ft.) in length, containing somewhat coarser and more prismatic crystals compared to the adjacent groundmass. They are oriented vertically and occur predominantly in the bottom half of the flow. The origins and dynamic behavior of vesicle cylinders are poorly understood; however they appear to represent an accumulation of exsolved magmatic gas bubbles which migrate upwards through the magma during the period when the cooling magma behaves as a Bingham plastic (i.e., possesses a finite yield strength, Walker 1987). --Paces 1988 66 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 71: Segregation cylinders standing up as resistant to wave washing, forming small mounds separated by a few feet. The flow is tilted about 20 degrees to the left in this view. Figure 72: Vesicle cylinder or segregation cylinder from Raspberry Island, showing its cylindrical shape in 3 dimensions. Here at Raspberry Island, exposures of vesicle cylinders (Figures 71, 72) show a fairly regular spacing, 1-3 m (3-10 ft) apart, and a marked variety of textures; some were evidently preserved almost as voids, while others are filled with material that closely resembles vesicular pegmatoid. An interesting aspect of the exposures here is the relationship between the ophitic textures of the flow and the vesicle cylinders: The grain size of oikocrysts seems to be diminished by the proximity to the vesicle cylinder. Vesicle cylinders (Goff, 1996) are found mainly in only two areas along this shoreline. This may reflect their restricted occurrence in a thin part (less than a few meters thick) of this flow. Based on limited field examination, this thin part seems to be in the lower part of the flow. The comparisons between this occurrence and written descriptions, by Paces (1988) of the PLV on the Keweenaw, by Marsh et al. (1991) of solidification in sheet-like basaltic bodies, and those from Hon et al. (1994) and Self et al. (1998), are illuminating. Also featured conspicuously along the E shore of Raspberry Island are slickenside surfaces. A study of the fault slickenfibers allowed Witthuhn-Rolf (1997) to use geometrical and statistical methods to define the kinematics of the closing of the rift (Figure 73). In Witthuhn-Rolf's study, 67 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift a) Figure 73: Equal area rose diagrams of the trend of slickensides on reverse faults on island along Rock Harbor, Isle Royale National Park (Witthuhn-Rolf, 1997). Mon EAST AND WEST CARIBOU INNER lULL, OUTER HILL AND DAVIDSON Equal Area +:..... = RASPBERRY AND EDWARDS STOKLEY BAY, TOOKER, SHAW AND SMITHWICK ISLANDS Figure 32: Left: Equal area rose diagrams of the trend of slickensides on (a) normal and (b) reverse faults on Isle Royale (Witthuhn 1993). Notice the similar trends that define the resolved shear stress on the faults. Right: Rose diagrams ofthe trends of slickensides on reverse faults measured on islands along the SE shoreline of Isle Royale (Witthuhn 1993). Figure 74: Epidote-coated slickenside surfaces along faults exposed in Raspberry Island lava flows. Raspberry and Edwards Islands offered one of the largest populations of measurements. The measurements revealed two consistent stress fields, for each limb of the syncline, that would satisfy the conditions envisioned for the opening and closing of the Midcontinent Rift. Most of the faults on Isle Royale, including both normal and reverse faults, trend NE. This suggests that the reverse faults represent reactivated normal faults. The orientation of reverse faults at Isle Royale differs significantly from the predominately N-S trending structures measured in the PLV on the Keweenaw Peninsula. 68 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 75: 4 cm thick vesicle sheet or pegmatoid layer within horizontally fractured section of basaltic lava flow at Raspberry Island. About two-thirds of the way along the shore of Raspberry Island, the exposures that occur are stratigraphically higher in the flow. Here the flow has a laminar structure that consists of fractures that are parallel to the bedding and spaced about 0.5-3 cm (0.2-1.2 in) apart. Within this part of the flow, vesicle cylinders are not seen, but small pegmatoid lenses (vesicle sheets: Figure 75) occur. Paces (1988) describes them: Pegmatoid horizons are similar to vesicle cylinders in that they consist of gas-rich, coarsely crystalline, granophyric material. However, they occur as discontinuous lenses and layers, typically 10 cm (4 in) to several meters thick, and are usually located between the flow top and most massive portion of the flow interior. Pegmatoids are best developed in thicker flows that have cooled slowly enough to allow in situ differentiation (Cornwall 1951; Lindsley et al. 1971). This material represents the last remaining volatile-rich liquid, which is injected into fractures oriented sub-parallel to the upperflow surface. Both vesicle cylinders and pegmatoid layers contain significant void space in the form of vesicles and gas pockets and contribute to the permeability of the lava flows. --Paces 1988 The origin of the pegmatoids is likely related to the process by which the vesicle cylinders were formed. However, for the pegmatoid origin, the rise of material in channels is limited by the 69 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift thermal gradient and by the associated solidification that happens above the zone of pegmatoids, so the material is blocked and accumulates in lensoid layers (Figure 75). It is possible that Keweenawan flows preserve the inflated nature of ponded flood basalts well because runout of inflated flows such as can occur on sloping volcanoes is prevented by the riftfilling geometry. Tookers and Davidson Islands Figure 76: Oblique Google Earth Image of Tookers I looking N. One of many small islands strung out along the south side of Rock Harbor, Tookers (Fig 76) has some nice exposures of lava flow tops on its south side. Flow tops are amygdaloidal and less resistant to weathering. Flow interiors are massive and featureless, except they nearly always have at least poorly-developed columns. Figure 77: Exposure of contact between two lava flows, showing a black massive, relatively fresh upper flow, in contact with a reddish altered amygdaloidal flow top. Photo from 7 Mile Point, Keweenaw Peninsula. Flow Top 70 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Figure 78: Wave-washed lava surface on SE corner of Davidson Island showing polygonal jointing pattern with 2-4 m diameter polygons. Such patterns may be seen on many ophitic flows on Isle Royale. Figure 79: Oblique Google Earth Image of Davidson I looking W. On Davidson Island (Fig 79) is the Boreal Research Center, a residence for researchers at Isle Royale. We will walk around this small island, visiting another exposure of the epiclastic sedimentary rocks and an exposure of a columnar-jointed, ophitic flow on the SE corner of the island (Figure 78). The wave-washed shoreline has exposed a surface perpendicular to the columns, which are 2-3 m (6-10 ft.) across. Large columns seem to be a regular feature of ophitic flows at Isle Royale. Mott Island We will stop at Mott Island (Figures 79, 80) to visit one of the best exposures of sedimentary units within the PLV, found at the SW end of the island, facing East Caribou Island near the Park headquarters complex. There are seven such units mapped by Huber (1973) in the Chippewa 71 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Harbor area. Most of them are remarkably constant in thickness and lithology throughout their lateral extents, which are 65 km (40 mi) or more. Figure 80 : Detail of Isle Royale Geologic Map (Huber, 1973) which shows part of Eastern Isle Royale including Mott Island. The brown colored unit is the interflow sediment we will visit. Paces (1988) reports the following about interflow sediments in the PLV: Occasionally, lava flows are separated by intervening sheets and lenses of terrigenous clastic sediment. Twenty two major interflow sedimentary horizons occur scattered throughout the PLV section and are described by Butler and Burbank (1929), White (1952), and Merk and Jirsa (1982). Interflow sedimentary beds vary in thickness from less than 1 cm (0.4 in) thick fine-grained siltstones filling fractures between flow top fragments to coarse boulder conglomerates over 100 m (330 ft.) thick locally. Typically, interflow sediments are poorly sorted, lithologically immature conglomerates and sandstones derived from a nearby volcanic source of some relief and deposited in an alluvial fan-type environment (Merk and Jirsa 1982). 72 �rough and ly toward gh, finally ce to form usands of Lake VolKeweenaw represent is volcanic e sedimened with the Lake Vole Copper nd other above the ported by basin from gins. This of streams, at the lavas f the basin, mes, reverover large t of a basin filled (fig. flows were oward the as filling by nwarping. was intered downslopes that bris to be y, with the c activity, mitted the er Harbor nger Kem a thick e the vol- enawan or osed along www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift A. Lava erupts near the center'Of the basin and spreads laterally toward the margins to form a sequence of. lava flows. Y 8. The basin subsides, and during a lull in volcanic activity gravels are swept into the basin and'spread out over the uppermost lava flow. C. Volcanic activity resumes, and the cycle starts over again. FLOOD BASALTS AND SEDIMENTS Figure 81 : Cross section of rift valley showing theaccumulation process of interbedding. 43) showing of lava from (Fig. fissure vents in the Center of the rift, sometimes by this sandstone, together with similar alternating with infilling sediments from sandstone exposed in the southwestern outside the rift (Huber, 1973). part of the basin (fig. 39). The gross synclinal form of the Keweenawan basin resulted from subsidence coincident with filling of the basin rather than later folding by squeezing. However, Keweenawan strata near the margins of the basin, as73 on the Keweenaw Peninsula and Isle Royale, were subsequently steepened Transportation was generally from the SE to NW*, or from basin margins towards the center of the subsiding graben (White 1952). Although the interflow sediments are volumetrically insignificant within the PLV (3% of the total lithologic volume) (Merk and Jirsa 1982; White 1971), they form distinct and relatively continuous stratigraphic marker horizons within an otherwise monotonous volcanic pile. The occurrence of occasional interflow sediments implies that rates of lava flow extrusion, sedimentation, and/or tectonic subsidence were not constant during the formation of the PLV. White (1960) shows that a subsidence-depositional equilibrium was established so that both lava flows and sediments were deposited on near-horizontal surfaces. Most lava flows were deposited directly on top of the underlying lava flow top indicating a more-or-less constant and relatively short repose period between eruptions. The infrequent presence of sedimentary beds between lava flows may indicate occasional hiatuses in magma extrusion, which allowed or alluvial fans to transgress out towards the center of the basin. Conversely, interflow sedimentary horizons may mark brief periods of increased depositional rates possibly related to episodic normal faulting and basin subsidence. --Paces 1988 *this quote refers to the Keweenaw, where Paces worked--on Isle Royale directions are reversed. �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Edison Fishery and the Lighthouse The Fishery (Figure 82) itself is a restored camp that is occupied each summer by a retired Lake Superior fisherman and his family; this man is employed by the Park to interpret what life was like here during the heyday of Isle Royale fishing camps, from before the establishment of the Park in 1936 until the sea lamprey invasion of the 1950s. Figure 82 Oblique Google Earth View of Edison Fishery and the Lighthouse looking SW. The lavas that underlie the site of the fishery and the lighthouse are a sequence of 45-50 ophitic flows, which occur between the Scoville Point flow (psp) and the overlying CHC. As we walk around the point we will see several flow tops exposed, good examples of cellular amygdaloids. This is an excellent place to find (but not to collect!) Isle Royale greenstone, a nodular, compact form of pumpellyite that is prized as a semi-precious gemstone (Huber 1983, see pp. 58-9). The geological purpose of stopping here is to look at the flow sections along the wave-washed shoreline, following it from this point to Tonkin Bay. We can also look at the amygdule mineral suite, which can be found on the pebble beaches. The amygdules of Isle Royale's flows contain a variety of secondary minerals, listed alphabetically (by Huber) as barite, calcite, chlorite, copper, datolite, epidote, laumontite, natrolite, prehnite, pumpellyite (chlorastrolite or “greenstone”), 74 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift quartz (agate), and thomsonite (see section on Amygdaloid). The prehnite is unusual in that it contains disseminated native copper inclusions and has a pink color, which has caused some to confuse it with thomsonite (Huber 1969). Overall, the assemblage is zeolite facies and prehnitepumpellyite facies, representing a slightly lower grade than much of the Keweenaw Peninsula area. This lower mineralization temperature may partially explain the lower abundances of native copper on Isle Royale than those found on the Keweenaw Peninsula. This metamorphic event reflects a large hydrothermal (hot, geothermal brine which was pumped through the porous flow tops and conglomerates of the Portage Lake Volcanics for years after the volcanism ended (Jolly, 1972). Figure 83: Oblique Google Earth View of Mt Franklin and Ojibway tower, looking SW. The view looks directly along the strike of the lava flows, which are dipping gently to the east. 75 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Franklin and Ojibway The Mount Franklin Trail begins 0.3 km (0.2 mi) W of Three Mile Campground (Figure 83). The trail immediately climbs a ridge supported by the Scoville Point Flow (psp), then levels off and descends. We will cross a boardwalk over a swamp and arrive at a valley where there is a junction with the Tobin Harbor Trail, 0.8 km (0.5 mi) from Three Mile Campground. We will continue on the Mount Franklin Trail, straight ahead, crossing the Tobin Creek swamp and then climbing a ridge underlain by the Tobin Harbor flow (pth). From here we will descend to cross another swamp and then begin the 300 ft. ascent of the Greenstone ridge. The entire swamp and ascent is underlain by the great Greenstone Flow (pg). At the top of the ridge there is a junction with the Greenstone Ridge Trail, which we will take left to go about 0.5 km (0.3 mi) to Mount Franklin, elevation 330 m (l080 ft.). Here there is a good view of the N side of the island, including Five Finger Bay, Lane Cove, and Amygdaloid Island, as well as of the Canadian Shoreline, including the Logan Sills and the Sleeping Giant. The Greenstone Flow is indeed the backbone of the island, forming the most prominent ridge all along; only at Blake Point, however, is a reasonably complete section through the flow exposed. The contact between the pegmatoid and the lower ophite units of the Greenstone is mainly located near the crest of the Greenstone ridge. The lower ophite underlies the N slope, which is a steep, anti-dip slope, and the pegmatoid armors the gentler dip slope to the S. Figure 84: Oblique Google Earth view of Ojibway Tower and Daisy Farm, looking E. 76 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Following the same trail, we will descend sharply to a wooded area and level off for about 0.4 km (0.25 mi) before climbing again. We will then reach the ridge crest and follow it for another 3.2 km (2 mi), with occasional outstanding views, to the Mount Ojibway tower. This structure was built in 1962 and was used initially as a fire tower. Now it is used for monitoring acid rain, along with other environmental monitoring. We can climb the tower stairs for full views of the surroundings, both to the N and S. From the tower we will descend to the Daisy Farm Campground via the Mount Ojibway Trail. (Figure 84). We will go down from the ridge to the first level spot and then begin to rise over a smaller ridge. The beginning of this small ridge is the approximate location of the top of the Greenstone flow; the ridge top and the dip slope to the S is underlain by the Tobin Harbor flow (pth). At the base of this ridge we will cross a swamp fed by Tobin Creek. Then we will ascend Ransom Hill, which has the Long Island Flow (pli) on its anti-dip (N) slope and the Edwards Island Flow (pei) on its dip slope (S) side, where there is some entablature jointing. From Ransom Hill, the trail descends to Daisy Farm Campground. Daisy Farm is located on the site of an old mining community, called Ransom, which was founded in 1847 with the clearing of land and the construction of a smelter. The mining prospects dimmed quickly, however, and the mining activity ended only two years later in 1849. Then, in 1866, all the buildings burned down. In later years, the place was the site of a sawmill, a garden that supplied vegetables to Rock Harbor Lodge, and a Civilian Conservation Corps (CCC) camp, which was a foundation for youth employment, developed by Roosevelt during the depression (Rakestraw 1965). What to take home After a several day journey, what are the earth sciences messages that stick with you? What are the globally significant issues that stand out? What is uniquely interesting about the place and time that is recorded in rocks here? What big ideas emerge from this geology? 1. Rodinia, a Proterozoic supercontinent, blanketed Earth’s mantle, and the higher heat flow of 1.1 billion years ago triggered huge volumes of hot magmatism from the mantle, first giving rise to ultramafic dike swarms, then basalts in huge quantities.  These dikes split the great supercontinent, but a nearby continental collision (Grenville) was apparently what prevented the formation of an ocean basin. 2. Large Scale Flood basalts occurred for a brief period, lasting only a few million years in the Keweenaw and Isle Royale.  These eruption rates, much higher than average, apparently were driven by a mantle plume. They are similar to other continental flood basalts and mafic large igneous provinces (LIPs) in these respects. There are volcanic, plutonic, and sedimentary elements to the mantle plume and rifting (see map, below). 3. Ponding of magma happened in a great crack—the midcontinent rift basin, locally called the Keweenaw Rift. Because lava solidifies by heat loss from the lower surface where it is contact with the cold ground and the upper surface where it is in contact with the air, thick lava flows cool much more slowly than thin ones, because the massive flow interiors, far 77 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift from the top and bottom of the flow, are shielded from heat loss. The Keweenaw Rift has flows as thick as 1200 ft, thicker than those found in other mafic LIPs. 4. The in situ differentiation within the largest lava flows may have occurred because of the existence of a large ponded magma body (not unlike a magma ocean) within the rift valley for perhaps up to a millennium. This results in pegmatite or dolerite horizons within the large flows, features that are not common in younger flood basalts. Vesicle cylinders and segregation cylinders are also conspicuous features of these ponded flows which occur in the lower parts of the flows, reflecting compaction of a dendritic mush consisting of ophitic crystals. 5. The hotspot (mantle plume head), along with the rifting it caused, created a big, elongate hole in the continent, that was partially filled with basalt and redbed sediments. This hole has persisted until now and it is this hole that coincides very closely to the position of Lake Superior. 6. An unexplained unique aspect of this rift situation is native copper mineralization. Though other rifts have all of the other mineral deposit types of the 1.1 Ga Lake Superior area, none has native copper. We are puzzled by this cosmic geochemical oddity. What happened to the sulfur usually found with chalcophile elements? 7. Fossils are difficult to find in Keweenawan rocks, generally, but cyanobacteria are conspicuous. Stromatolites within the rift basin here are associated with an oxidized ocean and an atmosphere that was holding at least some free oxygen. Following the redbeds of the rift were the multiple Snowball Earth events. Acknowledgements The opportunity to write a detailed guide to Isle Royale and to lead a field trip comes from the cooperation of many people. Lori Witting did the planning and financing issues for the trip. Mark Klawiter planned the food and field logistics. Bob Barron helped with numerous details. I would like to thank Liz Valencia and Greg Bickings of Isle Royale National Park for permitting and helping plan this field trip. Steve Roblee was an eager boat pilot. King Huber provided us with a complete set of his many publications about Isle Royale and also with lots of cheerful encouragement. Jim Paces, Tony Longo, and Rick Wunderman provided me with a lot of insight on the volcanic geology of Isle Royale. Kate Witthuhn-Rolf supplied some unpublished data. Discussions with Bruce Marsh and Angus Hellawell about solidification helped me to understand a little better what may have been going on inside Isle Royale's lava flows. Seth De Pasqual at IRNP provided LIDAR maps for the guide and explanations of them. Evgeniy Kulakov worked on the paleomagnetic information for us. George Robinson helped find some great mineral specimens to illustrate the zeolite facies 78 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift amygdaloids, and John Jaszsak helped with photographing them. Many researchers provided material for me to learn about and communicate about Isle Royale. There are so many crucial words and illustrations that are needed, and I tried to use as many as I could. Here are some of the names: Ted Bornhorst, Bill Cannon, Henry Cornwall, Jim DeGraff, Doug Elmore, Fraser Goff, John Green, Ken Hon, Wayne Jolly, Susanne Nicholson, Dick Ojakangas, Lauri Pesonen, Anthony Philpotts, Suzanne Schmidt, Steve Self, Dick Stoiber, George Walker, Walter White. Others are in the Bibliography. Ken VanDellen helped with editing the text and clarifying the English. References Cited • Basalt Volcanism Study Project, 1981, Basaltic volcanism on the terrestrial planets, Pergamon Press, Inc., New York, 1286 pp. • Bondre, N.R., R.A. Duraiswami, G. Dole (2004) Morphology and emplacement of flows from the Deccan Volcanic Province Bulletin of Volcanology, 66 , pp. 29–45 • Behrendt, J.C., A.G. Green, W.F. Cannon, D.R. Hutchinson, M. Lee, B. Milkereit, W.F. Agena, C. Spencer (1988) Crustal structure of the Midcontinent Rift System: results from the GLIMPCE deep seismic reflection profiles Geology, 16 , pp. 81–85. • Bornhorst, T.J., 1997, Tectonic context of native copper deposits of the North American Midcontinent Rift System, in Ojakangas, R.W., Dickas, A.B., and Green, J.C., eds., Middle Proterozoic to Cambrian rifting, central North America: Geological Society of America Special Paper 312, p. 127–136. • Bornhorst, T.J., and Brandt, D., 2009, Michigan’s earliest geology: The Precambrian, in Schaetzl, R., Darden, J., and Brandt, D., eds., Michigan Geography and Geology: New York, Pearson Custom Publishing, p. 24–39. • Bornhorst, T.J., and Lankton, L.D., 2009, Copper mining: A billion years of geologic and human history, in Schaetzl, R., Darden, J., and Brandt, D., eds., Michigan Geography and Geology: New York, Pearson Custom Publishing, p. 150–173. • Bornhorst, T.J., Paces, J.B., Grant, N.K., Obradovich, J.D., and Huber, N.K., 1988, Age of native copper mineralization, Keweenaw Peninsula, Michigan: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 83, p. 619–625. • Bornhorst, TJ and R Barron, 2011, Copper deposits of the western Upper Peninsula of Michigan, Geol Soc Amer Field Guide 24: 83-99. • Brannon, J.C. 1984, Geochemistry of successive lava flows of the Keweenawan North Shore Volcanic Group, Ph.D. dissertation, Washington University, St. Louis, MO, 312 pp. • Bresson, David, 2011 Large Igneous Provinces and mass extinctions. Scientific American, September 16, 2011 (http://blogs.scientificamerican.com/history-of-geology/2011/09/16/largeigneous-provinces-and-mass-extinctions/) 79 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift • Broderick, T.M., 1935, Differentiation in lavas of the Michigan Keweenawan, Geol. Soc. Am. Bull., v. 46, pp. 503-58. • Broderick, T.M., Hohl, C.D., and Eidemiller, H.N., 1946, Recent contributions to the geology of the Michigan copper district, Econ. Geol., v. 41, pp. 675-725. • Brown, A.C., 2006, Genesis of native copper lodes in the Keweenaw district, northern Michigan: A hybrid evolved meteoric and metamorphogenicmodel: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 101, p. 1437–1444. • Butler, B.S. and Burbank, W.S., 1929, The copper deposits of Michigan, USGS. Prof Pap., No. 144, 238 pp. • Cannon, W.F., and Phillips, B.A.M., 2007, Geologic and cultural history of the Grand Portage National Monument [field trip 2]: Institute on Lake Superior Geology, Annual Meeting, Lutsen, MN, Part 2 – Field Trip Guidebook, v. 53, pages 24-52. • Cannon, W.F., 1994, Closing of the Midcontinent rift: a far-field effect of Grenvillian compression, Geology, v. 22, pp. 155-8. • Cannon, W.E et al., 1989, The North American Midcontinent rift beneath Lake Superior from GLIMPCE seismic reflection profiling, Tectonics, v. 8, pp. 30532. • Cannon, W.F., Peterman, Z.E., and Sims, P.K., 1993, Crustal-scale thrusting and origin of the Montreal River monocline—A 35-km-thick cross section of the Midcontinent Rift in northern Michigan and Wisconsin: Tectonics, 12, p. 728–744, doi:10.1029/93TC00204. • Carmichael, I.S.E., Turner, EJ., and Verhoogen, J., 1974, Igneous Petrology, McGraw-Hill, New York. • Clark, J.A, Hendriks, M., Timmermans, T.J., Struck, C., and Hilverda, K.J., 1994, Glacial isostatic deformation of the Great Lakes region, Geol. Soc. Am. Bull., v. 106, pp. 19-31. • Cornwall, H.R., 1951, Differentiation in lavas of the Keweenawan series and the origin of the copper deposits of Michigan, Geol. Soc. Am. Bull., v. 62, pp. 159-202. • Crisp, J., and S. Baloga, 1994, The influence of crystallization and entrainment of cooler material on the emplacement of basaltic aa lava flows, J. Geophys. Res., 99: 11819-11831. • Davis, D.W. and Green, J.C., 1997. Geochronology of the North American Midcontinent Rift in western Lake Superior and implication for its geodynamic evolution. Canadian Journal of Earth Sciences, 35, 476-488. • Davis, D.W., and Paces, J.B., 1990, Time resolution of geologic events on the Keweenaw Peninsula and implications for development of the Midcontinent Rift system: Earth and Planetary Science Letters, v. 97, p. 54–64, doi:10.1016/0012-821X(90)90098-I. • DeGraff, J.M. and Aydin, A, 1993, Effect of thermal regime on growth increment and spacing of contraction joints in basaltic lava, J. Geoph. Res., v. 98, pp. 6411-30. • DeGraff, J.M., Long, P.E., and Aydin, A, 1989, Use of joint-growth directions and rock textures to infer thermal regimes during solidification of basaltic lava flows, J. Volc. and Geotherm. Res., v. 38, pp. 309-24. • Ding, X., Ripley, E.M., Shirey, S.B., and Li, C., 2012, Os, Nd, O, and S isotopic constraints on country rock contamination in the conduit-related Eagle Cu-Ni-(PGE) deposit, 80 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift • • • • • • • • • • • Midcontinent Rift System, Upper Michigan: Geochimica et Cosmochimica Acta, v. 89, p. 10-30. Elmore, R.D., 1984, The Copper Harbor Conglomerate: A late Precambrian fining-upward alluvial fan sequence in northern Michigan: Geological Society of America Bulletin, v. 95, p. 610–617, doi:10.1130/0016 -7606(1984)95<610:TCHCAL>2.0.CO;2. Elmore, R.D., Milavec, G.J., Imbus, S.W., and Engel, M.H., 1989, The Precambrian Nonesuch Formation of the North American Mid-Continent Rift, sedimentology and organic geochemical aspects of lacustrine deposition: Precambrian Research, v. 43, p. 191–213, doi: 10.1016/0301-9268(89)90056-9. Foster, J.W. and Whitney, J.D., 1851, Report on the geology of the Lake Superior land district, Washington, DC, AB. Hamilton, 400 pp. Goff, F.E., 1977, Vesicle cylinders in vapor-differentiated basalt flows, Ph.D. dissertation, University of California, Santa Cruz, CA, 83 pp. Goff, F. E.. 1996, Vesicle cylinders in vapor-differentiated basalt flows, J Volcanol Geoth Res 71: 167-185. Goodge, J. W., Vervoort, J. D., Fanning, C. M., Brecke, D. M. Farmer, G. L., Williams, I. S., Myrow, P. M., DePaolo, D. J., 2008, A Positive Test of East Antarctica-Laurentia Juxtaposition Within the Rodinia Supercontinent: Science, v. 321, p. 235-240. [PDF] Grand Portage National Monument, 1986, Grand Portage, National Park Service, U.S. Department of the Interior, GPO: 1986--491-414/20056. Green, J.C., 1982, Geology of Keweenawan extrusive rocks, Geo. Soc. Am. Mem., v. 156: 47-55. Green, J.C., 1989, Physical volcanology of mid-proterozoic plateau lavas: the Keweenawan North Shore Volcanic Group, Minnesota, Geol. Soc. Am. Bull., v. 101, pp. 486-500. Halls, H.C. and Pesonen, L.J., 1982. Paleomagnetism of Keweenawan rocks, Geological Society of America, Memoir, 156, 173-201. Hellawell, A, Sarazin, J.R., and Steube, R.S., 1993, Channel convection in partly solidified systems, Phil. Trans. R. Soc. Land., v. 345, pp. 507-44. • Holm, D.K., D.A. Schneider, S. Rose, C. Mancuso, M. McKenzie, K.A. Foland, K.V. Hodges 2007 Proterozoic metamorphism and cooling in the southern Lake Superior region, North America and its bearing on crustal evolution, Precambrian Research 157:106–126. • Hon, K. , J. Kauahikaua, R. Denlinger, K. Mackay Emplacement and inflation of pahoehoe sheet flows; observations and measurements of active lava flows on Kilauea Volcano, Hawaii Geological Society of America Bulletin, 106 (3) (1994), pp. 351–370 • Huber, N.K., 1969, Pink copper-bearing prehnite from Isle Royale National Park, Michigan, USGS. Prof Pap., No. 650-D, pp. D63-8. • Huber, N.K., 1973, Geologic map of Isle Royale National Park, Keweenaw County, Michigan, USGS. Map, No. 1-796. • Huber, N.K., 1973a, The Portage Lake Volcanics (middle Keweenawan) on Isle Royale, Michigan, U.S.G.S. Prof Pap., No. 754C, 32 pp. 81 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift • Huber, N.K, 1973b, Glacial and postglacial geologic history of Isle Royale National Park, Michigan, U.S.G.S. Prof Pap., No. 754-A, 15 pp. • Huber, N.K., 1983, The Geologic Story of Isle Royale National Park, U.S.G.S. Bull., No. 1309, 66 pp. • Jerram, D.A., M. Widdowson, 2005 The anatomy of continental flood basalt provinces: geological constraints on the processes and products of flood basalt volcanismLithos, 79 (3– 4) , pp. 385–405 • Hjartarson A. 1988. The great Thjorsa lava: Earth's largest Holocene lava flow. Natturufraedingurinn 58:1–16 • Jolly, W.T., 1974, Behavior of Cu, Zn, and Ni during prehnite-pumpellyite rank metamorphism of the Keweenawan basalts, northern Michigan: Economic Geology and the Bulletin of the Society of Economic Geologists,69, p. 1118–1125. • Karlstrom, K.E., Harlan, S.S., Williams, M.L., McLelland, J., Geissman, J.W., and Ahaell, K., 1999. Refining Rodinia: Geologic Evidence for the Australia Western U.S. (AUSWUS) connection for Proterozoic Supercontinent Reconstructions. GSA Today, v.9, n. 10, October, 1999 • Katterhorn, SA & CJ Schaefer, 2008, Thermal–mechanical modeling of cooling history and fracture development in inflationary basalt lava flows, Journal of Volcanology and Geothermal Research 170 (2008) 181–197 • Kern, A.N., Kulakov, E.V., Smirnov, A.V., Diehl, J.F., and K. Chamberlain, 2012. Paleomagnetism of the Coldwell Complex (Ontario, Canada): New Data and New Insights. American Geophysical Union Fall meeting, abstract GP21A-1131. • Keszthelyi, L., S. Self Some physical requirements for the emplacement of long basaltic lava flows Journal of Geophysical Research, 103 (B11) (1998), pp. 27,447–27,464 • Kilburn C. R. J. and G. Luongo 1993. Active Lavas - Monitoring and Modelling . UCL Press, London, 384pp. • Klewin, K.W., and Shirey, S.B., 1992, The igneous petrology and magmatic evolution of the Midcontinent Rift system: Tectonophysics, v. 213, p. 33-40. • Lacroix, A, 1928, Les Pegmatitoides des Roches Volcaniques a Facies Basaltiques, Ac. Sci. Paris Comptes Rendus, v. 187, pp. 321-6. • Lacroix, A, 1929, Les Pegmatitoides des Roches Volcaniques a Facies Basaltiques: A Propos de Celles du Wei-Tchang, Bull. Geol. Soc. China, v. 8, pp. 45-9. • Lane, AC., 1893, Geological report on Isle Royale, Michigan, Geol. Surv. ofMichigan, v. 6., pp. 1-265. • Lane, A.C., 1911, The Keweenaw series of Michigan, Michigan Geol. and Biol. Surv. Pub. 6, Geol. Ser. 4, v. 2, 983 pp. • Lane, A. C., and Seaman, A. E., 1907, Notes on the geological section of Michigan, Part 1. The pre-Ordovician: Journal of Geology, v. 15, p. 680-695. • Lindsley, D.H., Smith, D., and Haggerty, S.E., 1971, Petrography and mineral chemistry of a differentiated flow of Picture Gorge Basalt near Spray, Oregon, Carnegie Inst. of Washington, Yearbook 69, pp. 264-85 82 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift • Lofgren. G.E., 1980, Experimental studies on the dynamic crystallization of silicate melts, Physics of Magmatic Processes (ed. RB. Hargraves), Princeton Univ. Press, Princeton, NJ, pp. 487-551. • Long, P.E. and Wood, BJ., 1986, Structures, textures and cooling histories of Columbia River basalt flows, Geol. Soc. Am. Bull., v. 97, pp. 1144-55. • Longo, A.A., 1984, A correlation for a middle Keweenawan flood basalt: the Greenstone flow, Isle Royale and Keweenaw Peninsula, Michigan, M.S. thesis, Michigan Technological University, Houghton, MI, 198 pp. • Mangan, M. and B. D. Marsh. 1992. Solidification front fractionation in phenocryst -free sheet-like magma bodies. J. Geology, v. 100, p. 605-620. • Marsh, B.D., Gunnarsson, B., Congdon, R., and Carmody, R., 1991, Hawaiian basalt and Icelandic rhyolite: indicators of differentiation and partial melting, Geologische Rundschau, 80/2, pp. 481510. • McKee, B. and Stradling, D., 1970, The sag flowout: a newly described volcanic structure, Geol. Soc. Am. Bull., v. 81, pp. 2035-44. • Merk, G.P. and Jirsa, M.A., 1982, Provenance and tectonic significance of the Keweenawan interflow sedimentary rocks, Geol. Soc. Am. Mem., v. 156, pp. 97-105. • Nevanlinna, H., and Pesonen L.J., 1983. Late Precambrian Keweenawan asymmetric polarities as analyzed by axial offset dipole geomagnetic models, Journal of Geophysical Research, 88, 645–658. • Nicholson, S.W., Shirey, S.B., Schulz, K.J., and Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent Rift System basalts; implications for multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504-520. • Nicholson, S.W., and Shirey, S.B., 1990, Midcontinent Rift volcanism in the Lake Superior region; Sr, Nd, and Pb isotopic evidence for a mantle plume origin: Journal of Geophysical Research, v. 95, p. 10,851-10,868. • Ojakangas, R.W., Morey, G.B. and Green J.C., 2001. The Mesoproterozoic Midcontinent Rift System, Lake Superior region, USA, Sedimentary Geology, 141-142, 421-442. • Ojakangas, R.W., Morey, G.B., and Southwick, D.L., 2001, Paleoproterozoic basin development and sedimentation in the Lake Superior region, North America: Sedimentary Geology, v. 141–142, p. 319–341, doi:10.1016/S0037-0738(01)00081-1. • Paces, J.B., 1988, Magmatic processes, evolution and mantle source characteristics contributing to the petrogenesis ofMidcontinent rift basalts: Portage Lake Volcanics, Keweenaw Peninsula, Michigan, Ph.D. Dissertation, Michigan Technological University, Houghton, MI, 413 pp. • Paces, J.B. and Miller, J. D., Jr., 1993, Precise U-Pb ages of Duluth complex and related mafic intrusions, northeastern Minnesota: geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagmatic processes associated with the 1.1 Ga • Midcontinent rift system, J. Geoph. Res., v. 98, pp.13,997-14,013. 83 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift • Palmer, HC, 1970, Paleomagnetism and correlation of some Middle Keweenawan rocks, Lake Superior: Canadian Journal of Earth Sciences, v. 7, p. 1410-1436. • Pesonen L.J., and Nevanlinna, H, 1981. Late Precambrian Keweenawan asymmetric reversals, Nature. 294, 436-439. • Pesonen, L. J. and Halls, H.C., 1983. Geomagnetic field intensity and reversal asymmetry in late Precambrian Keweenawan rocks. Geophysical Journal. Royal Astronomical Society, 73, 241-270. • Pinkerton, H. & Wilson, L. (1994) Factors controlling the length of channel-fed flows, Bull. Volc., 56, 108. • Rakestraw, L., 1965, Historic mining on Isle Royale, reprinted in Borealis Isle Royale, Natural History Association, Houghton, MI. • Robertson, J.M., 1975, Geology and mineralogy of some copper sulfide deposits near Mount Bohemia, Keweenaw County, Michigan: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 70,1202–1224. • Robertson, W.A. and Fahrig, W.F., 1971. The great Logan paleomagnetic loop-the polar wandering path from the Canadian Shield rocks during the Neohelikian Era. Canadian Journal of Earth Sciences, 8, 1355-1372. • Rogan W, S Blake and I Smith 1996 In situ chemical fractionation in thin basaltic lava flows: examples from the Auckland volcanic field, New Zealand, and a general physical model JVGR 74: 89-99 • Ross PS, L Ukstins Peate, MK McClintock, YG Yu, IP Skilling, JDL White and BF Houghton, 2005, Mafic volcaniclastic deposits in flood basalt provinces: A review, Journal of Volcanology and Geothermal Research 145 (2005) 281–314 • Santin, S.F., 1969, Pegmatitoides in the horizontal basalts of the Lanzarote and Fuerteventura Islands, Series I, Bull. Volc., v. 33, pp. 989-1007. • Schneider, D., Holm, D. K., Boyle, C. O., Hamilton, M., and Jercinovic, M., 2004, Paleoproterozoic development of a gneiss dome corridor in the southern Lake Superior region, U.S.A.: In Whitney, Tessyier, and Siddoway (eds) Gneiss domes in orogeny: Geological Society of America Special Paper 380, p.339-357. • Schmidt, S.T. and Robinson, D. 1997. Metamorphic grade and porosity and permeability controls on mafic phyllosilicate distributions in a regional zeolite to greenschist facies transition of the North Shore Volcanic Group, Minnesota. Geol. Soc. Am. Bull. 109, 683-697. • Schmidt, P. W., and Williams G.E., 2003. Reversal asymmetry in Mesoproterozoic overprinting of the 1.88-Ga Gunflint Formation, Ontario, Canada: non-dipole effects or apparent polar wander?, Tectonophysics, 377,7–32, 2003. • Seaman, A.E. and Seaman, W.A., 1944, Geological column Lake Superior Region in general: Michigan Geol. Surv. Div., Progress Report No. 10. • Self, S., L. Keszthelyi, T. Thordarson The importance of pahoehoe Annual Review of Earth and Planetary Sciences, 26 (1998), pp. 81–110 • Stoiber, R.E., and Davidson, E.S., 1959, Amygdule mineral zoning in the Portage Lake Lava Series, Michigan copper district: Economic Geology, 54, p. 1250–1277, p. 1444–1460. 84 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift • Swanson, D.A., Wright, T.L., and Helz, RT., 1975, Linear vent systems (and estimated rates of magma production and eruption) for the Yakima basalt of the Columbia plateau, Am. J. Sci., v. 275, pp. 877-905. • Thordarson, T., and S. Self (1998), The Roza Member, Columbia River Basalt Group: A gigantic pahoehoe lava flow field formed by endogenous processes?, J. Geophys. Res., 103(B11), 27411–27445, doi:10.1029/98JB01355. • Tomkeieff, S.I., 1940, The basalt lavas of Giants Causeway district of Northem Ireland: Bull. Volcanol., v. 6, pp. 90-143. • Van Hise, C.R. and Leith, C.K., 1911, The geology of the Lake Superior region, U.S.G.S. Mon., No. 52, 381 pp. • Trubitsin, M. Kaban and M. Rothacher: "Mechanical and thermal effects of floating continents on the global mantle convection", PHYSICS OF THE EARTH AND PLANETARY INTERIORS (Vol. 171, S. 313-322). • Walker, G.P.L., 1987, Pipe vesicles in Hawaiian basaltic lavas: their origin and potential paleoslope indicators, Geol., v. 14, pp. 84-7. • White, W.S., 1952, Imbrication and initial dip in a Keweenawan conglomerate bed, J. Sed. Pet., v. 22, pp. 189-99. • White, W.S., 1960, The Keweenawan lavas of Lake Superior: an example of flood basalts, Am. J. Sci., v. 258-A, pp. 367-74. • White, W.S., 1971, Geologic setting of the Michigan copper district, In Guidebook for Field Conference. Michigan Copper District, Sept. 30 - Oct. 2, (ed. W.S. White), Soc. Econ. Geol., Michigan Technological University, Houghton, MI, pp. 3-17. • Whitmeyer S. J. and K E. Karlstrom 2007 Tectonic model for the Proterozoic growth of North America Geosphere 2007;3;220-259 • Witthuhn-Rolf, K.M., 1997, A structural analysis of the Midcontinent rift in Michigan and Minnesota, Geol Soc Amer Special Paper 312: 97-113. • Wolff, RG. and Huber, N.K., 1973, The Copper Harbor Conglomerate (middle Keweenawan) on Isle Royale, Michigan, and its regional implications, U.S.G.S. Prof Pap., No. 754-B, 15 pp. • Worster, M.G. and Huppert, H.E., 1993, The crystallization of lava lakes, J. Geoph. Res., v. 98, pp. 15,891901. Lat-Long Locations of this field trip You have noticed there is no road log for this trip. This is because there are no roads! Download all the locations from the web here: www.geo.mtu.edu/~raman/IsleRoyalekmz.zip These files will be readily ingested by Google Earth software or GPS software and provide precise locations for all the sites described here. A table of the Latitude and Longitude of all these sites is listed here so it can be used to manually transfer this information if needed. These values may be entered manually into GPS or Google Earth. 85 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Name Longitude Latitude 3 Mile CG -88.52960447 48.12410194 Amygdaloid Island Ranger St -88.65598008 48.13570657 Arch Amgd I -88.62558448 48.14889669 Belle Isle CG -88.58562501 48.15234621 Big Cols Davidson -88.51062061 48.1249503 Big Cols Rasp I -88.47841662 48.14023711 Big Cols Rasp -88.48350224 48.13787684 Blake Pt -88.42229232 48.19082848 Caribou Arch -88.56108894 48.09705948 Caribou CG -88.57214509 48.09498673 Cop Harb Cong RC -89.23205406 47.85168084 Crystal Cove -88.58980015 48.15869417 Daisy Farm CG -88.59552193 48.09214022 Davidson I -88.51535972 48.12257809 Duncan Bay CG -88.52185527 48.150598 Edison Fishery -88.58317221 48.08946992 Edwards Is -88.43527441 48.17172245 Gull Rocks East -88.26162826 48.26236504 Hill Pt -88.52528802 48.1655558 Johnson Is -88.58571927 48.14731944 Keyhole -88.61806043 48.14501207 L Louise -88.47250078 48.16924628 Lane Cove CG -88.5570814 48.14486573 Lighthouse -88.57937109 48.08979679 86 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Name Longitude Latitude Little Todd CG -88.92697185 48.02005966 Locke Pt -88.45901399 48.18450616 McCargoe CG -88.7082605 48.08740121 Merrit Lane CG -88.42972709 48.18442853 Minong Mine -88.72005096 48.08347491 Moose Skulls -88.59063351 48.08709128 Mott I Dock -88.54739095 48.10720599 Mott Sediment -88.55002491 48.10429157 Ollies Rocks -88.71228558 48.11692273 Ophite php Wash Hbr -89.17944981 47.93491098 Ophite pwi Wash Hbr -89.23071872 47.87582894 Passage Island Dock -88.35571791 48.23122681 Passage Light -88.36567255 48.22354584 Pickerel Cove -88.65241933 48.12402173 Pine Mountain -88.72816055 48.08439633 Porphyrite pgi Wash Hbr -89.21618244 47.88214454 Porphyrite pmp Wash Hbr -89.21962318 47.8702359 Porphyrite ph Wash Hbr -89.18438864 47.93089216 Porter I -88.44598813 48.17423934 Rasp I Dock -88.47534021 48.14220455 Rasp Seg Cyls -88.47477863 48.1405457 Raspberry Pegs -88.46879695 48.14351368 Red Rock Pt -88.45413695 48.17139189 -89.313 47.867 Rock of Ages Light 87 �www.geo.mtu.edu/~raman/SilverI/IRKeweenawRift Name Longitude Latitude Scoville Pt -88.44940521 48.16322165 Snug Harbor -88.4852324 48.14576228 South Rock -89.27218772 47.86125303 Susie Islands -89.5736758 47.96604403 Suzy's Cave -88.51477842 48.13207674 Todd Harbor CG -88.8219923 48.05083223 Tookers I -88.50329307 48.12941722 Trap pm Wash Hbr -89.2212717 47.90837977 Trap2 pm Wash Hbr -89.14971047 47.93373916 Voyaguer II Dock -89.65254479 47.96263767 Wendigo Mine -89.15127391 47.93227937 Wilson I -88.83672187 48.05654853 Windigo Dock -89.15820212 47.91194955 88 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Isle Royale: Keweenaw Rift Geology. Field Trip, 2013. Description An account of the resource Isle Royale: Keweenaw Rift Geology Physical Volcanology of Large Lava Flows Field Trip, Institute on Lake Superior Geology May 25-30, 2013 Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2013 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English https://digitalcollections.lakeheadu.ca/files/original/8154f0af057b119dd8208a23f584961e.pdf f4397eb22cabad582e48970a6ebf2063 PDF Text Text 69th ANNUAL MEETING Eau Claire, Wisconsin — April 24-25, 2023 INSTITUTE ON LAKE SUPERIOR GEOLOGY Part 1 — Program and Abstracts �Thank you to our sponsors! A SPECIAL THANK YOU TO OUR INDIVIDUAL CONTRIBUTORS: FREDERICK CAMPBELL, VAL CHANDLER, JIM DEGRAFF, THOMAS ERICKSON, TOM FITZ, DAVE GOOD, PAULA LEIER-ENGELHARDT, ALLAN MACTAVISH, BOB MAHIN, GORDON MEDARIS JR., JIM MILLER, STEVEN PINTA, TOD ROUSH, AND GERRY WHITE i �Proceedings of the 69th ILSG Annual Meeting – Part 1 69th ANNUAL MEETING INSTITUTE ON LAKE SUPERIOR GEOLOGY April 24-25th Eau Claire, Wisconsin HOSTED BY Rob Lodge, Esther Stewart, Carsyn Ames Co-Chairs University of Wisconsin- Eau Claire and Wisconsin Geological and Natural History Survey Proceedings - Volume 69 Part 1 – Program and Abstracts Compiled and edited by Carsyn Ames Cover Photos. Left— Photograph showing a group of men, women and children traveling through a forest north of Chippewa Falls, Wisconsin in a horse-drawn carriage, Chippewa Co., 1916. Center— Cross-bedding in basal Cambrian sandstone Eau Claire Co., 1919. Right — Outcrops of rhyolite schist along the north fork of the Eau Claire River, Eau Claire Co. 1919. ii �Proceedings of the 69th ILSG Annual Meeting – Part 1 69th INSTITUTE ON LAKE SUPERIOR GEOLOGY VOLUME 69 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD T RIP GUIDEBOOK Trip 1: PRECAMBRIAN GEOLOGY OF THE CHIPPEWA RIVER VALLEY Trip 2: WISCONSIN’S PALEOZOIC STRATIGRAPHY AND TOUR OF CRYSTAL CAVE Trip 3: PRECAMBRIAN GEOLOGY OF THE EAU CLAIRE RIVER VALLEY Trip 4: QUATERNARY GEOLOGY AND GEOMORPHOLOGY OF THE EAU CLAIRE REGION Reference to material in Part 1 should follow the example below: Grauch, V.J.S., Heller, Sam J., Stewart, Esther K., and Woodruff, Laurel G. 2023. Exploring the geology of the Midcontinent Rift under western Lake Superior using a preliminary velocity model of seismic line GLIMPCE C. in Ames C. (Ed.), Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 1 - Abstracts and Proceedings. v.69, part 1, p.3738. Published by the 69th Institute on Lake Superior Geology and distributed by the ILSG Secretary: Pete Hollings - ILSG Secretary Department of Geology Lakehead University 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Email: peter.hollings@lakeheadu.ca ILSG website: www.lakesuperiorgeology.org iii �Proceedings of the 69th ILSG Annual Meeting – Part 1 ISSN 1042-9964 Part 1: Program and Abstracts Table of Contents Institutes on Lake Superior Geology, 1955-2023 v Sam Goldich and the Goldich Medal vii Goldich Medal Guidelines ix Goldich Medalists and Goldich Medal Committee xi Citation for Goldich Medal Award to Peter Hollings xii Honoring the Pioneers of Lake Superior Geology xii Nomination for Thomas Benton Brooks, Pioneer of Lake Superior Geology xv Memoriams for Stephen Allard, Steven Hauck and Manfred Kehlenbeck xx Eisenbrey Student Travel Awards xxv Joe Mancuso Student Research Awards xxvi Doug Duskin Student Paper Awards and Award Committee xxvii Board of Directors and Session Chairs xxviii Field Trip Leaders and Guidebook Authors xxix Report of the 68th Annual Meeting xxx Technical Program xxxiv Poster Presentations xl Banquet Presentation xliii Abstracts 1-99 iv �Proceedings of the 69th ILSG Annual Meeting – Part 1 Institutes on Lake Superior Geology, 1955-2023 # Date Place Chairs 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 1955 1956 1957 1958 1959 1960 1961 1962 1963 1964 1965 1966 1967 1968 1969 1970 1971 1972 1973 1974 1975 1976 1977 Minneapolis, Minnesota Houghton, Michigan East Lansing, Michigan Duluth, Minnesota Minneapolis, Minnesota Madison, Wisconsin Port Arthur, Ontario Houghton, Michigan Duluth, Minnesota Ishpeming, Michigan St. Paul, Minnesota Sault Ste. Marie, Michigan East Lansing, Michigan Superior, Wisconsin Oshkosh, Wisconsin Thunder Bay, Ontario Duluth, Minnesota Houghton, Michigan Madison, Wisconsin Sault Ste. Marie, Ontario Marquette, Michigan St. Paul, Minnesota Thunder Bay, Ontario C.E. Dutton A.K. Snelgrove B.T. Sandefur R.W. Marsden G.M. Schwartz & C. Craddock E.N. Cameron E.G. Pye A.K. Snelgrove H. Lepp A.T. Broderick P.K. Sims & R.K. Hogberg R.W. White W.J. Hinze A.B. Dickas G.L. LaBerge M.W. Bartley & E. Mercy D.M. Davidson J. Kalliokoski M.E. Ostrom P.E. Giblin J.D. Hughes M. Walton M.M. Kehlenbeck v �Proceedings of the 69th ILSG Annual Meeting – Part 1 # 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 Date 1978 1979 1980 1981 1982 1983 1984 1985 1986 1987 1988 1989 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010 Place Milwaukee, Wisconsin Duluth, Minnesota Eau Claire, Wisconsin East Lansing, Michigan International Falls, Minnesota Houghton, Michigan Wausau, Wisconsin Kenora, Ontario Wisconsin Rapids, Wisconsin Wawa, Ontario Marquette, Michigan Duluth, Minnesota Thunder Bay, Ontario Eau Claire, Wisconsin Hurley, Wisconsin Eveleth, Minnesota Houghton, Michigan Marathon, Ontario Cable, Wisconsin Sudbury, Ontario Minneapolis, Minnesota Marquette, Michigan Thunder Bay, Ontario Madison, Wisconsin Kenora, Ontario Iron Mountain, Michigan Duluth, Minnesota Nipigon, Ontario Sault Ste. Marie, Ontario Lutsen, Minnesota Marquette, Michigan Ely, Minnesota International Falls, Minnesota 57 58 59 60 61 62 63 64 65 66 67 68 69 2011 2012 2013 2014 2015 2016 2017 2018 2019 2020 2021 2022 2023 Ashland, Wisconsin Thunder Bay, Ontario Houghton, Michigan Hibbing, Minnesota Dryden, Ontario Duluth, Minnesota Wawa, Ontario Iron Mountain, Michigan Terrace Bay, Ontario Meeting cancelled Virtual meeting Sudbury, Ontario Eau Claire, Wisconsin vi Chairs G. Mursky D.M. Davidson P.E. Myers W.C. Cambray D.L. Southwick T.J. Bornhorst G.L. LaBerge C.E. Blackburn J.K. Greenberg E.D. Frey & R.P. Sage J. S. Klasner J.C. Green M.M. Kehlenbeck P.E. Myers A.B. Dickas D.L. Southwick T.J. Bornhorst M.C. Smyk L.G. Woodruff R.P. Sage & W. Meyer J.D. Miller & M.A. Jirsa T.J. Bornhorst & R.S. Regis S.A. Kissin & P. Fralick M.G. Mudrey & Jr., B.A. Brown P. Hinz & R.C. Beard L. Woodruff & W.F. Cannon S. Hauck & M. Severson M. Smyk & P. Hollings A. Wilson & R. Sage L. Woodruff & J. Miller T.J. Bornhorst & J. Klasner J. Miller, G. Hudak, D. Peterson M. Jirsa, P. Hollings & T. Boerboom, P. Hinz & M.Smyk T. Fitz P. Hollings T.J. Bornhorst & A. Blaske J. Miller & M. Jirsa R. Cundari & P. Hinz J. Miller, C. Schardt & D. Peterson A. Pace, A. Wilson & T.J. Bornhorst L. Woodruff, W. Cannon & E.K. Stewart P. Hollings & M.C. Smyk Cancelled by the COVID-19 pandemic M. Jirsa, M. Smyk & P. Hollings R.M. Easton & W. Bleeker R. Lodge, E.K. Stewart, & C. Ames �Proceedings of the 69th ILSG Annual Meeting – Part 1 Sam Goldich and the Goldich Medal Sam Goldich received an A.B. from the University of Minnesota in 1929, a M.A. from Syracuse University in 1930, and a Ph.D. from the University of Minnesota in 1936. During World War II Sam worked for the U.S. Geological Survey in mineral exploration. In 1948, Sam returned to the University of Minnesota, and became Professor and Director of the Rock Analysis Laboratory the following year. He rejoined the U.S. Geological Survey in 1959 and was appointed as the first Branch Chief of the Branch of Isotope Geology. Sam returned to academia in 1964 when he went to Pennsylvania State University. He left PSU in 1965 and moved to the State University of New York at Stony Brook, where he stayed for 3 years. Restless yet again, he moved to Northern Illinois University in 1968 where he was a professor until his retirement in 1977. Sam’s final move was to Denver where he became an emeritus at the Colorado School of Mines. Sam died in 2000, less than a month before his 92nd birthday. In the late 1970s, Geological Society of America Special Paper 182, which included seminal geochronological studies by Sam Goldich and coworkers on the Archean rocks of the Minnesota River Valley, was nearing completion. At this time various ILSG regulars began discussing the possibility of recognizing Sam for his pioneering work on the resolution of age relationships and thus the geology of Precambrian rocks in the Lake Superior region. Three members, R.W. Ojakangas, J.O. Kalliokoski and G.B. Morey, presented the idea to the ILSG Board of Directors in 1978. The Board approved the creation of an award, provided funding could be obtained. It was suggested that collecting one or two dollars at registration for a dedicated account would provide resources for striking the medal. A general request was made to the ILSG membership for donations and Sam himself offered a challenge grant to match the contributions. In total $4,000 was collected and thus began the work of creating the Goldich Medal. The initial Goldich Award was presented to Sam by G.B. Morey in 1979 and consisted of a large paper proclamation. For the actual medal, G.B. Morey consulted with the foundry on production details, while Dick Ojakangas and Jorma Kalliokoski worked on the design of the award, suggesting that it be given for “outstanding contributions to the geology of the Lake Superior region.” Simultaneously, a committee of J.O. Kalliokosi, W.F. Cannon, M.M Kehlenbeck, G.B. Morey, and G. Mursky developed the Award Guidelines that were approved by the ILSG Board. By 1981 all the elements of the Goldich Award had come together, and the second recipient, Carl E. Dutton, Jr., received the Goldich Medal for 50 years of significant contributions to the understanding of the geology of the Lake Superior region. Since the beginning, the Awards Committee has consisted of individuals representing industry, government and academia, with each member of the Committee serving for three years. The medal is now awarded every year at the annual ILSG meeting. Reference: Morey, G.B. and Hanson, G.N. (editors). 1980. Selected studies of Archean gneisses and Lower Proterozoic rocks, southern Canadian Shield. Geological Society of America, Special Paper 182, 175 p. Prepared by various Goldich Medal Awardees, 2007 vii �Proceedings of the 69th ILSG Annual Meeting – Part 1 INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL viii �Proceedings of the 69th ILSG Annual Meeting – Part 1 Goldich Medal Guidelines (Adopted by the Board of Directors, 1981; amended 1999) Preamble The Institute on Lake Superior Geology was born in 1955, as documented by the fact that the 27th annual meeting was held in 1981. The Institute’s continuing objectives are to deal with those aspects of geology that are related geographically to Lake Superior; to encourage the discussion of subjects and sponsoring field trips that will bring together geologists from academia, government surveys, and industry; and to maintain an informal but highly effective mode of operation. During the course of its existence, the membership of the Institute (that is, those geologists who indicate an interest in the objectives of the ILSG by attending) has become aware of the fact that certain of their colleagues have made particularly noteworthy and meritorious contributions to the understanding of Lake Superior geology and mineral deposits. The first award was made by ILSG to Sam Goldich in 1979 for his many contributions to the geology of the region extending over about 50 years. Subsequent medallists and this year’s recipient are listed in the table below. Award Guidelines 1) The medal shall be awarded annually by the ILSG Board of Directors to a geologist whose name is associated with a substantial interest in, and contribution to, the geology of the Lake Superior region. 2) The Board of Directors shall appoint the Goldich Medal Committee. The initial appointment will be of three members, one to serve for three years, one for two years, and one for one year. The member with the briefest incumbency shall be chair of the Nominating Committee. After the first year, the Board of Directors shall appoint at each spring meeting one new member who will serve for three years. In his/her third year this member shall be the chair. The Committee membership should reflect the main fields of interest and geographic distribution of ILSG membership. The out-going, senior member of the Board of Directors shall act as liaison between the Board and the Committee for a period of one year. 3) By the end of November, the Goldich Medal Committee shall make its recommendation to the Chair of the Board of Directors, who will then inform the Board of the nominee. 4) The Board of Directors normally will accept the nominee of the Committee, inform the medallist, and have one medal engraved appropriately for presentation at the next meeting of the Institute. 5) It is recommended that the Institute set aside annually from whatever sources, such funds as will be required to support the continuing costs of this award. ix �Proceedings of the 69th ILSG Annual Meeting – Part 1 Nominating Procedures 1) The deadline for nominations is November 1. Nominations shall be taken at any time by the Goldich Medal Committee. Committee members may themselves nominate candidates; however, Board members may not solicit for or support individual nominees. 2) Nominations must be in writing and supported by appropriate documentation such as letters of recommendation, lists of publications, curriculum vita’s, and evidence of contributions to Lake Superior geology and to the Institute. 3) Nominations are not restricted to Institute attendees, but are open to anyone who has worked on and contributed to the understanding of Lake Superior geology. Selection Guidelines 1) Nominees are to be evaluated on the basis of their contributions to Lake Superior geology (sensu lato) including: a) importance of relevant publications; b) promotion of discovery and utilization of natural resources; c) contributions to understanding of the natural history and environment of the region; d) generation of new ideas and concepts; and e) contributions to the training and education of geoscientists and the public. 2) Nominees are to be evaluated on their contributions to the Institute as demonstrated by attendance at Institute meetings, presentation of talks and posters, and service on Institute boards, committees, and field trips. 3) The relative weights given to each of the foregoing criteria must remain flexible and at the discretion of the Committee members. 4) There are several points to be considered by the Goldich Medal Committee: a) An attempt should be made to maintain a balance of medal recipients from each of the three estates—industry, academia, and government. b) It must be noted that industry geoscientists are at a disadvantage in that much of their work in not published. 5) Lake Superior has two sides, one the U.S., and the other Canada. This is undoubtedly one of the Institute’s great strengths and should be nurtured by equitable recognition of excellence in both countries. x �Proceedings of the 69th ILSG Annual Meeting – Part 1 Goldich Medalists 1979 Samuel S. Goldich 1998 Zell Peterman 2016 Mark A. Jirsa 1980 not awarded 1999 Tsu-Ming Han 2017 Philip Fralick 1981 Carl E. Dutton, Jr 2000 John C. Green 1982 Ralph W. Marsden 2001 John S. Klasner 2018 Val W. Chandler 2019 Mark Severson 1983 Burton Boyum 2002 Ernest K. Lehmann 2020 not awarded 1984 Richard W. Ojakangas 2003 Klaus J. Schulz 2021 Alan MacTavish 1985 Paul K. Sims 2004 Paul Weiblen 2022 Terrence J. Boerboom 1986 G.B. Morey 2005 Mark Smyk 2023 Peter Hollings 1987 Henry H. Halls 2006 Michael G. Mudrey 1988 Walter S. White 2007 Joseph Mancuso 1989 Jorma Kalliokoski 2008 Theodore J. Bornhorst 1990 Kenneth C. Card 2009 L. Gordon Medaris, Jr 1991 William Hinze 2010 William D. Addison & 1992 William F. Cannon Gregory R. Brumpton 1993 Donald W. Davis 2011 Dean M. Rossell 1994 Cedric Iverson 2012 James D. Miller 1995 Gene La Berge 2013 Tom Waggoner 1996 David L. Southwick 1997 Ronald P. Sage 2014 Laurel Woodruff 2015 Rodney J. Ikola 2023 GOLDICH MEDAL RECIPIENT Peter Hollings Goldich Medal Committee Serving through the meeting year shown in parentheses. Steve Kissin (2018-2023*) Lakehead University (Committee Chair) Dorothy Campbell (2019-2024*) Ontario Geological Survey Dean Peterson (2022-2025) Big Rock Exploration *Terms of the committee members were extended 2 years because of the cancelation of the 2020 meeting and the logistical difficulties of voting during the 2021 virtual meeting. xi �Proceedings of the 69th ILSG Annual Meeting – Part 1 Citation for the Goldich Medal Recipient to Peter Hollings ILSG Members, it is our privilege to present the citation for this year’s recipient of the prestigious Goldich Medal to Dr. Peter Hollings. Pete received his Bachelor of Science with Honours in Geology from the Royal Holloway and Bedford New College, University of London in 1992. He continued as a postgraduate research assistant at Royal Holloway and Bedford New College until 1994 when he enrolled as a Ph.D. student at the University of Saskatchewan. He earned his Ph.D. in 1998 and his doctoral dissertation was titled “Geochemistry of the Uchi subprovince.” He had a one-year postdoctoral fellowship at Saskatchewan, followed by a two-year NSERC postdoctoral fellowship at the University of Tasmania. Pete joined the faculty at Lakehead University in 2001 as an Assistant Professor and in 2009 was promoted to full Professor, a title he continues to hold. Since 2013 Pete has been Director of the Centre of Excellence for Sustainable Mining and Exploration (CESME) at Lakehead University. He has served as Chair of the Department of Geology and as interim Dean of the Faculty of Science and Environmental Studies at Lakehead. Pete has been recognized for his research through several awards. In 2004 he and his coauthors were awarded the Julian Boldy Award by the Mineral Deposits Division of the Geological Association of Canada for an outstanding paper. In 2008 he was awarded the William Harvey Gross Medal by the Mineral Deposits Division of the Geological Association of Canada for significant contributions to the field of economic geology by a geoscientist under the age of 40. He was part of the team recognized by an award in 2012 and in 2014 by AMIRA International. In 2015 he was named the NSERC Distinguished Researcher for Lakehead University and in 2016 he was named the Lakehead University Research Chair in the NSERC/CHIR category. He received the Howard Street Robinson Medal from the Geological Association of Canada in 2017. In 2021, a paper on which he was co-author was awarded the 2020 Cameron-Hall Copper Medal for the most outstanding scientific publication in the journal Geochemistry: Exploration, Environment, Analysis (GEEA). Pete was awarded the NOHFC Industrial Research Chair in Mineral Exploration for a term from 2020 to 2025. Pete has an impressive professional record of publications and presentations. As of 2022, he has been first author or co-author of 145 refereed journal articles, 13 book chapters, 234 reports, 136 papers in refereed conference proceedings, and 87 abstracts in conference xii �Proceedings of the 69th ILSG Annual Meeting – Part 1 proceedings. While this is an impressive list of accomplishments, it is Pete’s ongoing contributions to our understanding of Lake Superior geology and to the Institute on Lake Superior Geology that make him a worthy recipient of the Goldich Medal. Pete has extensively conducted research on the geology of the Lake Superior region and the broader Superior Province. He has focused on both the Midcontinent Rift System (MRS) and Archean greenstone belts and their mineral resources. More than 30 of his published papers in refereed journals are on Lake Superior geology as well as about half of both his 30 first-authored conference proceedings and 29 first-authored refereed abstracts. He has contributed to more than 60 technical reports on Lake Superior geology. Of his 27 invited presentations, half have dealt with Lake Superior geology. Pete has a significant number of publications and presentations relevant to the discovery and utilization of natural resources in the Lake Superior region. Some of his numerous economic geology publications and presentations on topics outside of the Lake Superior region are also applicable to our regional geology. An area of emphasis in Peter’s research is the application of geochemistry and petrology to explore for ore deposits, including NiCu-PGE deposits (e.g., Lac des Iles Mine and the Thunder Bay North igneous complex). His other areas of interest include igneous geochemistry of the MRS, Archean greenstone belts and granites, the tectonic setting of komatiites, and Archean gold deposits. As the Director of CESME, he provides leadership in promoting the discovery of and environmentally responsible exploration for natural resources. Pete has also made contributions to understanding of the natural history and environment of the Lake Superior region as demonstrated by numerous publications focused on the timing and evolution of local rocks and mineral deposits. Pete’s research is firmly rooted in field work and uses geochemical and other data to test existing ideas and concepts and to develop new ones. He has successfully used local and regional geochemical data to provide evidence and/or implications for broader geological questions, such as atmospheric oxygen in the Precambrian, continental growth and lithospheric recycling, the Superior Province cratonic keel, and the earliest phases of Midcontinent Rift development. In addition to data-driven new ideas and concepts, Pete’s research efforts have resulted in development of new analytical approaches that can be applied to the Lake Superior region and beyond. As a Professor at Lakehead University, Pete is actively involved the education of geoscientists through classroom teaching and thesis supervision. He is committed to training and mentoring as evidenced by the large role students play in his research. He has supervised and co-supervised 37 honours undergraduate research projects and 32 Masters graduate student theses. Most of this student-focused research has involved Lake Superior geology. His former students now have senior positions with government and industry, and some have gone on to complete PhDs. Moreover, he supports and encourages students to attend and present their research at ILSG. xiii �Proceedings of the 69th ILSG Annual Meeting – Part 1 ILSG plays a significant role in Pete’s professional activities. He has authored/co-authored (many with his students) 75 ILSG abstracts (nearly 4 per year), six ILSG field trip guidebooks, and ILSG Special Publication #1, Field trip guidebook for the Slate Islands, Ontario. At his very first ILSG meeting in 2002, Pete co-authored an abstract and served on the Student Paper Awards Committee. Pete has Chaired or Co-Chaired four in-person annual meetings (Nipigon, 2005; International Falls, 2010; Thunder Bay, 2012; Terrace Bay, 2019) and the virtual meeting in 2021. He has served as the Secretary of the ILSG from 2003 to the present. As Secretary, he is responsible for email communications with the members of ILSG. As a member of the Board, he attends and chairs the annual Board meeting. In ILSG Board meetings he always considers and defends the best interests of Institute. Pete is the ILSG webmaster and played a key role in the current design of the ILSG website which he updates and maintains. Through his efforts, Lakehead University is the digital archive to all of the past ILSG proceedings and field trip guidebooks and provides open access of this content worldwide. A testament to the quality and accessibility of these documents was ILSG’s receiving the 2016 Outstanding Geologic Field Trip Guidebook Series Award by the Geoscience Information Society (GSIS), which Pete accepted on behalf of the Institute. The stature of ILSG in the regional, national, and international geological communities has been elevated because of the increased presence of ILSG on the worldwide web, in large part because of Pete’s efforts. Over the years, we have all witnessed Pete in action. He is collegial, easy to approach and gets along well with others, whether they be students, colleagues, or industry geoscientists. He is both a good listener and a good speaker. And he is open-minded. He has high personal standards and expects them to be reflected in the work of his students and research colleagues. Pete is truly enthusiastic about the geology of the Lake Superior region and about ILSG. Pete has made and continues to make substantial contributions to the field of geology and to the Institute on Lake Superior Geology. Pete has more than met the qualifications that are engraved on the Goldich Medal itself: “For outstanding contributions to the geology of the Lake Superior region.” We congratulate the 2023 Goldich Medalist, Peter Hollings. Citation by: Theodore J. Bornhorst, Goldich Medalist 2008 Mark C. Smyk, Goldich Medalist 2005 xiv �Proceedings of the 69th ILSG Annual Meeting – Part 1 Honoring the Pioneers of Lake Superior Geology (Adopted by the Board of Directors, 2016) Preamble At the suggestion of Gene LaBerge, the 2016 executive board agreed to implement a program to recognize historic pioneers in the understanding of geology in the Lake Superior region. Beginning with the 2017 annual meeting, nominations will be accepted from the membership for geologists whose work was conducted primarily before inception of the institute in 1955. Biographical sketches of those pioneers will be presented at future annual meetings so that all might appreciate the value of their contributions. Selection of nominees will be decided in part by the organizing committee of each year's annual meeting, in consultation with the Board, to ensure equitable geographic representation in the selection process. Award Guidelines 1) Nominations from the membership will be submitted via the Institute web site and forwarded to the Chair of the next Annual Meeting. The nominations will be no more than half a page in length and will summarise the contribution of the nominee. 2) The Organising Committee will select one or two individuals to be highlighted at the next Annual meeting and submit those names to the Board for approval. 3) The nominator will be requested to prepare a brief presentation to be given during the next annual meeting with a summary to be included in the Proceedings volume. 4) Unsuccessful nominations will be kept by the Secretary for two years and forwarded to the next meeting Chair; these nominations may be resubmitted at a later date. The Board will review this award every five years. Pioneers of Lake Superior Geology 2017 Douglass Houghton (1809-1845) 2018-20 not presented 2021 Newton Horace Winchell (1839-1914) 2022 Thomas Leslie Tanton (1890-1971) 2023 Thomas Benton Brooks (1836-1900) xv �Proceedings of the 69th ILSG Annual Meeting – Part 1 2023 Nomination for Thomas Benton Brooks Pioneer of Lake Superior Geology “During many years Major (T.B.) Brooks was the chief authority in the region on matters pertaining to geology, the ores and the mines of the iron region of Lake Superior.”1 Shortly after the U.S. Civil War Major Thomas Benton Brooks moved to the Marquette Iron Range. There over the course of less than a decade, he became the premier geologist, prospector, mining and civil engineer, and mining company executive of the region. During these formative years of the iron ore industry, when the Lake Superior region was providing about one-quarter of the iron ore used in the U.S., he was employed by the Iron Cliffs Company, the predecessor of the ClevelandCliffs Company, the Michigan and Wisconsin Geological Surveys, and served as a consultant to iron ore exploration and mining companies of the region. His contributions had a significant role in mapping the Precambrian geology and iron ranges of Michigan and Wisconsin and a lasting impact on the iron ore industry of the region. As stated by Prof. C.R. Van Hise, Brooks’ successor as the premier geologist of the Lake Superior region2: “Notwithstanding the immense advantage which it has been to have Brooks’ work as a foundation, it has taken many years of labor fairly to complete the structural story to which Brooks contributed important chapters. Only those who have labored in the Lake Superior region and who understand its peculiar difficulties can give Brooks credit for the remarkable work he did. His geological work is my ideal of what should be done in a new region of complex geology.” Thomas B. Brooks was born on June 15, 1836 in Monroe, NY, near the New Jersey border, and died nearby on November 22, 1900. In 1852 at the age of 16, he joined a surveying crew of the Erie Railroad and rapidly advanced from woodsman to instrument man. In 1853 he was employed with the New York Topographic and Geological Survey and then entered the Engineering Department of Union College of Schenectady, NY in 1856, graduating in 1858 in civil engineering. He remained at Union College as an instructor for a year and then took part in topographical surveys in New York, New Jersey, Pennsylvania and the U.S. Gulf Coast. In 1 Quoted from an article by Chas. A. Lawton in the Daily Mining Journal, November 29, 1900 entitled The Late Major Thomas Benton Brooks: Biographical Sketch of a Man Whose Name is Intimately Associated With the Early Development of Michigan’s Iron Mines. The Mining Journal, the predominant daily newspaper of Marquette, Michigan and the Northern Peninsula of Michigan, was founded in 1841. 2 As quoted by Bailey Willis of the U.S. Geological Survey in an obituary for Major Brooks in the Proceedings of the American Association for the Advancement of Science, New Series, Volume 13, No. 325(March 22, 1901), 460-462. xvi �Proceedings of the 69th ILSG Annual Meeting – Part 1 1860 he attended a series of lectures on geology given by Prof. J.P. Lesley former state geologist of Pennsylvania and Professor of Geology at the University of Pennsylvania. This was his only formal education in geology. He volunteered for the Union Army in 1861 and organized an engineering company that had a distinguished record during numerous Civil War campaigns. He retired from the Union Army in 1864 as a brevet colonel after being wounded in the battle of Denly’s Bluff, but referred to himself after the war as Major Brooks. In 1865 after leaving the Union Army he accepted a position with the Geological Survey of New Jersey where he conducted magnetic surveys with a dip needle to locate iron ores and was put in charge of mines and furnaces. Shortly thereafter, he was induced to take charge of the mines of the Iron Cliffs Company in the Marquette Iron Range as vice-president and general manager. He moved to Negaunee, Michigan, where his practical knowledge of geology and engineering, leadership skills, originality, keen powers of observation and deduction, and intense work ethic served him, the company, and the Lake Superior region well. This is where his extensive geological studies began and where he developed the instruments and methodology to exploit the iron ores of the Lake Superior region. He brought the dip needle to the Lake Superior region and was among or possibly was the very first, to use it in iron ore exploration and geologic mapping in the region. He also pioneered the dial (Sun) compass, which he modified for geologic use from the surveying solar compass developed by W.A. Burt. In 1869 he resigned from the Iron Cliffs Company and was given the responsibility of mapping and reporting on the Marquette Iron Range and was placed in charge of the Economic State Geological Survey of the district by the Michigan Geological Survey, essentially becoming the State Geologist of the Northern Peninsula. He received no salary for this position, but he was allowed to receive private funds from numerous iron ore companies and mines. Unfortunately, his intense work schedule took a toll on his health that caused him to leave Marquette with his family in the winter of 1872-73 for London, England and eventually Dresden, Germany, where he hoped to regain his health, but failed to do so. During this period he prepared reports on his iron range geologic studies for publication by the Michigan and Wisconsin Geological Surveys (Brooks, 1873 and 1880), articles on the geology of the region and magnetic surveying instruments and their use published in various journals including the American Journal of Science and Arts (Brooks and Pumpelly, 1872; Brooks, 1875), and co-authored the book “Iron Ores of Missouri and Michigan” (Pumpelly, Brooks, and Schmidt, 1876). During his years involved with the geology and ores of the Lake Superior region Major Brooks made numerous advances in the geological knowledge of the region that have served as a foundation for future studies and developed methods and instruments that proved useful for exploiting the ores of the region for many years. The following are a list of his major lasting accomplishments: • • He with the assistance of R. Pumpelly and R.D. Irving developed the dial (Sun) compass for geologic studies based on the principal of Burt’s surveying solar compass which together with the dip needle that he brought from the Geological Survey of New Jersey were used in the Lake Superior region for nearly a century to locate and outline iron-rich rocks and ores. His publications on these instruments led to their extensive worldwide use. He established procedures for conducting magnetic surveys for geological purposes in the Lake Superior region and methods of interpreting the observations of the surveys based on empirical studies. xvii �Proceedings of the 69th ILSG Annual Meeting – Part 1 • • • • • • • He was the first to describe the magnetic characteristics of the minerals and rocks of the Lake Superior region. He (Brooks, 1872a) recognized that magnetic anomalies observed in the area of non-magnetic Paleozoic (then Silurian) sedimentary rocks of the eastern part of the Northern Peninsula of Michigan were likely derived from the basement Precambrian rocks that crop out to the west. Accordingly, these anomalies could be used to trace the basement rocks and their structure beneath the sedimentary rocks. Furthermore, he realized that anomaly characteristics could be used to determine the depth to magnetic sources and thus, the thickness of the sedimentary rocks. In a similar manner he understood that perhaps the depth of Lake Superior could be determined from analysis of the lake magnetic anomalies. He founded the first assay facility for iron ores in the Lake Superior region in the city of Marquette which facilitated iron ore mining in the region. He conducted one of the first geological surveys of the Marquette, Menominee, Crystal Falls, and Gogebic Iron Ranges. He was the first to understand that the Marquette Iron Range occurs within a 75-km long syncline extending to the west from near Marquette, Michigan (Allen and Martin, 1922). He recognized the stratigraphic position of the copper-bearing rocks of the Northern Peninsula of Michigan and suggested the name Keweenawian (note his spelling) for the age of these rocks in American Journal of Science and Arts articles of 1872 and 1875. Subsequently, the term Keweenawan has been used for these rocks. He had an important role in developing safe, efficient methods of mining iron ores of the Lake Superior region (Brooks, 1972b). He was intensely interested in the education of his children and supported the studies of his son, Alfred Hulse Brooks, a famed geologist of the U.S. Geological Survey, Alaska Branch, who is honored by naming of the Brooks Range of Alaska after him. These are all significant contributions that have had a profound role in understanding of the geology of the Lake Superior region and the exploitation of its ores. They have largely gone unrecognized for the past century and a half, but they clearly distinguish Major Thomas Benton Brooks as a Pioneer of Lake Superior geology. References Allen, R.C., and Martin, H.M., 1922. A brief history of the Geological and Botanical Survey of Michigan. Michigan History Magazine, Volume VI, No. 44: 675-750. Lawton, C.A., 1900. The Late Major Thomas Benton Brooks: Biographical Sketch of a Man Whose Name is Intimately Associated with the Early Development of Michigan’s Iron Mines. The Daily Mining Journal, November 29, 1900. Pumpelly, Raphael, Brooks, T.B., and Schmidt, A., 1876. Iron Ores of Missouri and Michigan. G.P. Putnam’s Sons, New York: 624. Willis, B., 1901. Thomas Benton Brooks. Proceedings of the American Association for the Advancement of Science, Science, New Series, Volume 13, No. 325: 460-462. William J. Hinze, Purdue University xviii �Proceedings of the 69th ILSG Annual Meeting – Part 1 APPENDIX: PUBLICATIONS OF T.B. BROOKS Brooks, T.B., 1872a. On the use of the magnetic needle in mineral explorations on Lake Superior. Van Nostrand’s Eclectic Engineering Magazine (1869-1879), August 1, 1872; Volume 7, No. 44, American Periodicals: 161. Brooks, T.B., 1872b. An analysis of the cost and description of the methods of mining employed in the Marquette Iron Region, Lake Superior, Michigan. Transactions of the American Society of Civil Engineers, Volume XXXIV: 18. Brooks, T.B., and Pumpelly, R., 1872. On the age of the copper-bearing rocks of Lake Superior. American Journal of Science and Arts, Third Series, Volume III, No. XVIII: 428-432. Brooks, T.B., 1873. Geology of Marquette Iron Range, Geology of the Menominee Iron Range, and Geology of the Gogebic and Montreal Iron Ranges. Michigan Geological Survey, Volume 1, Chapters IV, V, VI, VII, and VIII, Part 1, Iron-Bearing Rocks: 117-243. Brooks, T.B., 1875. On the youngest Huronian rocks south of Lake Superior and the age of the copperbearing series. American Journal of Science and Arts, Third Series, Volume III, No. XI: 206211. Brooks, T.B., 1880. Geology of the Menominee Region. In Chamberlin, T.C. (ed.), Geology of Wisconsin, Volume 3, Part 7, Chapters 1, 2, and 3: 430-552. xix �Proceedings of the 69th ILSG Annual Meeting – Part 1 In Memoriam Stephen Allard Stephen Thomas Allard, 59, of Winona, MN, passed away on Friday, September 16, 2022. He was born May 2, 1963, in Manchester, New Hampshire and graduated from Manchester Central High School before going on to receive both his bachelor’s and master’s degrees from the University of New Hampshire, and his doctorate from the University of Wyoming. In 2002, Stephen moved to Winona, MN to begin his career as a professor at Winona State University. After serving for 19 years as a faculty member in the Department of Geoscience, Stephen retired from the university in December of 2021. During his tenure at WSU, Stephen served on several committees and taught 13 different courses drawing on his expertise in hard rock and structural geology. Stephen was dedicated to teaching and mentoring students through field-based research, leading courses and field trips throughout the United States, notably the many summers spent in the Black Hills of South Dakota. (modified from Hartford Courant newspaper) xx �Proceedings of the 69th ILSG Annual Meeting – Part 1 In Memoriam Steven A. Hauck This Fall the Institute on Lake Superior Geology lost a dedicated geologist and friend, Steve Hauck, who was a regular attendee of ILSG (since at least 1984) and worked on countless projects in the Lake Superior Region while employed at the Natural Resources Research Institute (NRRI) in Duluth, MN. During that time, Steve was a mentor to numerous geologists in the region throughout their early and continuing careers. Steve Hauck had just recently moved from Duluth to Euclid, OH where he passed away on October 6, 2022 at the age of 73. Steve as born on May 16, 1949, in Rochester, NY, where he graduated from Gates-Chili High School prior to attending Albion College where he earned a BS in geology. He enlisted in the US Army where he was trained as a Chinese translator and married fellow Albion student and the love of his life Barbara Horsley to whom he was married for 50 years. Steve loved to talk about geology on car trips and impressed his future father-in-law with his knowledge and enthusiasm. Steve later earned a MS degree in geology from the University of North Carolina before embarking on a geology career that eventually led him around the globe. He was first employed by Union Carbide in Grand Junction, CO, where he was responsible for exploration for uranium in the 4-corners region. While at Union Carbide he was also responsible for developing a world-wide exploration program in search of IOCG deposits (as they were later called) and visited many similar deposits including Olympic Dam, Pilot Knob and Pea Ridge in Missouri, and Kiruna iron deposits in Sweden to name a few. Steve’s first ILSG talk (1984) pertained to the distinguishing characteristics of these types of deposits and was titled “Comparison of Middle Proterozoic Iron Oxide Rich Ore Deposits, Mid-Continent, USA, South Australia, Sweden, and the Peoples Republic of China.” Steve was then hired as the second employee of the Minerals Division of the NRRI in 1984 as Research Director and Manager where he worked for over 30 years. He was initially responsible for building and equipping the division, focusing on economic geology, and initially hired graduate students from the University of Minnesota Duluth (UMD). During his tenure at the NRRI, Steve hired well over 30 UMD students (both undergraduate and graduate students) as well as many geologists in their early career years. Projects that he and his coworkers researched ranged from clay deposits in SW MN, to Cu-Ni and Fe-Ti deposits in the Duluth Complex, to the Biwabik Iron Formation, to geochemistry on a wide range of rocks spanning from the Archean to the Cretaceous. He worked closely with fellow geologists at the Minnesota Geological Survey, Minnesota Department of Natural Resources Lands and Minerals, and the U.S. Geological Survey, and collaborated with many geologists across the U.S. and overseas in academia and industry. xxi �Proceedings of the 69th ILSG Annual Meeting – Part 1 Steve was a Co-chair of the ILSG meeting for its 50th Anniversary in Duluth in 2004 and served on the Board of Directors for three years. Overall, Steve participated in three ILSG talks (one as primary author) and ten posters (four as primary author). Steve loved to talk about rocks and encouraged his co-workers to give talks and poster presentations at many of the ILSG meetings. Steve was an avid birder, cultivator of native plants, and shutterbug. He was predeceased by his youngest son, Davis, and his parents Arthur and Jean (Doron). He is survived by wife Barbara, son Steven (Danette), sisters Carlin Eagan (Daniel), Sandra Doron, and Mary McGuire (Mark), and two grandchildren Levi and Abigail. (modifed from Duluth Tribune newspaper) xxii �Proceedings of the 69th ILSG Annual Meeting – Part 1 In Memoriam Manfred Kehlenbeck Manfred was born in Bremen, Germany in 1937 to parents Emma and Theodor. This is where he spent his childhood, amidst the horrors of World War II, like so many of Europe's children. At age 14, Manfred immigrated with his parents to the U.S., landing in New York in July of 1952 and settling with relatives in Long Island until they could become established. Here he completed his high school education, then attended Hofstra University for his undergraduate degree. It was there that he was introduced to the science of geology, which became his lifelong interest and focus of his future education and career. It was also on Long Island that he met Elenore, who would eventually become his wife of 53 years. Manfred went on to Syracuse University in upstate New York to attain his M.Sc. in Geology and gain field experience in the beautiful Adirondack Mountains. And then, moving even further north, he attended Queen's University in Kingston, Ontario where he achieved his Ph.D. Since he has always planned to teach, he then accepted a position at the young Lakehead University in Thunder Bay, Ontario. Here he soon became fascinated with the Precambrian geology of the area and greatly enjoyed his teaching duties. He was a born teacher, winning Teacher of the Year awards both at Lakehead and in the Province. He served five terms as a Geology Department Chair, guiding the department into its M.Sc. program. His years at Lakehead were productive and happy ones. Upon his retirement, Manfred was able to expand on other interests and travel widely. In addition to trips in Canada, the U.S. and Germany, there were four “special” ones – professionally to Russia and China, and then the most fascinating ones, to the Arctic and Antarctic. His other areas of interest and hobbies were in watercolour and pen and ink drawings of local scenes, especially forests, lakes, rocks, and old buildings of NW Ontario xxiii �Proceedings of the 69th ILSG Annual Meeting – Part 1 and many views of Old Fort William. Many of his works hang in Thunder Bay homes. He became an avid gardener, curler and opera lover, and spent many hours volunteering for various causes. It was a happy and fulfilling retirement for Manfred and Elenore until his last illness and unexpected passing in the early morning hours of July 7, 2022 when he drew his last breath at the Thunder Bay Regional Health Sciences Centre after emergency surgery. Our thanks to the I.C.U. staff and especially to Katie and Michaela who were so kind and thoughtful during those last terrible hours, and to N.P. Crystal Kaukinen for the many years of care she had provided. Thanks also to all who have been so kind with phone calls, cards, offers to help, food and rides. Special thanks to Barb Morriss for always checking in, to Sam and Georgina Spivak for all the rides, and to Vince and Frieda DeSa who have been here for me everyday with their help and support – without them, I don't know how I would have survived this devastating time. Manfred was a good, kind, generous man, and loving and devoted husband. He is sorely missed. Auf Wiedersehen mein lieber Manfred. Published by The Thunder Bay Chronicle Journal on Aug. 13, 2022. xxiv �Proceedings of the 69th ILSG Annual Meeting – Part 1 Eisenbrey Student Travel Awards The 1986 Board of Directors established the ILSG Student Travel Awards to support student participation at the annual meeting of the Institute. The name “Eisenbrey” was added to the award in 1998 to honor Edward H. Eisenbrey (1926-1985) and utilize substantial contributions made to the 1996 Institute meeting in his name. “Ned” Eisenbrey is credited with discovery of significant volcanogenic massive sulfide deposits in Wisconsin, but his scope was much broader—he has been described as having unique talents as an ore finder, geologist, and teacher. These awards are intended to help defray some of the direct travel costs of attending Institute meetings, and include a waiver of registration fees, but exclude expenses for meals, lodging, and field trip registration. The number of awards and value are determined by the annual Chair in consultation with the Secretary and Treasurer. Recipients will be announced at the annual banquet. The following general criteria will be considered by the annual Chair, who is responsible for the selection: 1) The applicants must have active resident (undergraduate or graduate) student status at the time of the annual meeting of the Institute, certified by the department head. 2) Students who are the senior author on either an oral or poster paper will be given favored consideration. 3) It is desirable for two or more students to jointly request travel assistance. 4) In general, priority will be given to those in the Institute region who are farthest away from the meeting location. 5) Each travel award request shall be made in writing to the annual Chair, and should explain need, student and author status, and other significant details. Successful applicants will receive their awards during the meeting. xxv �Proceedings of the 69th ILSG Annual Meeting – Part 1 Joe Mancuso Student Research Awards The 2005 Board of Directors established the ILSG Student Research Fund with $10,000 US from the Institute’s general fund to encourage student research on the geology of the Lake Superior region. A minimum of two awards of $500 US each for research expenses (but not travel expenses) will be made each year. Students are expected to present their research orally or during a poster session at an ILSG meeting. The award winners will also be automatically eligible for the Eisenbrey Travel Awards. To allow the fund to grow, the Fund will receive one-half of any additional proceeds from each annual meeting, after all other commitments and expenses are covered. • The ILSG Board of Directors will be responsible for selecting a minimum of two awards each year. The ILSG Treasurer will issue the awards. • The ILSG Student Research Fund is available for undergraduate or graduate students working on geology in the Lake Superior region. • The applications are due to the ILSG Secretary by August 31st of each year. Awards will be made by October 1st of each year. • Names of the award recipients will be announced at the next annual meeting and posted on the ILSG website. • Details of the application process can be found on the ILSG web site. • The proposal will need to be signed by the researcher’s supervisor. The 2012 Board of Directors approved modification of the fund’s name, adding “Mancuso” to reflect the many contributions of Joseph Mancuso to the organization and sizeable donations made in his name. “Doc Joe,” as he was known by his students, taught geology for 36 years at Bowling Green State University, Ohio. He advised many graduate students in field-oriented research, and frequently brought them to Institute meetings. Joe was the 2007 Goldich Medalist. In fall 2022, the ILSG Board of Directors selected two students to be granted research funding of $1000.00 each from the Joe Mancuso Student Research Fund. The awardees were: Itai Bojdak-Yates Lawrence University Department of Geosciences TOPIC: Detrital zircon provenance study of Paleozoic sandstones from Wisconsin Lillian Glodowski University of Wisconsin- Eau Claire TOPIC: Petrogenesis of the Lynne Zn-CuPb Deposit, Oneida Co., Wisconsin Evan Weber University of Wisconsin- Eau Claire TOPIC: U/Pb Geochronology and Zircon Trace Element Geochemistry of the Pembine-Wausau Terrane of the Proterozoic Penokean Orogen, Wisconsin xxvi �Proceedings of the 69th ILSG Annual Meeting – Part 1 Doug Duskin Student Paper Awards Each year, the Institute selects the best of student presentations and honors the presenters with a monetary award. Funding for the award is generated from registrations of the annual meeting, and from generous donations to the fund in honor of Doug Duskin—an exploration geologist and longtime friend of the Institute. The 2012 ILSG Board of Directors approved adding Doug’s name to the award to acknowledge his contributions, and distribute those donations in a manner that would have pleased him. The Duskin Student Paper Committee is appointed by the Meeting Chair. Criteria for best student paper—last modified by the Board in 2001—follow: 1) The contribution must be demonstrably the work of the student. 2) The student must present the contribution in-person. 3) The Student Paper Committee shall decide how many awards to grant, and whether or not to give separate awards for poster vs. oral presentations. 4) In cases of multiple student authors, the award will be made to the senior author, or the award will be shared equally by all authors of the contribution. 5) The total amount of the awards is left to the discretion of the meeting Chair in conjunction with the Secretary, but typically is in the amount of about $500 US (increase approved by Board, 10/01). 6) The Secretary maintains, and will supply to the Committee, a form for the numerical ranking of presentations. This form was created and modified by Student Paper Committees over several years in an effort to reduce the difficulties that may arise from selection by raters of diverse background. The use of the form is not required, but is left to the discretion of the Committee. 7) The names of award recipients shall be included as part of the annual Chair’s report that appears in the next volume of the Institute. Student papers will be noted on the Program. xxvii �Proceedings of the 69th ILSG Annual Meeting – Part 1 Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected. The terms of Board members were extended 2 years because of cancellation of the 2020 meeting, and the difficulties of virtual voting by the membership during the 2021 meeting. Mike Easton, Chair (2022-2025) — Ontario Geological Survey Mark Smyk (2019-2024*) — Lakehead University Esther Stewart (2018-2023*) – Wisconsin Geological & Natural History Survey Peter Hollings — Secretary (2019-2024*) — Lakehead University Mark A. Jirsa — Treasurer (2022-2025) — Minnesota Geological Survey xxviii �Proceedings of the 69th ILSG Annual Meeting – Part 1 Field Trip Leaders and Guidebook Authors Field trips have been the mainstay of the ILSG since its inception 69 years ago. We want to give a special thanks to the field trip leaders and guidebook authors who volunteered their time and talent in carrying that tradition forward. 1) Precambrian geology of the Chippewa River Valley Rob Lodge- UW- Eau Claire Bob Hopper- UW- Eau Claire 2) Wisconsin’s Paleozoic stratigraphy and tour of Crystal Cave Carsyn Ames- Wisconsin Geological and Natural History Survey Esther Stewart- Wisconsin Geological and Natural History Survey William Batten- Wisconsin Geological and Natural History Survey Eric Stewart- Wisconsin Geological and Natural History Survey Ian Orland- Wisconsin Geological and Natural History Survey 3) Precambrian geology of the Eau Claire River Valley Rob Lodge- UW- Eau Claire Evan Weber- UW- Eau Claire (student) 4) Quaternary geology and geomorphology of the Eau Claire Region Doug Faulkner- UW- Eau Claire Elmo Rawling- Wisconsin Geological and Natural History Survey xxix �Proceedings of the 69th ILSG Annual Meeting – Part 1 REPORT OF THE 68th ANNUAL MEETING OF THE INSTITUTE ON LAKE SUPERIOR GEOLOGY The Ontario Geological Survey (OGS), with support from the Geological Survey of Canada (GSC), hosted the 68th Annual Institute on Lake Superior Geology on May 07 – 12, 2023 in the “Cavern” at Science North in Sudbury, Ontario. The meeting consisted of two days of technical sessions with pre- and post-technical session field trips. First, we would like to thank the meeting sponsors for their generous support, either through direct funding or in-kind support, namely: the Centre for Excellence and Sustainable Mineral Exploration in Thunder Bay, Gel Exploration Limited, the Northwestern Ontario Prospectors Association, Vale Canada, and the Ontario Geological Survey. We also thank the Individual Contributors to the Student Travel Scholarship fund: Mary Kay Arthur, Mike Beauregard, Ben Berger, Terry Boerboom, Jim DeGraff, Michael and Monica Easton, Dick Heglund, Joanna Hodge, Bob Mahin, Jim Miller, Dean Peterson, Mark and Laurie Severson, Al MacTavish and Graham Wilson. The 2022 meeting was the first in-person meeting held since the 2019 Terrace Bay meeting. An ILSG meeting questionnaire, which ran from January 20 to February 20, 2022, was key to shaping the format and venue of the meeting during a period of rapidly changing COVIDrelated regulations, with most responses favouring an in-person meeting. For technical reasons, a hybrid meeting was not possible. Total meeting registration was 80, including 12 students. This registration is about 80% of the attendance of the last two Sudbury area meetings (Sudbury 1997; Sault Ste. Marie 2006), and was a great turnout given the COVID-related travel restrictions still in place at the time of the meeting. Attendance from the United States was excellent, with attendance from the Sudbury area lower than expected, for unknown reasons. Despite the somewhat lower attendance, the technical program was nevertheless excellent, with a strong focus on Midcontinent Rift geology and mineralization in the Lake Superior region. In addition, four presentations focused specifically on Sudbury area geology. There was also time in the schedule for several impromptu presentations on a variety of topics on Wednesday afternoon prior to the announcement of the student awards. Proceedings Volume 68 was published in two parts. Part 1 – Program and Abstracts, compiled and edited by Michael Easton (OGS), contains 28 published abstracts for 21 oral and 8 poster presentations (one poster did not have an abstract). Students presented 5 oral and 5 poster presentations. Part 2 – Field Trip Guidebooks, also was compiled and edited by Michael Easton. It contains descriptions of three pre-meeting and two post-meeting field trips. Hard copies of the Abstract Volume and Field Trip Guidebooks for trip participants were printed by Johanne Roux and Carlo Castrechino (OGS) after it proved impossible to find a commercial printer who could produce the volumes in time for the meeting. Both volumes are available for download from the Institute on Lake Superior Geology website. Monica Easton is thanked for assisting in preparing the digital versions of both volumes. The 68th ILSG marked only the second time in the Institute’s long history that its annual meeting was held in Sudbury, the last time being in 1997. Since the discovery of distal ejecta xxx �Proceedings of the 69th ILSG Annual Meeting – Part 1 from the Sudbury impact in the western Lake Superior area in 2005, many members of the Institute had suggested that the time was right for another Sudbury meeting. The meeting location enabled organizers to offer five field trips that showcased a variety of Proterozoic rocks in the Sudbury area itself, as well as along the north shore of Lake Huron. Three field trips focussed on the geology and mineralization related to the Sudbury Structure, and the organizers wish to thank the local exploration companies that graciously provided information and access to their properties. Parts of the other two of the field trips had been offered at previous ILSG annual meetings (e.g., Sudbury 1997; Sault Ste. Marie 2006), but both greatly benefitted from the new mapping, research, discoveries and interpretations that had taken place in the intervening years. COVID-related shortages of rental vehicles and/or drivers led to pre-meeting trips being held over several days, which unexpectedly, provided more opportunities for attendees to take in several field trips if they wanted. All the field trips, and the meeting itself, were blessed with sunny weather and a minimum number of pesky insects. Total field trip participation was 96 (excluding leaders and volunteer drivers). A list of field trips is provided below (numbers correspond to trip numbers in the Guidebook volume): Pre-meeting field trips (and leaders) on Saturday, May 07; Sunday, May 8, and Monday, May 9. 5) A cross-section through the Huronian Supergroup at Elliot Lake, Ontario (Michael Easton, Ontario Geological Survey) (May 7) 2) Geology of the Grenville Front in the Sudbury area (Michael Easton, Ontario Geological Survey) (May 8) 1) Traverse across the Sudbury Impact Structure (Wouter Bleeker, Geological Survey of Canada, and Sandra Kamo, University of Toronto; Michael Lesher and Henning Seibel, Laurentian University) (Two-day trip, May 8 and May 9) Post-meeting field trips (and leaders) on Thursday, May12 3) Magmatism and brecciation in the Footwall Rocks of the southwestern Sudbury Structure (Caroline Gordon, Ontario Geological Survey; Carol-Anne Généreux, Laurentian University and Terrane Geoscience; and Brad Clarke, SPC Nickel Corporation) 4) An overview of the geology of the Sudbury Structure (Shirley Péloquin, Ontario Geological Survey) Many registrants attended the welcoming reception on Monday evening, which included an IMAX theatre presentation on “Dinosaurs of Antarctica”. Furthermore, the vast majority of registrants and invited guests attended the annual ILSG banquet on Tuesday night. Although a Homer Award overview presentation was given, no “recipients” were identified during the 2022 annual meeting, or in the previous 3 years! As always, a highlight of the post-banquet activities was presentation of the 2022 Goldich Medal. This year’s very deserving recipient was Terry Boerboom. The Goldich Medal citation was presented by Mark Jirsa, his colleague for many years. Mark described Terry’s contributions to the ILSG and to the greater understanding of Minnesota’s geology over several decades during his time as a student and his 35 years with the Minnesota Geological Survey. Terry is indeed a worthy recipient of this prestigious award. The 68th ILSG saw a return to the usual post-banquet guest speaker tradition. Andy Parmenter xxxi �Proceedings of the 69th ILSG Annual Meeting – Part 1 of the Canadian Nuclear Waste Management Organization (NWMO) travelled from Toronto to give an overview of NWMOs Geoscience site characterization of the Revell batholith in the Ignace area of northwestern Ontario. His talk provided detailed insights into the 3-D character of a Neoarchean granodioritic to granitic intrusion, based on detailed mapping and geophysical, seismic, and geochemical studies, as well as from multiple 1 km-long research cores obtained from the batholith. In 2022, the student paper committee had its usual difficult job of selecting the best among five excellent oral presentations and five poster presentations for the Doug Duskin Student Paper Awards. The committee awarded four prizes, with the best talk award going to Rebecca Price for her talk on “Mineralogy and Petrology of the Good Hope Carbonatite Complex, Marathon, ON” and the best poster award going to Khalid Yahia for his poster on “Geochemical and isotopic composition of Midcontinent Rift-related intrusions of the Thunder Bay North Igneous Complex, northwestern Ontario, Canada”. Runner-up prizes went to Audray Hinkenmeyer for her talk on “Characterizing Late Wisconsinan Rainy Lobe till from the Hudson Bay Lowlands to SW Minnesota: Insights on provenance and ice sheet behavior during Late Wisconsin glaciation” and to Katherine Langfield for her poster on “Slip Kinematics of the Hancock Fault in the Midcontinent Rift System, Keweenaw Peninsula, Michigan”. Eisenbrey Student Travel Grants were given to three students: Connor Caglitoti (Lakehead University), Katherine Langfield (Michigan Technical University), and Miles Harbury (University of Wisconsin, Milwaukee). The Institute’s Board of Directors met on Tuesday, May 10, 2022, and a brief overview of the meeting notes is provided below: 1. Accepted report of the Chairs for the 67th ILSG, Virtual Meeting; as published on the ILSG web site, and minutes of last Board meeting in May 2021. 2. Received, discussed, and accepted 2021-2022 ILSG Financial Summary. 3. Received, discussed, and accepted 2021-2022 report of the Secretary (Hollings). 4. Approved Michael Easton as on-going ILSG Board member 5. Discussed and approved renewal of Mark Jirsa as Institute Treasurer (end of term 2022). This was later approved by a vote of the membership. 6. Discussed and approved replacing Dan England as the “member from industry” on the Goldich Committee (end of term 2022) with Dean Peterson. 7. Approved Eau Claire as the site for the 69th annual ILSG meeting. The meeting will be hosted by Robert Lodge and Esther Stewart. 8. Reviewed and approved the guidelines for the Honouring the Pioneers of Lake Superior Geology with the charge that the document will be reviewed as needed. 9. Future meeting locations were discussed. Ted Bornhorst offered Houghton in 2024, Peter Hinz has offered Kenora as a future site and Mark Jirsa is keen to host the Mountain Iron meeting that was cancelled in 2020 because of the pandemic. In a subsequent discussion, Bernie Saini-Eidukat expressed a willingness to organize a meeting in St. Cloud. 10. The cost of insurance was discussed and it was agreed that the Board of Directors insurance and field trip insurance should be maintained for future meetings and that the costs would be included in the cost of each meeting. The fact that the Institute meets in both the US and Canada is an added complication. xxxii �Proceedings of the 69th ILSG Annual Meeting – Part 1 11. Jirsa advised the board of the donation of polar bear carvings from Mike Beauregard, and it was agreed that a silent auction would be held during the meeting with funds going to support student travel. Dan England later donated two samples with visible gold and, combined, these items raised $395 for the Eisenbrey fund 12. Bornhorst advised that there are a small number of hard copies missing from the MTU archives and that he will work to fill these. It was agreed that the ILSG would make a donation of $1 per member (minimum $100) each year to the library as a “thank you” for their efforts 13. The 68th ILSG meeting was a great success and we wish to thank all the people who contributed to that success, including staff of the Ontario Geological Survey who were pressed into action as editors, field trip leaders and drivers. Patty Cobin and Ted Bornhorst (A.E. Seaman Mineral Museum, Michigan Technological University) handled the premeeting registration. Ted also supplied the poster boards. Thanks go also to the staff at Science North who helped the meeting run smoothly as well as Bryston’s on the Park in Copper Cliff who provided a first-class banquet dinner, as well as lunches and snacks during the technical sessions. Michael Easton (OGS) and Wouter Bleeker (GSC) Co-Chairs, 68th Institute on Lake Superior Geology xxxiii �Proceedings of the 69th ILSG Annual Meeting – Part 1 TECHNICAL PROGRAM xxxiv �Proceedings of the 69th ILSG Annual Meeting – Part 1 TECHNICAL PROGRAM SUNDAY APRIL 23, 2023 All field trips begin and end at The Lismore Hotel 8:00 am - 5:00 pm PRE-MEETING FIELD TRIPS 1) Precambrian geology of the Chippewa River Valley Rob Lodge and Bob Hooper – UW- Eau Claire 2) Wisconsin’s Paleozoic stratigraphy and tour of Crystal Cave Carsyn Ames – Wisconsin Geological and Natural History Survey 4:00 pm - 10:00 pm Registration (Wilson Hall Lobby) 7:00 pm - 10:00 pm Welcoming Reception (Wilson Hall A/F) MONDAY APRIL 24, 2023 7:30 am – 11:30 am Registration (Wilson Hall Lobby) 8:00 OPENING REMARKS (Wilson B) Rob Lodge and Carsyn Ames, Co-Chairs, 2023 ILSG TECHNICAL SESSION I Session Chair: James DeGraff- Michigan Technological University * denotes a student eligible for Best Student Paper Award. To be eligible students must have graduated no more than one month before the ILSG meeting, be first author, and present the paper at the meeting. + denotes author that will present abstract, if different than the first author. 8:10 William J. Hinze, and +William Cannon 2023 Pioneer of Lake Superior Geology: Thomas Benton Brooks 8:30 Erika Vye and William Rose Geoheritage as an educational tool to explore relationships with land and water in the Keweenaw 8:50 William Rose New work developing Keweenaw geoheritage awareness 9:10 Matt Carter and Donald Elsenheimer Workshop Outcomes and Updates for the Minnesota Department of Natural Resource’s Drill Core Library 9:30 Dean Peterson On the Importance of Geologic Maps for Mineral Exploration 9:50 9:50 END OF TECHNICAL SESSION I COFFEE BREAK xxxv �Proceedings of the 69th ILSG Annual Meeting – Part 1 TECHNICAL SESSION II Session Chair: Ben Drenth- USGS and Amy Radakovich Block- Minnesota Geological Survey 10:00 Dana Peterson, Paul Bedrosian, and Carol Finn Subsurface characterization of the Duluth Complex and related intrusions from 3D modeling of gravity and magnetotelluric data 10:20 Paul Bedrosian, Tien Grauch, Laurel Woodruff, William Cannon, Benjamin Drenth, Esther Stewart, Dana Peterson, and James Jones Interpreted geophysical cross-sections through the Lake Superior region: Investigating three billion years of geologic history in sixteen lines of data 10:40 Tien Grauch, Sam Heller, Esther Stewart, and Laurel Woodruff Exploring the geology of the Midcontinent Rift under western Lake Superior using a preliminary velocity model of seismic line GLIMPCE C 11:00 Jennifer Smith, Victoria Tschirhart, Loughlin Tuck, Randy Enkin, and David Roy-Guay Exploring the application of full tensor magnetic gradiometry to better define conduit type NiCu-PGE targets 11:20 END OF TECHNICAL SESSION II 11:20-1:00 LUNCH BREAK and LSG BOARD OF DIRECTORS MEETING - lunches not provided to conference attendees- 11:20-1:00 Student Career Panel- (L.E. Phillips Memorial Public Library- 400 Eau Claire St. in the Riverview Room (Room 306)) TECHNICAL SESSION III Session Chair: Marcia Bjørnerud- Lawrence University 1:10 Wouter Bleeker, Jennifer Smith, Michael Hamilton, Sandra Kamo, Pete Hollings, Michael Easton, and Robert Cundari The Midcontinent Rift System: Neither triple junction nor failed rift? 1:30 Matthew Brzozowski, +Pete Hollings, Jing-jing Zhu, and Robert Creaser Contributions of diverse mantle sources during the early stages of Midcontinent Rift formation — Implications for a passive rifting model 1:50 *Daniel 2:10 *Katherine Langfield, 2:30 END OF TECHNICAL SESSION III Lizzadro-McPherson, James DeGraff and Ian Gannon Structural analysis and slip kinematics of the Keweenaw fault system between Bête Grise Bay and Gratiot Lake, Keweenaw County, Michigan James DeGraff, and Nolan Gamet Slip Kinematics of the Keweenaw and Hancock Faults within the Midcontinent Rift System, Upper Peninsula of Michigan xxxvi �Proceedings of the 69th ILSG Annual Meeting – Part 1 2:30 COFFEE BREAK TECHNICAL SESSION IV Session Chair: Pete Hollings- Lakehead University and 2023 Goldich Medalist 2:50 *Tianna 3:10 *Sam Ghantous, 3:30 *Blaize Briggs and Mary Louise Hill Quetico-Wabigoon Subprovince Boundary in the Superior Province north of Thunder Bay, Ontario, Canada 3:50 Margaret Upton, Phillip Larson, Allan MacTavish, and Peter Hinz Summary of the 2022 ILSG Field Trip to Iceland 4:10 END OF TECHNICAL SESSION IV 4:10 POSTER VIEWING - AUTHORS WILL BE PRESENT AT THEIR POSTERS 6:00 RECEPTION AND CASH BAR (Wilson Hall A/F) 7:00 Groeneveld, Peter Hollings, Wyatt Bain, and Lionnel Djon Petrography, geochemistry, and mineralization of the Archean Titan (Roaring River) intrusion, Northwestern Ontario Noah Phillips, Alex Lusk, Julie Newman, and Shaocheng Ji Are serpentine fault mirrors an indicator of seismic slip? A microstructural analysis ANNUAL BANQUET AND AWARDS (Wilson Hall A/F) SPEAKER: Curt Meine- Adjunct Professor at UW- Madison and Senior Fellow with the Aldo Leopold Foundation and Center for Humans and Nature IMAGINING “CONSERVATION GEOLOGY”: LESSONS FROM THE DRIFTLESS AREA TUESDAY APRIL 25, 2023 8:00 INTRODUCTORY REMARKS AND UPDATES (Wilson Hall B) Rob Lodge and Carsyn Ames, Co-Chairs, 2023 ILSG TECHNICAL SESSION V Session Chair: Allan MacTavish- Consulting Geologist and 2021 Goldich Medalist 8:10 *Justin Jonsson, Pete Hollings, Matthew Brzozowski, Wyatt Bain, and Lionnel Djon Petrogenesis of the mineralized horizons in the Offset and Creek zones, Lac des Iles Complex, N. Ontario 8:30 Pete Hollings, Jacob Hanley, Mark Smyk, Larry Heaman, and Brian Cousens Copper-rich melt inclusions from the St. Ignace Island Complex: Implications for magma mixing and mineralization xxxvii �Proceedings of the 69th ILSG Annual Meeting – Part 1 8:50 Alex Steiner, Dean Peterson, and Gabriel Sweet Magma Recharge and the distribution of Copper and Nickel in the Keweenaw Large Igneous Province 9:10 David Good Identifying regional exploration domains for Ni-Cu-PGE deposit types in the Midcontinent Rift 9:30 Julia Steenburg and Anthony Runkel Record of an Ancient Meteorite Impact Buried Beneath the Twin Cities, MN 9:50 COFFEE BREAK 10:10 Benjamin Drenth, Amy Radakovich Block, George Hudak, Kate Souders, and Stacy Saari Geophysical architecture of the Neoarchean Mentor anorthosite intrusive complex, northwestern Minnesota 10:30 Paul Weiblen The Use of Electric Pulse Disaggregation Technology to Recover Nickel Metal from Nickel Sulfide Ore Deposits 10:50 END OF TECHNICAL SESSION V 11:00 ADDITIONAL POSTER VIEWING – AUTHORS ARE ENCOURAGED TO BE AT THEIR POSTERS (Wilson C & D) 11:30-12:30 LUNCH BREAK - lunches not provided to conference attendees- TECHNICAL SESSION VI Session Chair: Laurel Woodruff- USGS and 2014 Goldich Medalist 12:40 *Margaret Upton, Howard Mooers, and Philip Larson Alteration Geochemistry Characterization and 3D Modeling of the Back Forty Volcanogenic Massive Sulfide (VMS) Deposit Stephenson, Upper Peninsula of Michigan, USA 1:00 Robert Lodge Re-evaluating the tectonics and metallogeny of terranes in the Paleoproterozoic Penokean Orogen, Wisconsin 1:20 William Cannon and Benjamin Drenth Eastward transition from banded iron-formation to ferruginous clastic rocks across the central Upper Peninsula of Michigan 1:40 Jamey Jones, William Cannon, Benjamin Drenth, and Paul O’Sullivan Provenance patterns and tectonic styles of ca. 2.3–1.8 Ga metasedimentary strata in northern Michigan based on regional mapping and detrital zircon U-Pb geochronology xxxviii �Proceedings of the 69th ILSG Annual Meeting – Part 1 2:00 *Audray Hinkemeyer, Howard Moores, and Phillip Larson Determining Provenance of Rainy Lobe Till using Geochemistry and Detrital Zircon Geochronology. 2:20 COFFEE BREAK 2:40 Gordon Medaris Jr. and Steven Driese Secular Changes in the Magnitude of Terrestrial Weathering 2:40 END OF TECHNICAL SESSION VI TECHNICAL SESSION VII Session Chair: Carsyn Ames- Wisconsin Geological and Natural History Survey 3:00 Caroline Rose Tips from a GIS Specialist: Moving maps to GeMS, and a utility for georeferencing quadrangles 3:20 Matthew Rehwald, Carsyn Ames, Sarah Bremmer, William Fitzpatrick, Eric Stewart, Bill Batten, and Stephen Mauel Mobile geologic mapping at the Wisconsin Geological and Natural History Survey 3:40 Roger Schulz Outcrop Scale Mapping Utilizing High-Accuracy GNSS with MnDOT’s Virtual Reference Station (VRS) Network: Minnesota Examples 4:00 Stephen Mauel, Eric Stewart, Matthew Rehwald, Esther Stewart, Carsyn Ames, Sarah Bremmer, and William Fitzpatrick 3D geologic mapping at the Wisconsin Geological and Natural History Survey 4:20 END OF TECHNICAL SESSION VII 4:20 BEST STUDENT PAPER AWARDS STUDENT TRAVEL AWARDS CLOSING REMARKS 4:40 END OF TECHNICAL SESSIONS WEDNESDAY APRIL 26, 2023 8:00am – 5:00pm POST-MEETING FIELD TRIPS Field trips begin and end at The Lismore Hotel 3) Precambrian Geology of the Eau Claire River Valley Rob Lodge and Evan Weber– UW- Eau Claire 4) Quaternary Geology and Geomorphology of the Eau Claire Region Doug Faulkner – UW- Eau Claire J. Elmo Rawling III– Wisconsin Geological and Natural History Survey xxxix �Proceedings of the 69th ILSG Annual Meeting – Part 1 POSTER PRESENTATIONS * denotes a student eligible for Best Student Paper Award. To be eligible students must have graduated no more than one month before the ILSG meeting, be first author, and present the paper at the meeting. *Zsuzsanna P. Allerton, Anita Hall, Françoise Roger, and Christian Teyssier Geochronology and Geochemical Analysis of the Giants Range Batholith in Northern Minnesota Carsyn Ames The Wisconsin Geological and Natural History Survey’s (WGNHS) 2020 and 2021 National Geological and Geophysical Data Preservation Program (NGGDPP) Projects *Ryan Barkley, Noah Phillips, and Pete Hollings The geologic setting, structural controls, and geochemical signature of the Eagle River Au deposit in Northwestern Ontario Marcia Bjørnerud, Buchholz, T., Falster, A.U, And Simmons, W.B. Deformation, metamorphism, fluid flow and pegmatite emplacement history of the post-1630 Ma Waterloo Quartzite of southern Wisconsin Amy Radakovich Block, Kate Souders, Benjamin Drenth, George Hudak, Stacy Saari, and Aaron Hirsch New geologic mapping in the Superior Province of northwestern Minnesota, USA: Pennington and Red Lake Counties *Itai Bojdak-Yates, Marcia Bjørnerud, David Malone, and Esther Stewart A revised provenance model for the Elk Mound Group in south-central Wisconsin based on detrital zircon analysis James DeGraff and William Rose Digital Image Capture and Database Compilation of Historic Mining Data from the Keweenaw Copper District, Michigan: A Progress Update Benjamin Drenth and William Cannon Geophysical mapping of the Great Lakes Tectonic Zone and surrounding Precambrian geology in the central Upper Peninsula, Michigan William Fitzpatrick and Eric Stewart Multiple overlapping features spatially associated with lead-zinc-copper mineralization in the Highland quadrangles, southwest Wisconsin, USA *Lillian Glodowski and Robert Lodge Characterizing volcanic host stratigraphy and syn-volcanic intrusions at the Lynne Zn-Pb-Cu deposit, Oneida Co., Wisconsin xl �Proceedings of the 69th ILSG Annual Meeting – Part 1 *Kaine Johnson and Robert Lodge Hydrothermal Alteration Facies of the Eisenbrey Zn-Cu Deposit, Rusk County, Wisconsin *Matthew Leahy and Robert Lodge Petrology and Geochemistry of the Paleoproterozoic Eau Claire Volcanic Complex, Eau Claire, WI *Francisca Nuñez-Ferreira, Lucas Zoet, and J. Elmo Rawling III Morphometry and formation process of eskers developed under the Chippewa Lobe of the Laurentide Ice Sheet *Jordan Peterzon, Noah Phillips, Peter Hollings and Lionnel Djon Fault zone architecture in mafic protoliths at the Lac des Iles mine, northwestern Ontario Caroline Rose, J. Elmo Rawling III, Eric Carson, John Attig, David Mickelson, William Mode, Mark Johnson, and Kent Syverson Quaternary Geology of Wisconsin at a scale of 1:500,000 (in review) Allison Severson, Eric Nowariak, and Phillip Larson Geology and geochemistry of the basal North Shore Volcanic Group and Midcontinent Rift Intrusive Supersuite, Cook County, MN, USA Eric Stewart, William Fitzpatrick, and Carsyn Ames Relay zones in weakly folded and faulted Paleozoic strata and their role localizing Mississippi Valley-type mineralization, southwest Wisconsin, USA *Madeline Taylor and Marcia Bjørnerud Deciphering the metamorphic and deformational history of the Hardwood Gneiss, Felch District, Michigan: Anomalously high-pressure rocks in the heart of the Penokean orogen *Evan Weber, Robert Lodge, and Jeffrey Marsh U/Pb geochronology and zircon petrochronology of Paleoproterozoic magmas from the Marshfield terrane Penokean Orogen, Wisconsin xli �Proceedings of the 69th ILSG Annual Meeting – Part 1 BANQUET PRESENTATION IMAGINING “CONSERVATION GEOLOGY”: LESSONS FROM THE DRIFTLESS AREA Curt Meine Adjunct Professor at UW- Madison and Senior Fellow with the Aldo Leopold Foundation and Center for Humans and Nature The field of conservation biology emerged in the 1980s when scientists became increasingly alarmed about the loss of biodiversity, and decided that they had a responsibility to put their science to work to address the issue. This required not only new interdisciplinary research, but new ways to put knowledge to work in our human and natural communities. Can we imagine a field of conservation geology that similarly seeks to integrate geological knowledge with history and culture, and addresses our concerns for our landscapes and for future generations? The Driftless Area provides ample examples and opportunities to explore that question. xlii �Proceedings of the 69th ILSG Annual Meeting – Part 1 ABSTRACTS 1 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Geochronology and Geochemical Analysis of the Giants Range Batholith in Northern Minnesota ALLERTON, Zsuzsanna1, HALL, Anita1, ROGER, Françoise2, and TEYSSIER, Christian1 1 Earth and Environmental Science Department, University of Minnesota, 150 Tate Hall, 116 Church St. SE, Minneapolis, MN 55455 2 Géosciences Montpellier, Université de Montpellier-Campus Triolet, c.c. 60 Place Eugéne Batallion 34090, Montpellier, Cedex 05, France The Giants Range Batholith (GRB) is a ~ 2.7-billion-year-old (2.7 Ga) granitic unit in northern Minnesota striking SW-NE from east of Ely to Grand Rapids (Figure 1). It is located N-NW of the 1.8 Ga Mesabi Iron Range and the 1.1 Ga Duluth Igneous Complex (DC). During emplacement, the GRB was situated at the southern edge of the Superior Craton, the Archean core of the North American Continent. At its eastern end the GRB is in contact with the Mesoproterozoic DC (1.1 Ga), which is the intrusive segment of the Mid-continent Rift System, and to the west the GRB flanks the Lower Member of the Ely Greenstone Formation. This project has two main goals: (1) better understanding the origin of the GRB; and (2) using the GRB to track the thermal and hydrothermal history of the rocks from the contact with the DC outward. The project included the compilation of existing data, such as geochronology and geochemistry, that have been collected to date on the GRB, based Figure 1. Simplified geologic map of Minnesota's arrowhead on Allison (1925), Griffin & Morey (1969), region showing the Giants Range Batholith in blue. Prior Viswanathan (1971), Boerboom & Zartman studies have been done in the area circled in red. The white dashed box shows the current and proposed area of this (1993), Boerboom (1994), and Southwick project. Modified from Jirsa, M.A., Miller jr., J.D., & Morey, (1994). The Minnesota Geological Survey G.B. (2008). (Jirsa, 2016) acquired U-Pb zircon age dates on selected samples. Only Boerboom & Zartman (1993) and Boerboom (1994) have completed trace element analysis. Their eight samples are from the central section of the GRB (red circle in Figure 1) and were collected along the northern margin. The samples from the GRB main body have not been analyzed for trace elements, and geochronological data are scarce. Our sampling campaign so far has concentrated on the northeastern part of the GRB (white dashed box in Figure 1) and builds on the work of Boerboom & Zartman (1993) and Boerboom (1994). Thin sections were cut and used in transmitted-light petrography to determine mineralogical composition, analyze textures, and identify accessory minerals for radiometric dating. Radiometric dating involved Laser-Ablation Inductively Coupled Plasma Mass-Spectroscopy (LA-CPMS) performed at the University of Clermont-Ferrand, France. We obtained zircon and titanite age dates for samples located within 1000 meters from the DC contact. The zircon grain separates from one sample produced a concordant age date of ~ 2690 ± 10 Ma. In-situ zircon analysis of another sample displays some discordia (Pb loss) that may be associated with hydrothermal alteration related to DC emplacement. The mounted titanite grains and in-situ analysis yielded ages (approx. 2450-2500 2 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Ma) that are consistently younger than zircons from the same samples. Current and future work include further sample collection in the study area (white dashed box in Figure 1) to obtain additional U-Pb dates on titanite and zircon to the thermal and hydrothermal history of GRB at the contact with the DC. The GRB samples also contain abundant apatite grains, some primary and some recrystallized, that will be dated using the U-Pb method to provide new data on the thermal and hydrothermal history of the GRB near the DC contact. Additionally, we will pursue acquiring bulk composition and trace element data in order to better understand the source of magma and the likely tectonic setting in which the GRB was emplaced. References Allison, I.S., 1925. The Giants Range Batholith of Minnesota. The Journal of Geology, 33(5): 488-508. https://doi.org/10.1086/623215. Boerboom, T.J. and Zartman. R.E., 1993. Geology, Geochemistry, and Geochronology of the Central Giants Range Batholith, Northeastern Minnesota. Canadian Journal of Earth Sciences, 30(12): 25102522. https://doi.org/10.1139/e93-217. Boerboom, T., 1994. Short Contributions to the Geology of Minnesota: Alkalic Plutons of Northeastern Minnesota; Report of Investigations 43. Minnesota Geological Survey, ISSN 0076-9177. Frost, B.R. and Frost, C.D., 2008. A geochemical classification for feldspathic igneous rocks. Journal of Petrology 49.11. Griffin, W.L. and Morey, G.B., 1969. Geology of the Isaac Lake Quadrangle, St. Louis County, Minnesota. Published in Cooperation with Minnesota Department of Iron Range Resources and Rehabilitation. Minnesota Geological Survey 5 P-8 Special Publication Series. University of Minnesota. Southwick, D.L., 1994. Short Contributions to the Geology of Minnesota: Assorted Geochronologic Studies of Precambrian Terranes in Minnesota: A Potpourri of Timely Information. Report of Investigations 43. Minnesota Geological Survey, ISSN 0076-9177. 3 �Proceedings of the 69th ILSG Annual Meeting – Part 1 The Wisconsin Geological and Natural History Survey’s (WGNHS)’s 2020 and 2021 National Geological and Geophysical Data Preservation Program (NGGDPP) Projects AMES, Carsyn1, GOTTSCHALK, Brad1, ROSE, Caroline1, SIBLEY, Dave1 1 Wisconsin Geological and Natural History Survey, University of Wisconsin- Madison, 3817 Mineral Point Rd. Madison, WI 53705 The Wisconsin Geological and Natural History Survey (WGNHS) received grants from the United States Geological Survey (USGS)’s National Geologic and Geophysical Data Preservation Program (NGGDPP) for FY2020 and FY2021. This program promotes the preservation and public accessibility of geoscience collections and data. Projects completed during the 2020 grant were 1) to preserve 130 boxes of hand samples, and 2) to convert 10 WGNHS maps to the standard Geologic Map Schema (GeMS) format. The majority of hand samples for this project came from a donation by Gene LaBerge (UW-Oshkosh) who worked extensively in and around Marathon County, and whose work resulted in a Marathon County bedrock map (LaBerge and Myers, 1983) published by WGNHS. The collection includes more than 1500 specimens from 583 separate outcrops. Successfully preserving these samples is of importance as Marathon Co. continues to urbanize and many of the outcrops these samples represent are being demolished due to land development. The 10 maps converted to GeMS format during the project include Pleistocene maps from northwestern Wisconsin and bedrock maps from southern and northeastern Wisconsin. Converting legacy maps to GeMS format is important because the digital use of WGNHS maps allows for wider and broader use by both internal and external stakeholders. Additionally, a survey of WGNHS external partners showed that a majority prefer digital versions of maps and data. Projects for the 2021 grant included 1) expanding the WGNHS data viewer’s capacity to deliver photos of bedrock cores, 2) digitizing borehole data from the Mineral Development Atlas (MDA)- a joint project between the USGS, United States Bureau of Mines (USBM), and state surveys of Wisconsin, Iowa, and Illinois- that gathered information related to metallic mineral exploration and mining in the lead-zinc district, and 3) photograph, log, and permanently archive seven cores from the Lynne Deposit, a volcangenic massive sulfide deposit in Oneida County. WGNHS’s data viewer, created in 2018, saw its capacity expanded to include photos of cores in their collection. The 2021 project used almost 300 donated Wisconsin Department of Transportation (WisDOT) cores to test pilot this new ability and results are available here: https://data.wgnhs.wisc.edu/data-viewer/. An additional part of this project included the correlation of 3300 scanned logs to the boreholes and updating location data for logs and cores. The MDA portion of the 2021 project focused efforts on mine workings in Lafayette Co., WI. Staff at the Survey geolocated more than 17,000 boreholes and corrected polygons for surface workings such as quarries, prospecting sites, and lead diggings. Lastly, WGNHS permanently archived seven cores from the Lynne Deposit, Oneida Co., WI in 2021. These cores were transferred to WGNHS’s samples repository from the University of Wisconsin- Eau Claire where they had been stored temporarily for student study. The cores (totaling approximately 2500 ft) were then logged and photographed by UW-Eau Claire students. These photos were also added to the WGNHS data viewer. References LaBerge, G., and Myers, P., 1983. Precambrian Geology of Marathon County, Wisconsin. Wisconsin Geological and Natural History Survey IC45: 1-88. https://wgnhs.wisc.edu/catalog/publication/000295/resource/ic45. 4 �Proceedings of the 69th ILSG Annual Meeting – Part 1 The geologic setting, structural controls, and geochemical signature of the Eagle River Au deposit in Northwestern Ontario BARKLEY, Ryan1, PHILLIPS, Noah1, HOLLINGS, Pete1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The Eagle River orogenic gold deposit is hosted in the Mishibishu greenstone belt of the western Superior craton, approximately 50 km west of Wawa, Ontario. The deposit, an active underground mine, has been in continuous production since 1995 and produced 1.485 (Moz) of Au through to the end of 2021 (SRK, 2022). The average grade is 9.7 g/pt and Au is primarily hosted in highly strained, milky white to grey quartz veins that dip to the north and strike east to west. The shear zones are hosted in an elliptical quartz diorite pluton, extending into iron rich mafic volcanic rocks. The Mishibishu greenstone belt is dominated by granitic plutons, mafic to felsic volcanics, and lesser amounts of metasedimentary packages. U-Pb zircon dates in the belt range from 2.6 to 2.8 Ga, indicating a Neoarchean environment (Keller, 1989). To understand the geological setting, structural controls, and the geochemical signature of the Eagle River deposit, we completed detailed structural field mapping, petrography, and whole rock geochemistry analysis of the rocks in and around the deposit. A total of 41 whole rock geochemistry samples were collected from the area north of the mine. Two suites were identified; suite one consists of calc-alkaline basalt, andesite, dacite, rhyolite, diorite, tonalite and granite. This suite is characterized by enriched La/Smn ratios of 2.06 to 6.83 and negative Nb anomalies (Nb/Nb* of 0.09 to 0.43), consistent with magmas formed in a supra-subduction environment. Suite two consists of tholeiitic basalt, andesitic-basalt and gabbro. This suite is characterized by flatter trace element patterns with La/ Smn ratios of 0.83 to 1.48 and minor Nb anomalies (Nb/Nb*of 0.40 to 0.85), consistent with primitive arc tholeiites (Fig. 1). Figure 1. Primitive mantle normalised diagrams for the calc-alkaline (blue) vs tholeiitic (red) rocks of the study area. Normalising values from Sun and McDonough (1989). Shear zones for this study appear to be ductile. The white to grey, boudinaged quartz veins are highly strained and flattened, indicating a ductile environment. Gold accumulates in areas of high strain. Quartz veins exhibit chessboard extinction patterns and lobate grain boundaries indicating that the veins have recrystallized through grain boundary migration dynamic recrystallization (Fig. 2; Stipp et al, 2002). Deformation therefore occurred at high temperatures in a low stress environment. 5 �Proceedings of the 69th ILSG Annual Meeting – Part 1 0.5 cm Figure 2. Recrystallization of quartz veins via grain boundary migration. References Keller, J., 1989. The evolution of the Mishibishu greenstone belt, near Wawa, Ontario. Electronic Theses and Dissertations. Stipp, M., Holger & Heilbronner, R., & Schmid, S., 2002. Dynamic recrystallization of quartz: Correlation between natural and experimental conditions. Geological Society London Special Publications. 200: 171-190. 10.1144/GSL.SP.2001.200.01.1. Sun, S.S., and McDonough, W.F., 1989. Chemical and Isotopic Systematics of Oceanic Basalts: Implications for Mantle Composition and Processes. In: Saunders, A.D., Norry, M.J., Eds., Magmatism in the Ocean Basins, Geological Society, London, Special Publications, 42: 313-345. S.R.K Consulting, 2022. 43-101 Eagle River Mine, Ontario, Canada, Wesdome Gold: 262. 6 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Interpreted geophysical cross-sections through the Lake Superior region: Investigating three billion years of geologic history in sixteen lines of data BEDROSIAN, Paul A.1, GRAUCH, V.J.S.1, WOODRUFF, Laurel G.2, CANNON, William F.3, DRENTH, Benjamin J. 1, STEWART, Esther K. 4, PETERSON, Dana E.1 and JONES, James V.5 1 U.S. Geological Survey, Building 20, MS 964, Denver Federal Center, Denver, CO 80225 U.S. Geological Survey, 2280 Woodale Drive, Mounds View, MN, 55112 3 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA 20192 4 Wisconsin Geological & Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 5 U.S. Geological Survey, 4210 University Drive, Anchorage, AK 990508 2 The northern midcontinent is a window into an Archean-Proterozoic continent, and the Mesoproterozoic Midcontinent Rift System (MRS) that nearly tore it apart. This complex tectonic collage has been largely unmodified during the last billion years yet is poorly exposed except in the Lake Superior region. The area is rich in mineral resources, including native and sedimentary copper deposits, iron formations, volcanogenic massive sulfide deposits, and nickel-copper-platinum-group element sulfide mineralization. In 2016, 2,710 line-km of airborne electromagnetic (AEM) and magnetic data were collected along sixteen regional transects spanning parts of three states and more than three billion years of geologic time (Bedrosian, 2019). The transects range from 100 to 300 km in length and cross parts of the Wisconsin Magmatic Terrane, the Penokean fold and thrust belt, the MRS, and the Archean Superior Province (Figure 1). Data modeling was challenging due to poor control on system height and the prevalence of induced polarization effects (Bedrosian et al., 2018). The final electrical resistivity models derived from the AEM data have been translated into interpreted geophysical cross-sections though an iterative, consensus building approach. A team was assembled with varied expertise in the geology, geophysics, and mineral resources of the MRS, Penokean, and Archean assemblages within the region. Over a period of two years, a series of interpretation sessions worked line-by-line through the transects, culminating in a workshop to synthesize and finalize interpretations. Geologic maps, potential-field data, and drill hole logs were examined alongside the AEM resistivity models and incorporated into the resulting interpretations. Constraints from seismic reflection and refraction studies, magnetotelluric models, geochronology, and detrital zircon studies were also considered where available. The resulting annotated geophysical cross-sections are a resource to be drawn and built upon for geologic and tectonic investigations. Some aspects these sections touch upon include: • Internal structure of the Animikie basin and the basal contact of the Duluth Complex • Geometry and deformation of MRS-flanking sedimentary basins • Structure of the MRS Ashland syncline • Geometry and extent of post-magmatic MRS clastics (Oronto and Bayfield Groups) • Geometry, internal variability and provenance of the Jacobsville Sandstone • Geometry and extent of Archean, Penokean, and MRS faults • Extent and dismemberment of Penokean-deformed metasedimentary units • Iron formations and Penokean structures along the early Proterozoic gneiss dome corridor • Patterns of reverse polarity dikes • Phanerozoic cover and underlying structure (e.g., eastern arm of the MRS) • Distribution, thickness, and variability in glacial cover 7 �Proceedings of the 69th ILSG Annual Meeting – Part 1 New insights and refinements are many but include (a) a restricted areal extent for Bayfield Group clastic rocks, (b) multiple distinct subunits within the Jacobsville sandstone, (c) a close stratigraphic relation between Penokean iron formations and conductive sulfide-rich metasediments, and (d) complex deformation and alteration of the main bowl Animikie basin. Figure 1. Location of AEM and magnetic profiles (magenta). Background geology is from a USGS MRS GIS compilation from published sources of the region. References Bedrosian, P.A., 2018. Geologic mapping and tectonic structure of the U.S. midcontinent via reconnaissance AEM, 7th Intl. Wksp on Airborne Electromagnetics, Kolding, Denmark: 4. Bedrosian, P., 2019. Multi-scale AEM and MT mapping of the Precambrian in Upper Michigan, Northern Wisconsin, and Eastern Minnesota, in Puumala, M., (ed.), Institute on Lake Superior Geology Proceedings, 51st Annual Meeting, Nipigon, Ontario, Part 1 - Abstracts and Proceedings. v.65, Part 1: 67. 8 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Deformation, metamorphism, fluid flow and pegmatite emplacement history of the post-1630 Ma Waterloo Quartzite of southern Wisconsin BJØRNERUD, M.1, BUCHHOLZ, T. 2, FALSTER, A. U.3, and SIMMONS, W. B.3 1 Geosciences Department, Lawrence University, Appleton Wisconsin 54911 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494 3 Maine Mineral & Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217 2 The Waterloo Quartzite, one of the upper Paleoproterozoic ‘Baraboo Interval’ quartzites of the southern Great Lakes region (Medaris et al., 2003), experienced a more complex structural history and higher-grade metamorphism (amphibolite facies) than any of the other quartzite units in this group. It is also distinctive in being intruded by bodies of granitic pegmatite. Natural outcrops of the Waterloo quartzite are limited, but a major quarry near the town of Waterloo (43.210 N, 88.450 W) provides three-dimensional exposures and access to a large volume of fragmented rock. This study is based on observations and samples taken at the quarry over several years as it was deepened and enlarged by blasting. The youngest detrital zircons in the Waterloo Quartzite date to 1634 Ma, younger than the 1710 Ma maximum depositional age of the Baraboo Quartzite, and an indication that sediment transport directions changed from southward to northward (modern coordinates) between the times of deposition of the Baraboo and Waterloo units (Schwartz et al., 2018). The protolith of the Waterloo quartzite was primarily pure quartz sandstone but also included pelites and quartz pebble conglomerates with clasts of jasper (Stewart, 2021). The earliest deformational feature in the Waterloo quartzite is a penetrative foliation (S1) defined by aligned grains of sub-mm muscovite in the pelitic layers; this muscovite has yielded an 40 Ar/39Ar age of 1452 +/- 7 Ma and has been interpreted as evidence of a pervasive fluid flow event that introduced potassium into the supermature sediments, in which K was originally absent (Medaris et al., 2003). In the quarry, the S1 foliation is nearly parallel to bedding; both surfaces strike toward the northeast (045° to 055°) and dip moderately (35°-55°) southeast, suggesting that the rocks lie on the SE limb of a tight NW-verging anticline. In places, mm- to cm-scale quartz veins with Ti-rich hematite masses on their margins lie parallel to the foliation and have fibers oriented perpendicular to the foliation. This points to another episode of fluid infiltration under a stress regime distinct from the one that formed the foliation. These early quartz veins are commonly folded and/or boudinaged. The S1 foliation is overprinted by porphyroblasts of andalusite, typically about 0.5 cm in size. Most of these have been altered to muscovite and/or kaolinite, indicating another episode of fluid infiltration. The kaolinite occurs mainly on the margins of the andalusite crystals, giving them a zoned appearance. In many specimens, the retrograded andalusites have a rusty red color that may be related to the presence of hematite in the kaolinitic rims (Geiger et al., 1982). The next structural feature to develop in these rocks are kink-like crenulations in pelitic horizons, seen abundantly in blocks in the quarry waste piles. At two sites where this crenulation cleavage (S2) was observed in place, it strikes N to NNW and dips steeply east. The geometry of the crenulations is strongly influenced by the presence of the andalusite porphyroblasts/ pseudomorphs, many of which have small, asymmetric pressure shadows of quartz and muscovite that seem to be related to the development of the crenulations. Sometime after the formation of the crenulation cleavage, the quartzite was intruded by Kfeldspar-dominated pegmatite dikes ranging in width from 1 cm to 3 m. Pegmatites have been known from the NW portion of the Waterloo Quarry for some years and have been discussed by 9 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Buchholz et al. (2016). The pegmatite dikes have sharp boundaries and granitic textures, with Kfeldspar, quartz and muscovite crystals of equal and uniform size. Blasting of a phyllitic horizon in the NE part of the quarry has recently exposed additional thin (< 5 cm wide) pegmatite dikes, with mmscale chilled margins at contacts with the host rock. Heavy mineral separates from these thin dikes contain fluorapatite, monazite-(Ce), ilmenite, columbite-(Mn), tantalite-(Mn), evidence of significant enrichment of Ta/Nb and Mn/Fe. In addition to the pegmatite dikes, coarse-grained pegmatite-like patches occur in the necks of boudinaged quartz veins and quartzose layers enclosed by phyllites, primarily in the NE corner of the quarry. Unlike the clearly igneous dikes, these patches have irregular boundaries with the host rock and their crystal size is variable. In thin section, K-feldspar and quartz in these patches show a micrographic texture. Muscovite in pelitic layers surrounding the boudins is coarser than in the rest of the rock, and rusty andalusite pseudomorphs are smeared and flattened in the vicinity of the boudins, suggesting that they had already been altered and softened by the time the boudins formed. Although they occur in the same area of the quarry, the pegmatite-like boudin patches do not seem to be physically connected to the pegmatite dikes. The patches are presumably older since they formed during the process of boudinage, while the dikes apparently postdate deformation. The pegmatite-like material in the boudin necks could either be hydro- thermal or produced by in situ melting related to influx of fluids or perhaps a local drop in pressure (mean stress) during boudinage. Examination of heavy mineral separates from the pegmatite-like bodies associated with boudins revealed fluorapatite and monazite-(Ce). One specimen of the boudin material contains small beryl crystals in a pocket-like void. This may be similar to beryl occurrences in regionally metamorphosed rocks in Austria (Franz et al, 1986), believed to have formed between 500- 550⁰C -- slightly higher than maximum metamorphic temperature estimates of 500⁰C for the Waterloo rocks (Medaris et al., 2003). Small crystals of greenish to light brown dravitic tourmaline are also present locally; analysis shows that these are Li-bearing, as are nearby muscovites. Additional mineral phases include chloritoid, spessartine garnet, fluorapatite, and gahnite, all pointing to the introduction of fluids with a rich mix of ions. Although the Waterloo quarry lies only 20 km in the across-strike direction from the south limb of the Baraboo syncline, it is not easy to correlate either the chronology or the orientations of structures at Waterloo with those in the more famous Baraboo Quartzite. Our observations from the Waterloo quarry suggest that the 1470-1450 Ma “Baraboo Orogeny” (Medaris et al., 2021) was a complex, multistage tectonic event whose details have not yet been fully documented. References Buchholz, T.W., Falster, A.U. & Simmons, W.B., 2016. Second Foord Pegmatite Symposium: 22-23. Franz, G., Grundman, G., & Ackermand, D, 1986. Tschermaks Min. Pet. Mitteilungen, 15: 167-192. Geiger, C., Guidotti, C. & Petro, 1982. Geoscience Wisconsin 6: 21-40. Medaris, L.G. & others, 2003. Journal of Geology, 111, doi:10.1086/373967 Medaris, L.G. & others, 2021. Geoscience Frontiers, 12, doi: 10.1016/j.gsf.2021.101174 Schwartz, J.J., Stewart, E.K. and Medaris, L.G., Jr., 2018. ILSG Proceedings, 64: 93–94. Stewart, E.K., 2021. Wisconsin Geological & Natural History Survey Map 508. Stewart. E.K., Brengman, L. & Stewart, E.D., 2021. Journal of Geology, 129, doi:10.1086/713687. 10 �Proceedings of the 69th ILSG Annual Meeting – Part 1 The Midcontinent Rift System: Neither triple junction nor failed rift? BLEEKER, Wouter1, SMITH, Jennifer1, HAMILTON, Michael2, KAMO, Sandra2, HOLLINGS, Pete3, EASTON, Michael4, and CUNDARI, Robert5 1 Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8; wouter.bleeker@canada.ca Jack Satterly Geochronology Lab., University of Toronto, 22 Ursula Franklin St., Toronto, ON M5S 3B1 3 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 4 Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5 5 Ontario Geological Survey, 435 James Street South, Thunder Bay, ON P7E 6S7 2 The Midcontinent Rift System has often been described in terms of i) a failed intracontinental rift system; with ii) a basic ‘triple junction’ architecture, the three arms of the triple junction being represented by the SW arm of Lake Superior, the SE arm of Lake Superior, and a less developed rift structure reaching up into the Lake Nipigon area. Here we challenge both views. Although it is certainly true that on a local scale, i.e. the North American midcontinent, the rift system failed and inverted, it is likely that on a more global scale the system did not fail but led to ocean opening at ca. 1103 Ma, i.e. the waning phase of the “Early Magmatic Stage”3 of the Midcontinent Rift. This led to a global reorganization of plate stresses. The majority of robust structural indicators suggest that this early stage rifting, initiated at ca. 1110 Ma and waning towards ca. 1103 Ma, was oriented on an NW-SE axis or trend (present orientation), from Lake Nipigon to the SE arm of Lake Superior and beyond. The significant gradient in rifting and lithospheric stretching, from Lake Nipigon (minor rifting followed by sagging) to the SE rift arm (major rifting), requires that the rotation pole for this early phase of rifting was situated to the northwest, somewhere in northwest Ontario. At larger distances from this rotation pole, up to 90° of arc away(?), to the southeast (present orientation), lithospheric spreading may have reached ~1000 km and thus likely led to ocean opening. This early phase of rifting with its NW-SE axis came to a close with a marked hiatus of ~4-5 Myr (the “Magmatic Hiatus”), represented in most sections by a distinct unconformity of conglomerates and more shallow dipping basalt flows on top of older, more steeply dipping basalt flows. When rifting resumed, after this significant hiatus, it opened up the SW arm of the Midcontinent Rift organized on a SW-NE trending rift system. This phase was accompanied by the “Main Magmatic Stage” and was initiated at 1099-1098 Ma, the emplacement age of the Duluth Complex (e.g., Paces and Miller, 1993). The marked gradient in rifting and lithospheric stretching on this SWNE rift system, with major crustal stretching in the central part of Lake Superior, and less stretching farther to the southwest, indicates that the rotation pole for this younger phase of rifting was situated well to the southwest, perhaps in Texas or on the future western margin of Laurentia. This SW-NE rift system shows marked jogs, and may have stepped over to the south, through the eastern arm of Lake Superior, and continued to the northeast in the area now obscured by final accretion and collision of the Grenville orogen at ca. 1 Ga. As for the first phase of rifting, we observe locally (in North America) only one proximal end (relative to the rotation poles) of the larger rift systems—systems that may well have been global in scale. Clearly the later SW-NE rift system is distinct from the earlier NW- SE rift system, with completely different rotation poles, and rift axes that are essentially perpendicular to each other. Relevant to the early NW- SE rift phase, the extent of the rifting and the location of its rotation pole, are occurrences of carbonatite complexes in northwest Ontario, and large diabase sills at the base 3 We use here the magmatic stage terminology of Miller and Nicholson (2013) but with modified age boundaries. 11 �Proceedings of the 69th ILSG Annual Meeting – Part 1 of the Athabasca Basin (the Moore Lakes sills), the latter with an age exactly equivalent to those of the Nipigon diabase sills (see Bleeker et al., 2020 and references therein). At a larger scale, there are 1108 Ma magmatic provinces on several other continents (e.g., the Umkondo sills in South Africa; Hanson et al., 2004). And relevant to the younger SW-NE rift phase is the major diabase sill magmatism of the SW USA Diabase Province at 1095-1085 Ma (e.g., Bright et al., 2014; Heaman and Grotzinger, 1992). Clearly, we need to zoom out to develop a broader understanding of the Midcontinent Rift System and see it in a more global context. The GSC-funded project to refine our knowledge of this major rift system started with an attempt to better define the ages (both precision and accuracy) of some of the major events and many of the mineralized intrusions. We currently have ~30 U-Pb samples in various stages of progress and some early results were reported in Bleeker et al. (2020) and Smith et al. (2020). Several others will be discussed as part of this presentation. Our initial focus was to resolve many of the problematic age ‘outliers’, the majority of which were based on extrapolations from sparse and discordant data. Most of these outliers are now gone. Based on our current data and review of the published literature, major age divisions may be summarized as follows: Initiation: ca. 1111-1110 Ma, as best defined by the large Echo Lake subvolcanic layered intrusion (a robust zircon age, reported in Cannon and Nicholson, 2001). Early Magmatic Stage: 1110-1103 Ma, with emplacement of regional diabase sill complexes (ca. 1108-1106 Ma) following early rift intrusions and regional plateau basalt building (1110-1107 Ma). The Tamarack intrusion, still organized on a NNW-SSE dyke-like system, is ca. 1104 Ma. Hiatus: 1103-1099 Ma, in many places marked by an angular unconformity. Main Magmatic Stage: 1099-1092 Ma, initiated with emplacement of the Duluth Complex and later characterized by the very extensive flood basalts of the Portage Lake Volcanic Group. Late Magmatic Stage: 1092-1084 Ma, waning volcanism and intercalated rift-fill sediments. Sagging and Rift-Fill Stage: 1084 to ca. 1060 Ma, final rift fill sedimentation, Oronto Group. References Bleeker, W. et al., 2020. The Midcontinent Rift and its mineral systems: Overview and temporal constraints of Ni-Cu-PGE mineralized intrusions. GSC Open File 8722: 7–35. DOI: org/10.4095/326880. Bright, R.M. et al., 2014. U-Pb geochronology of 1.1 Ga diabase in the southwestern United States: Testing models for the origin of a post-Grenville large igneous province. Lithosphere, 6:135–156. Cannon, W.F. and Nicholson, S.W., 2001. Geology map of the Keweenaw Peninsula and adjacent area. U.S. Geological Survey, Geological Investigations Series, Map I-2696, scale 1:100 000. Heaman, L.M. and Grotzinger, J.P., 1992. 1.08 Ga diabase sills in the Pahrump Group, California: Implications for development of the Cordilleran miogeocline. Geology, 20: 637–640. Miller, J.D. and Nicholson, S.W., 2013. Geology and mineral deposits of the 1.1 Ga Midcontinent Rift in the Lake Superior region – An overview. Precambrian Research Center Guidebook 13-1:1–50. Paces, J.B. and Miller, J.D., 1993. Precise U‐Pb ages of Duluth complex and related mafic intrusions, northeastern Minnesota: Geochronological insights to physical, petrogenetic, paleomagnetic, and tectonomagmatic processes associated with the 1.1 Ga midcontinent rift system. Journal of Geophysical Research: Solid Earth, 98: 13 997–14 013. Smith, J.W. et al., 2020. Timing and controls on Ni-Cu-PGE mineralization within the Crystal Lake Intrusion, 1.1 Ga Midcontinent Rift. GSC Open File 8722: 37–63. 12 �Proceedings of the 69th ILSG Annual Meeting – Part 1 New geologic mapping in the Superior Province of northwestern Minnesota, USA: Pennington and Red Lake Counties BLOCK, Amy Radakovich1, SOUDERS, A. Kate2, DRENTH, Benjamin J.3, HUDAK, George J.4, SAARI, Stacy M. 5, HIRSCH, Aaron C.1 1 Minnesota Geological Survey, 2609 Territorial Road, St. Paul, MN 55114 U.S. Geological Survey, PO Box 25046, MS 963, Denver Federal Center, Denver, CO 80225 3 U.S. Geological Survey, PO Box 25046, MS 973, Denver Federal Center, Denver, CO 80225 4 Natural Resources Research Institute, 5013 Miller Trunk Highway, Duluth, MN 55811 5 Minnesota Department of Natural Resources, 1525 3rd Ave E, Hibbing, MN 55746 2 The Earth Mapping Resources Initiative (Earth MRI) is a partnership between the USGS and state geological surveys/science agencies that funds data collection and geologic mapping in order to better characterize areas of potential critical mineral resources. Earth MRI recently funded a highresolution airborne geophysical survey (Allen Langhans and Drenth, 2023; Fig. 1, blue box) and acquisition of new geochronologic (Souders, A.K., in review), petrologic, and geochemical data in a part of the Superior Province in northwestern Minnesota that is prospective for numerous criticalmineral-producing systems. Previous geologic mapping of the area (Jirsa et al., 1999; Jirsa et al., 2011) was limited by an absence of outcrop, limited drill hole data, and only one geochronologic age. Newly acquired data support ongoing bedrock mapping across a large area (Fig. 1, red box); This map highlights the geology of Pennington and Red Lake Counties (Fig. 1, orange box). The map area comprises three conterminous subprovinces of the Archean Superior Province which are situated in unusually close proximity to one another; in the map area, the Quetico metasedimentary province pinches to as little as 5 km of thickness in map view where it separates the Wabigoon and Wawa volcanoplutonic subprovinces on either side. Ages from a biotite tonalite in the Snake River batholith (ca. 2758 Ma), a diorite in the Grygla pluton (ca. 2771 Ma), and a biotite-hornblende tonalite in the Red Lake Falls pluton (ca. 2701 Ma) (Souders, in review) define multiple Neoarchean episodes of intermediate to felsic intrusive activity within the Wabigoon subprovince. Interpretation of the improved aeromagnetic data suggests a revised, more southerly position of the Wabigoon-Quetico subprovince boundary, as well as modifications of numerous other geologic contacts across the map area. Finally, new geochronologic ages confirm an Archean (ca. 2737 Ma) age for the Mentor Anorthosite Intrusive Complex (MAIC) (Souders, in review), and new geophysical interpretations reveal that the MAIC is as much as twice as large and much more structurally complex than previously thought (Drenth et al., this volume). Both findings regarding the MAIC are consistent with what is known of other Archean anorthosites in the Superior Province (Sotirou & Polat, 2020; Polat et al., 2018). Work in the larger Earth MRI mapping area is ongoing. Additional geochronologic data will shed light on the depositional history and timing of mineralization of volcanic strata in both the Wawa and Wabigoon subprovinces. Geochemical analyses will supplement petrographic observations and help refine tectonic provenance of all rock units. A new geologic map of the entire area (Fig. 1, red box) will be completed, and a comprehensive mineral potential model will better assess the potential for critical minerals in the area. 13 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Pennington Red Lake Figure 1. Generalized subprovince map of the Superior Province in northwest Minnesota, USA, showing the location of both the recent geophysical survey (blue outline), ongoing new mapping (red outline) for the EarthMRI project, and the map area for this poster (orange outline). References Allen Langhans, A.D., and Drenth, B.J., 2023. Airborne magnetic and radiometric survey, northwestern Minnesota, 2021: U.S. Geological Survey data release, https://doi.org/10.5066/P97D2JJE. Drenth, et al., this volume. Jirsa, M.A., Chandler, V.W., and Runkel, A.C., 1999. M-092 Bedrock geologic map of northwestern Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/973. Jirsa, M.A., Boerboom, T.J., Chandler, V.W., Mossler, J.H., Runkel, A.C., and Setterholm, D.R., 2011. Geologic map of Minnesota, bedrock geology: Minnesota Geological Survey State Map S-21, scale 1:500,000. Polat, A., Longstaffe, F.J., and Frei, R., 2018. An overview of anorthosite-bearing layered intrusions in the Archaean craton of southern West Greenland and the Superior Province of Canada: implications for Archaean tectonics and the origin of megacrystic plagioclase: GEODINAMICA ACTA, v. VOL. 30, NO. 1: 84–99, https://doi.org/10.1080/09853111.2018.1427408. Sotiriou, P., and Polat, A. 2020. Comparisons between Tethyan anorthosite‐bearing ophiolites and Archean anorthosite‐bearing layered intrusions: implications for Archean geodynamic processes: Tectonics, v. 39, 35. Souders A.K., in review. U-Pb Geochronology of the Mentor Anorthosite Intrusive Complex (MAIC) and Regional Plutonic Units. U.S. Geological Survey Data Release. https://doi.org/10.5066/P9WMD477. 14 �Proceedings of the 69th ILSG Annual Meeting – Part 1 A revised provenance model for the Elk Mound Group in south-central Wisconsin based on detrital zircon analysis BOJDAK-YATES, Itai S.1, BJØRNERUD, Marcia1, MALONE, David H.2, and STEWART, Esther K.3 1 Department of Geosciences, Lawrence University, Appleton, WI, 54911, United States Department of Geography, Geology, and the Environment, Campus Box 4400, Illinois State University, Normal, IL, 61790, United States 3 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Rd, Madison, WI, 53705, United States 2 The Late Cambrian Elk Mound Group consists of three sandstone formations deposited in a shallow tropical sea: the Mount Simon, Eau Claire, and Wonewoc formations, in ascending order. The formations underlie much of the upper Midwestern United States in vast, thin sheets, which thicken toward the Illinois Basin further south. These formations have long fascinated geologists due to their extraordinary physical and chemical maturity, but they have often eluded explanation thanks to those same qualities. Recent studies have employed detrital zircon (DZ) U-Pb analysis to constrain the sources of the sand, and workers have begun to build regional provenance models that describe the origins of the sand and the routes it took to arrive at its present location. Our study builds upon these models with new samples from the Mount Simon Sandstone, a quartz arenite deposited in terrestrial and shoreface environments (Dott et al. 1986). We analyzed samples from outcrops of nonmarine deposits high on the Wisconsin Arch near Wisconsin Dells, WI, as well as a drill core taken 26 miles east of the Dells (the Triemstra core, near Belle Fountain, WI). We place these samples in the context of previous DZ work in this area, especially a study by Konstantinou et al. (2014). The formations of the Elk Mound Group are poorly defined in central Wisconsin, and the samples reveal a transition from Mesoproterozoic source provinces towards Late Archean source provinces as one moves up section and to the west (Figure 1). This transition is understood to represent a shift from sediments derived from the more local Wolf River Batholith (ca. 1470 Ma) and Penokean orogenies (ca. 1830 Ma) to more distal sediments derived from the Superior Province (ca. 2650 Ma). However, other sedimentary basins such as the Animikie and Huronian basins and the Midcontinent Rift may have contributed sediments as well. The physical maturity of the sand grains supports a recycled origin, as multiple cycles of weathering and erosion would have been necessary to produce such rounded grains (Dott, 2003). Sedimentological details of the sandstone reveal additional information about shifts in provenance. A pair of samples from the Wisconsin Dells area (upper Chapel Gorge and lower Mirror Lake) show relatively high proportions of Penokean-age sediments. The sedimentology of the outcrops sampled records a transition from a dune environment to a braided river environment, and these rivers may have brought sediment from the Penokees. Additionally, the proportional increase in Archean-age sediments correlates with a rise in sea level as one rises through the Elk Mound Group. This correlation suggests that local sediment sources were drowned by sea level transgressions, while the distal Superior Province remained high enough to continue eroding and contribute sediment to a shallow sea already rich in Archean-age sand. Paleocurrent indicators derived from optical borehole image logs from wells across central Wisconsin add to the regional provenance picture with evidence of predominant currents flowing toward the west and southwest, giving some indication of the more immediate source and final transport of these sediments. 15 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. DZ data from six samples gathered in central Wisconsin, organized in ascending order through the section and from east to west. (The Triemstra sample is the oldest and furthest east; the Wonewoc sample is the youngest and furthest west.) The Chapel Gorge samples came from the east bank of the Wisconsin River about 1.5 miles north of Wisconsin Dells. The Mirror Lake samples came from the northwest shore of Mirror Lake about 3.5 miles south of Wisconsin Dells. The Wonewoc sample was collected near Wonewoc, WI, about 21.5 miles west of Wisconsin Dells, and was analyzed by Konstantinou et al. (2014). References Dott Jr., R.H., Byers, C.W., Fielder, G.W., Stenzel, S.R., and Winfree, K.E., 1986. Aeolian to marine transition in Cambro-Ordovician cratonic sheet sandstones of the northern Mississippi Valley, USA. Sedimentology, 33: 345-367. Dott Jr., R.H., 2003. The Importance of Eolian Abrasion in Supermature Quartz Sandstones and the Paradox of Weathering on Vegetation-Free Landscapes. The Journal of Geology, 111(4): 387-405. Konstantinou, A., Wirth, K.R., Vervoort, J.D., Malone, D.H., Davidson, C., and Craddock, J.P., 2014. Provenance of Quartz Arenites of the Early Paleozoic Midcontinent Region, USA. The Journal of Geology, 122: 201-216. Lovell, T.R., and Bowen, B.B., 2013. Fluctuations in Sedimentary Provenance of the Upper Cambrian Mount Simon Sandstone, Illinois Basin, United States. The Journal of Geology, 121: 129-154. 16 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Quetico-Wabigoon Subprovince Boundary in the Superior Province north of Thunder Bay, Ontario, Canada BRIGGS, Blaize1, and HILL, Mary Louise1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The boundary zone between the Quetico and Wabigoon subprovinces is a complex zone of deformation and metamorphism that historically has been described as a fault, change in metamorphic grade and/or change in lithology. This boundary zone is exposed along Highway 527 within a roughly 23km stretch of highway. At the south end of this zone the DeCourcey Lake outcrop is a strongly foliated, mylonitic gneiss containing quartz, feldspar, garnet, sillimanite, muscovite, and biotite with pegmatites and boudinaged quartz veins that is interpreted to be part of the Quetico subprovince. The north end of the zone is marked by weakly foliated Max Lake conglomerate that displays primary sedimentary textures and is interpreted to be part of the Wabigoon subprovince. Cataclasis was discovered 9.8km north of the DeCourcey Lake outcrop and marks a sharp change from the high-grade amphibolite to granulite facies Quetico lithologies south of the cataclasite to subgreenschist to greenschist facies Wabigoon lithologies to the north. This cataclasite is characteristic of brittle deformation and evidence for a fault that has not been reported in previous studies. The fault is mapped parallel to the foliation of the cataclasite (Fig. 1). This cataclasite is interpreted to be a boundary fault marking the abrupt transition between the Quetico and Wabigoon subprovinces along Highway 527. Figure 1. Map of study area along Highway 527 showing sampled outcrops and new subprovince boundary. Thirteen outcrops along Highway 527 were mapped and sampled for microstructural analysis. 17 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Thin sections created from these samples were used to identify deformation microstructures in quartz and feldspar. Feldspar deformation microstructures are particularly useful and can be used as a proxy for temperature since feldspar needs higher temperatures than quartz to deform internally. Identifying deformation regimes for feldspar is important as most of the rocks within the study area are dominantly composed of quartz and feldspar. Metamorphic Grade Low Temperature 400 ℃ Textures/Deformation Structures -Patchy undulose extinction - Fracturing and cataclasis -Angular grains -Grain size faults with bent cleavage plane/twins Low-Medium 400-500 ℃ Medium 450-600 ℃ High 600 ℃ -Internal fracturing (minor dislocation glide) -Bulging recrystallization (BLG) -Tapered deformation twins -Bent twins -Undulose extinction -Deformation & kink bands -Core & mantle texture -Fine grain recrystallization/uniform grain size -Micro-kinking -Less abundant deformation twins -Sub-grain rotation (SGR) -Bulging recrystallization (BLG) -Core and mantle texture -Myrmekite along foliation planes Table 1. Feldspar deformation structures and corresponding temperatures/metamorphic grade based on descriptions from Passchier and Trouw (2005). References Passchier, C.W. and Trouw, R.A.J., 2005. Microtectonics, Second Edition. Springer. Berlin, New York: 366. 18 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Contributions of diverse mantle sources during the early stages of Midcontinent Rift formation — Implications for a passive rifting model BRZOZOWSKI, Matthew1,2, HOLLINGS, Pete1, ZHU, Jing-jing3, CREASER, Robert4 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada British Columbia Geological Survey, 1810 Blanshard Street, Victoria, BC V8T 4J1 Canada 3 State Key Laboratory of Ore Deposit Geochemistry, Institute of Geochemistry, Chinese Academy of Sciences, 99 Lincheng West Road, Guiyang, Guizhou Province 550081, PR China 4 Earth & Atmospheric Sciences, University of Alberta, 116 Street & 85 Avenue, Edmonton, AB T6G 2R3, Canada 2 It is generally accepted that the Midcontinent Rift System (MRS) and associated magmatism originated as a result of the impingement and melting of the Keweenaw Plume beneath the crust ca. 1.1 Ga (Hutchinson et al. 1990). This interpretation is based largely on Sm–Nd and Re–Os isotope data, and the need for a heat source to explain the large volumes of magma generated (Cannon 1992; Nicholson et al. 1997; Shirey 1997). This view has recently been challenged, however, given the long duration of magmatism associated with the MRS (Hollings and Heggie 2014) and paleomagnetic evidence that is indicative of rapid plate motion during the formation of the MRS (Swanson-Hysell et al. 2014). Alongside these ambiguities are uncertainties in the nature of the sources that fed the MRS with magma (e.g., plume vs. subcontinental lithospheric mantle)? Clarifying these ambiguities has remained challenging given that many of the earliest magmas in the MRS were variably contaminated by crustal material (e.g., the Nipigon sills), masking potential contributions from distinct mantle sources. Development of a robust genetic model for the early history of the MRS and the critical mineral resources associated with this magmatism requires a firm understanding of these contributions. To address this, we integrated new Os isotope data of Initiation (>1,109 Ma), Early (1,109–1,104 Ma), and Hiatus (1,104–1,098 Ma) stage rocks with variations in their bulk-rock trace-element and Nd isotope geochemistry (Brzozowski et al. 2023). Early MRS rock suites are characterized by highly variable γOsi values of -10 to 3857, with Early Stage melts exhibiting the greatest variability (-10 to 3857) and Initiation Stage melts exhibiting the smallest variability (4 to 50). Given that the γOsi values do not correlate with La/Sm, Gd/Yb, and MgO, this variability could not be due to variable degrees of partial melting, retention of garnet in the mantle, or fractional crystallization, respectively. Several of the rock suites of interest exhibit elevated Th/Nb–Th/La and radiogenic εNdi–Sri values that are indicative of crustal contamination and/or contributions from a subcontinental lithospheric mantle (SCLM) source. Based on numerical modeling, the radiogenic εNdi and γOsi values recorded by the mafic–ultramafic intrusions and sills are indicative of their crystallization from hybrid melts (enriched SCLM-derived melt > plume-derived melt) that assimilated <10% crustal material during emplacement (Fig. 1). In contrast, the melts that fed the diabase sills and subaerial lavas likely originated from depleted portions of the Keweenaw Plume based on their variably negative to positive γOsi values, and were contaminated during emplacement (Fig. 1). Although contamination can explain the range of εNdi values exhibited by the rock suites, it cannot independently account for the range of γOsi values because i) this would require unrealistically high degrees of contamination and ii) not all of the rock suites were contaminated (cf. Wolfcamp Basalt). Rather, it is likely that fractionation of sulfide liquid and/or Os-bearing platinum-group minerals also contributed to this variability. Together, these results indicate that i) not all of the rock suites in the MRS crystallized from plume-derived melts, ii) melt contributions from the SCLM were greatest during the early stages of rift formation, and iii) the MRS likely initiated passively, with plume impingement being a coincidence that provided the energy and material necessary for voluminous 19 �Proceedings of the 69th ILSG Annual Meeting – Part 1 magmatism. Figure 1. Variation in γOsi and εNdi in hybrid magmas generated by mixing of melts from various mantle reservoirs. The numbers along the mixing curves are the mixing percents. The numbers in rounded boxes are the εNdi values of the contaminant. References Brzozowski M.J., Hollings P., Zhu J-J., Creaser R.A., 2023. Osmium isotopes record a complex magmatic history during the early stages of formation of the North American Midcontinent Rift — Implications for rift initiation. Lithos: 436–437:106966. Cannon W.F. 1992. The Midcontinent rift in the Lake Superior region with emphasis on its geodynamic evolution. Tectonophysics 213: 41–48. Hollings P., Heggie G. 2014. Rethinking the Midcontinent Rift–puncturing the ‘Plume Paradigm’. In: 60th Institute on Lake Superior Geology. Hibbing, Minnesota, pp 57–58. Hutchinson D.R., White R.S., Cannon W.F., Schulz K.J., 1990. Keweenaw hot spot: Geophysical evidence for a 1.1 Ga mantle plume beneath the Midcontinent Rift System. J Geophys Res 95: 10869. Nicholson S.W., Schulz K.J., Shirey S.B., Green J.C., 1997. Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development. Can J Earth Sci 34: 504–520. Shirey S.B. 1997. Re-Os isotopic compositions of Midcontinent rift system picrites: implications for plume – lithosphere interaction and enriched mantle sources. Can J Earth Sci 34: 489–503. Swanson-Hysell N.L., Burgess S.D., Maloof A.C., Bowring S.A., 2014. Magmatic activity and plate motion during the latent stage of Midcontinent Rift development. Geology 42: 475–478. 20 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Eastward transition from banded iron-formation to ferruginous clastic rocks across the central Upper Peninsula of Michigan CANNON, W. F.1, DRENTH, Benjamin J.2 1 U.S. Geological Survey, MS 954, Reston, VA 20192; 2U.S. Geological Survey, Denver, CO 80225 The classic Paleoproterozoic iron-formations of the Lake Superior iron ranges are predominantly banded cherty chemical sedimentary rocks characterized by centimeter-scale interbedding of chert and various iron minerals. New observations from legacy iron exploration drill cores that sampled Precambrian rocks below Paleozoic sediments to the east of the exposed iron ranges in the Upper Peninsula of Michigan show that highly ferruginous fine-grained clastic sedimentary rocks are predominant in that area, and that true cherty iron-formation is a subordinate component of the ferruginous sedimentary section. Most of our information is derived from a collection of cores from proprietary exploration holes held by Cleveland-Cliffs Iron Company, who has allowed us to examine, sample, and describe the rock units. Those holes were drilled to test five large-amplitude magnetic anomalies (Figure 1). Cores from four additional anomalies that are publicly available at the Michigan Geologic Sample Repository were also studied. Figure 1. Reduced to pole aeromagnetic anomaly map showing anomalies sourced in sub-Paleozoic Precambrian basement, names assigned to each magnetic anomaly, and drill holes used in this study. Crosshatched pattern is the area of Paleozoic cover. The ferruginous clastic rocks examined in this study are generally laminated at centimeter- to millimeter-scale and range from fine-grained quartzite to siltstone. Most are even-bedded. Laminae alternate between quartzo-feldspathic and ferruginous; some of the latter are nearly 100% iron minerals. Average iron mineral content of individual short core segments is as much as 50% by visual estimates. All are metamorphosed to varying degrees, but unambiguous relict clastic textures are preserved widely. The combination of textures and mineral content leaves no doubt that these are 21 �Proceedings of the 69th ILSG Annual Meeting – Part 1 clastic rocks that accumulated very anomalous concentrations of iron. Figure 2. A. Whole thin section of thinly interlaminated fine sandstone and siltstone from the LaBranch deposit. Light layers are quartzo-feldspathic fine sandstone. Darkest layers are nearly all magnetite. B. Crossed polars view of quartz, microcline, biotite, and magnetite in fine sandstone from the Gladstone deposit. C. Same view as B in reflected light showing numerous magnetite grains. Bright partial rims on some grains are martite. Figure 3. Schematic section of approximately 100 kilometers showing the transition from Vulcan Ironformation in the west, as exposed on the Menominee Range (Bayley et al., 1966) and Felch Trough (James et al., 1961), to ferruginous clastic-dominated sedimentary rocks in areas covered by Paleozoic sediments in the east. We interpret these ferruginous clastic rocks as the lateral equivalent of the Menominee Group, which includes the major banded iron-formations of the Menominee and other iron-ranges of the western Upper Peninsula of Michigan. They record a gradation from the purely chemical and clasticstarved true banded iron-formations to the west, to a more shoreward facies where fine clastic sedimentation predominated and overwhelmed slow precipitation of chert beds. Intermittent periods of diminished clastic input allowed sporadic deposition of layers of cherty banded iron-formation, some of which are granular, indicating deposition in shallow water. These relationships show that the lateral disappearance of true banded iron-formations resulted from suppression of chemical chert precipitation by the input of fine-grained clastic sediments. However, intense iron deposition persisted into this more proximal fine-clastic-dominated facies resulting in abundant ferruginous clastic rocks. References Bayley, R.W., Dutton, C.E., and Lamey, C.A., 1966. Geology of the Menominee Iron-bearing District, Dickinson County, Michigan, and Florence and Marinette Counties, Wisconsin: U.S. Geological Survey Professional Paper 513: 96. James, H.L., Clark, L.D., Lamey, C.A., and Pettijohn, F.J., 1961. Geology of Central Dickinson County, Michigan: U.S. Geological Survey Professional Paper 310: 176. 22 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Workshop Outcomes and Updates for the Minnesota Department of Natural Resource’s Drill Core Library CARTER, Matt J.1 and ELSENHEIMER, Donald2 1 Minnesota Department of Natural Resources, 1525 3rd Ave E, Hibbing, MN 55746 2 Minnesota Department of Natural Resources, 500 Lafayette Rd, Saint Paul, MN 55155 The Minnesota Department of Natural Resources (DNR) provides public access to more than one million meters of drill core from over 9,000 locations across the state at its Hibbing Drill Core Library (DCL). This archive opened in 1967 and has been an invaluable resource for bedrock mapping, mineral exploration, and research, including numerous ILSG presentations. In November 2022, the DNR convened a workshop to gather stakeholder input on DCL policies and procedures (Carter et al., 2023). A need to update these policies and procedures was identified by DNR staff after conducting a 2022 inventory of DCL holdings and determining its current storage capacity, an assessment of projected core submissions, participation in a National Geological and Geophysical Data Preservation Program (NGGDPP) data management workshop, and a review of the policies and procedures of the United States Geological Survey (USGS) and peer repositories. Workshop participants were affiliated with the mining/mineral exploration industry, government agencies, academic institutions, and consulting firms. Feedback was gathered through exercises and participant surveys that focused on the mission of the DCL, prioritization of storage for various materials, sampling and related policies, and desirable enhancements to DCL databases. DNR staff used input from the workshop and reviewed the mission statements from the USGS, the DNR, and peer repositories to create a mission statement for the DCL. DCL curational decisions on what to add or retain in its collection have not previously been constrained by storage capacity. Given anticipated core submissions, participants were encouraged to consider submission and retention priorities, even with a planned addition of a fourth DCL storage building. It was recommended that prioritization should be given to materials that are costlier to replace, are more difficult to access (present and future) as well as complete (i.e., non-skeletonized) diamond drill hole cores that have economic and/or geologic significance. Suggestions were made in favor of retaining pulp and reject samples derived from bedrock core, while acknowledging the potential for the materials to degrade over time. It was suggested that unless surficial materials (e.g., outcrop, glacial sediments) have historical significance or were from areas with restricted access then they should be given a low storage priority. Participants encouraged the DNR to consider strategies that might optimize storage capacity or lower retention costs, such as standardized containers for unconsolidated materials or off-site storage of lower priority samples within the collection. Established DCL policies for facility visits, sampling protocols, and derivative thin sections and dataset submissions are comparable to peer repositories. While reviewing these policies, workshop participants expressed concerns about missing or oversampled intervals and suggested improving communication about the allowable sample size based on the proposed analyses. It 23 �Proceedings of the 69th ILSG Annual Meeting – Part 1 was generally accepted that samples, unused materials, and thin sections should be returned within a year. Yet, it was recognized that multi-year projects may need accommodations, that regular communication and updates must be provided by visitors who want to retain materials for over a year, and that consequences need to be established and enforced for those that do not follow policies. In general, the DCL could improve the communication of its policies to ensure visitors are better able to follow them. Participants also offered ideas on enhancing online access to DCL holdings and associated datasets. These included improving the accuracy of some drill hole collars as well as the link between historical and other publicly available data to drill holes. Digital images of cores boxes were also desirable and the DNR is conducting a pilot program to evaluate digital image collection. The importance of the DCL and the value it offers to researchers, the local mining and mineral exploration community, and the citizens of Minnesota was emphasized by workshop participants. DNR staff are currently using workshop feedback and relevant policies and procedures at peer repositories to make preliminary curational decisions that support the DCL’s mission on topics such as storage prioritization, development of operational policies, enhancements to associated databases, and future decision-making. Discussions and input on preliminary policy ideas at venues such as ILSG will help craft a published update to DCL policies and procedures. References Carter, M.J., Elsenheimer, D. and Arends, H., 2023. Minnesota Minerals Coordinating Committee Drill Core Library Workshop. Minnesota Department of Natural Resources, Lands and Minerals Division, OFR 411: 57. 24 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Digital Image Capture and Database Compilation of Historic Mining Data from the Keweenaw Copper District, Michigan: A Progress Update DeGRAFF, James1 and ROSE, William1 1 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 U.S.A. The Michigan copper rush starting at Copper Harbor in 1843 (Fig. 1) led to 150 years of mining that produced ~7.5 x 106 MT of copper (Bornhorst, T.J. and Barron, R.J., 2011), attracted ~100,000 persons from 40 countries, and profoundly influenced understanding of Lake Superior geology, advances in mining technology, and the region’s pattern of life. Many companies invested significantly in trenching, coring, and mining operations that generated an enormous body of geologic information. USGS efforts in the 1940s and 1950s to map bedrock geology and to assess mineral resources have compiled much of this information as bedrock geology maps with supporting cross sections and reports. Though available online in various formats, these map products are the tip of an iceberg of original detailed source data that is not easily accessed. Significant exploratory drilling that postdates map publication has not been utilized for later geologic investigations because of the same difficulty of access. Paper records and microfiche that decay with time are stored at various locations, which further complicates their use. Many groups could benefit from improved access to this vast amount of information. Therefore, we began a ‘skunk-works’ project to identify and gather information into a digital image repository, to extract it into tabular databases, and to explore how to make it available to scientists, industry, land-use planners, and the general public. Figure 1. Michigan’s copper mining district with generalized bedrock geology. Figure modified from (Bornhorst, T.J. and Barron, R.J., 2011). Numbered field trip stops generally define the extent of copper mining and exploration between 1843 and present. Limited mining also occurred on Isle Royale just off the map to the 25 �Proceedings of the 69th ILSG Annual Meeting – Part 1 north. The initial phase of the project was to identify sources, access data, establish procedures, and demonstrate feasibility. We started with drill holes, trenches, and mine openings posted on USGS geology maps of the Keweenaw Peninsula. Features were symbolized in Google Earth from georegistered maps, assigned unique codes, and recorded with their data in tables having a common layout (Stage 1). Derivative tables contain data unique to a class, such as azimuth and inclination of drill holes found on core logs (Stage 2). Data captured up to this stage are useful for positioning and orienting features on maps and in subsurface models. Stage 3 captures geologic data as a function of location in a feature, e.g., distance along a drill hole. Such information, available from core descriptions at the Keweenaw National Historical Park (Keweenaw National Historical Park, 2016), the USGS/Denver Archives (White, W.S., 1985), old reports and plates, often requires careful transcription to extract it from image records. Other potential sources of such mining data include early reports of the Michigan Geological Survey, university archives, and private collections. Besides preserving and making these data available to others in an easy-to-access format, we hope to build subsurface models that can benefit research, mineral exploration, and land-use planning (Fig. 2). Figure 2. Possible uses of the database once it is further developed. Acknowledgements: We thank Ted Bornhorst (MTU), Jeremy Mason (KNHP), Bill Cannon (USGS), and Jenny Stevens (USGS) for making us aware of and facilitating access to the two archives that currently are being digitally captured and tabulated. This work is possible because of the foresight of many late geologists who gathered and preserved the original paper records. References Bornhorst, T.J. and Barron, R.J., 2011. Copper deposits of the western Upper Peninsula of Michigan, in Miller, J.D., Hudak, G.J., Wittkop, C., and McLaughlin, P.I., eds., Archean to Anthropocene: Field Guides to the Geology of the Mid-continent of North America: Geological Society of America Field Guide 24: 83– 99, doi:10.1130/2011.0024(05). Keweenaw National Historical Park, 2016. Calumet & Hecla Records – 00019/004.02.01.03-007 Microfiche Drill Core Log Library: Calumet, Michigan, U.S. Department of the Interior, National Park Service, on microfiche (accessed August 2016). White, W.S., 1985. “Unpublished diamond drillhole core logs”: U.S. Geological Survey, Field Records Collection, Boxes 282: 287-290. 26 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Geophysical mapping of the Great Lakes Tectonic Zone and surrounding Precambrian geology in the central Upper Peninsula, Michigan DRENTH, Benjamin J.1, CANNON, William F.2 1 2 U.S. Geological Survey, PO Box 25046, MS 973, Denver Federal Center, Denver, CO 80225 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA 20192 The Great Lakes Tectonic Zone (GLTZ) forms the boundary between the Wawa-Abitibi subprovince (north side) and Minnesota River Valley subprovince (south side) within the Archean Superior Province. The GLTZ is concealed for all of its 1100 km length, except south of Marquette in the central Upper Peninsula of Michigan (Sims, 1991; Sims and Day, 1993). Near KI Sawyer, it is exposed as a NW-striking, 2.3 km wide mylonite zone along a strike length of about 11 km, with a mylonitic foliation that dips steeply to the SW (Sims, 1993). The location extent of the GLTZ is unknown to the east where it is concealed beneath Paleozoic sedimentary rocks. We use legacy aeromagnetic data (Daniels et al., 2009) in combination with modern aeromagnetic data (Drenth and Brown, 2020) and ground gravity data to geophysically characterize the GLTZ and map its eastward extent under cover and map additional nearby covered Precambrian tectonic elements. Discontinuous NW-striking aeromagnetic gradients observed over the mylonite zone are interpreted to be produced by structurally juxtaposed rocks with varying magnetizations, and such relations are observed locally in outcrops. Mapping of similar gradients across the region shows that they are widely distributed, but have highest concentration within 3 km of the center of the GLTZ. Gravity data show a steep regional gradient along the GLTZ trend, which is likely produced by the juxtaposition of a dense greenstone belt on the north against lower-density gneisses and granites on the south. Using the distribution of aeromagnetic gradients, broader aeromagnetic patterns, and the regional gravity gradient, the GLTZ is interpreted to extend about 55 km under cover to the east, where it changes to an E-W strike and possibly NE strike (Fig. 1). Interpretations are less detailed and less certain east of the area covered by high-quality aeromagnetic data. Interpreted Paleoproterozoic features have similar strike as the GLTZ. This includes an undated dike swarm and an elongated trough of variably magnetic and dense Paleoproterozoic strata that extends from the Gwinn district southeast under Paleozoic cover. The trough is truncated on its southeastern margin by an interpreted extension of the Norway Lake fault. The GLTZ is terminated on the east by broad aeromagnetic and gravity highs produced by rocks of the buried eastern arm of the 1.1 Ga Midcontinent Rift. The intersection of the rift and the GLTZ is the location of a change in the strike of the rift from crudely N-S north of the GLTZ to NW south of the GLTZ. References Daniels, D.L., Kucks, R.P., Hill, P.L., and Snyder, S. L., 2009. Michigan magnetic and gravity maps and data: a website for the distribution of data: U.S. Geological Survey Data Series 411: http://pubs.usgs.gov/ds/ds411. Drenth, B.J., and Brown, P.J., 2020. Airborne magnetic survey, Iron Mountain-Chatham region, central Upper Peninsula, Michigan, 2018: U.S. Geological Survey data release, https://doi.org/10.5066/P91EF3CI. Sims, P.K., 1991. Great Lakes tectonic zone in Marquette area, Michigan - implications for Archean tectonics in north-central United States: U.S. Geological Survey Bulletin 1904-E: 17. Sims, P.K., 1993. Structure map of Archean rocks, Palmer and Sands 7.5-minute quadrangles, Michigan, showing Great Lakes tectonic zone: U.S. Geological Survey Miscellaneous Investigations Map I-2355, 1:24,000 scale. Sims, P.K., and Day, W.C., 1993. Great Lakes tectonic zone -- revisited: U.S. Geological Survey Bulletin 1904- 27 �Proceedings of the 69th ILSG Annual Meeting – Part 1 S, 11 p. Figure 1. Preliminary interpretations. 28 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Geophysical architecture of the Neoarchean Mentor anorthosite intrusive complex, northwestern Minnesota DRENTH, Benjamin J.1, BLOCK, Amy Radakovich2, HUDAK, George J.3, SOUDERS, A. Kate4, SAARI, Stacy5 1 U.S. Geological Survey, PO Box 25046, MS 973, Denver Federal Center, Denver, CO 80225 Minnesota Geological Survey, 2609 Territorial Road, St. Paul, MN 55114 3 Natural Resources Research Institute, 5013 Miller Trunk Highway, Duluth, MN 55811 4 U.S. Geological Survey, PO Box 25046, MS 963, Denver Federal Center, Denver, CO 80225 5 Minnesota Department of Natural Resources, 1525 3rd Ave E, Hibbing, MN 55746 2 The ca. 2737 Ma (Souders, 2023) Mentor anorthosite intrusive complex (MAIC) lies near the northern margin of the Wawa subprovince of the Archean Superior Province, in an area of northwestern Minnesota where the Wawa, Quetico, and Wabigoon subprovinces are juxtaposed in close proximity (Fig. 1). The rocks of interest are entirely concealed by 10s to >100 m of unconsolidated Quaternary sediments and localized Cretaceous strata and saprolite. The MAIC comprises a large volume of megacrystic anorthosite, with a lesser volume of oxide-rich gabbros. The gabbros are known, from a single borehole intersection at ~70 m depth, to be enriched in vanadium (see http://minarchive.dnr.state.mn.us), and have further potential for chromium and titanium mineralization. New interpretations are based on data from an Earth Mapping Resources Initiative (MRI)-sponsored aeromagnetic survey flown in 2021 and pre-existing ground gravity data, constrained by approximately ten boreholes in the area. The anorthosite is weakly magnetized and dense, with a mean measured density of 2940 kg/m3, producing a 10-60 mGal gravity high. Pervasive epidote alteration is a suggested explanation for the high density of the anorthosite (the density of unaltered anorthite is 2730 kg/m3). The oxide-rich gabbros are strongly magnetized, producing aeromagnetic anomalies as large as 6000 nT, making them readily mappable across the complex. New geophysical interpretations (Fig. 1) suggest that the MAIC is significantly broader in extent than previously interpreted (Jirsa et al., 1999) and can be traced along strike for approximately double its originally interpreted length. The MAIC covers an area of about 640 km2 along a strike length of about 85 km, and forward modeling suggests a depth extent as great as 7 km. The MAIC is here interpreted to be the largest known anorthosite complex in the Superior Province, as measured by preserved extent in map view (cf. Sotiriou and Polat, 2020). The MAIC is observed in drill core to intrude a package of basalt flows at its northwest boundary and is itself intruded by multiple low-density felsic plutons that produce 10-20 mGal, 4-20 km wide gravity lows. The large felsic pluton along the southeastern margin of the MAIC is dated at 2702 ± 6.5 Ma (Souders, 2023), and is here called the Fertile pluton after the nearby town. This tectonomagmatic setting is consistent with other anorthosite complexes of the Superior Province, that commonly intrude packages of mafic volcanic flows and are themselves commonly intruded by felsic plutons (e.g., Polat et al., 2018). Disrupted trends and patterns of geophysical anomalies indicate that the MAIC was variably deformed, likely via both faulting and folding, in a complex fashion. 29 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Preliminary geophysical interpretations of geology surrounding of the Mentor anorthosite intrusive complex and surrounding area. Inset shows location of study area. References Jirsa, M. A., Chandler, V. W., and Runkel, A. C., 1999. M-092 Bedrock geologic map of northwestern Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/973. Polat, A., Longstaffe, F. J., and Frei, R., 2018. An overview of anorthosite-bearing layered intrusions in the Archaean craton of southern West Greenland and the Superior Province of Canada: implications for Archaean tectonics and the origin of megacrystic plagioclase: GEODINAMICA ACTA, v. 30, 1:84–99. https://doi.org/10.1080/09853111.2018.1427408. Sotiriou, P., and Polat, A. 2020. Comparisons between Tethyan anorthosite‐bearing ophiolites and Archean anorthosite‐bearing layered intrusions: implications for Archean geodynamic processes: Tectonics, v. 39: 35. https://doi.org/10.1029/2020TC006096. Souders A.K., 2023. U-Pb Geochronology of the Mentor Anorthosite Intrusive Complex (MAIC) and Regional Plutonic Units: U.S. Geological Survey data release. https://doi.org/10.5066/P9WMD477. 30 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Multiple overlapping features spatially associated with lead-zinc-copper mineralization in the Highland quadrangles, southwest Wisconsin, USA FITZPATRICK1, William, and STEWART1, Eric 1 Wisconsin Geological and Natural History Survey, University of Wisconsin-Madison Division of Extension, 3817 Mineral Point Road, Madison, WI, 53705 Several features of the Paleozoic bedrock units of the Upper Mississippi Valley (UMV) leadzinc district have been spatially correlated with sulfide mineralization in the Sinnipee Group including folds and faults (e.g. Heyl et al., 1959) and paleovalleys in the base St. Peter unconformity surface (e.g. Mai and Dott, 1985). The significance of these features in creating fluid pathways with sufficient flow to explain the temperature anomalies associated with ore deposition has been justified by the modeling of Arnold et al. (1996). New detailed 1:24,000 scale mapping of two quadrangles in the Highland area created a detailed structural and stratigraphic framework for this area at the northernmost margin of the UMV district, with the results providing a case study allowing the precise geometry of factors such as fold zones and paleovalleys relative to lead zinc mineralization to be revealed. Numerous E-W and N-S trending fold zones with amplitudes of 20-60 ft were identified during mapping of the Highland quadrangles (Fig. 1). Lead-zinc mineralization as defined by the digitized mineral development atlas (MDA) mine maps (Pepp et al., 2019) is clustered on the margins of the synclines, most commonly found on gently sloping ramps below the crest of adjacent structural highs. The largest deposits in the Highland area are spatially associated with pit zones where the base Platteville drops for an additional 40-80 ft below the trough of the synclines over restricted elliptical areas. These pit zones are the site of the steepest folding observed in the mapped area, and may have been important for compromising the integrity of the overlying Maquoketa formation, providing a fluid pathway for migrating brines through this regionally important aquitard (Arnold et al., 1996). Numerous paleovalleys filled with St. Peter formation were identified during mapping (Fig. 2), with the largest in the southeast and southwest corners of the quadrangles mapped continuing down to the Jordan formation with the Prairie du Chien group entirely removed. By removing the Prairie du Chien group, these paleovalleys provide connectivity between the thick, lower Cambrian sandstone aquifer and the upper St. Peter aquifer, allowing large volumes of migrating brines to migrate upward in section towards the favorable ore host units in the Sinnipee Group (Arnold et al., 1996). In the Highland district, the likely flow paths from these paleovalleys to the places where the Maquoketa aquitard was compromised at the pit zones directly correspond to areas with known lead-zinc mineralization. References Arnold, B.W., Bahr, J.M., and Fantucci, R., 1996. Paleohydrology of the upper Mississippi valley zinc-lead district: Society of Economic Geologists Special Publication, no. 4: 378-389. https://doi.org/10.5382/SP.04.28. Heyl, A.V., Jr., Agnew, A.F., Lyons, E.J., Behre, C.H., Jr., and Flint, A.E., 1959. The geology of the Upper Mississippi Valley zinc-lead district: U.S. Geological Survey Professional Paper 309: 310 p., 24 pls., https://doi.org/10.3133/pp309. Mai, H., and Dott, R.H., Jr., 1985. A subsurface study of the St. Peter sandstone in southern and eastern Wisconsin: Wisconsin Geological and Natural History Survey Information Circular 47: 35 p., 2 pls., https://wgnhs.wisc.edu/catalog/publication/000297. Pepp, K., Siemering, G., and Ventura, S., 2019. Digital atlas of historic mining activity in southwestern Wisconsin, 40 p., https://learningstore.extension.wisc.edu/products/digital-atlas-of-historic-miningfeatures-and-potential-impacts-in-southwestern-wisconsin. 31 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. 10ft structure-contour map for the base of the Platteville formation with interpreted fold axes marked by red lines with black outlines with arrows denoting synclines and anticlines. Green polygons mark surface diggings and blue polygons mark underground mine workings from the MDA data digitized by Pepp et al., 2019. Figure 2. Cross section running E-W through the southern part of the Highland quadrangles. Large black arrows mark likely flow paths for mineralizing fluids ascending from St. Peter paleovalleys (Oa) to pit zones which locally breach Maquoketa aquitard. Are serpentine fault mirrors an indicator of seismic slip? A microstructural analysis 32 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Are serpentine fault mirrors an indicator of seismic slip? A microstructural analysis GHANTOUS, Sam1, PHILLIPS, Noah 1, LUSK, Alex 2, NEWMAN, Julie 3, & JI, Shaocheng 4 1 Department of Geology, Lakehead University, Thunder Bay, ON, Canada Department of Geology & Geophysics, Texas A&M University, College Station, TX, USA 3 United States Geological Survey, Denver, CO, USA 4 Department of Civil, Geological and Mining Engineering, École Polytechnique, Montréal, QC, Canada 2 Fault mirrors are smooth, sheened surfaces along a fault plane. An array of microstructures may produce a fault mirror which each have respective formation mechanisms and associated slip velocities. Fault mirrors in certain compositions may be an indicator of ancient earthquakes with seismic slip velocities, but not all fault mirrors are associated with seismic slip. We study the microstructures of two serpentine mirror surfaces, which have not yet been described in the literature, to determine their formation mechanisms and to assess whether they serve as indicators of paleo-seismic slip. One sample is a medium green mirror surface from a late normal fault cutting dunites from the Twin Sisters complex, Washington State, USA. The second mirror surface is pale green and cuts a serpentinite from the Thetford Mines ophiolite in Quebec, Canada. Both fault mirrors have slickenlines on their surfaces indicating that they formed during slip. The mirror surface from the Twin Sisters complex consists of a ~2 micron thick, potentially amorphous, low asperity serpentine layer which may have formed during seismic slip. The mirror surface from the Thetford Mines ophiolite consists of a ~0.5 centimeter-thick layer which is composed of radiating serpentine microcrystallites which are ~ 1 micrometer in length and 10’s to 100’s of nanometers in width. These serpentine microcrystallites are interpreted to have crystallized from a serpentine gel phase during slip. While we hypothesize that these samples are both indicative of seismic slip, similar structures may form if serpentine gels crystallize during aseismic creep. Serpentine fault mirrors may represent paleo-seismic slip, but a microstructural examination of the mirror surface is required to establish a seismic origin. Figure 1. SEM photomicrographs of radiating serpentine microcrystallites from the Thetford Mines ophiolite fault mirror. 33 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Characterizing volcanic host stratigraphy and syn-volcanic intrusions at the Lynne Zn-Pb-Cu deposit, Oneida Co., Wisconsin GLODOWSKI, Lillian N. 1, LODGE, Robert W.D. 1 1 Department of Geology, University of Wisconsin-Eau Claire, 101 Roosevelt Avenue, Eau Claire, WI 54701 The Lynne Zn-Pb-Cu deposit in Oneida County, Wisconsin is one of several volcanogenic massive sulfide (VMS) deposits located within the understudied Paleoproterozoic (1.8-1.9 Ga) Penokean Volcanic Belt (PVB). The PVB formed as the Marshfield and Pembine-Wausau terranes collided and accreted onto the Superior Craton during the Penokean orogeny (Schulz and Cannon, 2007). VMS deposition in Wisconsin has been interpreted to be associated with continental back-arc rifting in a submarine environment. However, little data is available on the deposit-level at Lynne and other deposits in the PVB to test this model. Volcanic and tectonic variability in VMS forming environments and the effect of basement inheritance on metallogeny are important for district-scale exploration. This study constrains the volcanic and tectonic setting at the Lynne deposit via trace element systematics and aims to improve regional metallogenic models in the PVB. Historically, the Lynne deposit was subdivided by Adams (1996) based upon their relative stratigraphic position to the ore horizon into upper and lower “Rhyolite”, “Dacite”, and “Volcaniclastic (VCS)” with mineralized zones occupying the lower VCS unit (Figure 1A). This study relogged seven drill holes from the Lynne deposit and sampled for petrographic and geochemical analyses. The new geochemical data presented in this study reveals there are no petrochemical differences between the upper and lower host strata (Figure 1B). There were also no petrochemical differences observed between the volcanic host rocks and the intruding footwall granodiorite. Therefore, the rocks in this study have been subdivided based simply upon composition and petrography. The volcanic rocks which host the Lynne deposit are comprised primarily of medium to dark grey felsic to intermediate lapilli and crystal tuff. The sedimentary rocks at the Lynne deposit are observed to be very fine-grain, dark grey siltstones with thin parallel laminations and are assumed to be volcanically derived. The Lynne deposit is intruded by a pluton of medium-grained granodiorite which disrupts the lower massive sulfide lenses. The granodiorite appears in a variety of colors ranging from pink and orange to grey and white. Smaller mafic and felsic dikes also crosscut the Lynne deposit. The mafic dikes are dark grey to green with a fine-grain mafic matrix and feldspar phenocrysts. Felsic dikes are commonly light to medium grey with a fine-grain felsic matrix. The geochemical data indicates that VMS deposition at the Lynne deposit occurred in a bimodal-felsic petrochemical assemblage consistent with a continental setting. The shared FII-type lithogeochemistry of the felsic volcanic rocks, granodiorite pluton, and felsic dikes suggests these rocks formed under similar extensional, shallow crustal conditions and originated from the same magmatic system. Combined with the lack of a metamorphic aureole around the pluton, the intruding footwall granodiorite is likely the syn-volcanic intrusion which eventually intruded its own volcanic pile (Galley et al., 2003). Improved geochemical and petrographic data on the Lynne deposit will allow for more accurate and improved models which can be compared to other deposits throughout the PVB and around the world. 34 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. A) Geologic cross section of the Lynne deposit highlighting host stratigraphy, bore hole traces, approximate sample locations, and mineralized zones. Modified from Kennedy (1997). B) Rock type classification diagram of the Lynne. Diagram from Pearce (1996). References Adams, G.W., 1996. Geology of the Lynne base-metal deposit, north-central Wisconsin, U.S.A., in LaBerge, G.L., ed., Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume: Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, Cable, WI, v. 42, part 2: 161179. Galley, A.G., 2003. Composite synvolcanic intrusions associated with Precambrian VMS-related hydrothermal systems: Mineralium Deposita, v. 38: 443–473. Kennedy, L.P., 1997. Summary geologic and geotechnical report for the Lynne project Oneida County, Wisconsin, U.S.A., Unpublished report of Noranda Minerals Wisconsin Corp.: 26. Pearce, J.A., 1996. A users guide to basalt discrimination diagrams, Trace Element Geochemistry of Volcanic Rocks: Applications for Massive Sulphide Exploration. Geological Association of Canada, Short Course Notes 12: 79-133. Schulz, K.J. and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research, 157: 4-25. 35 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Identifying regional exploration domains for Ni-Cu-PGE deposit types in the Midcontinent Rift GOOD, David1 1 Department of Earth Sciences, Western University, London, ON N6A 5B7 Canada A new classification strategy for Midcontinent Rift basalts and associated gabbro and ultramafic rocks is proposed, the main objective being to identify magmatic suites associated with known Ni-Cu-PGE occurrences and their spatial distribution across the rift. The study is based on the idea that units with similar incompatible trace element signatures formed under similar conditions in a similar mantle source region and had been subjected to similar contamination or fractionation processes. Elements used in this study are REE, Th, Nb, and Zr. The approach taken is to identify point cloud clusters (magmatic suites) on contoured point density plots for REE represented by ‘lambda’ parameters which emphasize slope and curvature of REE patterns. The resultant groups are checked in Gd/Yb vs. Th/Nb and Gd/Yb vs. La/Sm diagrams which identify influence by crustal contamination or clinopyroxene fractionation, respectively. Melts produced in a metasomatised mantle source are a special case and are distinguished from contaminated melts in a Zr-Th-La diagram. The data set comprises a total of 1815 samples, 343 of which are basalt, from 70 mafic units. Data are carefully screened for discrepancies and extreme outliers removed. Results indicate a total of eight distinct magmatic suites (Groups 1 to 8). The groups are not listed in stratigraphic order because many units appear simultaneously, and a few are active for long time periods during the MCR event. Highlights of the study with respect to Ni-Cu-PGE mineralized intrusions include: a) Group 1 includes the Current, Seagull and Thunder Intrusions and the Lower Suite basalts of the Osler Volcanic Group; b) Group 2 is the most voluminous and includes the Duluth, Tamarack and Crystal Lake deposits, the Pigeon, Cloud and Arrow intrusions, and basalts of the Greenstone Flows, Upper Suite at Black Bay (OVG) and Upper Groups A and B at Mamainse Point; c) The Eagle deposit is intermediate between Groups 2 and 5 but overlaps the field for all flows in Lower Mamainse Point Group A; d) Group 7 includes the Two Duck Lake (Marathon deposit), Abitibi Dykes and metabasalt unit 3a; e) Group 8 includes the Geordie Lake deposit, Wolfcamp basalt, Copper Island dykes and a few of the Pukaskwa dyke swarm; and f) Groups 3 and 4 are not, as yet, associated with mineralized intrusions and includes the Nipigon sills and basalts of the Centre and Upper Suites of OVG. A map of the Midcontinent Rift showing regional domains for each Group is presented, highlighting the extent of igneous rock domains for each of the known Ni-Cu-PGE deposit types, and their locations relative to the central axis of the MCR. 36 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Exploring the geology of the Midcontinent Rift under western Lake Superior using a preliminary velocity model of seismic line GLIMPCE C GRAUCH, V.J.S.1, HELLER, Sam J.2, STEWART, Esther K.3, and WOODRUFF, Laurel G.4 1 U.S. Geological Survey, MS 973, Federal Center, Denver, CO 80225 U.S. Geological Survey, MS 939, Federal Center, Denver, CO 80225 3 Wisconsin Geological & Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 4 U.S. Geological Survey, 2280 Woodale Drive, Mounds View, MN, 55112 2 Seismic-reflection data were collected in the 1980s as part of the Great Lakes International Multidisciplinary Program on Crustal Evolution (GLIMPCE) to investigate the 1.1 Ga Midcontinent Rift System (MRS). GLIMPCE Line C crosses western Lake Superior from north to south shores (Fig. 1 inset). Many previous workers have interpreted the MRS in Line C as an asymmetric central graben filled with 10–20 km of subaerial basalt flows, overlain by 7-10 km of sedimentary section, and underlain by magmatic underplating. The central graben was interpreted to have formed from extensional normal faults, later reactivated as high-angle reverse faults. The northern part of Line C crosses over a prominent gravity low called the Grand Marais Ridge (GMR; Fig. 1 inset), previously interpreted as an Archean granitic basement high. Line C interpretations are commonly shown on a section plotted against two-way travel time along with a crudely estimated depth scale. We are undertaking a more rigorous approach by developing a detailed velocity model for time to depth conversion. The modeling for Line C is guided by velocities resulting from a pre-existing seismic refraction study, intervals defined by seismic horizons, and correlation with velocity models from neighboring seismic-reflection lines. Velocities are verified using common-reflection point gathers from pre-stack depth migration. Several salient points about the MRS can be gleaned from the preliminary velocity model alone (Fig. 1). The north and south sides of the model are dissimilar, reflecting the disparate geology of the north and south shores. On the south side, we identify an outline reminiscent of a bird (Fig. 1) that helps focus discussion without implying any geologic significance. Aided by a land-based seismic line near the southeast end of Line C, we can tentatively identify the geologic units under the lake within the bird outline (Fig. 1) and interpret a sag basin rather than a graben. The basin contains inferred Porcupine Mountains Volcanics (PM; 6.1 km/s), Portage Lake Volcanics (PLV; 5.9 and 6.5 km/s), with older, possibly reversed magnetic polarity, volcanic units at the base (6.9 km/s). A thick gabbroic sill (6.8 km/s) is inferred within the PLV section. We interpret the truncated PLV (5.9 km/s) and PM (6.1 km/s) intervals at the bird’s head to represent an eroded cliff face of the tilted northern limb of the sag basin. Sheet-like mafic intrusions (7.1 km/s) arise from the lower crust/upper mantle (7.2 km/s) and diverge upwards, following the geometry of the central sag basin. The interpretation that these 7.1 km/s units represent discontinuous or only partially evident magmatic feeder zones is based on their high velocities and sheet-like forms, which in part are constrained by neighboring industry seismic sections. The sedimentary section above the sag basin includes the Oronto Group (3.4, 4.7, 5.2, and 5.6 km/s) and likely Bayfield Group (3.0 km/s). An angular unconformity between Oronto Group (5.6 km/s) and underlying PM (6.1 km/s) at the bird’s head indicates the north limb of the sag basin was tilted prior to deposition. In contrast, the units on the south limb appear conformable. Using aeromagnetic patterns that lead from the north shore into the lake, we tentatively identify a highly reflective (not shown) 4.7 km/s interval as rhyolites of the upper northeast sequence of the North Shore Volcanic Group (NSVG). This unit is interpreted to be angularly unconformable with overlying sedimentary rocks of the same velocity (4.7 km/s). The 5.6 km/s and 6.5 km/s intervals 37 �Proceedings of the 69th ILSG Annual Meeting – Part 1 beneath the interpreted rhyolites are likely older NSVG volcanic rocks that form a carapace over the GMR. The 6.1 km/s velocity of the GMR corroborates its interpretation as a granitic basement high. The model indicates that a 5.6 km/s unit (NSVG?) dives below the bird outline to depths below 15 km. Whether this unit is connected to deeper parts of the sag basin or separated by faulting is obscured by the 7.1 km/s sheet-like intrusions. The sedimentary section on the north side of the model tilts to the south, unconformably overlies volcanic rocks (4.7 and 5.2 km/s) and is truncated by the overlying 3.0 km/s interval (Bayfield Group or equivalent). The sedimentary package on the north side collectively has lower velocities and is thinner than the sedimentary package on the south side. It is unclear if the northern section is correlative with the Oronto Group or represents less consolidated, younger rocks, possibly eroded from the NSVG or basalts at the bird’s head. Identification of velocity intervals and their relations at and under the bird’s head are key to understanding the tectonomagmatic picture but remain somewhat obscure. Suffice to say for Line C that magmatism and syn-magmatic subsidence played a greater role in the origins of the MRS than previously realized. Moreover, unconformable relations within the sedimentary package may be evidence of multiple post-magmatic tectonic events. Figure 1. Preliminary velocity model for GLIMPCE Line C showing velocity intervals in km/s. Inset map shows Line C in relation to the Grand Marais Ridge and neighboring seismic lines in Lake Superior. The white dashed line outlines a bird-like pattern to guide discussion. Velocities near the bird’s head are interfingered only to provide a smooth transition for the depth migration; lines are drawn to better represent the form of the depthconverted seismic horizons, which are not shown for simplicity. Vertical exaggeration=2. 38 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Petrography, geochemistry, and mineralization of the Archean Titan (Roaring River) intrusion, Northwestern Ontario GROENEVELD, Tianna1, HOLLINGS, Peter1, BAIN, Wyatt1, DJON, Lionnel2 1 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada, Impala Canada, 69 Yonge Street, Suite 700 Toronto, ON M5E 1K3, Canada The Archean Titan intrusion, formerly known as the Roaring River mafic intrusion, is one of several mafic-ultramafic complexes in northwestern Ontario that are currently the focus of ongoing PGE exploration. The Titan intrusion is located ~145 km North of Thunder Bay, Ontario, in the Winnipeg River terrane of the western Superior Province and is part of the Roaring River Complex (Figure 1). The Titan intrusion was identified as an underexplored area during a lake sediment survey in 2000 (Ontario Geological Survey). In the following five-year period, there were several periods of prospecting and soil surveys carried out in the area as well as one diamond drilling project, which aimed to determine the extent of the Titan intrusion within the Roaring River Complex and assess the potential for economic Ni-Cu-PGE mineralization. This early exploration revealed petrologic and geochemical similarities between the Titan intrusion and the mineralized mafic-ultramafic rocks in the Lac des Iles (LDI) Complex, which lies ~60 km to the south of the Titan intrusion (Figure 1). The LDI Complex is the largest of a series of mafic and ultramafic intrusions known as the LDI suite, all within the Marmion terrane, and hosts the world-class LDI palladium mine. An unpublished U-Pb age for zircons from the Titan intrusion yielded an age of 2690 ± 3.2 Ma, broadly coeval with the LDI Complex, dated at 2689 ± 1.0 Ma (Heaman and Easton, 2006). Outcrop across Titan is sparse, due to the presence of pervasive glacial till and Proterozoic diabase sills. Samples were collected in the summer of 2021 and analyzed for whole rock and PGE geochemistry, sulphur and Sm-Nd isotope analysis, and detailed petrographic characterization. The intrusion consists of a mix of lithologies, ranging from pyroxenites to gabbros to leucogabbros. The lithologies are distributed throughout the intrusion and suggest a simple magma body, where one pulse of magma underwent fractional crystallization within a closed system. Sulphide mineralization is generally confined to pyrite and chalcopyrite, though inclusions of pyrrhotite were observed occasionally. Sulphide mineralization is typically fine-grained and disseminated, though larger blebs do occur, usually of either pyrrhotite or chalcopyrite. The pyrite is considered to be a hydrothermal phase, likely formed from secondary precipitation while the larger blebs of pyrrhotite are considered to be a primary magmatic phase. The Titan intrusion is characterized by enriched LREE’s and fractionated HREE’s, with negative Nb, Zr, Hf, and Ti anomalies (Figure 2). Titan samples have a range of (La/Sm)N from 0.7 to 3.8, a range of (Gd/Yb)N from 2.3 to 7.4, and a range of Nb/Nb* values from 0.02 to 0.47. The geochemistry behavior is consistent with formation in a supra subduction zone setting, which fits with the regional setting of the Winnipeg River and Marmion terranes during this time period (~2.74-2.69 Ga). Only small amounts of crustal material appears to have been incorporated, based on εNd values of 0.70 to 1.82, compared to an estimated depleted mantle at 2.7 Ga which would have a εNd value of +3. Titan appears to be a simple intrusion when compared to intrusions of similar size in the LDI suite and many of the similarities between Titan and the LDI suite appear to occur from regional characteristics of the area in this time period (~2.74-2.69 Ga). 39 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. (left) A regional geology map of the western Superior Province highlighting the approximate locations of the Titan intrusion, the Lac des Iles Complex, and the city of Thunder Bay, modified from Stott et al., 2010. Figure 2. (below) Primitive mantle normalized spider plot, showing representative values for oceanic island basalts (OIB), continental arc, oceanic arc, and the span of values for Titan. Concentrations normalized to primitive mantle from Sun and McDonough (1989) OIB from Sun and McDonough (1989), continental and oceanic arcs from Kelemen et al. (2014). References Heaman, L.M. and Easton, R.M., 2006. Preliminary U/Pb geochronology results: Lake Nipigon Geoscience Initiative. Ontario Geological Survey, Miscellaneous Release-Data 191. Kelemen, P.B., Hanghøj, K., Greene, A.R., 2014. One View of the Geochemistry of Subduction-Related Magmatic Arc, with an Emphasis on Primitive Andesite and Lower Crust. Treatise on Geochemistry, vol. 4: 749-806. Ontario Geological Survey., 2000. Garden-Obonga Lake Area Lake Sediment Survey: Gold and PGE Data; Open File Reports 6028: 76. Stott, G.M., Corkery, M.T., Percival, J.A., Simard, M., Goutier, J., 2010. A Revised Terrane Subdivision of the Superior Province, in Summary of Field Work and Other Activities, 2010. Ontario Geological Survey, Open File Report 6260: 20-1 to 20-10. Sun, S.S., McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. Geological Society, London, Special Publications, vol. 42: 313-345. 40 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Determining Provenance of Rainy Lobe Till using Geochemistry and Detrital Zircon Geochronology. HINKEMEYER, Audray M.1, MOOERS, Howard D.1, and LARSON, Phillip C.2, O’SULLIVAN, Paul B.3 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Duluth, MN 55812 Vesterheim Geoscience, PLC, Hibbing, MN 3 GeoSep Services, 1521 Pine Cone Road, Moscow, Idaho 83843 2 Till of the Late Wisconsin Rainy lobe (RL), which emanated from the Labradoran sector of the Laurentide ice sheet, is exposed at the surface from SW Minnesota to the extreme NE part of the State. The RL advanced to its maximum limit in southwestern Minnesota well prior to the Last Glacial Maximum (ca. 27-30 ka BP) and retreated into Ontario by 17.9 ka BP. This till exhibits dramatic spatial and temporal changes in provenance from the Hewitt till of SW Minnesota to the Independence till in the NE. While texture, fabric, and physical properties are similar, lithologic changes include a decrease in carbonate and greywacke of the Omarolluk Fm. with an increase in mafic rocks of the Duluth Complex as the ice retreated. The observed change in lithology reflects changes in the mean transport length (MTL) of the till. The MTL is the average distance of transport defined by the indicator lithology abundance. The Hewitt till has a mean transport length of > 1000 km, whereas the Brainerd and Independence tills have mean transport lengths of approximately 400 and 100 km, respectively (Berthold, 2015). Two models have been proposed to explain the lithological differences (particularly carbonate) in RL tills. Goldstein (1989) postulated that the downglacier increase in carbonate in the Hewitt till was the result of progressive incorporation, by regelation or deformation, of older underlying till that was rich in carbonate. However, Goldstein also postulated an accretionary origin for the Wadena drumlins, which would imply continuous deposition rather than erosion. This subglacial erosional vs. depositional paradox remains unresolved. Larson (2008) concluded that the changes in sedimentology and landforms record systematic changes in provenance related to changing basal boundary conditions in the interior of the LIS. As the RL advanced early in the last glacial cycle, a continuous till sheet composed of sediment from Hudson Bay and the Hudson Bay lowlands (HBL) extended to SW MN. As the ice approached its maximum limit, much of this till sheet was then eroded exposing Canadian Shield bedrock along the central portion of the flow path (Fig. 1). Early in this phase of glaciation, the sediments reflect long-distance transport from Hudson Bay, and later phases reflect increased proportions of felsic shield lithologies and Duluth Complex rocks. These two models of Rainy lobe till sedimentology are evaluated using mixing models, till matrix geochemistry, and detrital zircon geochronology. The tills underlying the Hewitt till are typically finer textured and contain significant concentrations of Cretaceous age carbonates and shales. Therefore, a multicomponent mixing model is developed to examine sedimentological variability by incorporation of older, underlying tills (e.g. Goldstein, 1989). To evaluate the model of Larson (2008), which implies long vs. short transport distances, twenty-eight samples collected along a transect from SW to NE Minnesota, and six samples collected from the HBL, were processed and sent for geochemical analysis. Fifteen of these samples were processed and analyses for detrital zircon geochronology using laserablation, ICPMS. 41 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Results of a 48-element analytical suite along with latitude, longitude, and depth were run through a principal component. The first 3 factors were retained for analysis. Factors 1 and 3 distinguished mafic vs felsic igneous rock geochemical signatures and carbonate content, respectively. Factor 1, felsic vs. mafic lithologies, can be used as a proxy for MTL and shows locally vs distally derived lithologies. Factor 3 distinguishes tills based on carbonate content. Core SLL (Independence till) plots positively on factor 1 indicating a short MTL. Core CSS (Brainerd till) represents an intermediate MTL, while cores UMRB and TG (Hewitt till) SW of the Wadena drumlin field have the longest MTL. In addition, the samples with the longest MTL plot in high carbonate space, positive on Factor 3. Detrital zircon age populations represented on probability density plots show that the shortest MTL Figure 1. Factor 1 (MTL) vs. factor 3 (carbonate content). samples have the highest signature of local 1.1 Ga MidContinent Rift zircons. A Kolmogorov-Smirnoff (K-S) test statistically compares age populations and determines if they are statistically different. Results from the K-S test reveal that HBL ages are statistically similar to samples from central Minnesota (core CSS). The mixing model, indicates that the Hewitt till is not a mixture low-carbonate RL till and older underlying tills. Geochemistry, and detrital zircon analyses support the model of Larson (2008). Early deposits of the RL in SW Minnesota are geochemically similar to the high-carbonate HBL samples, indicating a distal provenance. This similarity is also observed in the detrital zircon results from the K-S test. Subsequently younger deposits lose the HBL signature and start to incorporate more felsic craton and eventually mafic signatures of the Mid-Continent rift system. References Berthold, A.J., 2015. Surface Boulder Concentrations of the Late Wisconsinan Rainy Lobe, Minnesota, USA. M.S. Thesis, University of Minnesota Duluth: 48. Goldstein, B.S., 1985. Stratigraphy, sedimentology, and late-Quaternary history of the Wadena drumlin region, central Minnesota: Minneapolis, University of Minnesota, Ph.D. dissertation: 216. Larson, P.C., 2008. Quantification of Glacial Sediment Erosion, Entrainment and Transport Processes and Their Implications for the Dynamic History of the Laurentide Ice Sheet. Ph.D. Dissertation, University of Minnesota: 76. 42 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Copper-rich melt inclusions from the St. Ignace Island Complex: Implications for magma mixing and mineralization HOLLINGS, Pete1, HANLEY, Jacob2, SMYK, Mark1,3, HEAMAN, Larry4, and COUSENS, Brian5 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada 2 Department of Geology, Saint Mary’s University, 923 Robie Street, Halifax, NS, B3L 2Y5 Canada 3 Ontario Geological Survey, Ministry of Mines, Suite B002, 435 James St. South, Thunder Bay, ON P7E 6S7 Canada 4 Department of Earth & Atmospheric Sciences, University of Alberta, 126 Earth Sciences Building, Edmonton, AB, T6G 2E3, Canada 5 Ottawa-Carleton Geoscience Centre, Department of Earth Sciences, Carleton University, 1125 Colonel By Drive, Ottawa. Ontario, K1S 5B6, Canada The St. Ignace Island Complex (SIC) comprises volcanic and intrusive rocks that were emplaced the upper portions of Midcontinent Rift-related, ca.1008 Ma Osler Group volcanic rocks (Davis and Sutcliffe 1985; Fig. 1). The St. Ignace Island complex is a ~26 km2 stock with a core of quartz-feldspar-phyric rhyolites and dacites and an outer ring of anorthosite and gabbro (Sutcliffe and Smith 1988; Giguere 1975). The petrology and geochemistry of the SIC has been described by Smyk et al. (2006) and Hollings et al. (2023). The pink to grey, felsic rocks at the center of the complex are quartz-phyric, with rare pyroxene and feldspar phenocrysts. Textures at a variety of scales show evidence of the mingling and mixing of partially crystallized mafic and felsic liquids in SIC rocks. Mafic and felsic liquids may be incipiently mixed, resulting in partially disaggregated mafic enclaves hosted in a felsic matrix. With progressive mixing, the felsic volcanic domains in the rock become darker and phenocrysts of quartz and alkali feldspar appear embedded in the mafic Figure 1. (A) Map of upper Great Lakes. (B) Regional domains. In the most intensely mixed geology of the St. Ignace Island complex. Age data samples, small, mafic crystalline clots are (black stars) from Davis and Sutcliffe (1985) and Davis dispersed throughout a felsic matrix, and as and Green (1997). (C) Geological map of St. Ignace rare mafic enclaves, consisting of only a thin Island, modified after Giguere (1975). rind of mafic rock surrounding coarsegrained plagioclase phenocrysts. Well-preserved silicate melt inclusions (MI), many completely glassy, were observed in quartz, clinopyroxene and some plagioclase phenocrysts from the felsic and mafic rocks of the SIC, representing some of the oldest unrecrystallized silicate melt inclusions recognised to date. Melt inclusions from quartz from the felsic rocks are broadly rhyolitic in composition whereas those from 43 �Proceedings of the 69th ILSG Annual Meeting – Part 1 plagioclase in the mafic rocks range from basalt to basaltic andesite. The melt compositions are interpreted to represent the end-member liquids in the system with direct evidence of mixing of the two. Concentrations of Cu and Ag (in both mafic and felsic MI), and Mo (in felsic MI), are up to an order of magnitude higher in both the mafic and felsic MI than in continental crust and the host bulk rock concentrations. We propose that the melt inclusions have preserved pre-eruptive metal tenors that were subsequently modified by sulfide saturation, degassing, or post-solidus hydrothermal alteration. The elevated Cu and Ag contents are similar to those noted in arc-related and extremely oxidized early Midcontinent Rift-related rocks and may account for the world-class volcano-sedimentary-hosted Cu(Ag) deposits within the Rift as well as the presence of small, porphyry-style deposits. References Davis, D.W., and Green, J.C., 1997. Geochronology of the North American Midcontinent rift in western Lake Superior and implications for its geodynamic evolution; Canadian Journal of Earth Sciences, v.34: 476488. Davis, D.W., and Sutcliffe, R.H., 1985. U-Pb ages from the Nipigon plate and northern Lake Superior; Geological Society of America Bulletin, v.96: 1572-1579. Giguere, J.F., 1975. Geology of St. Ignace Island and adjacent islands, District of Thunder Bay; Ontario Division of Mines, Geological Report 118: 35. Hollings, P., Hanley, J., Smyk, M., Heaman, L., and Cousens, B., 2023. The ~1.1 Ga St. Ignace Island complex, Northern Ontario, Canada: Evidence for magma mixing and crustal melting in the generation of Midcontinent Rift-related bimodal magmas and implications for regional metallogeny. Journal of Petrology, in review. Smyk, M., Hollings, P., and Heaman, L., 2006. Preliminary investigations of the petrology, geochemistry and geochronology of the St. Ignace complex, Midcontinent Rift, Northern Lake Superior, Ontario. In Wilson, A.C. (ed.), Proceedings and Abstracts, Institute on Lake Superior Geology 52nd Annual Meeting, Proceedings Volume 52, Part 1 – Program and Abstracts, 61-62. Sutcliffe, R.H., and Smith, A.R., 1988. Geology of the St. Ignace Island volcanic-plutonic complex; Summary of Field Work and Other Activities, Ontario Geological Survey, Miscellaneous Paper 141: 368-371. 44 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Hydrothermal Alteration Facies of the Eisenbrey Zn-Cu Deposit, Rusk County, Wisconsin JOHNSON, Kaine, P. 1, and LODGE, Robert W.D. 1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI This study focuses on the hydrothermal alteration zones surrounding the volcanogenic massive sulfide (VMS) Eisenbrey Zn-Cu deposit in Rusk County, northwestern Wisconsin. The Eisenbrey deposit is hosted within the Paleoproterozoic Pembine-Wausau terrane and is a part of the Penokean volcanic belt, along with many other VMS deposits including the Crandon, Lynne, and Flambeau deposits. The goal of this research is to develop a petrographic and geochemical categorization of alteration types and complete a geochemical mass balance to produce specific alteration trends. Data collected on the hydrothermal alteration at the Eisenbrey deposit is being compared with other Wisconsin VMS deposits to produce a better depositional framework for VMS mineralization. The Penokean Orogen is the culmination of various accretionary events and volcanism. The Penokean Orogen began around 1.88 Ga along the southern margin of the Superior Craton. The collision and subsequent accretion of the Pembine-Wausau terrane resulted in subduction moving to the south and began back arc basin development. Most VMS deposits within the Penokean volcanic belt formed within this back arc extensional environment (Shultz and Cannon 2007). Arc magmatism continued until roughly 1.85 Ga. when an Archean crustal fragment, known as the Marshfield terrane, accreted to the Pembine-Wausau terrane & Superior Craton. VMS systems are characterized by volcanic-sedimentary hosted massive sulfide deposits that form at or near sea floor. Formation is associated with convection of metal rich hydrothermal fluids rising through the crust and mobilizing elements. These deposits are commonly poly-metallic with common mineralization of Zn-Cu-Pb-Ag-Au rich sulfides. Hydrothermal alteration in VMS environments results in mobilization of major elements during modification of primary minerals. The style of alteration varies based on the volcanic setting and fluid chemistry, but commonly are noted by gains in MgO, Fe2O3, K2O, and/or SiO2 and losses in Na2O and CaO (Galley et al., 2007). The Eisenbrey deposit (Figure 1) is relatively poorly understood. Regional metamorphism at the Eisenbrey deposit is lower amphibolite facies and has completely recrystallized the alteration zone at the deposit. Eisenbrey deposit is the only known VMS occurrences associated with Algoma-type iron formation and formed within the “Main Arc Sequence” (DeMatties, 2022). Therefore, improving our understanding of the Eisenbrey hydrothermal system can aid in identifying new exploration criteria in non-typical VMS environments for the Penokean Orogen. Samples of the hydrothermal alteration zone at the Eisenbrey deposit were analyzed across twelve drill holes from both the structural hanging wall and footwall to the ore horizon. These samples were initially divided into alteration mineral assemblages based on petrography. Alteration types include chlorite-cordierite-anthophyllite, quartz-anthophyllite-biotite, quartz-white mica, quartz-biotite. These alteration types were then characterized using major and trace element geochemistry and mass balance calculations. The alteration at Eisenbrey has notable gains in Fe2O3 and MnO; with losses in SiO2, MgO, and Na2O. This contrasts alteration at Flambeau, which has gains in K2O and SiO2 (Lodge et al., 2022). 45 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Representative cross-section of the Eisenbrey deposit with representative photomicrographs of common alteration types (right). I. shows Quartz-Anthophyllite alteration (T-22), II. shows Chlorite-Cordierite alteration (T-40), III. Shows Quartz-White Mica alteration (T-22) References DeMatties, T.A., 2022. Exploration-resource assessment of productive felsic volcanic centers in the paleoproterozoic penokean volcanic belt of northern Wisconsin, Michigan and East-central Minnesota, USA: Ore Geology Reviews, v. 141: 104489. Galley, A.G., Hannington, M.D., and Jonasson, I.R., 2007. Volcanogenic massive sulphide deposits, in Goodfellow, W.D., ed., Mineral Deposits of Canada: A Synthesis of Major Deposit-Types, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods: GAC-MAC, Special Publication No. 5: 141-161. Lodge, R.W.D., Lemke, T.C., Blotz, K.E., 2022. Using Ore Petrography and Geochemical Mass Balance to Constrain the Hydrothermal Environment at the Paleoproterozoic Flambeau Cu-Zn-Au Deposit, Wisconsin, USA. Society of Economic Geology, Society of Economic Geologist Annual Meeting Proceedings, Denver, CO, paper P2.15. Schulz, K.J., and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157: 4–25. 46 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Provenance patterns and tectonic styles of ca. 2.3–1.8 Ga metasedimentary strata in northern Michigan based on regional mapping and detrital zircon U-Pb geochronology JONES, Jamey1, CANNON, William F.2, DRENTH, Benjamin J.3, and O’SULLIVAN, Paul4 1 U.S. Geological Survey, Alaska Science Center, Anchorage, AK U.S. Geological Survey, Geology Energy Minerals Science Center, Reston, VA 3 U.S. Geological Survey, Geology, Geophysics, and Geochemistry Science Center, Denver, CO 4 GeoSep Services LLC, Moscow, ID 2 Detrital zircon U-Pb data from ca. 2.3–1.8 Ga metasedimentary successions in northern Michigan are used to test regional stratigraphic correlations and yield key insights into provenance and tectonic styles along the southern Superior craton. Circa 2.3–2.2 Ga Chocolay Group turbiditic strata and quartzite record initial rifting and basin formation along the southern Superior margin. Unimodal ca. 2.7–2.6 Ga age populations were derived from abundant Archean batholiths in the surrounding region. Distinctive ca. 2.3 Ga populations are rare but present in some samples, but the source(s) of these grains is not well understood. Chocolay Group detrital zircon data are very similar to upper Huronian Supergroup strata to the east and with other global ca. 2.3–2.2 Ga glaciogenic successions. The ca. 2.1 Ga Dickinson Group contains bimodal ca. 2.9 and 2.7 Ga age populations in the East Branch Arkose and Solberg Schist that are distinctive in the region and suggest a mixture of recycled 2.3 Ga Chocolay Group quartzite and more diverse regional Archean basement sources. Minor ca. 2.1 Ga grains indicate derivation from nearby plutonic sources or eroded volcanic equivalents of the same age, consistent with magmatism, regional uplift, and final rifting of the southern Superior craton ca. 2.1. After a ca. 100 Ma hiatus, the Ajibik and Siamo Formations of the ca. 1.90–1.85 Menominee Group have unimodal ca. 2.7–2.6 Ga age populations that suggest continued derivation from ca. 2.7– 2.6 Ga batholiths and (or) recycling of older underlying strata. The Goodrich Formation of the basal Baraga Group (ca. 1.85–1.83 Ga) shows similar patterns. A provenance shift to prominent ca. 1.85 Ga populations occurs in turbiditic strata of the Michigamme Formation (upper Baraga Group), indicating arrival of the outboard Wisconsin magmatic terrane to the south. Michigamme strata record basin evolution between the southern Superior Province and the exotic terrane as it approached and collided during the ca. 1.87–1.83 Ga Penokean orogeny, but the relative role of Penokean versus younger ca. 1.78–1.76 Ga tectonism in regional folding and metamorphism remains uncertain. Additional mapping and geochronology focused on Michigamme strata will better constrain regional depositional ages, facies relationships, and tectono-metamorphic patterns. 47 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Petrogenesis of the mineralized horizons in the Offset and Creek zones, Lac des Iles Complex, N. Ontario JONSSON, Justin1, HOLLINGS, Peter1, BRZOZOWSKI, Matthew1, BAIN, Wyatt1, DJON, Lionnel2 1 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Impala Canada, 69 Yonge Street, Suite 700 Toronto, ON M5E 1K3 Canada The Lac des Iles Complex is a Neoarchean (2.69 Ga; D.W. Davis cited in Stone et al., 2003) polyphase mafic-ultramafic complex located in the Marmion terrane of the Superior Province, 85 km north of Thunder Bay, Ontario, Canada. The intrusive complex can be subdivided into two discrete subcomplexes: the ultramafic-dominated North Lac des Iles Complex and the mafic-dominated South Lac des Iles Complex (SLDIC). The SLDIC has been subdivided into four intrusive series, termed the gabbronorite, breccia, norite, and diorite series (Decharte et al., 2018). To date, economic Pd-rich mineralization has been discovered in both the breccia and norite series, and occurs proximal to the contacts between the breccia and gabbronorite series and between the breccia and norite series. The objectives of this study are to i) evaluate the mechanisms of formation of the mineralized horizons near the contact between the breccia and norite domains in the Offset and Creek zones of the SLDIC, ii) evaluate the role that crustal contamination played in this process, and iii) assess the tectonic setting in which the SLDIC formed. The breccia and norite series are both composed of varitextured, brecciated, and equigranular leucocratic-melanocratic norites and gabbronorites, and their altered equivalents. The breccia series contains a greater proportion of brecciated and varitextured rocks, while the norite series contains a greater proportion of equigranular rocks. All pre-alteration lithologies are essentially plagioclaseorthopyroxene cumulates with varyingly minor quantities of interstitial clinopyroxene, biotite, magnetite, chalcopyrite, pentlandite, and pyrrhotite. Variable degrees of hydrothermal alteration are indicated by the presence of tremolite-actinolite and talc (after pyroxenes), chlorite and sericite (after plagioclase), and pyrite (after pyrrhotite). Although the breccia and norite series are mineralogically similar, the breccia series is generally more leucocratic (i.e., higher plagioclase/pyroxene ratio) than the norite series. Neodymium isotopic evidence indicates that the Offset and Creek Zone magmas were crustally contaminated. ɛNd values of 19 analyzed samples range from +0.38 to -3.47 (median = -2.13), which is consistently more negative than the ɛNd value of +2.24 expected in an uncontaminated mantlederived magma that crystallized at 2.69 Ga. The crustal contaminant that imparted the negative ɛNd values is unlikely to be the tonalitic gneiss that hosts the SLDIC, as the ɛNd value of one reported tonalitic gneiss sample is -1.77 (Brugmann et al., 1997). The lack of correlation between ɛNd and geochemical or spatial variations suggests that variable crustal contamination was not the cause of the geochemical variability observed within the Offset and Creek Zones. Samples from both the breccia and norite series have similar trace-element chemistry, including enriched LILE/LREE patterns, flat HREE patterns, and pronounced negative Nb anomalies. Although these characteristics can be caused by assimilation of crustal material, it is more likely that they are the result of formation of the parental magma in a magmatic arc. Evidence for this interpretation includes low Nb/Yb ratios, high Ba/Th ratios, low Th content, and the lack of correlation between geochemical variability and Nd isotopic variability. Evidence from S isotopes of sulfide minerals and whole-rock geochemistry suggests that the addition of crustal S was not necessary in the formation of the Pd-rich mineralization within the Offset and Creek zones. δ34S values of 54 crystals from 17 samples range from -0.37‰ to +3.28‰ VCDT 48 �Proceedings of the 69th ILSG Annual Meeting – Part 1 (median = +1.11‰), with values from 52 of 54 crystals falling in the expected range of mantle-derived sulfur (0 ± 2‰; Seal, 2006). Based on the association of low Cu/Pd ratios with high Pd values, Offset and Creek zone ores formed at high R factors, which were likely high enough to cause the PGE enrichment without incorporation of crustal sulfur. The higher degree of Pd enrichment in the Offset Zone compared to the Creek Zone was likely due to a greater amount of sulfide liquid in the Offset Zone that also underwent higher R factors; the distribution of sulfide liquid and magma flow may have been influenced by primary structural constraints on the geometry of the intrusion. No evidence was found for significant low-temperature remobilization of chalcophile elements, including the PGEs. The compositional variability observed within the breccia and norite domains suggests that both domains formed via multiple pulses of compositionally similar magma. The proximity of mineralization to the interpreted feeder conduits suggests that the distribution of mineralization is largely the result of PGMs/Pd-rich pentlandite crystallizing as the magma transitioned from the feeder structure outwards into the periphery of the intrusive complex. This process may have repeated several times as successive magma pulses infiltrated the partially crystallized intrusive complex, resulting in the redistribution of ores in brecciated zones. References Brugmann, G.E., Reischmann, T., Naldrett, A.J., and Sutcliffe, S.H., 1997. Roots of an Archean volcanic arc complex: the Lac des Iles area in Ontario, Canada. Precambrian Research, vol. 81: 223-239. Decharte, D., Hofton, T., Marrs, G., Olson, S., Peck, D., Perusse, C., Roney, C., Taylor, S., Thibodeau, D., and Young, B., 2018. Feasibility study for Lac des Iles mine incorporating underground mining of the Roby Zone. North American Palladium, NI 43-101 Technical Report: 435. Seal, R.R., 2006. Sulfur isotope geochemistry of sulfide minerals. Reviews in Mineralogy and Geochemistry, vol. 61: 633-677. Stone, D., Lavigne, M.J., Schnieders, B., Scott, J., and Wagner, D., 2003. Regional geology of the Lac des Iles area, in Summary of Field Work and Other Activities 2003. Ontario Geological Survey, Open File Report 6120: 15-1 to 15-25. 49 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Slip Kinematics of the Keweenaw and Hancock Faults within the Midcontinent Rift System, Upper Peninsula of Michigan LANGFIELD, Katherine1, DeGRAFF, James1, GAMET, Nolan1 1 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, Houghton, MI, USA The Keweenaw fault is a major compressional structure along the center of the Keweenaw Peninsula and positioned near the southern edge of the Midcontinent Rift System (MRS). The smaller Hancock fault connects with the hanging wall of the Keweenaw fault and, together, the two faults define a thrust slice. The MRS formed ~1.1 billion years ago when a major extensional event split a significant portion of the ancient North American continent across the Upper Midwest. The rifting produced large volumes of basaltic lava, roughly ending with the Portage Lake Volcanics that have an exposed thickness of 3-5 km along the Keweenaw Peninsula (1). A common interpretation of the Keweenaw fault is that it originally formed as a normal fault during MRS extension and then inverted to become a reverse fault during a post-rift compressional event, most likely the Grenville Orogeny (2,3). Another interpretation is that the Keweenaw and Hancock faults are parts of a detached fault system that was initiated during the Grenville Orogeny (4). Until a few years ago, ideas about these and similar faults in the region considered only dip slip with an either normal or reverse sense of motion. Recent bedrock mapping and measurements of faultslip lineations, however, have revealed a significant component of right-lateral strike-slip on the Keweenaw fault system near its northeastern end which is about twice the magnitude of north-side-up reverse slip (5, 6). To clarify the slip kinematics of this region we utilized bedrock mapping and fault slip measurements between Hancock and Mohawk, MI to clarify the geometry and slip kinematics of the NE-trending Keweenaw and Hancock faults and to relate their characteristics here to what is observed along the more easterly trending portion of the fault system previously studied (Fig. 1). Rose diagrams of slickenlines rakes found along the Hancock and Keweenaw Faults show that both faults have roughly equal dip-slip versus strike-slip components (Fig. 2). This bimodal distribution of rake data differs from previous EDMAP projects, possibly due to the overall curvature of the Keweenaw Peninsula. The strike-slip to dip-slip component ratio was 2:1 (Mueller, 2021). The resulting map from this project indicates that the Keweenaw Fault isn’t a single fault trace, but instead connected fault segments (Fig. 3) The updated map and cross-section from this project proposes a new model for the Keweenaw Fault system kinematics. Acknowledgements This project was funded by the U.S Geological Survey’s EDMAP program under Award No. G21AC10681. This funding was matched by the Department of Geological and Mining Engineering and Sciences of Michigan Technological University, as well as sponsorship by the Michigan Geological Survey. Funding was also provided by the ILSG Student Research Fund for work done in the Quincy Mine, as well as an award by the Michigan Space Grant Consortium. Thanks goes to Tom Wright for access to the Quincy Mine. Additionally, we thank Ian Gannon, Breeanne Heusdens, Jack Hawes, Braxton Murphy, and Dillon Breen for fieldwork assistance. 50 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Map showing geology of the Keweenaw Peninsula. The boxes show the areas for the previous and current EDMAP project. (Cannon and Nicholson, 2001). Figure 2. Rake histograms showing the distribution of low and high angle rake on the Keweenaw fault (A) and Hancock fault (B). Arrows indicate mean rake of each dataset. Figure 3. Updated bedrock geologic map and legend of study area. References Cannon, W.F., and Nicholson, S.W., 2001. Geologic Map of the Keweenaw Peninsula and Adjacent Area, Michigan, U.S. Geological Survey, 1:100000 scale. Cannon, W.F., 1994. Closing of the Midcontinent rift ‒ A far-field effect of Grenvillian compression: Geology, v. 22: 155-158. Bornhorst, T.J., 1997. Tectonic context of native copper deposits of the North American Midcontinent Rift System: in Ojakangas, R.W., Dickas, A.B., and Green, J.C. (eds.), Middle Proterozoic to Cambrian Rifting, Central North America: Boulder, Co, GSA Special Paper 312: 127-136. DeGraff, J.M. and Carter, B.T., 2022. Detached structural model of the Keweenaw fault system, Lake Superior region, North America: Implications for its origin and relationship to the Midcontinent Rift System: Geological Society of America Bulletin, https://doi.org/10.1130/B36186.1. Tyrrell, C.W., 2019. Keweenaw Fault Geometry and Slip Kinematics – Bête Grise Bay, Keweenaw Peninsula, Michigan [M.S. thesis]: Houghton, Michigan, Michigan Technological University: 30. Mueller, S.A., 2021. Structural Analysis and Interpretation of Deformation Along the Keweenaw Fault System West of Lake Gratiot, Keweenaw County, Michigan, Open Access Master’s Thesis, Michigan Technological University 51 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Petrology and Geochemistry of the Paleoproterozoic Eau Claire Volcanic Complex, Eau Claire, WI LEAHY, Matthew D.1, LODGE, Robert W.D.1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA 1 The 1.8 Ga Eau Claire Volcanic Complex (ECVC) is located in the northwestern portion of Wisconsin primarily exposed in the Eau Claire River valley. The complex is part of the Marshfield terrane of the Penokean Orogen which developed along the southern margin of the Superior craton (Schulz & Cannon, 2007). Following the accretion of a juvenile ocean island arc, now known as the Pembine-Wausau terrane (PWT), with the southern margin of the Superior craton, opposing subduction zones closed the ocean between the accreted PWT and MT resulting in coeval magmatism on both terranes prior to collision around 1850 Ma. The origin of the MT is uncertain but is believed to be a small Archean craton that is either a rifted fragment of the Superior Province (Zi et al., 2021) or Wyoming Province (Malone et al., 2019). The suture between these terranes is the Eau Pleine Shear Zone. Paleoproterozoic subduction-related volcanism began to develop along MT’s northern margin, resulting in arc volcanism and back-arc spreading with associated calc-alkaline felsic magmas (DeMatties, 2022). This volcanism continued until the terrane collided with the subduction trench, resulting in a major compressional event along the Superior craton (Sims et al., 1989; Shultz and Cannon, 2007). This comprehensive interpretation of the tectonic setting fits well with the eastern portion of the MT where rocks are more abundantly exposed. However, the lack of outcrop exposure due to extensive Cambrian sedimentary strata has restricted research and mineral exploration in western parts of the orogen (DeMatties, 2022). This includes the ECVC, which is based on geophysical data, and has high potential for supergene-enriched VMS-style mineralization (DeMatties, 2022). The main objective of this study is to map and sample volcanic, metamorphic, and intrusive packages of the ECVC exposed along the North Fork of the Eau Claire River (Figure 1A) and Chippewa River for whole-rock geochemistry and petrographic analysis. Trace element geochemical data can be used to determine magmatic and tectonic settings of these rocks and improve regional tectonic models for the ECVC and MT. Twenty-four samples were analyzed for major elements via XRF and trace elements via ICPMS. Rock classifications were given in the field, reevaluated during petrographic analysis, and grouped into suites based on geochemistry. The majority of the suites were separated into four main categories: felsic gneiss (Figure 1B), mafic gneiss (Figure 1C), amphibolite (Figure 1D), and granitoid (Figure 1E). Each suite was diagnosed with a tectonic signature using multiple trace element diagrams. Th/Yb versus Nb/Yb displayed geochemical characteristics of deep crustal recycling for the majority of the samples, related to the active subduction that occurred during the advancement of the MT. The only suite that differs from this trend is the amphibolite group, which has a lower Th-Yb-Nb concentration, insinuating magma-crustal interactions with the protolith basalt. A tectonic classification tertiary diagram using La-Y-Nb solidified the theory that calc-alkaline arc magmatism dominated the MT region, while the amphibolite suite trends towards a more tholeiitic arc composition. This interpretation is backed by a magmatic affinity diagram as well using Th-Yb-Zr-Y percents. 52 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. (A) Regional map of the Eau Claire River with the North Fork and relative location in Wisconsin, (B) Poorly exposed bedrock of a felsic gneiss, (C) Isoclinal folded trondhjemite at Hamilton Falls trending eastwest, (D) Elongated pipe vesicles on an amphibolite outcrop near Knights Pool, (E) Intrusive contact between pegmatite and amphibolite. References DeMatties, T. A. (2022). Exploration-resource assessment of productive felsic volcanic centers in the Paleoproterozoic penokean volcanic belt of northern Wisconsin, Michigan and East-central Minnesota, USA. Ore Geology Reviews, 141, 104489. https://doi.org/10.1016/j.oregeorev.2021.104489 Malone, S.J., Nicholson, K.N., and Dowling, C.B., 2019, Preliminary geochemistry on the Marshfield Terrane, west-central Wisconsin: Geological Society of America Abstracts with Programs, doi: 10.1130/abs/2018am-322316. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4–25, doi: 10.1016/j.precamres.2007.02.022. Sims, P.K., Schmus, W.R., Schulz, K.J., and Peterman, Z.E., 1989, Tectono-stratigraphic evolution of the early Proterozoic Wisconsin magmatic terranes of the Penokean orogen: Canadian Journal of Earth Sciences, v. 26, p. 2145–2158, doi: 10.1139/e89-180. Zi, J.-W., and al., et, 2021, Refining the Paleoproterozoic tectonothermal history of the Penokean orogen: New U-Pb age constraints from the pembine-wausau terrane, Wisconsin, USA: doi: 10.1130/gsab.s.14700069. 53 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Structural analysis and slip kinematics of the Keweenaw fault system between Bête Grise Bay and Gratiot Lake, Keweenaw County, Michigan LIZZADRO-McPHERSON, Daniel1, DeGRAFF, James1, and GANNON, Ian2 1 2 Department of Geological and Mining Engineering Sciences, Michigan Technological University, 630 Dow Environmental Sciences, 1400 Townsend Drive, Houghton, MI 49931 USA The Keweenaw fault is perhaps the most important geologic structure on the Keweenaw Peninsula, with an estimated 7-11 km (1) of reverse slip juxtaposing Cu-bearing volcanic strata of the ~1.1 Ga Portage Lake Volcanics above ~1.0 Ga Jacobsville Sandstone. The fault has been interpreted as a riftbounding normal fault later inverted by compressional pulses of the Grenville Orogeny (2) and, more recently, as part of a detached thrust fault system unrelated to an earlier normal fault (1). The fault is shown on published maps as a nearly continuous fault trace whose sinuosity implies multiple fault segments and complex slip dynamics. Recent mapping has revealed that the Keweenaw fault at its most northeastern exposure on land is better characterized as a network of interconnected, left-stepping fault segments with easterly strike and exhibiting a 2:1 ratio of dextral strike slip to reverse slip (3). This project focused on the eastern half of a 2019-2020 EDMAP project (Fig.1) to map the Keweenaw fault system between Bête Grise Bay and Gratiot Lake. New mapping combined with structural and fault-slip analyses produced a revised bedrock geology map (Fig. 2) and a 3D-model (Fig. 3) that better constrain the geometry of the fault system, revealing folds and fault-bounded blocks in the main fault’s footwall. Analyses of fault slip data indicates a strike-to-dip slip ratio of 1.7:1 and a local shortening direction of 083°-263°. Slip along faults is a function of their strike relative to the shortening direction. Eastward transport of faultbounded blocks relative to the distal footwall was facilitated by mostly strike slip on longer EWtrending faults and reverse slip on shorter NEtrending faults, coupled with layer-parallel detachments along weak layer boundaries. The fault network defines a complex multistranded Figure 1. Bedrock geology of the Keweenaw transpressional system with overall dextral strike Peninsula (4), showing the 2017-2018 (grey box) slip and north-side-up reverse slip. Footwall folds and 2019-2020 (green box) EDMAP study areas. in Jacobsville strata adjacent to mostly strike-slip faults are considered to be cogenetic drag folds that formed during the Rigolet phase of the Grenville orogeny. These findings are consistent with recent mapping projects adjacent to the study area and investigations that relate far-field compressive pulses of the Grenville Orogeny to deformation of Keweenawan strata. Acknowledgements Funding provided by the USGS EDMAP program (Award No. G19AC00140) with a matching contribution from the Department of Geological and Mining Engineering and Sciences, Michigan Technological University and additional support from the Keweenaw Community Forest Company. Sponsored by the Michigan Geological Survey. 54 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 2. Keweenaw fault system between Lac La Belle and Gratiot Lake. Deer Lake fault block is in the Keweenaw fault’s footwall between the lakes. Cross-sections shown by thin black lines labeled A - F. Figure 3. Cross-section B-B' showing modeled hanging-wall and footwall structural and stratal relationships across the Deer Lake fault block. References DeGraff, J.M. and Carter, B.T., 2023. Detached structural model of the Keweenaw fault system, Lake Superior region, North America: Implications for its origin and relationship to the Midcontinent Rift System: Geological Society of America Bulletin, v. 51, no. 1: 449–466. Cannon, W.F., Green, A.G., Hutchinson, D.R. et al., 1989. The North American Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling. Tectonics, v.8:305-332. Tyrrell, C.W., 2019. Keweenaw Fault Geometry and Slip Kinematics – Bête Grise Bay, Keweenaw Peninsula, Michigan: Michigan Technological University, M.S. thesis: 30. Cannon, W.F. and Nicholson, S.W., 2001. Geologic Map of the Keweenaw Peninsula and Adjacent Area, Michigan. U.S. Geological Survey, Map I-2696, Scale 1:100,000. 55 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Re-evaluating the tectonics and metallogeny of terranes in the Paleoproterozoic Penokean Orogen, Wisconsin LODGE, Robert W.D.1 1 Department of Geology & Environmental Science, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA The tectonic model for the development of the Penokean orogen was synthesized in a classic paper by Schutz and Cannon (2007) that compiled decades of mapping, sedimentology, U/Pb geochronology, and geophysical surveys. The orogen started at ca. 1880 Ma with the accretion of Pembine-Wausau terrane, an oceanic arc complex, onto the margin of the Superior Province. A subduction flip after accretion resulted in overprinting continental arc volcanism and rifting (Figure 1A) until the collision a collision of an Archean crustal block, known as the Marshfield terrane, at ca. 1850 Ma. Several undeformed intrusions, interpreted as post-tectonic intrusions, constrain the end of the Penokean orogen at ca. 1835 Ma (Figure 1B). Perhaps the most important event during the orogen was the formation of the ~150 million tonnes of volcanogenic massive sulfide (VMS) deposits in the Pembine-Wausau terrane at ca. 1875 Ma (Sims et al, 1989; Quigley, 2016). This event was widespread across multiple VMS deposits. This presents a clear episode of submarine rifting and was assigned to a period of continental back-arc tectonism by Shultz and Cannon (2007). This is supported by the presence of inherited Archean zircons at the Lynne and Back Forty VMS deposits (Quigley, 2016) indicating the presence of Archean crust during the formation of Pembine-Wausau magmas. However, new U/Pb data has documented a second VMS forming event at ca. 1835 Ma at the Back Forty (Quigley, 2016) and Eisenbrey (Weber and Lodge, 2022) VMS deposits. Recognition of this extensional event has led to an alternate tectonic model wherein back-arc extension reactivated multiple times during ridge subduction (Zi et al., 2021). One of the principal issues that needs to be resolved with the classic Penokean tectonic model is the regional setting of Penokean VMS mineralization. VMS deposits formed in continental settings have different petrochemical associations than those formed in oceanic settings. New lithogeohemical data from mafic and felsic rocks at several VMS deposits (Flambeau, Eisenbrey, Lynne, Wolf River) suggest that most of the deposits hosted in rocks that are consistent with oceanic settings, while some suggest a continental setting. This suggests that the continental setting for the VMS mineralization does not apply to all deposits and that the extent of Archean basement needs to be better defined. Zircon petrochronology provides a mechanism to better resolve the nature of continental basement and its influence on metallogeny by providing a link between age of magmatism and tectonic setting and/or crustal inheritance. Once again, some deposits within the Pembine-Wausau terrane provide evidence for Archean basement, while others do not. However, in the process of discovering new ages, we also discovered that VMS forming environments continued until ca. 1835 Ma in a juvenile, oceanic setting. It was also discovered that some of the rocks from the Eau Claire volcanic complex of the Archean Marshfield terrane were mantle-derived, oceanic magmas that were ~1875 Ma with no evidence for Archean inheritance and seems eerily similar magmas from the Pembine-Wausau terrane. While Penokean magmas are known to intrude Archean rocks in the Black River Falls region of Wisconsin (Weber and Lodge, 2022), they clearly show Archean inheritance. As the hunt for domestic critical minerals makes its way to Wisconsin, the Penokean terranes and their metallogenic setting needs to be re-evaluated. 56 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 2. Illustration of tectonic models proposed by Shultz and Cannon (2007) and various new petrochemical or zircon petrochronology datasets that highlight some inconsistencies in the model. REFERENCES Quigley, A., 2016. Setting of the volcanogenic massive sulfide deposits in the Penokean Volcanic belt, Great Lakes region, USA: Unpublished M.S. thesis, Colorado School of Mines: 95. Schulz, K.J., and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157: 4-25. Sims, P.K., Van Schmus, W.R., Schulz, K.J., and Peterman, Z.E., 1989. Tectonostratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean orogen: Canadian Journal of Earth Sciences, v. 26: 2145-2158. Weber, E.M., and Lodge, R.W.D., 2022. New U/Pb Geochronology from the Proterozoic Penokean Orogen, Wisconsin: Implications for VMS Metallogeny. Society of Economic Geology, Society of Economic Geologist Annual Meeting Proceedings, Denver, CO, paper P5.10. Zi, J.-W., Sheppard, S., Muhling, J.R., and Rasmussen, B., 2021. Refining the Paleoproterozoic tectonothermal history of the Penokean Orogen: New U/Pb age constraints from the Pembine-Wausau terrane, Wisconsin, USA: Geological Society of America Bulletin, v. 134: 776-790. 57 �Proceedings of the 69th ILSG Annual Meeting – Part 1 3D geologic mapping at the Wisconsin Geological and Natural History Survey MAUEL, Stephen1, STEWART, Eric1, REHWALD, Matthew1, STEWART, Esther K. 1, AMES, Carsyn1, BREMMER, Sarah1, and FITZPATRICK, William1 1 Wisconsin Geological and Natural History Survey, University of Wisconsin-Madison Division of Extension, 3817 Mineral Point Road, Madison, WI, 53705 The Wisconsin Geological and Natural History Survey has constructed a preliminary 14-county 3-D geologic data model across southern Wisconsin. The model was constructed primarily from well construction reports (WCRs) that have been refined for several WGNHS projects, as well as data from the Mineral Development Atlas, borehole geophysics, and data from previous mapping performed at various scales. Well Construction Reports (WCRs) from digital and analog sources were assembled in a GIS geodatabase. The land surface elevation for each well was extracted from a DEM, and the elevation was then used to “hang” each well’s downhole lithology. By displaying and exaggerating the data in 3D, the different lithologies were carefully selected and assigned to geologic formations. Prior to interpolation, statistical outliers were identified, inspected, and edited when appropriate. The elevation for each formation contact was used to interpolate a raster. The resultant raster was inspected to identify obvious outliers, and after the outliers were edited or removed, a “final” raster of each contact was generated. The formation contact rasters can be intersected with a bedrock elevation raster to produce a geologic map. New data can be added to the model when available, and a new updated map can be generated. The products derived from this type of 3D geologic modelling are useful to the public in many applications. Harmful minerals or metals dissolved in groundwater are a realistic concern in Wisconsin, and determining the geologic formation in which a well terminates can help to avoid or resolve water quality issues. 3D geologic modeling can help to inform decision making about land use and land practices, land conservation, zoning and planning, identification of natural hazards, and the construction & engineering of wells, roads, railways, and buildings. 58 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Secular Changes in the Magnitude of Terrestrial Weathering MEDARIS, L. Gordon Jr.1, and DRIESE, Steven G.2 1 2 Department of Geoscience, University of Wisconsin–Madison, Madison, WI 53706 Department of Geosciences, Baylor University, Waco, TX 76798 In a recent investigation of paleosols in the Lake Superior region, the magnitudes of weathering in six Proterozoic paleosols were found to be less than those in four Phanerozoic paleosols and four modern soils (Medaris et al., 2022). However, in view of this relatively small database, the apparent age distinction in the magnitudes of weathering might be spurious, and thus we have expanded the database to test the veracity of secular changes in the magnitude of terrestrial weathering. Twenty-one first-cycle paleosols in igneous and metaigneous rocks with well-characterized and relatively homogenous protolith compositions were selected for comparison. These paleosols occur world-wide, vary in age from 100 Ma to 2960 Ma, and have protolith compositions ranging from gabbro to granite. This expanded database confirms that the magnitude of weathering in Phanerozoic paleosols and modern soils is greater than that in Precambrian paleosols. Potassium metasomatism is a common phenomenon in paleosols (Rye and Holland, 1998), and among the 17 Cambrian and Precambrian paleosols investigated here, 14 experienced potassium metasomatism, which is recorded by the presence of neoblastic muscovite, illite, or microcline. The effect of such K-metasomatism is illustrated in a plot of Al2O3-(CaO*+Na2O)-K2O, where compositional trends for modern soils and unmetasomatized paleosols are oriented subparallel to the AC*N join (Fig. 1A), and those for K-metasomatized paleosols are rotated towards the K apex (Fig. 1B). (A) (B) Figure 1. Protolith compositions and paleosol trends in the system, Al2O3-(CaO*+Na2O)-K2O. A: Modern soils and paleosols without K-metasomatism; B. Paleosols with K-metasomatism. In K-metasomatized paleosols, the amount of K2O removed by weathering is unknown, but may be estimated by comparison to an average for the depth variations of K2O and Na2O in modern soils, for which: (% change K2O) / (% change Na2O) = – 1.40z3 + 0.95z2 – 0.31z + 0.75 where z is normalized depth. Following this approach, the removal of K2O is estimated to be 47 ± 4% for the combined Cambrian and Precambrian paleosols and observed to be 54 ± 21% for the Cretaceous paleosols and 47 ± 21% for modern soils (Fig. 2A). Interestingly, no correlation exists between the percentage of K2O removed and age (or protolith composition; not shown). In contrast, the total 59 �Proceedings of the 69th ILSG Annual Meeting – Part 1 addition of K2O to the weathered profiles, expressed in terms of Depth-Normalized Mass Flux, progressively increases with decreasing age, i.e. 0.74 ± 0.29 at 2960 Ma, 0.96 ± 0.18 at 2450 Ma, 1.04 ± 0.52 at 1600-2200 Ma, and 1.62 ± 0.36 at 500 Ma (Fig. 2B). (A) (B) Figure 2. A: Percentages of K2O removed from soils and paleosols; B: Total K2O added to paleosols, expressed as Depth-Normalized Mass Flux (DNMF). The percentage removal by weathering for the sum of SiO2, CaO, Na2O, and K2O(est or meas) progressively increases from Archean (17.3±1.5%) to Proterozoic (21.0±3.7%) to Cambrian (25.1±3.1%) to Cretaceous (37.1±10.8%) paleosols. In comparison, the percentage of mass removed from five modern soils is 36.0±3.7%, which lies within the values for the Cretaceous paleosols. We suggest that the greater magnitude of weathering in Phanerozoic soils compared to Proterozoic ones is due to higher concentrations of organic acids during Phanerozoic soil formation, which resulted from the emergence of sparse cryptophytes in biological soil crusts in Cambrian time and subsequent greening of the continents with vascular plants from Devonian time to the present. Figure 3. Percentages of the total mass of SiO2, CaO, Na2O, and K2O removed from modern soils and paleosols. References Medaris et al., 2022. Journal of Geology, v. 130, in press. Rye & Holland, 1998. American Journal of Science, v. 298: 621-672. 60 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Morphometry and formation process of eskers developed under the Chippewa Lobe of the Laurentide Ice Sheet NUÑEZ-FERREIRA, Francisca1, ZOET, Lucas1, and RAWLING III, J Elmo 2 1 2 Department of Geoscience, University of Wisconsin-Madison, Madison, WI, 53705 Wisconsin Geological and Natural History Survey, University of Wisconsin‐Madison, Madison, WI, 53705 Eskers are an important indicator of paleo subglacial hydrologic conditions and a good alternative to direct glaciological observations because they are one of the few landforms that record those processes. Esker morphology and sedimentology is useful to gain insight into how sediment transport relates to subglacial hydrology along channels, which in consequence provides understanding on ice dynamics. However, large discrepancies in the formation mechanisms of eskers still exist and there are even fewer attempts to investigate the influence of soft bed conditions on this process. To address this, we analyzed the morphometry and distribution of eskers formed under the Chippewa Lobe of the Laurentide Ice Sheet (Figure 1). This includes mapping the sinuosity and spatial distribution with 2m resolution LiDAR, comparing these to sediment thickness derived from a water well data base, and examining the sediment sequence of one large esker exposed to sand and gravel extraction (~20 m tall) (Figure 2). The LiDAR analysis revealed a direct relation between sinuosity and length of eskers formed in soft bed conditions, with a mean of 1.07 that is very similar to eskers formed under hard bed conditions. Eskers spacing over the soft bed of the Chippewa Lobe appear closer than over hard beds in Canada (e.g Storrar et al, 2014). The spacing of eskers decrease when the ice margin retreats, meaning that melt rates increase (Boulton et al, 2009; Hewitt, 2011). Moreover, the relation between the distribution of eskers and till thickness indicates that eskers formed preferentially over thin layers of sediment, specifically near 18 meters for the Chippewa Lobe. The results from the grain size distribution of the large esker showed that the critical shear stress changed nonmonotonically throughout the formation of the esker. As such, we can assume that the water velocity or depth of the channel likely changed sporadically with time while the esker formed. Figure 1. Distribution of eskers formed under the Chippewa Lobe during the Last Ice Age. 61 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 2. Location of the selected esker for sediment analysis. The yellow start shows the location of a pit where the samples were extracted for the analysis. References Boulton, G.S., Hagdorn, M., Maillot, P.B., & Zatsepin, S., 2009. Drainage beneath ice sheets: groundwater–channel coupling, and the origin of esker systems from former ice sheets. Quaternary Science Reviews, 28(7-8): 621-638. Hewitt, I.J., 2011. Modelling distributed and channelized subglacial drainage: the spacing of channels. Journal of Glaciology, 57(202): 302-314. Storrar, R.D., Stokes, C.R., & Evans, D.J., 2014. Morphometry and pattern of a large sample (> 20,000) of Canadian eskers and implications for subglacial drainage beneath ice sheets. Quaternary Science Reviews, 105: 1-25. 62 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Subsurface characterization of the Duluth Complex and related intrusions from 3D modeling of gravity and magnetotelluric data PETERSON, Dana E. 1, BEDROSIAN, Paul A. 1 and FINN, Carol A. 1 1 U.S. Geological Survey, MS 973, Federal Center, Denver, CO 80225 The Mesoproterozoic Duluth Complex and related intrusions in northeastern Minnesota make up the second largest exposed mafic intrusive complex in the world, second only to the Bushveld Complex in Africa. It is one of the major plutonic components of the Midcontinent Rift System and hosts a variety of copper-nickel sulfide and platinum-group-element deposits. Given the complex geology of the area, 3D modeling is necessary to provide a complete picture of the variable densities and geometries of intrusive suites throughout the Duluth Complex as well as their extent at depth. In this study, we use Bouguer gravity data collected over the past ~60 years and magnetotelluric data collected in 2019 to create new 3D models of density, resistivity, and subsurface structure of the region constrained by geologic data. We use the results of these models to calculate the total volume of the Beaver Bay Complex, Duluth Complex, and onshore North Shore Volcanic Group, and estimate preliminary intrusion and emplacement rates using age estimates from Swanson-Hysell et al. (2021). We model both thickness and density of intrusive and volcanic rocks in the region using Oasis GMSYS-3D. The igneous layer in our starting model is 11 km thick with a constant density of 2,941 kg/m3. Other surfaces in the model include topography, near surface glacial deposits, a high-density lower crustal layer, and the base of the crust. We start our inversion by allowing the basal surface of the igneous units to vary and then invert for density within the igneous layer, within a range of 2,630-3,180 kg/m3. Our gravity modeling indicates that intrusive and volcanic rocks reach a maximum thickness ~23 km, or half the crustal column, with densities ranging from ~2,730-3,030 kg/m3 and a mean density of 2,940 kg/m3. The thickest, highest density areas of the model are beneath the Beaver Bay Complex and other mapped diabase intrusions. We interpret the two thickest areas in our gravity model as feeder zones for the Beaver Bay intrusive complex and possibly also for the Duluth Complex, in-line with interpretations arising from previous gravity studies in the area (Allen, 1994; Miller et al., 2002). Preliminary volume estimates from 3D gravity modeling indicate the present-day Duluth Complex, Beaver Bay Complex, and onshore volcanic rocks constitute ~92,100 km3 of igneous material. We calculate the volume of separate mapped units by extending the mapped geologic boundaries at the surface to depth within our 3D model. Three major geologic groups each comprise ~30% of this total volume: 1) the North Shore Volcanic Group, 2) diabase units of the Beaver Bay Complex and intrusions to the northeast and southwest of it, and 3) the Duluth Complex Layered and Anorthositic series. The older Early gabbro series and Felsic series of the Duluth Complex make up the remaining ~10% volume. 206Pb/238U zircon ages for the Anorthositic and Layered series from Swanson-Hysell et al. (2021) indicate that rocks of these units were emplaced contemporaneously over a period of 500,000 ± 260,000 years, suggesting an emplacement rate of ~0.06 km3/year, assuming a constant rate on magma input. Using recently acquired magnetotelluric data, we invert for resistivity in the study area using ModEM (Kelbert et al., 2014). Our magnetotelluric model highlights an arcuate low resistivity anomaly at depths of ~9-20 km, westwardly adjacent to the high-density and high resistivity feeder zones (Figure 1). This anomaly may represent a plane of weakness along which magma intruded to form the Beaver Bay Complex and the Duluth Complex. Low resistivities in this case would be attributed to sulfide or graphitic mineralization that developed along the contact between intruding magma and country rock. These resistivities are also similar to values observed in the Paleoproterozoic metasedimentary rocks of the Animikie Basin, located to the southwest of the Duluth Complex. The 63 �Proceedings of the 69th ILSG Annual Meeting – Part 1 spatial and temporal relationship between the Animikie Basin and Duluth Complex raises a tantalizing hypothesis that the basin structure may have preferentially localized magma intrusion. If this was the case, entrainment of conductive metasedimentary rocks of the Animikie Group along a pre-existing fault could also explain the arcuate low resistivity anomaly observed adjacent to the highly resistive feeder zones. Figure 1. Depth slice through the 3D magnetotelluric resistivity model at ~14 km depth. Black dashed line is the surface extent of the Duluth Complex and related intrusive and volcanic rocks. White lines are 5 km contours of Duluth Complex thickness extracted from our gravity model, starting from 10 km. Cyan line is the outline of Lake Superior. Acknowledgements Any use of trade, firm, or product names is for descriptive purposes only and does not imply endorsement by the U.S. government. References Allen, D.J., 1994. An integrated geophysical investigation of the midcontinent rift system: Western Lake Superior, Minnesota, and Wisconsin. PhD thesis: Purdue University, West Lafayette, Indiana: 267. Kelbert, A., Meqbel, N., Egbert, G.D. and Tandon, K., 2014. ModEM: A modular system for inversion of electromagnetic geophysical data. Computers & Geosciences, 66:40-53. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., Wahl, T.E., 2002. RI-58 Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/58804. Swanson-Hysell, N.L., Hoaglund, S.A., Crowley, J.L., Schmitz, M.D., Zhang, Y., & Miller Jr, J.D., 2021. Rapid emplacement of massive Duluth Complex intrusions within the North American Midcontinent rift. Geology, 49(2): 185-189. https://doi.org/1.1130/G47873.1. 64 �Proceedings of the 69th ILSG Annual Meeting – Part 1 On the Importance of Geologic Maps for Mineral Exploration PETERSON, Dean1 1 Big Rock Exploration, 2505 West Superior Street, Duluth, MN 55806 The basis for most types of geologic investigations is fundamentally rooted in geologists’ observations and interpretations made of landscapes, exposed rocks, and surficial materials in their natural habitat: “in the field”. Coherent geologic maps, which may take many years to create, represent assembled collections of observations in context of space and geologic time, requiring teams of geologists who are usually employed by federal or state/province geological surveys. The outcomes of these concerted efforts in the field are published geologic maps at various scales. It is these works of publicly funded geologic mapping that form the foundation upon which mineral exploration programs and mineral resource developments are built (Figure 1). These early endeavors are key components in national goals to define domestic resources of critical minerals. In decades past, many university geology students in the USA (including the author) were employed as summer interns assisting geological survey geologists in the bedrock geologic mapping of 1:24,000 scale quadrangles. This type of early professional experience can have profound implications for the careers of these students. Student knowledge gained includes the understanding of what it takes to systematically map bedrock exposures and structural zones, categorize the various rock types into lithologic map units, write out detailed descriptions of these map units, generate geologic cross sections and correlation diagrams, and putting all of these components together into a map that the geologic survey will subsequently publish. In today’s mineral industry, geologic maps are largely digital compilations of publicly available regional/district scale GIS data (downloaded and/or digitized from geological survey websites) merged/overlain with detailed industry geologic mapping of prospects and/or project areas. For the most part, the mineral industry quickly compiles digital data into geologic databases and is seemingly always searching for new ways to quickly capture data in the field digitally. The ease with which the mineral industry can generate digital geologic map products today can be good, bad, or ugly. The state of such geologic map outcomes by industry entities rests largely on the knowledge and experience of the company geologists. The US Geological Survey’s (USGS) Earth Mapping Resources Initiative (Earth MRI) program is a partnership of the USGS, the Association of American State Geologists (AASG) and other governmental, Tribal, and private-sector entities to update the nation’s surface and subsurface mapping to improve our knowledge of the geologic framework in the United States and to identify areas that may have the potential to contain undiscovered critical mineral resources. In November 2021, the US government passed the Infrastructure Investment and Jobs Act, one outcome of which is an investment of $320 million into Earth MRI to develop a better understanding of sustainable mineral production and mine waste options. An industry appeal to Earth MRI programs is to reinvigorate the education of future professional geologists by employing hundreds of geology student interns in upcoming geologic mapping projects. 65 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. The mineral development trapezoid. 66 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Fault zone architecture in mafic protoliths at the Lac des Iles mine, northwestern Ontario PETERZON, Jordan1, PHILLIPS, Noah1, HOLLINGS, Peter1, DJON, Lionnel2 1 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1, Canada Impala Canada, 69 Yonge Street, Suite 700 Toronto, ON M5E 1K3, Canada Faults are important geologic structures that host earthquakes and serve as permeable pathways through the upper crust. From an economic perspective, faults may transport and trap mineralized fluids. In turn, trapped mineralization may be offset or remobilized by later faulting. Fault zones are complex structures that produce an array of fault rock fabrics and architectures. Fault zone architecture typically consists of three components: 1) a fault core where most of the slip has been accommodated, 2) a damage zone bounding the fault core where fracture density increases with proximity to the fault core, and 3) an undeformed and less altered protolith. (Faulkner et al., 2010). Permeability is significantly enhanced in damage zones due to the high density of fractures and is diminished in fault cores due to the presence of clay-rich fault gouges. Faults may therefore act as conduits or barriers for fluid flow depending on the proportion of fault core to damage zone (i.e., the fault zone architecture; Caine et al., 1996). Fault zone architecture has been well studied in felsic to intermediate protoliths but studies on mafic protoliths are lacking. Here, we examine late faults within the mafic Lac des Iles complex to characterize fault zone architecture in mafic protoliths. The Lac des Iles complex is a series of mafic-ultramafic intrusive bodies occurring within the Marmion terrane of the Superior Province. The complex has been dated at 2689 ± 1.0 Ma and was emplaced into a ~3.01 – ~2.68 Ga granite-greenstone terrane (Djon et al., 2018). The Lac des Iles mine, owned and operated by Impala Canada, is a working Pt-Pd mine which is classified as a structurally controlled magmatic sulfide deposit. Extensive Ni-Cu-PGE mineralization has been offset by two late reverse faults in the high-grade zones (>4 g/t Pd): the Camp Lake fault and the Offset fault. A depletion in Pt-Pd mineralization is observed surrounding the late Camp Lake fault which extends ~180m into the hanging wall and ~145m into the footwall. Five drill holes that cross the late faults were logged and sampled in detail, with a fracture density counting program conducted systematically in the hanging wall and footwall. Fracture density increases as a power law function with proximity to the fault core and correlates with alteration. Tonalite has a higher fracture density and fracture density decay rate than gabbronorites near the fault. Fracture density and hematite/epidote alteration are more intense in the damage zone when faults cut through tonalite than when faults cut through gabbro. Fault cores in tonalite display a range of textures, from chlorite-rich gouges to fault breccias with calcitic matrix, while fault cores in gabbro only display chlorite-rich gouges. In this study, felsic protoliths have a higher fracture density than mafic protoliths indicating that fluid flow was more effective in felsic protoliths which may have contributed to depleted mineralization. This implies that host rock lithology strongly affects fault zone structure, including alteration assemblages, fracture densities, and permeabilities. 67 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Simplified regional map of the Lac des Iles intrusive complex (Djon et al., 2018). Figure 2. (A): Fracture density data from a single drill hole displaying an increase in fractures with proximity to faulting. (B): Underground exposure of the Camp Lake Fault at Lac des Iles. (C): Schematic of a typical fault zone architecture with corresponding cartoons of typical fracture density and permeability across the fault (Faulkner et al., 2010). References Caine, J.S., Evans, J.P., and Forster, C.B., 1996. Fault zone architecture and permeability structure. Geology, 24 (11): 1025-1028. Djon, M.L., Peck, D.C., Olivo, G.R., Miller, J.D., and Joy, B., 2008. Contrasting Style of Pd-rich Magmatic Sulfide Mineralization in the Lac des Iles Intrusive Complex, Ontario, Canada. Economic Geology, 113 (3): 741-767. Faulkner, D.R., Jackson, C.A.L., Lunn, R.J., Schlische, R.W., Shipton, Z.K., Wibberley, C.A.J., and Withjack, M.O., 2010. A review of recent developments concerning the structure, mechanics and fluid flow properties of fault zones. Journal of Structural Geology, 32 (11): 1557-1575. 68 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Mobile geologic mapping at the Wisconsin Geological and Natural History Survey REHWALD, Matthew1, AMES, Carsyn1, BREMMER, Sarah1, FITZPATRICK, William1, STEWART, Eric1, BATTEN, William1 and MAUEL, Stephen1 1 Wisconsin Geological and Natural History Survey, University of Wisconsin-Madison Division of Extension, 3817 Mineral Point Road, Madison, WI, 53705 The Wisconsin Geological and Natural History Survey (WGNHS) currently collects and analyzes data using a number of mobile applications for different purposes. Field data collection has become a tool for collection of new data and the verification of existing data. It has allowed the survey to create an automated pipeline to capture photos, notes, as well as record location information into one central location for a respective project. We had 4 objectives to implement while incorporating mobile applications. The application had to be 1) easy to use, 2) efficient, 3) easy to update, and 4) capable of displaying many datasets in the field. At the WGNHS we utilize mobile field applications for the collection of new data and the verification of existing map data. A mobile field application has the advantage of making many different data sets available to the user in the field within the flexible scale of a mobile GIS application. The incorporation of other mobile applications (FieldMove Clino) for data collection can increase efficiencies and are a vital to aid interpretations. Mobile field applications allow for field reconnaissance from almost anywhere. When considering a large project with a lot of data, increasing efficiency in field mapping techniques without compromising quality is important. Automating much of the data collection and data transfer eliminates the need for individuals to spend time cataloging digital pictures, copying field notes, and uploading field data. It’s a great advantage to be able to easily update or add additional map layers and data, and to see the data already collected. A visual display of data collection progress is useful in time management and project planning. The ease of which an application can be updated consumes time and affects project budget. Ease of use is also important, Accessibility and technical expertise should not be barriers to data collection. The ease of use of a mapping application has positive impacts the project participants, the project budget, and the project output. 69 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Diagram of the flow of geologic data from the source to and from the mobile application. Managing the data allows for customization of the functionality and the display of the data. 70 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Quaternary Geology of Wisconsin at a scale of 1:500,000 (in review) ROSE, Caroline1, RAWLING III, J. Elmo1, CARSON, Eric C.1, ATTIG, John W.1, MICKELSON, David M.1, MODE, William N.2, JOHNSON, Mark D.3, and SYVERSON, Kent M.4 1 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 University of Wisconsin–Oshkosh Department of Geology, 645 Dempsey Trail, University of Wisconsin– Oshkosh, Oshkosh, WI 54901 3 Department of Earth Sciences, University of Gothenburg, Gothenburg, Sweden 4 University of Wisconsin–Eau Claire Dept. of Geology, 145 Phillips Hall, University of Wisconsin-Eau Claire, Eau Claire, WI 54702 2 In 2023 the Wisconsin Geological and Natural History Survey staff expect to publish a new statewide compilation map of Quaternary geology at a scale of 1:500,000. A preliminary version is presented here by the principal cartographer. Pre-existing statewide coverages of the surficial geology are limited to Chamberlin’s 1881 map of Quaternary formations and Hadley and Pelham’s 1976 map of glacial deposits at 1:500,000, which differentiates only six map units. No modern compilation of the surficial geology of the state at a scale of 1:500,000 or larger has been completed before. Figure 1. Statewide Quaternary geologic mapping in Wisconsin: Left: Chamberlin’s 1881 “Quaternary Formations of Wisconsin”. Center: Hadley and Pelham’s 1976 “Glacial Deposits of Wisconsin”. Right: Draft polygons of 1:500,000 scale surficial geologic map being compiled by WGNHS geologists. This effort began in 2019 due to a one-time funding opportunity from the US Geological Survey’s National Cooperative Geologic Mapping Program. Authors compiled previous mapping at 1:100,000 scale for 44 of Wisconsin’s 72 counties, along with partial mapping at the 1:100,000 scale and/or mapping at the 1:250,000 scale for 13 additional counties. Some areas had no prior mapping available at detailed scales. New map units have been developed for the 1:500,000 scale and are divided into glacial and nonglacial sediment that is characterized by lithology and subdivided by geomorphology. Glacial deposits are mapped at the formation level following the WGNHS Lexicon of Pleistocene Stratigraphic Units. We use color hue to differentiate among the various glacial formations by source areas with green groupings derived from the Superior basin, blue groupings from the 71 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Michigan basin. We assign the darkest colors to the strings of moraines and hummocky till marking the extent of glacial lobes of the most recent Wisconsin Glaciation. Nonglacial Quaternary units are generally shown in warm colors, including modern alluvium, colluvium, lake deposits, and meltwater stream deposits, with small pockets of terraces which are highlighted along major river valleys. The Driftless Area in southwestern Wisconsin shows the dendritic patterns of eroding colluvium along branching alluvial tributaries with windblown silt on the uplands. Some large deposits of organic sediment and areas of exposed or thinly covered bedrock are included at this scale. This map layout is being produced entirely in ArcGIS Pro, which is a relatively new layout process for our office. We are organizing the GIS data according to the USGS standard Geologic Map Schema (“GeMS”), and we make use of this data structure to draw the unit description text in the Explanation of Map Units (legend) directly from a table in the geodatabase using a dynamic text element. This saves us from the extra work of synchronizing the layout text with the database text. Although ArcGIS Pro does not natively offer an easy solution for geologic map legends, we have been able to find a series of work-arounds to achieve the desired legend layout. References Chamberlin, T.C., 1881. General map of the Quaternary formations of Wisconsin, plate 2 of Atlas of the Geological Survey of Wisconsin: [Madison, Wisc.], Wisconsin Geological Survey, scale approximately 1:960,000. Hadley, D.W., and Pelham, J.H., 1976. Glacial deposits of Wisconsin—Sand and gravel resource potential: Wisconsin Geological and Natural History Survey Map M061: 19 p., 1 pl., scale 1:500,000, https://wgnhs.wisc.edu/catalog/publication/000385 [Previously Map 10.]. Acomb, L., Attig, J.W., Baker, R.W., Brownell, J., Clayton, Lee, Fricke, C., Frolking, T.A., Frye, J.C., Hemstad, C., Jacobs, P.M., Johnson, M.D., Knox, J.C., Leigh, D.S., Mason, J.A., McCartney, M.C., Mickelson, D.M., Mode, W.N., Muldoon, M.A., Need, E.A., Schneider, A.F., Simpkins, W.W., Socha, B.J., Syverson, K.M., Willman, H.B., 2011. Lexicon of Pleistocene Stratigraphic Units of Wisconsin: Wisconsin Geological and Natural History Survey Technical Report 001: 180. 72 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Tips from a GIS Specialist: Moving maps to GeMS, and a utility for georeferencing quadrangles ROSE, Caroline1 1 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 USA The USGS has recently been requiring that geologic mapping deliverables use their new standard database format, called the Geologic Map Schema, or “GeMS.” The Wisconsin Geological and Natural History Survey (WGNHS) began converting geologic maps into the GeMS format four years ago. I will offer a brief overview of GeMS and will use the GIS data for the Geology of LaCrosse County map (available for download on our website) to demonstrate how GeMS captures the components of a geologic map in geodatabase format. We have created several documents to facilitate the process of migrating maps into GeMS, and we have made them available in this Github repository: https://github.com/wgnhs/gems. My advice to anyone beginning this process is to first consult our “Workflow Overview” document for a high-level summary of the steps. When completing the GeMS-specified attributes, the “Quick-Reference Sheets” are a convenient arrangement of the GeMS documentation, with each layer or table printed on a separate reference sheet, to put focus on one layer or table at a time. To help verify that a GeMS database is complete, the “GeMS Fields Checklist” is designed to help in confirming completion of GeMS attributes. Two of our documents address the process of authoring metadata for a GeMS geodatabase. The document titled “Metadata For GeMS Maps - Step by Step in ArcCatalog” is a guide to starting FGDC metadata in ArcCatalog before using the USGS-provided metadata script. The “Metadata Summary for GeMS Fields” is a reference to show where GeMS attributes appear in the FGDC metadata, as produced by the metadata script. All of these documents are housed on our github page, along with other resources such as python scripts and slides from various presentations. We are making it a priority to share these with other GeMS users; we hope these resources are useful to other organizations working through the process of converting maps into GeMS. I will also briefly summarize how we have involved the GeMS format in our map layouts in ArcGIS Pro by drawing from the Description of Map Units table to automatically lay out the legend using Dynamic Text elements. In the second half of this talk, I will give an overview of a semi-automated utility for georeferencing maps, especially USGS quadrangles. The software is called QuadG+ and was developed by USGS and University of Wisconsin collaborators to build the Historical Topographic Map Collection. It is available for free download at https://geography.wisc.edu/quad-g/ and has proven useful to Wisconsin survey staff for georeferencing maps with field notes. 73 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. The free software QuadG+ automatically detects the corners and other control marks in a scan of a USGS quadrangle References Burt, James E., Jeremy White, Gregory Allord, Kenneth Then, A-Xing Zhu, 2022. Quad-G+: Automated Georeferencing of Scanned Map Images User Manual Version 2.13 December 2022. University of Wisconsin – Madison. Accessed March 27, 2023. https://uwmadison.app.box.com/s/tkccw1j5u3ensn2e10hrl1eiek78z9r6/file/1125666147300. 74 �Proceedings of the 69th ILSG Annual Meeting – Part 1 New work developing Keweenaw geoheritage awareness ROSE, William1 1 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 U.S.A. Telling Keweenaw Geostories in ~ten minutes. Old stories of Keweenaw geohistory have been made into web-based illustrated summaries meant to fill awareness of geoheritage from literature sources. About 8-15 minutes long with ~20 illustrations, these stories tell about the Ontonagon Boulder, Douglass Houghton, Louis Agassiz, Jane Schoolcraft and Hiawatha, Pasties and Keweenaw Miners, Big Annie and the 1913 Strike, Discovery of the Keweenaw Fault, the Green Rock at Copper Harbor, Ben Franklin and Lake Superior, and the Discovery of the C&H Conglomerate. The stories may be viewed online (https:// vimeo.com/showcase/9801619). They show how local history is guided by geology. They are intended to supplement local and statewide awareness and pride. Bringing the Boulder Home to the UP. The Ontonagon Boulder was a legendary float of native copper 75 �Proceedings of the 69th ILSG Annual Meeting – Part 1 which was on the west branch of the Ontonagon River until 1847 (https://vimeo.com/showcase/9801619/ video/785968264). The word of mouth of this unusual precious rock led to widespread interest, but it was difficult to move. Dispute over the ownership of the Boulder was spirited, and eventually it ended up in Washington DC at the Smithsonian Mineral Science Museum. The Boulder is considered a sacred object by Ojibway (Erik Redix, 2017, American Indian Quarterly, 41 (3)). Repatriation of the boulder to the UP was applied for, but rejected by the Smithsonian in 2000. UP residents and tourists have no access to this iconic legend. Currently (for decades) the boulder resides out of public view. We propose a loan of the boulder to allow it to visit museums such as Cranbrook, Univ of Michigan and the AE Seaman Mineralogical Museum, partner of the Keweenaw National Historic Park Building a Statue of a feminist labor leader. Anna Klobuchar Clemenc was a feminist labor leader in Calumet during the miners’ strike of 1913 (https://vimeo.com/showcase/9801619/video/748833299). She had fame for her leadership of labor parades when she wrapped herself in the American Flag to inhibit the violent confrontations. Tall and homely, “Big Annie” used her personnage advantageously and allied with the Western Federation of Miners. She worked with Mother Jones and with the women’s vote efforts in Washington. She was the first member of the Michigan Women's Hall of Fame. The Michigan legislature has officially named June 17 as “Big Annie Day”. A bronze life-sized statue of her is planned for permanent display in Red Jacket, outside of the Calumet Opera House and one block away from the Italian Hall. For more info: https://www.facebook.com/profile.php?id=100090193837168 76 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Outcrop Scale Mapping Utilizing High-Accuracy GNSS with MnDOT’s Virtual Reference Station (VRS) Network: Minnesota Examples SCHULZ, Roger1 1 Big Rock Exploration Geologic mapping has long been utilized to visualize the underlying geology of a region. An important tool used in geologic mapping are those that resolve the mapper’s locations at a given time. The tools used to locate a mapper have advanced greatly since the time of pace and compass, chains, and grids. With that advancement comes ever more accurate location data. One of the most common modern mapping tools utilizes satellite networks to send a signal from which location data is calculated on a consumer grade handheld GPS unit. While handheld GPS’s are useful in mapping moderate to small scale (e.g., 1:5000 or 1:24,000), the accuracy limitations of these units are not capable of resolving outcrop-size maps (e.g., <1:250 scale). Given the limited outcrop in places like the Lake Superior region, it is necessary to extract all possible data from a given outcrop, lending greater importance to small-scale maps. Attaining a level of location accuracy needed for such outcrop scale mapping requires additional real-time corrections of satellite data. The Global Navigation Satellite System (GNSS) encompasses three major satellite networks operated by the USA (GPS), Russia (GLONASS), and the EU (Galileo). When utilized within the GNSS framework, it is possible to have reliable satellite coverage anywhere in the world, a requirement for accurate location data. GNSS functions via one-way communication of radio waves from the satellites to a receiver that calculates distance from the satellite to the receiver. Distance calculations based on the speed of the signal (c) and the time differential (Δt) between the signal being sent then picked up by the receiver (D = c • Δt). To triangulate the position of the observer, this calculation must be solved by multiple satellites. This results in positional data that is generally accurate to 10m in the horizontal, at best. The reason for the inaccuracy is that the atmosphere interferes with the speed of the signal resulting in a delay. It is possible to achieve more accurate data by correcting for this differential delay using established ground-based networks. Differential correction using Virtual Reference Station (VRS) utilizes base stations at control monuments that continually collect positional data generating an average position that can be used to determine the degree of atmospheric delay. When used in a network of base stations, the average atmospheric delay for an area can be determined. The regional delay, or differential, can be communicated to a handheld unit over an internet connection, thereby eliminating the effect of atmospheric delay. Positions can then be determined to centimeter-scale accuracy, a requirement of mapping outcrop scale features. MNDOT has implemented a statewide Virtual Reference Station (VRS) network with over 140 base stations over control monuments whose purpose is to correct for the atmospheric delay and generate high-accuracy GNSS datasets.2 This network is free to use for anyone. Figure 1 below is a case study from South Pass, Wyoming where a trench was mapped at 1:250 using a Trimble Geo 7x. The trench this study area contained auriferous quartz veins and barren quartz veins anastomose along a pair of sheared faults separated by several meters and are connected ladder veinlets. Without the decimeter-scale accuracy of the corrected positional data, it would not have been possible to accurately locate the geology, geochemical samples, or structural data within the trench and the adjacent outcrops. Such an approach could be extremely useful in visualizing complex intrusive outcrops in the Duluth Complex, tracing of the contacts of lava flows and interflow sediments along the shore of Lake Superior, and veins and stockworks within Archean rocks. 77 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Trench Mapping and Sampling for Relevant Gold Corp at the Golden Buffalo Project - South Pass, Wyoming. by Big Rock Exploration LLC References GNSS Timing and Atmospheric Interferences: How GNSS Is Solving These Problems. Global GPS Systems, 24 Jan. 2023, https://globalgpssystems.com/gnss/gnss-timing-and-atmospheric-interferences-how-gnss-issolving-these-problems/. Land Management. MnCORS Network - Land Management - MnDOT, https://www.dot.state.mn.us/surveying/cors/index.html. Understanding RTK VRS Networks. Global GPS Systems, 24 Jan. 2023, https://globalgpssystems.com/gnss/understanding-rtk-vrs-networks/. 78 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Geology and geochemistry of the basal North Shore Volcanic Group and Midcontinent Rift Intrusive Supersuite, Cook County, MN, USA SEVERSON, Allison R.1, NOWARIAK, Eric S.1, LARSON, Phillip C.2 1 Minnesota Geological Survey, Department of Earth and Environmental Sciences, University of Minnesota-Twin Cities, MN, USA 2 Vesterheim Geoscience PLC, Duluth, MN, USA Northeastern Minnesota preserves complex relationships between Mesoproterozoic volcanic flows and comagmatic gabbroic to granophyric intrusive rocks associated with the ca. 1.1 Ga Midcontinent Rift System (MRS), as well as Paleoproterozoic metasedimentary rocks. Over the last two years, bedrock mapping of nine 1:24K quadrangles in northeastern-most Minnesota (Fig. 1) has elucidated some of these relationships between the Rove Formation, Logan sills, Puckwunge sandstone, reversely polarized North Shore Volcanic Group (NSVG), and gabbroic and granophyric rocks of the Midcontinent Rift Intrusive Supersuite. Results described herein are based on field and thin section observations, and associated geochemistry, which will be compiled and published as part of the Minnesota Geological Survey’s County Geologic Atlas Series. Volcanic rocks lie conformably on top of the Puckwunge sandstone in the eastern map area (Fig. 1). In the western part of the map area, the Crocodile Lake Gabbro (CLG) is in contact with the Paleoproterozoic Rove Formation to the north, with the Rove being highly deformed, metamorphosed, and partially melted proximal to the contact. South of the CLG, is the coeval Cucumber Lake Granophyre (CLGp), which is in contact with the Grand Portage Lavas (GPL), Esther Lake Lavas (EL), and Hovland Lavas (HL) of the NSVG to the south. The NSVG youngs from north to south, and transitions from mafic to more felsic from north to south which is most evident in the transition from the GPL to the overlying EL (Fig. 1). Geochemically, this sequence evolves along a strong tholeiitic trend (Fig. 2). Lithologic and geochemical patterns suggest the >1108 Ma GPL, EL, and HL were likely sourced from a long-lived, evolving magma. The basal GPL amygdaloidal basalt preserves 5 - 75 cm long pillows with somewhat enigmatic siliceous, carbonate, and glassy selvages that also preserve hyaloclastic and perlitic textures. These flows are geochemically primitive and contain abundant altered olivine, pyroxene, and oxide phenocrysts. The pillowed basal unit grades into thick, massive to ophitic basaltic and basaltic andesite flows of the EL. The transition from the GPL is also marked by a change in trace element geochemistry from an enriched mantle to a more depleted mantle signature. The base of the HL consists of a package of strongly glomeroporphyritic, amygdaloidal andesites and basaltic andesites transitioning to porphyritic rhyolite and icelandite. Porphyritic basaltic to andesitic lavas in the westernmost map area also preserve pillow structures, but these flows vary in thickness and extent, suggesting aqueous subbasins within the HL volcanic basin. Intercalated throughout the HL are abundant dikes and sills of ultraphyric diabase containing 15-60% of >5 mm plagioclase phenocrysts within a basaltic, locally ophitic very fine-grained groundmass. These intrusives are interpreted to be hypabyssal and locally cut across volcanic stratigraphy. Though these dikes and sills are endemic to the area, temporal relationships between these intrusives, the surrounding volcanics, and the Brule-Hovland Gabbro are unknown. The ca. 1107 Ma CLG and the CLGp comprise some of the earliest known rocks within the intrusive Duluth Complex. Basal gabbroic cumulates of the CLG grade into dioritic-monzonitic rocks of the Crocodile Lake “Mixed Zone”, below the contact with the overlying CLGp. This Mixed Zone is typified by complex dikes and plagioclase cumulate rocks, rich in micrographic interstitial felsic mesostasis. Abundant quench textures and pegmatitic zones, as well as distinct geochemical patterns 79 �Proceedings of the 69th ILSG Annual Meeting – Part 1 suggest the Mixed Zone is a “cap” to the CLG rather than a gradual transition to the CLGp. REE patterns and Eu anomalies within these coeval intrusives suggest liquid immiscibility between mafic and felsic components of the source magma may have played a significant role in their genesis (Fig 3.). Other intrusive gabbroic rocks include the texturally varied Brule-Hovland Gabbro, which cross-cuts the HL. Figure 1. Regional geologic map of northeastern Cook County, MN. Ongoing partially USGS-funded STATEMAP projects outlined with bold lines. Generalized geology is from MGS miscellaneous map series M-119. Figure 2. AFM diagram of volcanic rocks. Figure 3. Chondrite-normalized REE diagram of Crocodile Lake and Cucumber Lake intrusives, based on Sun and McDonough, 1989. 80 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Exploring the application of full tensor magnetic gradiometry to better define conduit type NiCu-PGE targets SMITH, Jennifer1, TSCHIRHART, Victoria1, TUCK, Loughlin2, ENKIN, Randy1, and ROYGUAY, David3 1 Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8 Defence Research and Development Canada, Ottawa 3 SBQuantum,Sherbrooke, QC, J1H 1Z1 2 Magmatic Ni-Cu-PGE sulfide deposits are often associated with small conduit- or chonolithtype intrusions. These deposit types are notoriously challenging exploration targets owing to: 1) their small size, 2) lack of alteration halo or distal footprint, 3) complex and variable morphology, and 4) unpredictable depositional sites of sulfides (Barnes 2023). Furthermore, mafic rocks commonly retain significant remanent magnetization which, if not detected, can result in inaccurate modelling and targeting of these deposits. With a significant increase in the global production of Ni forecasted for the transition to a low-CO2 future, these deposit types will likely become an increasingly important source of Ni, both in Canada and globally. With fewer new discoveries being made, despite increased exploration expenditure, new methods and knowledge are needed to facilitate successful exploration at the regional and deposit scales and to ultimately secure a stable Ni supply. Historically, exploration has traditionally relied on geophysics (gravity, magnetics, electromagnetics), to identify potential mafic and/or ultramafic host intrusions, with airborne magnetic surveying dominating due to its low cost, and ability to survey vast areas rapidly and systemically. Although there is incredible value in Total Magnetic Intensity (TMI) data there are numerous limitations to this approach (e.g. non-uniqueness, scalar measurements, can’t distinguish remanence from induced field). The full tensor magnetic gradiometry (FTMG) technique, which measures the full magnetic gradient tensor at each measurement point, overcomes many of the limitations of TMI data. Advantages of FTMG include: (a) superior resolution of near-field sources, (b) enhanced detectability at low-magnetic latitudes, (c) automatic removal of the regional field and diurnal variations, and (d) additional target information from a single flight line. FTMG can provide a more complete picture of the subsurface magnetic properties and improved discrimination between magnetic sources. This leads to improved imaging of complex structures, more accurate models of the subsurface, and improved understanding of geological processes. While quantum FTMG is in use by industry, practicalities relating to the system hamper its widespread deployment. Currently, existing quantum FTMG relies on SQUID technology for large scale airborne surveying. The application of SQUID technology has shown great benefits due to the enhanced sensitivity and fidelity of the system. However, these systems typically weigh ~270 kg and require extremely low sensor temperatures, making them impractical for ground and uncrewed aerial vehicle (UAV) surveying. These limitations have warranted the development of a complimentary ground and UAV quantum FTMG system such as the diamond-based quantum magnetometer in development by SBQuantum. This rugged and compact system leverages quantum properties of nitrogen vacancy (NV) centres in a diamond to provide highly accurate, quantum-based FTMG measurements. The Geological Survey of Canada (GSC) is in the early stages of establishing a new collaborative partnership with Defence Research and Development Canada (DRDC), SBQuantum, and numerous other industry and academic partners. The aim of this partnership and wider project is to derisk quantum magnetic gradiometer use across Canada with the purpose of facilitating widespread adoption by the Canadian exploration industry, academia and the military. This will be achieved 81 �Proceedings of the 69th ILSG Annual Meeting – Part 1 through the field testing and validation of the ruggedized quantum FTMG system developed by SBQuantum. As part of this project, SBQ’s quantum magnetic gradiometer will be deployed on several Canadian critical mineral systems, allowing comparison with traditional airborne and/or ground total magnetic field systems and non-quantum FTMG systems. As part of this, a detailed study will be undertaken on the Ni-Cu-PGE bearing Escape Lake intrusion in northern Ontario, which presents as a complicated magnetic signal that is strongly affected by remanent magnetization and associated with the 1.1. Ga Midcontinent Rift. With conventional total field geophysical methods unable to address the challenging features which are often characteristic of small, conduit-type magmatic sulfide deposits, this case study will explore the use and application of quantum FTMG in the context of improving targeting of conduit type Ni-Cu-PGE deposits. This study will be the first to generate publicly accessible quantum FTMG data over critical mineral deposits in Canada and will act to improve exploration capacity by validating tools useful for critical metal deposits whose complex geophysical expressions are not easily resolved by traditional geophysical techniques. The increased accuracy of these quantum technologies, which map the magnetic field at an enhanced scale, provide the ability to resolve the complexity of these deposits. Providing enhanced tools to facilitate exploration and delineate deposits better will aid with the identification of new Canadian deposits of critical metals needed for the lower carbon and digitized economy supply chain. This will aid Canada’s Critical Minerals Strategy set forth in the 2022 Federal Budget. References Barnes, S.J., 2023. Lithogeochemistry in exploration for intrusion-hosted magmatic Ni–Cu–Co deposits. Geochemistry: Exploration, Environment, Analysis, 23(1): geochem2022-025. 82 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Record of an Ancient Meteorite Impact Buried Beneath the Twin Cities, MN STEENBERG, Julia R. 1, and RUNKEL, Anthony C. 1 1 Minnesota Geological Survey, 2606 W. Territorial Rd., St. Paul, MN 55114 USA An impact crater is proposed in the southeast part of the Twin Cities metropolitan area, 11 miles (18 km) south of St. Paul within an area with significant residential and industrial development. The crater lies within a predictable package of Paleozoic sedimentary rocks in the Twin Cities structural basin where near its center includes 14 formations with a total thickness of about 1,200 feet (365 meters) (Mossler, 2008; Mossler, 2013). Paleozoic formations are characterized by widespread layers of sandstone, shale, and carbonate deposited in shallow seas during the Cambrian and Ordovician Periods (500 to 450 Ma). They are underlain by Mesoproterozoic (1,100 Ma) sedimentary and volcanic rocks of the Keweenawan Supergroup associated with the Midcontinent Rift. Paleozoic rocks in this area have limited exposure along the Mississippi and Minnesota River bluffs, roadcuts, and rock quarries, but elsewhere are buried beneath a variable thickness of Quaternary glacial sediments. Without extensive exposures, a variety of subsurface datasets are used for bedrock mapping including core, drill cuttings, geophysical logs, passive seismic stations, and driller’s descriptions from water well records. While mapping the bedrock geology of Dakota County, an area of discordance with the surrounding Paleozoic stratigraphy was observed in geologic cuttings samples, and corroborated with additional cuttings, geophysical logs and water wells driller’s records. Drill samples reveal as much as 575 ft (175 m) of anomalous sandstone, siltstone and shale with some intervals containing abundant cloudy and fractured quartz sand grains. The samples are from an area entirely buried by several hundred feet of glacial deposits within a deep buried channel carved into the surrounding bedrock layers adjacent to the Mississippi River near the city of Inver Grove Heights. Beneath the anomalous sequence of strata and in additional samples near the site, local Cambrian and Mesoproterozoic stratigraphic layers are recognized but are out of the usual stratigraphic order and in places entirely overturned. Microscopic investigation has resulted in the detection of shocked metamorphic features including planar deformation features (PDFs) in the fractured quartz grains, confirming the impact origin of this structure (Fig. 1). As such, this area is referred to as the Pine Bend Impact Structure (PBIS) (Steenberg, in prep). Based on the available geologic data near the site and current models of crater formation from similarly sized structures in layered sedimentary target rocks we interpret this feature to be a complex crater, approximately 4 km wide with an apparent central uplift and possible terraced rims (Grieve, 1991). The total disturbed area may be as large as 9 square miles (23 square kilometers). Based on published crater- to- meteor size ratios, the size of the meteor is estimated to be several hundred meters in diameter (Grieve and Pilkington, 1996). Due to its location, within a buried bedrock valley, the upper sequence of this structure has been removed by erosion, making it difficult to precisely date the impact. It may be as old as Late Cambrian (~490 Ma), having occurred during or after deposition of the Jordan Sandstone based on the age of the overturned strata and the apparent lack of carbonate from the overlying Prairie du Chien Group in the samples. We have also collected a pebble with PDFs from strata approximating the Jordan-Prairie du Chien contact in an outcrop about 10 kilometers from the crater. If this pebble is ejecta from the PBIS, it also supports a latest Cambrian or very Early Ordovician age of impact. This would make the PBIS older than known craters in surrounding states which are Ordovician and younger (French et al., 2004; French et al., 2018). The dynamic nature of our planet has left us with a small sample size of terrestrial impact structures, nearly 200 confirmed impact structures are currently recognized on Earth (Gottwald et al., 2020). Although Minnesota has known impact debris from the Sudbury Impact Structure, this would 83 �Proceedings of the 69th ILSG Annual Meeting – Part 1 be Minnesota’s first documented crater, giving us a rare opportunity to better understand the important geological and biological effects of meteorite impact events on Earth. Figure 1. Photomicrographs of mounted quartz sandstone rock chips from a cuttings sample, sample depth is 525 feet. A- Two sets of planar features and feather features. B- One set of decorated planar deformation features. Photos by L. Ferriere, Natural History Museum, Vienna, Austria. References French, B.M., Cordua, W., and Plescia, J.B., 2004. The Rock Elm meteorite impact structure, Wisconsin: Geology and shock-metamorphic effects in quartz. GSA Bulletin, 116: 200–218. French, B.M., McKay, R.M., Liu, H.P., Briggs, D.E.G., and Witzke, B.J., 2018. The Decorah structure, northeastern Iowa: Geology and evidence for formation by meteorite impact. GSA Bulletin, 130: 2062– 2086. Gottwald, M., Kenkmann, T., and Reimold, W.U., 2020. Terrestrial impact structure. In: TheTan-DEM-X Atlas, Part 1 and 2, Friedrich Pfeil, Munich, Germany. Verlag Dr. Grieve, R.A.F., 1991. Terrestrial impact: the record in the rocks. Meteoritics, 26: 175–194. Grieve, R.A.F., and Pilkington, M., 1996. The signature of terrestrial impacts. AGSO Journal of Australian Geology and Geophysics, 16: 399-420. Mossler, J.H., 2008. Paleozoic stratigraphic nomenclature for Minnesota. Minnesota Geological Survey Report of Investigations RI-65: 76, 1 pl. Mossler, J.H., 2013. Bedrock geology of the Twin Cities ten-county metropolitan area, Minnesota. Minnesota Geological Survey Miscellaneous Map M-194: scale 1:125,000. Steenberg, J.R., in prep. Bedrock geology, pl. 2 of Steenberg, J.R., project manager, Geologic atlas of Dakota County, Minnesota. Minnesota Geological Survey County Atlas C-57: 6 pls., scale 1:100,000. 84 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Magma Recharge and the distribution of Copper and Nickel in the Keweenaw Large Igneous Province STEINER, Alex1, PETERSON, Dean1, SWEET, Gabriel1 1 Big Rock Exploration, 2505 W. Superior Street, Duluth, MN 55806. The Keweenaw large igneous province (LIP) was formed over a protracted period of magmatism that emplaced Cu-Ni-PGE bearing mafic to ultramafic intrusions along the arcuate MidContinent Rift system, thus creating one of the largest critical mineral resources in North America. The magmatic activity associated with the Keweenaw LIP has been divided into a series of tectonomagmatic stages extending from at least 1115 Ma to 1080 Ma. Of the six stages of formation, significant orthomagmatic sulfide deposits were formed during Stage 1 (plume impact stage; 11151110 Ma), Stage 2 (early stage; 1110-1105 Ma), and the Stage 4 (the main stage; 1101-1094 Ma). Stage 1 and 2 intrusions in the Minnesota and Michigan are Ni-rich while those of stage 4 in the Duluth Complex are considerably more copper-rich. Here we discuss a potential mechanism of copper enrichment via magma recharge-evacuation-fractional crystallization (REAFC) where the parameters of differentiation are based upon a conceptual model for the formation of continental LIPs. It has been recognized that continental LIPs form in a series of phases that reflect the conditions of magma generation and differentiation prior to the eruption and eventual formation of continental flood basalts (Jerram and Widdowson, 2005). Early phases of LIP formation are dominated by more primitive lavas, that pass through a magma plumbing system that is immature and inefficient at differentiating magmas (Steiner et al., 2021). However, the magmatic plumbing system of the most voluminous eruptive phase is mature and capable of differentiating magmas to a considerable degree. The key difference between these two periods is the amount of magma recharge, which has a profound impact on the geochemical composition of the resultant magmas where compatible elements become buffered and incompatible elements become enriched (Lee, Lee and Wu, 2014). The relative Cu-enrichment of mineralized Stage 4 intrusions compared to earlier Ni-rich Stage 1 and 2 intrusions may be explained by several mechanisms. Mechanisms such as sulfide upgrading and high-R factors have been recognized as important contributors to Cu-rich mineralization (Peterson and Boerst, 2013). However, recent chemo-stratigraphic examinations of Keweenaw LIP lavas from the Keweenaw Peninsula have demonstrated that REAFC processes are controlling erupted lava compositions during the eruption of Stage 4 lavas (Davis et al., 2021). To test the effect of REAFC on the proportions of Ni and Cu that may be available to form an orthomagmatic sulfide deposit, REAFC geochemical modelling utilizing the equations of Lee et al. (2014) were performed on a generalized basaltic composition (MgO = 10%, Ni = 250 ppm, Cu = 116 ppm (Prinz, 1967)). Figure 1 demonstrates the liquid line of descent for Cu, Ni, and MgO during REAFC differentiation and pure fractional crystallization. During pure fractional crystallization, MgO and Ni behave compatibly, gradually decreasing in concentration with continued differentiation while Cu gradually increases. However, during REAFC differentiation, both Ni and MgO become buffered while Cu becomes decoupled, increasing in concentration while Ni and MgO remain the constant. The consequence of this decoupling is that Cu can become considerably enriched relative to Ni, thereby producing a magma that contains greater than anticipated Cu concentrations compared to pure fractional crystallization. When this Cu-enriched magma reaches sulfur saturation, the subsequent sulfide magma would have a greater abundance of Cu to scavenge, resulting in the Cu-rich sulfide deposits observed in in the Duluth Complex. 85 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. REAFC calculations (Lee, Lee and Wu, 2014) for a generic basalt. Model parameters are recharge/evacuation = 0.43, assimilation = 0.07, fractional crystallization = 0.5. Crystallizing phases were olivine (25%), plagioclase (65%), and clinopyroxene (15%). References Davis, W.R. et al., 2021. Geochemical, petrographic, and stratigraphic analyses of the Portage Lake Volcanics of the Keweenawan CFBP: implications for the evolution of main stage volcanism in continental flood basalt provinces, Geological Society, London, Special Publications: SP518-2020–221. doi:10.1144/SP518-2020-221. Jerram, D.A. and Widdowson, M., 2005. The anatomy of Continental Flood Basalt Provinces: geological constraints on the processes and products of flood volcanism, Lithos, 79(3): 385–405. doi:https://doi.org/10.1016/j.lithos.2004.09.009. Lee, C.-T.A., Lee, T.C. and Wu, C.-T., 2014. Modeling the compositional evolution of recharging, evacuating, and fractionating (REFC) magma chambers: Implications for differentiation of arc magmas, Geochimica et Cosmochimica Acta, 143: 8–22. doi:10.1016/j.gca.2013.08.009. Peterson, D. and Boerst, K., 2013. Twin Metals Minnesota’s Maturi Deposit, in Cu-Ni-PGE Deposits of the Duluth Complex, Geology and Development: Precambrian Research Center, Workshop on the Copper, Nickel, Platinum Group Element Deposits of the Lake Superior RegionOctober 6-13, 2013, Field Trip Guidebook: 45–57. Prinz, M., 1967. Geochemistry of basaltic rocks: trace elements. In’, Basalts, 1: 271–333. Steiner, R.A. et al., 2021. Initial Cenozoic Magmatic Activity in East Africa: New Geochemical Constraints on Magma Distribution within the Eocene Continental Flood Basalt Province, Geological Society, London, Special Publications: SP518-2020–262. doi:10.1144/SP518-2020-262. 86 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Relay zones in weakly folded and faulted Paleozoic strata and their role localizing Mississippi Valley-type mineralization, southwest Wisconsin, USA STEWART, Eric1, FITZPATRICK, William1, and AMES, Carsyn1 1 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Road, Madison, WI, 53705 Folds and faults have long been known to play a role in localizing Mississippi Valley-type zinclead mineralization in the historic Upper Mississippi Valley base metal district (UMVD) of southwestern Wisconsin. However, a simple correlation between mineralization and map-scale structures is overly simplistic since it does not explain why mineralization often occurs only along isolated portions of folds and faults. New 1:24,000 scale geologic mapping as part of the United States Geological Survey Earth Mapping Resources Initiative (EarthMRI) in the Stitzer region of the northern UMVD was initiated to improve understanding of the relationship between folds, faults, and mineralization. The Mineral Point anticline is the dominant structure in the Stitzer area (Figure 1). It is an asymmetric, northwest-trending gentle fold with a maximum amplitude of around 180 feet. The fold deforms platform Cambrian and Ordovician siliciclastic and carbonate strata, and contains several doubly plunging segments. Structural highs along the fold (Figure 1) correspond to aeromagnetic anomalies (Daniels and Snyder, 2002). Deformation bands in sandstone are common along the more steeply dipping northeast limb of the fold. The asymmetry of the fold and the correspondence of structural highs to aeromagnetic anomalies suggest the Mineral Point anticline is a forced fold, forming from thrust reactivation of a buried Precambrian fault. At depth near the Precambrian basement, the segments of the Mineral Point anticline probably transition into fault segments. Simple 2D kinematic modeling suggests contraction is highest near the base of the overlying folded section. If deformation bands accommodate some of the contraction in the basal siliciclastic sequence, then significant numbers of deformation bands are probably present low in the Paleozoic section. Mineralization and historic mining are heavily concentrated where two segments of the Mineral Point anticline overlap, and a third smaller anticline terminates (Figure 1). The area between the overlapping segments of the Mineral Point anticline is interpreted to represent the area above a relay zone between thrust segments. As mineralizing brines approached the Mineral Point anticline from the south, flow was probably altered due to the abundance of impermeable deformation bands. Flow conduits developed in the relay zone between fault-fold segments, focusing the brines upward and concentrating mineralization. References Carlson, J., 1961. Geology of the Montfort and Linden Quadrangles, Wisconsin, in Geology of parts of the Upper Mississippi Valley zinc-lead district. U.S. Geological Survey Bulletin 1123–B: 95– 138, 2 pls. Daniels, D. and Snyder, S., 2002. Wisconsin aeromagnetic and gravity maps and data; a web site for distribution of data. U.S. Geological Survey Open-File Report 2002-493. Taylor, A., 1964. Geology of the Rewey and Mifflin quadrangles, Wisconsin, in Geology of parts of the Upper Mississippi Valley zinc-lead district. U.S. Geological Survey Bulletin 1123–F: 279–360, 2 pls. West, W., 1971. Geologic map of the Ellenboro quadrangle, Grant County, Wisconsin. U.S. Geological Survey Geologic Quadrangle Series 959: 1 pl. 87 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Figure 1. Simplified structure contour map of the base of the Ordovician Platteville Formation. Mines are concentrated in the SE portion of the map near the junction of three anticline-syncline pairs. Additional data sources include the Mineral Development Atlas, Carlson (1961), West (1971), and Taylor (1964). 88 �Proceedings of the 69th ILSG Annual Meeting – Part 1 Deciphering the metamorphic and deformational history of the Hardwood Gneiss, Felch District, Michigan: Anomalously high-pressure rocks in the heart of the Penokean orogen TAYLOR, Madeline1 and BJØRNERUD, Marcia1 1 Geosciences Department, Lawrence University, Appleton Wisconsin 54911 The Neoarchean Hardwood Gneiss is a mafic granulite with minor metapelites, exposed over an area of about 6 km2 between the towns of Foster City and Hardwood, MI, 8 km southeast of the eastern end of the Paleoproterozoic “Felch Trough” (James, 1961). The area lies at the heart of the ca. 1.85 Ga Penokean orogen and within the superimposed Yavapai-age (1.75 Ga) ‘gneiss dome corridor’ (Drenth et al., 2021). In contrast to the primarily felsic gneisses of the region, which contain inherited zircons with ages of 3.8-3.5 Ga, the Hardwood Gneiss is mostly mafic and yields no zircons older than 2.7 Ga (Ayuso et al., 2018). Zircons from the Hardwood also record a period of growth between 2.2 and 1.9 Ga, which does not correspond to any known thermal events in the region (Cannon et al., 2018). Most notably, the Hardwood experienced much higher-pressure metamorphism than any other rocks in the region. Using a variety of geo- thermometers and -barometers, Peterson & Geiger (1990) concluded that the mafic rocks underwent two distinct metamorphic events, the first, ‘M1’, at 8.2-11.6 kbar and ca. 770°C, and a another, ‘M2’, at 6-10 kbar and 610-740°C, while the pelites experienced only the second. Maximum pressure estimates for the nearby Peavy metamorphic node, in contrast, are < 5 kbar (Attoh and Klasner, 1989). It is difficult to explain how the Hardwood complex, with its distinctive geochronologic and metamorphic signatures, came to be incorporated into the Penokean orogen. This study presents detailed field, petrographic and microstructural observations that may help constrain the origin and history of the Hardwood Gneiss. Peterson & Geiger (1990) identified three compositional units in the Hardwood: metabasite, amphibolite, and metapelite. The metapelite, a garnet-biotite schist, is clearly a distinct unit, exposed in the western end of the outcrop area, but our work suggests that the amphibolite and metabasite are both part of a heterogeneous igneous complex that included anorthositic, gabbroic and noritic horizons – perhaps a Neoarchean layered mafic intrusion. If this complex was of mantle origin, it could explain the absence of older Archean zircons. In addition to their compositional variety, the metamafic rocks display a wide range of metamorphic and deformational textures. In outcrop, they have a strong, apparently mylonitic, foliation that dips mainly NE but is somewhat variable in orientation and may be folded. In thin section, microstructures show that a combination of brittle and plastic deformation mechanisms contributed to the intense fabric. In plagioclase-rich horizons, the feldspars tend to be the largest crystals, apparently surviving as porphyroclasts. These show both cataclastic fracturing and highly distorted twins, a combination usually interpreted to indicate that deformation took place at mid-crustal depths and temperatures of ca. 500°, the brittle-plastic transition for feldspars. These intensely deformed rocks show evidence of only limited, and heterogeneous, recrystallization, either dynamic or static. This suggests that deformation was brief and that the rocks cooled quickly after deformation ceased. Garnet-bearing horizons within the mafic complex display especially remarkable textures. Clusters and trains of garnets, apparently broken -- and in some cases, shattered – are engulfed in a very fine-grained (<0.01mm) feldspathic matrix. The unusual shapes of some of the garnet fragments – including crescents and splinters – may indicate seismic fragmentation (Hawemann et al., 2019). The largest garnet fragments tend to have inclusion-free cores and ‘spongy’ poikilitic rims, while smaller fragments are commonly poikilitic throughout, with a notable 89 �concentration of opaque inclusions. In some cases, the edges of the small garnet fragments are so diffuse that they cannot be seen in plane light. Peterson & Geiger (1990) interpreted the poikilitic rims and small inclusion-rich garnets as records of a second metamorphic event, but we speculate that these are resorption features rather than overgrowths. ‘Spongy’ or ‘amoeboid’ poikilitic rims are known to form around granulite-facies garnets during the introduction of external fluids (Baxter et al., 2017), or when garnets are engulfed in pseudotachylyte melts (Austrheim et al., 1996). Because they form under disequilibrium conditions, such resorption rims are unlikely to yield reliable P-T results, and this could account for the large range of P-T conditions Peterson & Geiger (1990) suggested for their ‘M2’ metamorphic event. Given the shattered nature of the Hardwood garnets, we tentatively speculate that the very fine-grained material in which they occur could represent coseismic fault rock – either (devitrified) pseudotachylyte or/and ultracataclasite flushed with seismically-pumped fluids. In this interpretation, the Hardwood complex would have experienced only one high-P/T metamorphic event, followed by mylonitization, cataclasis and seismic faulting. If the quasi- brittle deformation of the feldspars – which seems to be part of the same event that shattered the garnets -- can be interpreted as occurring at ca. 500°, the deformation would have had to happen well after the granulite-facies event. However, feldspar plasticity can be suppressed in very dry rocks (e.g. Bjørnerud & Austrheim, 2004), so it is also possible that the seismic event(s) occurred under high-temperature conditions and possibly soon after the granulite facies metamorphism that formed the inclusion-free garnets. Whether any of these events occurred during the Penokean orogeny remains unclear. One possible constraint on the timing of the main foliation-forming event comes from the occurrence of an unfoliated mafic within a feldspathic layer in the Hardwood Gneiss on the south bank of the East Branch of the Sturgeon River. If this sill could be dated or linked geochemically with known mafic magmatic units in the region, this would establish the youngest possible deformation age for the Hardwood Gneiss. The Hardwood pelites are classic garnet-biotite schists, with asymmetric quartz-vein boudins and garnet ‘tails’ that suggest normal-sense shear along the NE-dipping foliation. Low- angle normal faulting would be the most efficient way to juxtapose deep crustal rocks like the Hardwood Gneiss against the shallower units that surround it. But the pelites, which represent supracrustal material and record amphibolite rather granulite-facies conditions, lie on the western edge of the Hardwood outcrop area, so top-to-the-east normal slip does not help explain how the high-pressure mafic units were brought up from depth. The area between the Felch Trough and the Niagara Fault is among the most structurally complex of parts of the Penokean/Yavapai orogen, with many anastomosing faults of different generations (Drenth et al., 2021). The orientations of structures within the Hardwood complex have almost certainly been altered since their formation by later faulting and tilting. For now, the Hardwood Gneiss remains a micro- terrane of unknown provenance within the Penokean orogen. References Attoh, K. & Klasner, J., 1989. Tectonics 8: 911-933. Austrheim, H., Erambert, M., & Boundy, T., 1996. Earth & Planetary Science Letters 139: 223-238. Ayuso, R., et al., 2018. Institute on Lake Superior Geology Proceedings 64: 7-8. Baxter, E., Caddick, M., & Dragovic, B., 2017. Rev. Min. & Geochem. 83, 469–533. doi: 10.2138/rmg.2017.83.15 Bjørnerud, M. & Austrheim, H., 2004. Geology 32: 765-768. Cannon, W.F., Schulz, K., Ayuso, R. & Mroz, T., 2018. ILSG Field Trip Guidebook 64: 1-38. Drenth, B., Cannon, W.F., Schulz, K., & Ayuso, R., 2021. Precam. Res. 369. doi: 10.1016/j.precamres.2021.106205 Hawemann, F., et al., 2019. Solid Earth 10: 1635-1649. doi: 10.5194/se-10-1635-2019 James, H., Clark, L., Lamey, C., & Pettijohn, F., 1961. USGS Professional Paper 310. Peterson, J. & Geiger, C., 1990. Journal of Geology 98: 273-281. 90 �Alteration Geochemistry Characterization and 3D Modeling of the Back Forty Volcanogenic Massive Sulfide (VMS) Deposit Stephenson, Upper Peninsula of Michigan, USA UPTON, Margaret1, MOOERS, Howard1, LARSON, Phillip2 1 Department of Earth and Environmental Sciences, University of Minnesota - Duluth, 1114 Kirby Drive, 102 Heller Hall, Duluth, MN 55812 2 Cleveland-Cliffs Hibbing Taconite Company. Hibbing, MN 55746 The Gold Resources Back Forty zinc-and-gold-rich polymetallic volcanogenic massive sulfide (VMS) deposit is located near Stephenson in the Upper Peninsula of Michigan. In general, VMS deposits are created in submarine environments when heated seawater circulates through oceanic crust and precipitates base and precious metals at or near the seafloor due to both cooling and neutralization of the ore fluid. In the process, host rock mineralogy and geochemistry are modified by both downwelling and upwelling hydrothermal fluids, which produces distinct alteration mineral assemblages and metasomatic changes within the host rock (Shanks and Thurston, 2012; Galley et al., 2007). Alteration mineral assemblages and their spatial distribution can be used to unravel the geochemical evolution of the system and help locate mineralization. The relationship between host rock and alteration mineralogy is not well understood or documented at the Back Forty Deposit but essential for understanding its genesis. This study 1) identifies the alteration mineral assemblage present at the Back Forty Deposit using lithogeochemistry results; 2) calculates the elemental gains and losses associated with hydrothermal alteration; 3) develops a working method for immediate qualitative alteration values from core logging; and 4) creates a model of the alteration zonation in coordination with the existing stratigraphy and mineralization. Core from nine drill holes (~ 2,950 meters), were logged to observationally identify alteration mineral assemblages, intensity, and their textural characteristics. The deposit, hosted in felsic pyroclastic rocks, shows mostly sericite alteration, which was used to establish an alteration intensity scale of 1-4 (1: weak, 5: intense). Major alteration mineral assemblages observed were sericite ± silica ± chlorite. Sericite alteration is pervasive throughout the deposit (2.5-3.5) with silica alteration intensity ranging from 1-2 and a few areas of silica flooding (3.5-4.5). Weak to moderate (1.5-2.5) chlorite alteration occurred throughout the deposit within the host rhyolite crystal tuff units as spotty chlorite coarse-grained agglomerations. Lithogeochemistry (1,300 count) was evaluated using the alteration box plot (Large et al., 2001) and the isocon mass balance method (Grant, 1986) (fig. 1), which are essential in quantitative assessment of chemical changes associated with alteration mineral assemblages and their spatial distribution, and the identification of hydrothermal fluid pathways and mineralization vectors within the deposit. In addition to using isocon results, alteration box plot results were modeled based on sericite, chlorite, and total alteration. The production of cross sections based upon this numerical modeling identify the alteration mineral zonation and its relative extent; this model is evaluated to determine the relationship between massive sulfide mineralization and alteration intensity. From these results, downhole core logging of alteration assigned numeric values (“quick log”) is evaluated as a method to make fast-paced exploration decisions while awaiting longer lead-time lithogeochemical results. By leveraging the process and combination of core logging for alteration mineralogy and intensity paired with geochemical analysis, it may be possible to determine the origin direction of hydrothermal fluid flow associated with mineral deposition and aid in future exploration efforts to locate additional mineralization on the Back Forty Deposit property. 91 �Results from this study show the sericite alteration is most significantly related to Zn and Cu mineralization, whereas the chlorite alteration is most associated with Ag, Au, and Pb. Distinctly depleted species associated with sericite alteration include Ba, Sr, Na2O, Rb; with chlorite commonly depleted in Br, Ba, Sr, Na2O. Least v. Intense Sericite Alteration 50 45 45 Be Au Ge Zn Ga U Cu 40 35 More Altered Cs Dy Ce La Ag Ni Cr 20 Tl Hf Sn Sc TiO2 10 MgO As Yb Tl Mo Zr Pr Pr Hg Cd U Ir 15 Lu Sc 10 Hf V TiO2 Te Ba Tb MnO Br Eu Ga Cs Co Nd K2O Be In 20 Er Ho Y Cr2O3 Nb Ge y = 0.996x R² = 0.996 Na2O CaO 10 15 20 25 30 5 NdYb Sm Br Rb Gd Dy Tm Eu Bi Ba AL2O3 K2O 35 40 45 50 Cr Ta Sr y = 1.067x R² = 0.999 Na2O CaO BaO Re 0 Least Altered Th MgO P2O5 0 Lu SiO2 La Ce Ir Ho Sr 5 Ni Pb In Cd 0 25 BaO Tb V Co Bi Gd Er Hg Fe2O3 5 Rb Pb Y MnO Sb Cu Fe2O3 Sm P2O5 15 Tm Cr2O3 Ta Se W 30 W Zr 25 Zn 35 SiO2 Nb 0 40 Th 30 Least v. Intense Chlorite Altered 50 5 10 15 20 25 30 35 40 45 50 Least Altered Figure 1. ISOCON plot of selected elements used to compare elemental gains and losses between least and most altered samples. Isocon line of best fit is defined by relative immobile. Species above the isocon line are enriched; below are depleted (Grant, 1986). References Aquila Resources (now Gold Resources), data current as of April 2021. Galley, A., Hannington, M., Jonasson, I., 2007. Volcanogenic Massive Sulphide Deposits. Geological Survey of Canada, Special Publication 5: 141-161. Grant, J. A., 1986. The Isocon Diagram: A Simple Solution to Gresens' Equation for Metasomatic Alteration. Economic Geology, v. 81: 1976-1982. Large, R. R., Gemmell, B.J., Paulick, H., 2001. The Alteration Box Plot: A Simple Approach to Understanding the Relationship between Alteration Mineralogy & Lithogeochemistry Associated with Volcanic-Hosted Massive Sulfide Deposits. Economic Geology, v. 96: 957-971. Shanks, W.C.P., Thurston, R., 2012. Volcanogenic Massive Sulfide Occurrence Models. USGS Scientific Investigations Report 2010–5070–C: 363. 92 �Summary of the 2022 ILSG Field Trip to Iceland UPTON, Margaret1, LARSON, Phillip2, MACTAVISH, Allan3, HINZ, Peter4 1 Department of Earth and Environmental Sciences, University of Minnesota - Duluth, 1114 Kirby Drive, 102 Heller Hall, Duluth, MN 55812 2 Cleveland-Cliffs Hibbing Taconite Company. Hibbing, MN 55746 3 AGC GeoConsulting, 777 Red River Road, Thunder Bay, ON P7B IJ9 4 Retired, Ontario Ministry of Energy, Northern Development and Mines, Thunder Bay, ON During Summer of 2022 (July 26-August 9), a group of 16 people set out to tour the diverse and awe-inspiring geology of Iceland, led by ILSG representative geologists Phil Larson, Peter Hinz, and Allan MacTavish. The 15 day trip held many surprises for all involved, including the worst stretch of weather Phil has experienced in Iceland (of 11 visits!) as well as a once-in-a-lifetime experience to see a volcanic eruption. In addition to the trip leaders, the group of 16 people included 3.5 professional geologists, 1.5 graduate students, one retiree, one Goldich Medal laureate, and 9 members of the Minnesota Geological Society. Stops throughout the trip focused on a wide range of topics: • Volcanism, both historical and the 2021 Geldingadalir eruption; Figure 1. Photo credit: Tom Hart • Icelandic cuisine, lore, and the historical and cultural evolution of the nation; • Environmental geochemistry of subsurface and near-surface processes; • Volcanic flows, igneous petrology for mineralogy and volcanic textures; • Geothermal energy and its uses; • Hydrology and hydrologic events related to glaciers and volcanics; • Geomorphology as it relates to ecology, volcanics, and glaciers • Wind, water, and glacial erosional features; glacial nomenclature By special arrangement, the trip was scheduled to overlap with the onset of the 2022 Meradalir eruption (fig.1). An advance party made a midnight scouting foray to the vent site before a Force 13 gale descended on the island. 93 �This presentation summarizes the highlights from the trip (fig.2). Featured locations include the Fagradalsfjall eruptions on the Reykjanes Peninsula, the Vestmannaeyjar Islands, climbing atop and viewing the Laki fissure from above, trekking to the highlands to view Askja and its pumice fields, the Jökulsárgljúfur canyon and scablands, free roadside hákarl stands, being lowered into a dormant magma chamber, a sampling of geothermal pools, plus the many epic waterfalls along the way! Figure 2. Generalized map of the trip route. 94 �GEOHERITAGE AS AN EDUCATIONAL TOOL TO EXPLORE RELATIONSHIPS WITH LAND AND WATER IN THE KEWEENAW VYE, Erika1 and ROSE, William2 1 Great Lakes Research Center, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 2 Geoheritage is an evolving field in the United States that considers the protection, interpretation, and management of geologic features with significant scientific, educational, cultural, or aesthetic value (Brocx & Semeniuk, 2007; Geological Society of America, 2017; National Park Service & American Geosciences Institute, 2015; Reynard & Brilha, 2017). Geoheritage strongly emphasizes the importance of the varied personal values people have for geologic features and the wide-ranging relationships we have with landscapes. As such, geoheritage is an effective geoscience communication tool affording place-based learning experiences that nurture our sense of place, deepen our Earth science literacy, and inspire stewardship and protection of our place. This presentation explores geoheritage education and outreach initiatives in Michigan’s Keweenaw Peninsula for both formal and informal learning communities. The Keweenaw Peninsula sits at the heart of the Midcontinent Rift and is renowned for the world’s largest accessible native copper deposit and Lake Superior, the largest freshwater lake on Earth. These geologic processes and features have fostered varied human relationships with the landscape, including the oldest metal workings in the Western Hemisphere and the European immigration wave of 18401910 triggered by the Copper Boom. This intersection of deep time, industrial, and cultural heritage has been the focus of teacher professional learning institutes and student internship experiences that explore the compelling geoheritage of our place. These programs: a) focus on complex environmental issues rooted in Earth systems processes of importance within the community, b) emphasize strong community partnerships that bring together varied values and perspectives of our place; c) explore other ways of knowing about the dynamic and interconnected geologic and human stories that serve as the foundation of the landscape’s past, present, and future through equitable knowledge exchange, and d) elevate Earth science literacy for educators and students by connecting the underpinning geology to current environmental issues with wide-ranging impacts in our communities today such as cultural identity, subsistence uses, recreation, and sense of place. The geologic formations of the Midcontinent Rift are beautifully exposed in the Keweenaw for researchers, teachers, students, and geotourists. As the Keweenaw shifts from an extractive industrial economic past, geoheritage initiatives support a future based on education, conservation, and sustainable tourism. Current initiatives in our community include a) the development of geotourism opportunities - Keweenaw Geotours, b) strong partnerships with local conservation groups to maintain access to world-class geosites that provide outstanding Earth science learning opportunities, and c) exploration of recreational opportunities including the concept of a shoreline hiking trail following the high water mark of Lake Superior. Geoheritage education and outreach opportunities help foster a culture of stewardship, increase Earth science literacy, and provide opportunities to share our varied relationships with land and water. 95 �Figure 1. Teachers and students explore the geoheritage of the Keweenaw by land and water. References Brocx, M. and Semeniuk, V., 2007. Geoheritage and geoconservation - history, definition, scope and scale. Journal of the Royal Society of Western Australia, 90: 53-87. Geological Society of America 2017. GSA Position Statement: Geoheritage. Retrieved from: https://www.geosociety.org/documents/gsa/positions/pos20_Geoheritage.pdf. National Park Service and American Geosciences Institute 2015. America’s Geologic Heritage: An Invitation to Leadership. NPS 999/129325. National Park Service, Denver, Colorado. Reynard, E. and Brilha, J., 2017. Geoheritage: Assessment, protection, and management. Elsevier, ISBN: 9780128095317. 96 �U/Pb geochronology and zircon petrochronology of Paleoproterozoic magmas from the Marshfield terrane Penokean Orogen, Wisconsin WEBER, Evan1, LODGE, Robert W.D.1, MARSH, Jeffrey2 1 Department of Geology and Environmental Science, University of Wisconsin-Eau Claire, Phillips Hall Eau Claire, WI 54701 2 Department of Earth Sciences, Laurentian University, 933 Ramsey Lake Rd, Sudbury, ON P3E 6H5, Canada This study presents U-Pb, Hf-Lu, and trace isotopic element data from zircons obtained from volcanic and intrusive rocks from the Paleproterozoic Penokean magmas within the Marshfield terrane in northern Wisconsin. The Penokean Orogen hosts both the Proterozoic Pembine-Wausau and the Archean Marshfield terranes. The Eau Pleine Shear Zone is interpreted as the paleosuture zone between these two terranes (Sims et al., 1989). Both terranes host volcanic and intrusive rocks that were formed during the Penokean orogen. The Pembine-Wausau terrane is a juvenile arc system that was developed through subduction during the Penokean orogen that accreted against the Superior craton. The volcanic rocks in this terrane are tholeiitic and calcalkaline in nature (Schulz and Cannon, 2007). The Marshfield terrane is thought to be an accreted fragment of an Archean craton that collided with the Pembine-Wausau terrane and the Superior craton (Klier, 2019). The Marshfield terrane is mainly comprised of gneisses that underlie Early Proterozoic volcanic rocks (Sims et al., 1989), but due to poor exposure of these rocks this terrane is poorly understood. This study aims to provide a better understanding of the volcanic terranes in the region to improve regional models of the southern portion of the Penokean orogen. Samples were collected from Big Falls and other locations within the Eau Claire volcanic complex, as well as from granites and gneisses exposed in Black River Falls. Zircons from these samples were then analyzed at Laurentian University (Sudbury, Ontario, Canada) via Split-Stream Laser Ablation Inductively Coupled Plasma Mass Spectrometer (LASS-ICP-MS) to obtain U-Pb, HfLu, and trace isotopic element data. Results reveal complex age relationships and basement architectures. The Big Falls gneiss, part of the Eau Claire volcanic complex lying south of the Eau Pleine Shear Zone (Fig. 1), resulted in an interpreted U-Pb age of 1874.7±2.1 Ma (Fig. 1) which is consistent with VMS-forming events in the Pembine-Wausau terrane. Zircon trace element geochemistry from the Eau Claire volcanic complex indicate rocks formed from a hydrated but reduced melt. This melt may have occurred in a back-arc setting where decompression occurred in a metasomatized mantle, which is characteristic of back-arc signatures. Hf-Lu isotopic data from the Eau Claire volcanic complex show the rocks here lack an Archean inheritance. The data from the Eau Claire volcanic complex was compared to a granite intrusion in Black River falls and both the Eisenbrey and Lynne VMS deposits in the Pembine-Wausau terrane. Based on Hf-Lu data, the Black River Falls granite showed inheritance of basement, which is expected based on field relationships with Archean rocks from the Marshfield terrane. The Eisenbrey and Lynne deposit have juvenile signatures which is characteristic of an oceanic arc system. According to trace isotopic element data, the VMS deposits also formed from a more oxidized and hydrated melt which is a similar geodynamic setting seen in the Eau Claire volcanic complex. Since we would expect basement inheritance in the Eau Claire volcanic complex, these results question what is known about the Marshfield terrane and its relationship to the Penokean. 97 �Figure 1. Geologic map of Eau Claire and Chippewa Falls area highlighting the location of Big Falls alongside a concordia diagram plotting the age of Big Falls at 1874.7±2.1 Ma. Cathodoluminescence imaging of individual zircons are also shown with their corresponding ages. References Brown, B.A., 1988. Bedrock Geology Map of Wisconsin (Regional Map Series: West-Central Sheet), University of Wisconsin-Extension Geological and Natural History Survey, Scale: 1:250,000. Klier, J.J., 2019. The Marshfield Terrane: Redefinition of Origin Through Zircon Geochronology and Geochemistry [MSc thesis]: Ball State University: 115. Schulz, K.J. and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research 157: 4-25. Sims, P.K., Van Schmus, W.R., Schulz, K.J., and Peterman, Z.E., 1989. Tectonostratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean orogen: Canadian Journal of Earth Science, v. 26: 2145-2158. Zi, J.-W., Sheppard, S., Muhling, J.R., and Rasmussen, B., 2021. Refining the Paleoproterozoic Tectonothermal History of the Penokean Orogen: New U-Pb Age Constraints from the Pembine-Wausau terrane, Wisconsin, USA: GSA Bulletin, v. 134: 776–790. 98 �The Use of Electric Pulse Disaggregation Technology to Recover Nickel Metal from Nickel Sulfide Ore Deposits WEIBLEN, Paul1 1 Minnesota Geological Survey (Retired), 2609 West Territorial Road, St. Paul, MN 55114 All metals, except for the noble metals like gold, occur in nature as metal sulfides. The chemical process “Plat Sol”1 can be used to recover nickel metal from nickel sulfide ores. The demand for nickel metal has increased dramatically due to the need for nickel metal for electric vehicle batteries. Elon Musk, always ahead of the curve, has signed an agreement with Talon Metals to be the sole recipient of any nickel metal they produce. Similarly, the Biden Administration is encouraging a transition from fossil Fuel vehicles to electric vehicles.2 However, a particle size of less than a millimeter is required for the feed to the Plat Sol process. Electric Pulse Disaggregation Technology3 provides much more efficient and less expensive method than conventional crushing and grinding for reducing the particle size of ore samples. Figure 1 provides details on the disaggregater. Inside the 3D printed gray cap on the left below is a stainless steel hemisphere with a pointed electrode projecting upward. On the right, is a black 3D printed cap with an electrode like the one above. When the two caps are screwed together, a sphere is formed. The electrodes are separated ~ 5 mm forming a spark gap. The two hemispheres are filled with water and inch-sized sample fragments. When the 50KV power supply is turned on the discharge across the spark gap vaporizes the water, which in turn separates different minerals along their grain boundaries. Examples of “zapped” Talon Metals nickel sulfide ore will be shown. Figure 1. Image of the disaggregator set up. References Google “Plat Sol” https://www.whitehouse.gov/briefing-room/statements-releases/2021/08/05/fact-sheet-president-bidenannounces-steps-to-drive-american-leadership-forward-on-clean-cars-and-trucks/ https://www.researchgate.net/project/Electric-pulse-disaggregation-and-hydroseparation-for-mineral-processing *** Abstract Withdrawn*** 99 � https://digitalcollections.lakeheadu.ca/files/original/0b0e110830002a4043b2d65b71c5ca51.pdf 2cbbb59065186547c55e7817dcd81e3e PDF Text Text 69th ANNUAL MEETING Eau Claire, Wisconsin — April 24-25, 2023 INSTITUTE ON LAKE SUPERIOR GEOLOGY Part 2 — Field Trip Guidebooks �Thank you to our sponsors! A SPECIAL THANK YOU TO OUR INDIVIDUAL CONTRIBUTORS: FREDERICK CAMPBELL, VAL CHANDLER, JIM DEGRAFF, THOMAS ERICKSON, TOM FITZ, DAVE GOOD, PAULA LEIER-ENGELHARDT, ALLAN MACTAVISH, BOB MAHIN, GORDON MEDARIS JR., JIM MILLER, STEVEN PINTA, TOD ROUSH, AND GERRY WHITE i �Proceedings of the 69th ILSG Annual Meeting – Part 2 69th ANNUAL MEETING INSTITUTE ON LAKE SUPERIOR GEOLOGY April 24-25th Eau Claire, Wisconsin HOSTED BY Rob Lodge, Esther Stewart, Carsyn Ames Co-Chairs University of Wisconsin- Eau Claire and Wisconsin Geological and Natural History Survey Proceedings - Volume 69 Part 2 – Field Trip Guidebooks Compiled and edited by Rob Lodge Cover Photos. Upper — Photograph of E.O. Ulrich taking notes in the field describing the Cambrian Mount Simon Formation in the Chippewa Falls region in 1913. Lower — Photograph of geologists E.F. Bean and E.C. Edwards fording the Eau Claire River at Morrison’s Ford in 1919. iii �Proceedings of the 69th ILSG Annual Meeting – Part 2 69th INSTITUTE ON LAKE SUPERIOR GEOLOGY VOLUME 69 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD T RIP GUIDEBOOK Trip 1: PRECAMBRIAN GEOLOGY OF THE CHIPPEWA RIVER VALLEY Trip 2: WISCONSIN’S PALEOZOIC STRATIGRAPHY AND TOUR OF CRYSTAL CAVE Trip 3: PRECAMBRIAN GEOLOGY OF THE EAU CLAIRE RIVER VALLEY Trip 4: QUATERNARY GEOLOGY AND GEOMORPHOLOGY OF THE EAU CLAIRE REGION Reference to material in Part 2 should follow the example below: Lodge and Hooper, 2023. Precambrian geology of the Chippewa River Valley: A transect through the western Marshfield Terrane. in Lodge, R.W.D. (Ed.), Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 2 – Field Trip Guidebooks. v.69, part 2, p.1-26. Published by the 69th Institute on Lake Superior Geology and distributed by the ILSG Secretary: Pete Hollings - ILSG Secretary Department of Geology Lakehead University 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Email: peter.hollings@lakeheadu.ca ILSG website: www.lakesuperiorgeology.org ISSN 1042-9964 iv �Proceedings of the 69th ILSG Annual Meeting – Part 2 Part 2: Field Trip Guidebooks Table of Contents Page Field Trip 1: Precambrian geology of the Chippewa River valley: A transect through the western Marshfield Terrane Field Trip 2: Wisconsin’s Paleozoic stratigraphy and tour of Crystal Cave Field Trip 3: Precambrian Geology of the Eau Claire River Valley: Re-discovering the Eau Claire Volcanic Complex Field Trip 4: Quaternary Geology and Geomorphology of the Eau Claire Region v 1 27 48 71 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Field Trip 1 – Precambrian geology of the Chippewa River Valley: A transect through the western Marshfield Terrane Robert W.D. Lodge and Robert L. Hooper Department of Geology & Environmental Science, University of Wisconsin-Eau Claire, Eau Claire, Wisconsin 54701 of the 18.2 Mt Back Forty VMS deposit in Michigan, easing of the Wisconsin sulfide mining moratorium in 2017, and a recent national push for securing critical mineral resources. However, this has also highlighted the lack of modern datasets, notably lithogeochemistry, on these deposits that could be used to further our knowledge of the mineral-forming systems in the VMS belt. The Pembine-Wausau terrane has received most of the historic and recent attention since it hosts approximately 150 million tonnes of known VMS mineralization. However, little attention has been given to the Penokean volcanic deposits that overprinted the Marshfield Terrane that are presented in this guidebook. DeMatties (2022) recognized the gap in knowledge for these Penokean volcanic deposits within the Marshfield Terrane, also called the Eau Claire Volcanic Complex, and highlighted their exploration potential. Introduction The erosional outliers of Precambrian bedrock in the Chippewa River Valley represent the southernmost extent of the Canadian Shield before it is completely covered by Paleozoic sedimentary strata. The rocks exposed here are part of the Paleoproterozoic Penokean Orogeny, a collisional orogen that resulted from the accretion of the Pembine-Wausau and Marshfield terranes onto the (present-day) southern margin of the Superior Province. This region was last visited by members of the Institute of Lake Superior Geology in 1980 when a field trip through the region was conducted by Paul Myers (Myers et al., 1980). Since this time, there has been ‘new’ U/Pb data collected by the USGS (Sims et al. 1989) and others (Van Wyke et al, 1997; Klier, 2019), regional syntheses of the Penokean volcanogenic massive sulfide (VMS) mineralization (DeMatties 1989; 1994; 2018; 2022), maps published by government surveys (Mudrey et al, 1987; Brown 1988), and orogenwide tectonic model (Shultz and Cannon, 2007) that is being revisited based on new U/Pb data (Zi et al., 2021). After forty years of advancing our knowledge of the Penokean Orogen, it is worth touring again. The portion of the Marshfield terrane that is visited in this guidebook is well known, but grossly understudied and much of its regional context is unknown. Students from the University of Wisconsin-Eau Claire have been visiting many of the locations in this guidebook for decades to learn how to map and describe rocks in the field, measure structures and interpret geologic histories, and learn the basic mechanics of field work. Faculty, students, and alumni from Eau Claire consider these outcrops classic. This guidebook will (re-)introduce these rocks and provide an updated view on their context to the Marshfield terrane and Penokean Orogen. Ongoing research in this region hopes to expand the lithogeochemistry and zircon petrochronology database to better delineate the The Penokean Orogen is perhaps best known for hosting numerous VMS deposits. The passing of the “Prove-it-first” law, or sulfide mining moratorium, in 1997 effectively shut down mineral exploration and mining in Wisconsin. One of the most complete descriptions of several deposits was published by the Institute of Lake Superior Geology (LeBarge, 1996). More recently, the mineral exploration industry has been reinvigorated because of the 2002 discovery 1 �Proceedings of the 69th ILSG Annual Meeting - Part 2 geodynamic evolution and crustal architecture of this region. Determining the presence or absence of Archean basement throughout the Marshfield terrane will help refine terrane boundaries and improve our understanding of the metallogeny of the region to assist in future mineral exploration efforts. in a suprasubduction zone setting and are now structurally juxtaposed along the southern edge of the Archean Superior Province during the earliest phases of forming the Columbia, or Nuna, supercontinent (LaBerge and Myers, 1984; Sims et al., 1989; Schulz and Cannon, 2007). The orogen is host to at least 150 million metric tonnes (Mt) of VMS and associated mineralization (DeMatties, 1994, 2018) but remains one of the more poorly understood and underexplored mineral districts in North America. Regional Geology The Paleoproterozoic Penokean Orogen (ca. 1.8 Ga) in the Lake Superior region (Figure 1) is a classic Precambrian orogenic belt comprised of dominantly sub-marine volcanic rocks and associated plutons. The Penokean rocks formed The Penokean Orogen has been divided into the Interior and Exterior domains. The Exterior Figure 1 – Geologic map of the major tectonic assemblages and major structures of the Penokean Orogen. Notable and important abbreviations that are important for this guidebook are EPSZ, Eau Pleine shear zone; NFZ, Niagara fault zone. Figure from Shultz & Cannon (2007). 2 �Proceedings of the 69th ILSG Annual Meeting - Part 2 domains are sutured to the Superior Craton by the Niagara fault zone (Figure 1). The Exterior domain consists of passive margin, rift, and forearc basin sediments and Archean crustal blocks from the Superior Province that were folded and faulted in the foreland part of the orogen. The Interior Domain consists of two accreted terranes, the Pembine-Wausau and Marshfield terranes, that are sutured by the Eau Pleine Shear zone (Figure 1). The PembineWausau terrane is a composite accreted oceanic arc overprinted by continental-margin magmatism and hosts numerous VMS deposits and occurrences (DeMatties, 1994; Shultz & Cannon, 2007) (Figure 2). The Marshfield terrane is composed of Archean crustal fragments of unknown origin that were overprinted by Penokean magmas during the orogen (Figure 2) and is described in more detail in the section to follow. Shultz and Cannon (2007) synthesized tectonic events during the Penokean Orogeny based on a detailed compilation of lithologic, structural, sedimentological, isotopic, and geochronological datasets. This classic model proposed that an oceanic arc, now referred to as the PembineWausau terrane, collided with the southern Superior Province around 1880 Ma. Following a subduction flip from south-directed to northdirected subduction, continental arc magmatism and back arc extension followed until about 1850 Ma when convergence with an Archean crustal block, known as the Marshfield Terrane accreted to the southern edge of the Wausau- Pembine Terrane along the Eau Pleine Shear Zone (ESPZ). Sedimentation related to this convergence in a foreland basin setting continued until about 1835 Ma. The end of the Penokean orogen was constrained by a series of undeformed posttectonic plutons dated at 1830 Ma which stitched the terranes. Figure 2 – Schematic tectonic evolution of the Penokean Orogen provided by Shultz and Cannon (2007) based on geophysical, sedimentological, and geochronological complications. contradictory data came when Quigley (2016) obtained a high-precision U/Pb zircon age of 1832.98 ± 0.52 Ma from a rhyolite at the Back Forty deposit via CA-ID-TIMS. The other analyzed VMS deposits across the PembineWausau terrane by Quigley (2016) provided consistent U/Pb zircon ages ca. 1875 Ma and supported the Shultz and Cannon (2007) tectonic model. Additional U/Pb zircon ages from volcanic units (Beecher Formation) and plutonic rocks (Dunbar Gneiss, Newingham Tonalite) in the eastern part of the orogen by Zi et al. (2021) However, this classic tectonic model for the evolution of the Penokean orogen has recently been re-evaluated in light of new U/Pb data obtained throughout the orogen. The first 3 �Proceedings of the 69th ILSG Annual Meeting - Part 2 supported the younger extensional tectonic event proposed by Quigley (2016). These new ages resulted in a revised Penokean tectonic model where long-lived northward subduction along a continental margin with repeated extensional and contractional regimes in response to retreat and advance of the subducting oceanic plate (Figure 3). Weber and Lodge (2022) obtained a U/Pb age of 1831.4 ± 2.0 Ma on the dacite unit hosting the Eisenbrey deposit in the western part of the orogen, suggesting that this second VMS forming event was widespread. A summary of the geochronology is presented in Figure 4. part of the Marshfield terrane and lie immediately south of the Eau Pleine Shear Zone. Current tectonic models suggest that the Marshfield Terrane represents an Archean microcontinent of uncertain origins (Sims et al., 1989; Schulz and Cannon, 2007; Zi et al., 2021). Some of the earliest work on the terrane by Sims et al. (1989) noted only eight Archean U/Pb ages from isolated outcrops along the Wisconsin, Black, and Chippewa Rivers, many of which were compiled from unpublished sources. One of those was the gneiss exposed at Jim Falls (Stop 3 in this guidebook) which was dated at 2522 ± 22 Ma. Current tectonic reconstructions usually have Paleoproterozoic volcanic rocks in the Marshfield terrane being deposited on Archean basement at about 1870–1860 Ma. The Paleoproterozoic volcanic sequence is referred to as the Eau Claire Volcanic Complex by DeMatties (2018; 2022) and are preserved only as erosional remnants. The Eau Claire Volcanic Complex consists principally of an interlayered sequence of felsic to mafic volcanic rocks, dacite porphyry, and a variety of clastic and chemical sedimentary rocks (Sims et al., 1989). Some conglomerates contain granitic gneissic clasts that were interpreted to be Archean (Myers et al. 1980), but no definitive ages were determined on the clasts. Throughout the Marshfield terrane there are various Paleoproterozoic intrusions of gabbro, diorite, and tonalite. These have U/Pb ages of 1835-1865 Ma (Sims et al., 1989; Van Wyck and Johnson, 1997; Weber and Lodge, 2022). Otherwise, our knowledge of the Archean basement of the Marshfield terrane and associated Paleoproterozoic volcanic rocks remains as sparse as the outcrop exposures. Figure 3 - Schematic illustration of the revised tectonic model of the Penokean Orogen. Figure is from Zi et al. (2021). Abbreviations: NF—Niagara fault zone; EPSZ—Eau Pleine shear zone.Marshfield Terrane The study of the Marshfield terrane remained stagnant until new U/Pb and Lu-Hf isotopic data from zircons was published as a masters thesis (Kleir, 2019). The new isotopic data in the Marshfield Terrane collected from the Chippewa and Yellow River valleys will be presented at various stops on this field trip. In our opinion, one of the most significant results was that the “Archean” rocks from the Jim Falls region of Marshfield Terrane This guidebook visits field sites from the northwestern exposures of rocks interpreted to be 4 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 4 - Time-space plot for the tectonic components of the Penokean Orogen. Plot is from Zi et al. (2021). See citation for references on data sources. Sims et al. (1989) is a metasedimentary rock that has a significant proportion of Paleoproterozoic zircons (Kleir, 2019). While the data clearly indicates the presence of Archean rocks in the sedimentary source region, the sedimentary provenance does not require that Archean rocks represent the basement architecture in the northern part of the Marshfield Terrane. U/Pb ages from the northern part of the Marshfield Terrane collected in the Chippewa and Yellow River areas are interpreted as Paleoproterozoic in age (~1.83-1.88 Ga) and Hf isotopies indicate a juvenile source (without Archean contributions). This finding raises questions about the extent of the Archean basement in the Marshfield Terrane and consequently, the basement architecture in 5 �Proceedings of the 69th ILSG Annual Meeting - Part 2 the region. Preliminary geochemistry from Klier (2019) and our ongoing studies in the region show interesting trends that will help distinguish petrogenetic processes. Figure 5 highlights the REE trace element characteristics of these deposits and their application to each stop in subsequent sections below. thermometry determined temperatures between 719-769°C (Hannack and Radwany, 2018). A rutile U/Pb age of 1835 Ma from Sims et al. (1989) in the Eau Claire Volcanic Complex (Big Falls – Fieldtrip 3 in this volume) may indicate the timing of metamorphism since new zircon U/Pb age from the same region provided a crystallization age of ~ 1875 Ma (Weber and Lodge, 2022). Regional metamorphism in this region is at lower to upper amphibolite facies. Hornblendeplagioclase thermo-barometry from gneisses in the Chippewa River valley indicate metamorphism occurred at temperatures between 606-646°C and pressures between 5.74-6.64 Kbar (Hafften and Radwany, 2018). A sample of amphibolitic gneiss from the Eau Claire Volcanic Complex using the edenite-richterite Field Trip Stops The overall objective to this guidebook is to tour the Precambrian exposures of the Marshfield terrane along a southwest-northeast transect as exposed in the Chippewa River Valley. Starting within the city of Chippewa Falls, the trip will work its way to the northwest along the river and presumably get closer to the terrane boundary at the Eau Pleine Shear Zone. Stops 1-4 and 6 are all within the Marshfield Terrane, whereas Stop 5 is considered the southernmost exposure of the Pembine-Wausau Terrane. Fieldtrip stops are summarized in Figure 6. New data have us questioning what we know about the Marshfield Terrane. For example: Where exactly is the northern boundary of the Marshfield terrane in the Eau Claire region, and how much of the Marshfield Terrane, as currently defined, has an Archean basement architecture? Most of the locations in this guidebook are at the downstream side of hydro-electric dams. These areas are prone to sudden flooding and the upmost caution and careful planning should be used prior to visiting these locations. In addition, rocks here are uneven and slippery especially when wet. To access larger sections of outcrops, low water conditions or ladders (temporary bridges) may be required. In addition, all locations in this region may contain poisonous plants (e.g. nettle, poison ivy) and black-legged ticks that can transmit diseases. While this is unlikely to be a concern in early spring during the 2023 ILSG conference, future users of this manual should plan appropriately. Figure 5 - Trace element diagrams from the rocks in the Chippewa River valley region. Data from Cornell is preliminary data from ongoing projects whereas the remainder is from Klier (2019). (A) Classification diagram from Pearce (1996) showing protolith compositions. (B) Primitive mantle-normalized rare earth element diagram using normalizing values from Sun and McDonough (1989). 6 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 6 - Regional geology of the Chippewa Falls and Eau Claire region showing fieldtrip stops and approximate location of the Eau Pleine Shear zone. Rocks to the south of the Shear Zone are interpreted to be part of the Marshfield Terrane, whereas rocks to the north are part of the Pembine-Wausau terrane. Figure compiled from Mudrey et al. (1987) and Brown (1988). 7 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Stop #1 – Nonconformity at Irvine Park Claire and Chippewa Falls region are some of the southernmost exposures of the crystalline basement in the Lake Superior region before it disappears beneath the undeformed Paleozoic sedimentary strata. This is one of the many exposures of the “Great Unconformity” that is present throughout this region. Details of this unconformity in this region is described in detail in the most recent ILSG guidebook presented in Eau Claire (Chan et al. 1991) and is summarized below. Lat: 44.9542° Long: -91.3972° Precambrian Unconformity The Precambrian- Cambrian boundary is represented by a highly variable surface in the mid-continent area. In west-central Wisconsin, the Precambrian surface forms an extensive planation surface with a regional SW dip of less than 1 degree. Archean iron formations in the Black River Falls region and Proterozoic quartzites throughout the state, most famously the Baraboo Syncline, form isolated monadnocks on the peneplain. The peneplain was mantled by a layer of paleosols as much as several hundred feet thick. In some areas, Cambrian rocks directly rest upon barren, moderately weathered Precambrian rocks. Considering the low paleolatitude of the continent during the Cambrian, deep weathering of the Precambrian surface must have occurred before the Upper Cambrian deposition. The Precambrian basement, however, shows variable degrees of weathering and the weathering is complicated by potassium metasomatism overprinting associated with Silurian and Devonian K-rich basinal brines (Lui, 1997; Lui et al., 2003). Potassium metasomatism along the unconformity is responsible for the development of illite, interlayered I/S and authigenic Kfeldspar in both saprolites and in the Cambrian rocks in the Chippewa River Valley. Where the Precambrian is mafic (gabbros, amphibolites and gneisses) the potassium metasomatism commonly produces a distinctive bright bluegreen clay (celadonite) seen at Stop 2 on this field trip and at Big Falls (Fieldtrip 3 in this guidebook). These potassic brines have been The outcrop described at this stop is located along the east bank of Duncan Creek within Irvine Park in Chippewa Falls. Upon entering the park, drive north on Irvine Park Drive past the zoo and bison enclosure until the inter-section with Bear Den Road. There is ample parking in this area near the intersection that crosses Duncan Creek to the east and find the small foot trail that leads northward to the outcrop. Potential hazards include poisonous plants and ticks, but they are unlikely to be a problem in early spring. There is also uneven and potentially wet ground. The purpose of this location is to highlight the conditions that are impeding the study of the Precambrian bedrock in the region. The Eau 8 �Proceedings of the 69th ILSG Annual Meeting - Part 2 implicated in the formation of MVT deposits in the Tri-state region (Aleinkoff et al., 1993). At this location (and numerous others) where the Cambrian formations are in contact with Precambrian plutonic rocks, the basement is heavily spheroidal weathered (Photo 1, Figure 7) and is generally deeply altered to kaolinite saprolite and then metasomatically altered to illite I/S and authigenic Kspar. Mt. Simon Formation sandstone and conglomerate fill the wedges among the spheroids. In some areas such as Little Falls and Big Falls in Eau Claire County (Fieldtrip 3, this volume) or Rock Dam in Jackson County, Cambrian sandstones rest upon Proterozoic amphibolite and meta-rhyolites that are only partially altered. Mount Simon Formation is a coarse-grained, medium to thick-bedded quartz pebble conglomerate. The topographic relief on the Precambrian surface was probably only a few meters as sedimentary channels are typically less than 1 meter deep. The presence of trace fossils (rusophycus and Climactichnites, or trilobite burrows/tracks; Photo 2) and planar and bipolar cross-bedding (Photo 3) suggest a littoral or shallow marine tidal flat environment of deposition for the lower part of the formation. The upper part of the Mt. Simon Formation contains feldspathic quartz arenite with smallscale ripple bedding, brachiopod fragments (lingula sp.), and worm trails (planolites) The Photo 2 - Climactichnites fossil from the lower Mt. Simon Formation collected along the Chippewa River near downtown Eau Claire. Climachtinites trace fossils are typical of tidal flats in the Cambrian. Field of view is ~1m across. Photo 1 – Irvine Park outcrop photograph showing nonconformity between Cambrian Mt. Simon Formation (above) and Paleoproterozoic trondhjemite (below). Photo courtesy of Scott Clark (UW-Eau Claire Cambrian Mount Simon Formation The sediment above the unconformity consists largely of Upper Cambrian Mt. Simon Formation deposited on the mid-continent region of North America during the Dreisbachian transgression. The Mount Simon Formation is a fine to coarsegrained, moderately to well sorted, quartz arenite with a local basal conglomerate. The Mount Simon Formation varies in thickness 40 to 180 meters in the Upper Mississippi Valley. Locally, in the Chippewa Valley area, the lower part of the Photo 3 - Cross-bedded conglomerate and sandstone of the Cambrian Mount Simon Formation in the Irvine Park area, Chippewa Falls. Photo courtesy of Scott Clark (UW-Eau Claire). 9 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 7 – Conceptual field sketch of the unconformity at Irvine Park. Figure from Chan et al. (1991). Stop 2 – Penokean and Mid-Continent Rift Intrusions at Lake Wissota Dam grain size distribution shows a generally fining upward sequence. Precambrian Intrusion Lat: 44.9429° Long: -91.3425° The Paleoproterozoic biotite tonalite (trondhjemite) showing spheroidal and saprolitic weathering at this location is interpreted to be similar to the larger trondhjemite intrusion that underlies much of the Chippewa River valley. The trondhjemite can be more easily observed below the Chippewa Falls hydroelectric dam in downtown Chippewa Falls and below the Wissota hydroelectric dam (Stop 2 in this guidebook). The biotite trondhjemite at Chippewa Falls Hydro was dated by Van Schmus (1980) at 1,840 ± 15 Ma. Saprolites like the one exposed here are characteristic of areas of prolonged tropical to subtropical weathering on a granitoid bedrock surface of low relief. The saprolite at this outcrop contains high clay content and angular quartz and feldspar. The alteration intensity increases approaching the Cambrian Mt Simon Formation. This outcrop is located on the downstream side the dam on Lake Wissota. Drive to the end of 74th Avenue in Chippewa Falls and park in the Chippewa Rod and Gun Club & Marina. From here, you can walk southward along the access road to the dam (about 1 km). There are several places to cross the small steam to access the largest part of the outcrop. The largest potential hazard at this location is the stream crossing and uneven, wet walking area. A ladder or other temporary structure might be required to assist in crossing the stream if water levels are high. This outcrop highlights some of the magmatic history in this region. The majority of the exposure here is a Paleoproterozoic biotite tonalite (trondhjemite) that has local pods and dikes or medium gray biotite tonalite and alkali feldspar granite pegmatite. The outcrop is intruded by at least three gabbroic dykes associated with the mid-continent rift. The largest 10 �Proceedings of the 69th ILSG Annual Meeting - Part 2 of which is clearly visible in arial view (Figure 8). The entire area is covered by thin outwash gravels and silts that varies with seasonal flooding events. Some of the tonalite near the Chippewa River displays the same spheroidal weathering seen at Irvine Park so this location is just below the Great Unconformity and displays some of the same associated potassic alteration along faults and joints seen in Irvine Park (Stop 1). The potassic alteration is responsible for most of the pink color seen in outcrop. Figure 8 - (top) Generalized geology of the Wissota Dam region. Figure modified from Myers et al. (1980). (bottom) Aerial view of the outcrops with the mid-continent rift highlighted. Image obtained from Google Earth. tonalite for rocks with higher mafic concentrations. The oldest, abundant rock at Wissota Dam is a weakly foliated hornblende, biotite trondhjemite composed of oligoclase (50%), quartz (30%), microcline (5%), biotite (10%), and 5% hornblende with common accessory euhedral to subhedral titanite (Photo 4). Weak foliation strikes Nl5°W and dips steeply east. This is intruded by small dykes and masses of medium-grained, medium-grey hornblendebiotite tonalite (± epidote) that locally contains Granitoid Intrusive Suite Most of the Paleoproterozoic intrusive igneous rocks at Wissota are quartz diorites or tonalites with various proportions of hornblende and biotite. For clarity, and to be consistent with the terminology used by previous geologists working in the Chippewa River Valley, on this field trip we refer to the lightest colored tonalites (color index 15 or less) as trondhjemite and reserve 11 �Proceedings of the 69th ILSG Annual Meeting - Part 2 trondhjemite is an FI-type felsic rock with strongly depleted HREE (Figure 5) representing deep crustal melting (Hart et al. 2004). The trondhjemite is cut by east-northeast-trending pegmatite veins and pods and quartz ± pyrite) and epidote veinlets. Potassic alteration especially along any fractured surfaces the trondhjemites produces a pink color in outcrop (Photo 5). Minor cataclastic fault zones cut the granitoid intrusions with leftlateral displacement. A thin, branching discordant sheet of foliated biotite trondhjemite approximately 1-3 meters wide and trends N55°W. Drag folded foliation in the enclosing rocks indicates left-lateral displacement. Photo 4 – Photographs of main lighter colored tonalite phase at Wissota Dam. (A) Photomicrographs in plane-polarized light showing feldspar grains are lightly weathered with opaque rims around titanite. In cross-polarized light, quartz grains show moderate undulatory extinction. Photo from Klier (2019). B) Field photograph of biotite tonalite (trondhjemite) showing medium-grained, equigranular texture. Feldspars weather pink in color and mafic phases tend to be recessively weathered. lenticular xenoliths of banded amphibolite. The tonalite pods show no grain size diminution and sometimes have irregular shapes suggesting that some tonalites may be enclaves of earlier phases of the trondhjemite. In other cases, the tonalites are clearly dykes crosscutting the trondhjemite. The tonalite dikes tend to be unaltered with vitreous dark-brown biotite (~25%) and lack foliation. All minerals in the foliated tronhjemite show internal fracturing and dislocation, and contain quartz grains with undulatory extinction, display grain boundary migration and dynamic quartz recrystallization (Klier, 2019). The Photo 5 – Photographs of pegmatite and associated alteration at Wissota Dam. (A) Thin quartz-epidote veining and potassic alteration surrounding conjugate fracture sets adjacent to pegmatite. (B) Coarse grained texture of the pegmatite. Feldspar crystals can be as large as 5-7 cm in size. 12 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Slickenside fault surfaces elsewhere in this outcrop have similar strike and dip with the slickensides plunging 5°NW. of the largest dyke and the mineral chemistry was examined using SEM-EDS and optical petrography. The main dyke has an aphanitic chilled margin a few cm wide along both the north and south sides and the grain size consistently coarsens toward the middle of the dike into a medium grained olivine gabbro (Photo 7A). A prominent set of joints perpendicular to the cooling surface along the dike walls are interpreted as columnar joints and these are especially prominent on the south side below the power lines. More pronounced columnar joints are also seen in one of the smaller (2m wide) dikes along the northwest side of the area next to the Chippewa River. No internal contacts are apparent at the outcrop scale suggesting that this large dike represents a single cooling unit of magma intruded into the upper crust. West of the power lines the dike is cut by two faults, one left lateral strike slip fault with a few meters of displacement and a low angle reverse fault with well-developed chlorite slickensides and extensive alteration including chlorite and hematite, and calcite filled tension fractures. Mid-Continent Rift Dykes Three diabase dykes related to mid-continent rift extension intrude the granitoids (Photo 6) exposed below the Wissota Dam spillway and the largest dike is an ENE trending (~N65E) olivine tholeiite that averages 47m in width. The large dyke has a notable sharp and chilled margin. Unpublished data from the dyke at this location and others along the Chippewa River indicate a tholeiitic composition that shows slightly more Mg-enrichment trends on AFM diagrams in comparison to other parts of the dyke swarm in the region. Ongoing student-faculty research at the University of Wisconsin-Eau Claire is examining the composition of the large dyke at Wissota Dam. Samples were collected as a cross section The chilled margins consist of very finegrained plagioclase with variable compositions ranging from An35 to An63, in a devitrified-glass matrix crowded with submicron Fe-Ti oxides and sparse sub-calcic augite (Average cpx = [Mg.68Fe.60Ca.55Al.09Ti.02Mn.01] [Si1.91Al.09O6]). Within two meters of the north side of the diabase the dike contains single crystals of labradorite up to 10 cm across apparently sourced from a deeper magma chamber and transported (floated?) upwards during intrusion of the dike. Locally the chilled margin is altered to chlorite and very fine grained bright blue-green celadonite (Photo 7D), alkali-feldspar and dark red biotite. Five samples collected from the central 35 m of the dike all consists of an olivine gabbro with a well-developed ophitic texture (Photo 7B). The mineralogy from the center includes both titaniferous augite (1-2wt% TiO2) and titanaugite (>2wt% TiO2) oikocrysts with pink and lavender Photo 6 – Photographs showing sharp, chilled margin of mid-continent rift gabbro with Paleoproterozoic trondhjemite intrusion. 13 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Photo 7 – Photographs of mid-continent rift gabbroic dyke at Wissota Dam. (A) Outcrop photo showing fractured and weathered surface of dyke. Weathered surface shows medium-grained texture. (B) photomicrograph in crosspolarized light (40x) showing ophitic texture. (C) Photomicrograph in plane-polarized light (40x) showing aggregate of euhedral to subhedral olivine (ol) crystals (Fo 40-45) in plagioclase and cpx matrix where cpx as pinkish purple pleochroism typical of the titanaugite composition. (D) Photomicrograph in plane-polarized light (100x) showing greenish blue celadonite (cel) replacing biotite (bt) in the transitition zone between the chilled margin and dikes central olivine gabbro. pleochroism, laths of normally zoned plagioclase with labradorite cores (An55-65) and thin rims of andesine (An30-35) and unusually large aggregates of euhedral to subhedral olivine containing over 50 individual olivine crystals (Photo 7C). The augite and biotite show little variation across the dyke but the olivine becomes progressively more Fe-rich towards the south with an average of Fo43 in the north to Fo35 near the southern contact. The opaque minerals are primarily ilmenite with titaniferous magnetite lamellae often rimmed with a highly titaniferous reddish orange biotite. In the transition zone between the chilled margin and the center 30 m of the dyke much of the biotite is replaced (altered) with celadonite K(Mg,Fe2+)(Al,Fe3+)[Si4O10](OH)2 with a brilliant blue-green color in plane polarized light (Photo 7D). Unusual olivine aggregates (glomerocrysts?) occur throughout the central 35m of the dyke and often consist of more than 50 crystals (Photo 8). In some of the olivine aggregates the minerals show at least some crystallographic alignment. Olivine within individual aggregates have a very narrow range of chemistry. In one aggregate 16 grains were analyzed and the average composition was Fo47.5±0.2(2σ); this standard deviation is about the 14 �Proceedings of the 69th ILSG Annual Meeting - Part 2 magma with limited chemical variation and could be produced by turbulent flow (synneusis) agglomeration or by a high degree of undercooling and ripening of dendritic olivine. It seems very likely that this dike was an active conduit for magmas reaching the surface to produce MCR lava flows even though Chippewa Falls is almost 200 km south of the main MCR rift axis. Geochemical results which are pending should help further constrain the system. Stop 3 – Gneisses and Pegmatites at Jim Falls Photo 8 – Photomicrograph in cross-polarized light of olivine aggregate (glomerocrysts) showing consistent orientation of olivine crystals in the cluster. Gray crystals all have an optic axis almost perpendicular to the section. Magnification 100X. Lat: 45.0549° Long: -91.2734° same size as the analytical error for EDS analysis on olivine. Olivine aggregates have been described from several basaltic conduits where they have been attributed to differential crystal movement during turbulent flow in an active conduit such as at Kilauea (Helz, 1987) or as xenocrysts extracted from a deforming cumulate. However, there is no reference to aggregates with such a large number of crystals. The texture and chemistry of the aggregates at Wissota come closest to matching olivine aggregates collected from lava flows at La Reunion (Welsch et al., 2013) which they ascribe to rapid dendritic crystal growth and ripening under a high degree of undercooling (-ΔT > 60°C) from low viscosity basalts. Petrographic Interpretation: The dyke is sourced from a lower-level fractionated magma chamber of enriched basalt (E-MORB or alkali olivine parent) as evidenced by the olivine composition (~Fo40), plagioclase (An60) and titanaugite/ilmentite modal mineralogy. Plagioclase zoning from An60 cores to An30 rims suggests at least limited reaction with wall rocks. As a fractionated magma it seems likely that the ascending magma contained phenocrysts of both olivine and plagioclase that were kinetically fractionated by turbulent flow resulting in a chilled margin largely devoid of phenocrysts. The olivine aggregates form in equilibrium with This outcrop is located within the spillway of the hydroelectric dam near the community of Jim Falls. About 500 m north from the intersection of 15 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Highway 178 and County Highway Y (the main road into the community of Jim Falls), there will be an old, abandoned bridge that that used to be the access bridge to the community. There is ample parking in front of this bridge. On the south side of the old bridge is a small foot path that leads down to the outcrops along the river. These outcrops are smoothly polished from the flooding at the dam. They are uneven and quite slippery when wet. If water levels are high, there are also outcrops immediately downstream of the spillway to the north. This location has intensely folded amphibolite and biotite quartzofeldspathic gneiss that is intruded by granitoid intrusions and associated pegmatites (Figure 9). Intense shearing and metamorphism results in little preserved primary texture within the gneisses and amphibolites. On the east bank of the river is a gabbroic dyke associated with the mid-continent rift. Figure 10 - Tera-Wasserburg concordia diagram of biotite quartzofeldspathic gneiss from Jim Falls. A wide spread of ages is suggestive of a detrital origin. Figure from Klier (2019). sedimentary rocks. Trace element geochemistry from Klier (2019) was inconclusive in determining volcanic protolith because of low Ti abundance. Preliminary petrography from student projects at the University of Wisconsin-Eau Claire and Kleir (2019) seem to suggest that amphibolites are rather rare. Other geochemical results from ongoing research at the University of Wisconsin-Eau Claire are pending. The outcrops at this stop are one of the original “Archean” exposures of the Marshfield terrane that was described in in Sims et al. (1989), but new data in the region casts doubt on that original interpretation. The gneisses at this location were assigned a U/Pb age of 2522 ± 22 by Sims et al. (1989). Little description of the data was provided in that original reference and the date itself was referenced as unpublished data from personal communication. Klier (2019) resampled the gneiss from the region and analyzed zircons using LA-ICPMS. The resulting data clearly shows a large spread of ages and a significant portion of those are Paleoproterozic in age. There are clearly older sources of detritus for these meta-sedimentary rocks, some as old as 2841 Ma. However, the dominant source of detritus was Paleoproterozoic (Figure 10). Based on this new data, the Jim Falls region is not obviously an Archean crustal fragment. A biotite quartzofeldspathic gneiss was sampled by Klier (2019) for U/Pb geochronology. Based on recent field work, this rock appears to be the dominate lithology that exists in the immediate region around and under the bridge. Kleir (2019) describes the rock containing classic mylonitic textures and is comprised of quartz (60%), alkali feldspar (25%), biotite (15%), and trace zircon (Photo 9A). There are prominent bands of porphryoblastic quartz and cryptocrystalline biotite. Biotite is also present rarely as larger “destroyed” grains. Quartz has undulatory extinction and has undergone grain boundary migration recrystallization. Some feldspar grains display domino-type fragmented porphyroclastic textures. Portions of feldspar grains have diminished to sericite. Weakly Amphibolites and Gneisses Myers et al. (1980) interpreted the amphibolites and gneisses in this region to be derived from mafic volcanic rocks and associated 16 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 9 – Precambrian geologic map of Chippewa River near Jim Falls. Figure digitized from Myers et al. (1980). 17 �Proceedings of the 69th ILSG Annual Meeting - Part 2 isoclinally folded amphibolite occur in the granitic rocks. Granitoids and Pegmatites Granitic rocks range in composition from trondhjemite to alkali feldspar granite. Pegmatite dike intrusion occurred at several stages of "granite" intrusion. Mineralogy includes alkali feldspar (65%) quartz (32%), plagioclase (3%), and trace zircon.. The grains are subhedral to anhedral with intergrowths and granophyric textures occasionally present. Quartz grains appear stretched and strained and have undergone either subgrain rotation recrystallization or grain boundary migration recrystallization (Klier, 2019). The older granitic rocks are foliated and locally mylonitized. Shearing and boudinage of pegmatite stringers transposed them into oblique concordance with lamination in the enclosing rocks. A rough correlation can be made between relative age and concordance of veinlets. thinly laminated amphibolite was intruded by granite so that lenticular slices of the amphibolite were dragged en echelon away from the wall (Photo 10). The coarse granite pegmatite intruded under stress contains en echelon fractures filled with very coarse quartz. Photo 9 – (A) Photomicrograph in plane-polarized light of biotite quartzofeldspathic gneiss showing sericite-altered feldspar crystals and pronounced dynamic recrystallization of matrix. Foliation defined by elongation of grains and alignment of biotoite. Photo from Klier (2019). (B) Outcrop photo showing gneiss intruded by boudinaged granitic dykes. Gneissic layering is very fine and difficult to see in this photo. chlortizied biotite bands define foliation (Photo 9A). Garnetiferous hornblende gneiss and schist are folded with high-amplitude isoclinal folds with a persistent ENE strike. Small (F2) folds plunge gently east-northeast. These are folded F1 isoclinal folds, and a few hinges can be found in the outcrop. Some of the granitic pegmatites appear to be folded as well or are slightly boudinaged (Photo 9B), suggesting that pegmatites intruded prior to F2 or where exploiting layering within the folded gneisses and amphibolites during emplacement. Xenoliths of Photo 10 – Typical intrusive relationship between pegmatite surrounding amphibolites and gneisses. Small pegmatite veinlets and en echelon fracturing along margins is common resulting in lens-shaped gneissic fragments. 18 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Stop #4 – Amphibolites and Gneisses at Cornell Dam outcrops almost continuously for 4 km down the river. The amphibolite could also be classed as a gneissic, mafic hornblende tonalite or hornblende gneiss. This is also one of the few areas that this trip visits that you can potentially see primary depositional features! Immediately below the dam, the rock is a fine-grained amphibolite with elongate bulbous inclusions that appear to stretched pillows (Photo 11A) that contain local irregular to lens-shaped quartz-epidote nodules (Photo 11B). Lat: 45.1625° Long: -91.1596° The amphibolite is composed of subhedral to anhedral, lensoidal hornblende clusters (54%) with coarse, lensoidal porphyroclasts of twinned plagioclase (28%) and fine-grained quartz. Banding in the amphibolite Is cut by lenticular segments of granite and quartz veinlets. Garnets This outcrop is located within the spillway of the hydroelectric dam near the community of Cornell. About 500 m southwest from the bridge into Cornell on Highway 178 is the Wisconsin Department of Natural Resources Ranger Station where there is ample parking. Just south of the Ranger Station is a small road (called Pine Point Road) that leads toward the water. There are foot trails and gated roads (accessible by foot) that lead toward the outcrops at the dam and by the river. These outcrops are uneven but are generally dry and easily traversed under normal river conditions. If water levels are high, outcrops can also be visited on the shoreline above the dam near the Municipal Works buildings in Cornell. Photo 11 – Flow-like features in the amphibolites at Cornell Dam. (A). Streched pillow-like structures with cm-scale darkened pillow margins. (B) Irregular quartz-epidote nodules that are common in submarine or hydrothermally-altered submarine flows. Myers et al (1980) described this location as a laminated (foliated) garnet amphibolite that 19 �Proceedings of the 69th ILSG Annual Meeting - Part 2 in the amphibolite tend to be moderately poikioblastic with minor rotational features. The distribution of garnet clusters shows little relation to banding. Trace element chemistry of these rocks show clear tholeiitic trends and flat REE patterns on normalized diagrams (Figure 5). Further downstream from the dam, the rock becomes notably lighter in color and there appears to be a lower percentage of amphiboles. These rocks share similar trace element patterns (Figure 5) and are interpreted to be genetically related. The reason for the change in texture may be due to increase structural modification and gneissic banding development. The outcrops are also intruded by mafic dykes that clearly cross-cut the dominant foliation (Photo 12). These dykes trend N40°W and are approximately 30-50 cm in apparent thickness with no obvious chill margin. Since these dykes cross-cut the structural fabric, they are assumed to be related to the mid-continent rift. However, no petrography or chemistry has been done to confirm this hypothesis. of the river, there is a small vehicle parking area and footpath that leads to the dam on 260th Avenue about 100 m west of the intersection with County Highway M. This trail will take you to the dam and carefully navigate to the north bank of the river downstream of the dam. To access the south bank of the river, drive to the end of Irvine Avenue before it turns into a private driveway. There are numerous small foot paths that will lead to the south bank of the river. Photo 12 – Gabbroic dyke intruding through amphibolites at Cornell Dam. Stop 5 – Amphibolites and Deformed Diorite at Holcombe Dam The outcrops at this location are considered part of the Pembine-Wausau terrane, or is it? While the location of the Eau Pleine Shear Zone becomes problematic in this region, Sims et al. (1989) consider the Jump River Shear Zone the northern boundary of the Marshfield Terrane. Magnetic lineaments mark this shear zone and extend it close to these outcrops. Depending on Lat: 45.2251° Long: -91.1289° As time permits, each side of the river at Holcombe Dam has different rock types to examine, but are vastly different approaches to see them. To visit the outcrops on the north bank 20 �Proceedings of the 69th ILSG Annual Meeting - Part 2 the map, the shear zone lies just north or just south of the outcrops at this location (e.g. Mudrey et al., 1987). The rocks at this stop are a gneissic quartz diorite intrusion and an amphibolite schist (Figure 11). So, Marshfield or Pembine-Wausau terrane? The newest geochronology from the region is inconclusive. The rock exposed on the southern bank of the river below the dam is a foliated amphibolebiotite schist (Photo 13). Klier (2019) describes this rock as banded at the microscopic scale. Quartz grains are well banded, fairly subhedral to anhedral and feature undulatory extinction. Their boundaries are somewhat irregular and indicative of bulging recrystallization. Amphibole grains are hornblende to tremolite. Biotite appears as brown to light green grains and typically feature Photo 13 – Photomicrograph in plane-polarized light of amphibole-biotite schist with trace amounts of epidote in a quartzofeldspathic matrix. Figure from Klier (2019). Figure 11 – Precambrian geologic map of the Holcombe Dam region. Unit Abbreviations: qd: quartz-diorite, bgn: banded gneiss, ams: amphibolite schist. Figure from Myers et al. (1980). 21 �Proceedings of the 69th ILSG Annual Meeting - Part 2 shear bands with cryptocrystalline biotite crosscutting crystals. Klier (2019) obtained a U/Pb zircon age via LA-ICPMS of 1858 ± 1.0 Ma (Figure 12). This age does not definitively put the rocks in this region in Marshfield or PembineWausau terrane. There does not appear to be any Archean inherited zircons, as one might expect if Penokean magmas were overprinting an Archean crustal block. Photo 14 – Outcrop of quartz diorite on north bank of Chippewa River at Holcombe Dam. quartz, with minor amounts of biotite, muscovite and epidote. The quartz diorite contains two types of inclusions: hornblende rich ultramafic inclusions and spotted mafic inclusions. Ultramafic inclusions occur along the northwest portion of the exposed quartz diorite, generally less than 0.5 meters in length, although one is at least 2 meters long. Ultramafic inclusions are composed of 7585% hornblende and 11—13% biotite with a small amount of plagioclase (meta pyroxenites?). Chlorite occurs as an alteration product of biotite and less commonly of hornblende and can compose more than 20% of the rock. Figure 12 – Tera-Wasserburg concordia diagram of amphibolitic schist on the south bank of the Chippewa River near Holcombe dam. Figure from Klier (2019). Stop 6 – Tonalites and quartz diorites at Cadott Bridge On the north bank of the river, the outcrop is predominately a synkinematic quartz diorite (Photo 14; Figure 13). The quartz diorite is a medium-grained, dark to medium grey rock with rusty weathering surfaces. It is faintly foliated and has white discontinuous bands and lenticles which are more quartz rich than the rest of the rock. Quartz diorite is composed of plagioclase (32-51%), quartz (11-31%) and mafic minerals (12-33%). Mafic minerals range from entirely hornblende to entirely biotite. The quartz diorite is cut by medium-grained granite pods with migmatitic contacts and by finer-grained dykes with sharp contacts. The granite is a pink, faintly foliated rock which locally contains porphyritic microcline grains reaching 1 cm in size. Granitic rocks consist of plagioclase, microcline and Lat: 45.9535° Long: -91.1508° This outcrop shows is easily accessible under the Main Street bridge in Cadott. Just north of the bridge near the Main Street-Yellow Street intersection there is a parking area on the north bank of the river. From this parking area, there are footpaths that lead to the waters edge. The predominant rock type here is foliated biotite quartz diorite to biotite tonalite (Figure 14) composed of plagioclase (An25-35, 30-55%), quartz (10-40%), hornblende (0-25%) and biotite (0-15%) (Myers et al. 1980). Mafic minerals are partly replaced by chlorite (of several varieties), epidote, and sericite. Magnetite (1-5%) is a by- 22 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 13 – Detailed outcrop map of the quartz diorite gneiss on the north bank of the Chippewa River at Holcombe Dam. Figure from Myers et al. (1980). Figure 14 – Precambrian geologic map of the region downstream of Cornell Dam. Figure modified from Myers et al. (1980). 23 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Photo 15 – Mylonitized biotite tonalite at the Cadott Bridge on the Yellow River. Photo 16 – Photograph of elongated xenoliths (?) of chlorite-rich metavolcanic rocks in biotite tonalite at Cadott Bridge on the Yellow River. product in the chloritization of hornblende. The tonalites have been mylonitized (Photo 15) and locally recrystallized and contain lenticular xenoliths of chlorite and epidote-rich metavolcanic(?) rock (Photo 16). The older cataclastic foliation is axial-planar to isoclinally folded pegmatite, aplite, and quartz layers. These rocks are cut by a pervasive N65-75°W trending foliation and mylonitic shear zones. Acknowledgements Despite decades of regular visits from groups from the University of Wisconsin-Eau Claire, the most extensive detailed maps and rock descriptions were provided by Paul Myers and collaborators in the 1980 ILSG guidebook (Myers et al, 1980). There is some recent research activity in the region, but between new but pending analyses and future ambitions, the descriptions and maps provided in that ILSG guidebook are the most detailed and accurate for the region. A lot of the geologic descriptions have been updated and figures have been digitized while adding new data and insights where available. West of the Yellow River bridge, foliated biotite tonalite encloses angular xenoliths of hornblende tonalite or amphibolite containing strongly deformed aplite and pegmatite stringers. Isoclinally folded quartz, aplite, and pegmatite veinlets exist as angular xenoliths in a lighter biotite tonalite. 24 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Hafften, D., and Radwany, M., 2018, Geothermobarometry of a Precambrian amphibolite from Cornell WI: Proceedings of the Institute on Lake Superior Geology 64th Annual Meeting, Iron Moutain, Michigan, p. 45-46. In addition, the authors of this guidebook would like to thank the countless undergraduate and graduate students that have worked on these outcrops and have continued to inspire new work in the region. Specific acknowledgement is deserving to Matt Leahy and his efforts in digitizing figures and compiling geochemistry for the guidebook. Hannack, G., and Radwany, M., 2018, HornblendePlagioclase thermometry of the Eau Claire River Complex, western Wisconsin: Proceedings of the Institute on Lake Superior Geology 64th Annual Meeting, Iron Mountain, Michigan, p. 47-48. References Aleinikoff, J.N., Walter, M., Kunk, M.J., and Hearn, P.P., Jr., 1993, Do ages of authigenic K-feldspar date the formation of Mississippi Valley–type PbZn deposits, central and southeastern United States? Pb isotope evidence: Geology, v. 21, p. 73– 76. Hart, T. R., Gibson, H. L., and Lesher, C. M., 2004, Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic massive Cu-Zn-Pb sulfide deposits: Economic Geology, v. 99, p. 1003-1013. Helz, R. T. 1987, Diverse olivine types in the lava of the 1959 eruption of Kilauea Volcano and their bearing on eruption dynamics. USGS Professional Paper 1350, p. 691-722. Brown, B. A., 1988, Bedrock geology of Wisconsin, west-central sheet, Wisconsin Geological and Natural History Survey Map 87–11b. Chan, L. S., Myers, P. E., and Hay, R. L., 1991, Features and significance of the PrecambrianCambrian contact in western Wisconsin. Institute of Lake Superior Geology 37th Annual Meeting, Eau Claire, Wisconsin, Field Trip Guidebook 2, 17 p. Klier, J. J., 2019, The Marshfield Terrane: Redefinition of origin through zircon geochronology and geochemistry: Unpub. M.S. thesis, Ball State University, 115 p. LaBerge, G. L., 1996, Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume, Proceedings of the 42nd Annual Meeting of the Institute on Lake Superior Geology, Cable, Wisconsin. DeMatties, T. A., 1989, A proposed geologic framework for massive sulfide deposits in the Wisconsin Penokean volcanic belt: Economic Geology, v. 84, p. 946-952. LaBerge, G. L., and Myers, P. E., 1984, Two early Proterozoic successions in central Wisconsin and their tectonic significance: Geological Society of America Bulletin, v. 95, p. 246-253. DeMatties, T. A., 1994, Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview: Economic Geology, v. 89, p. 1122-1151. Lui, J., 1997, K-Metasomatism in Uppermost Precambrian Rocks in West-Central, Wisconsin and Southeastern, Missouri. Unpub. PhD thesis, University of Illinois. 227p. DeMatties, T. A., 2018, Effects of paleoweathering and supergene activity on volcanogenic massive sulfide (VMS) mineralization in the Penokean Volcanic Belt, northern Wisconsin, Michigan and east-central Minnesota, USA: Implications for future exploration: Ore Geology Reviews, v. 95, p. 216-237. Lui, J. Hay, R. L., Deino, A. and Kyser, T. K., 2003, Age and origin of authigenic K-feldspar in uppermost Precambrian rocks in the North American Midcontinent. Geological Society of America Bulletin, v. 115, p. 422-433. DeMatties, T. A., 2022, Exploration-resource assessment of productive felsic volcanic centers in the Paleoproterozoic Penokean Volcanic Belt of northern Wisconsin, Michigan and east-central Minnesota, USA: Ore Geology Reviews, v. 141, article 104489. Myers, P. E., Cummings, M. L., and Wurdinger, S. R., 1980, Precambrian geology of the Chippewa Valley, Wisconsin, Institute of Lake Superior Geology 26th Annual Meeting, Eau Claire, Wisconsin, Field Trip Guidebook 1, 123 p. 25 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Mudrey, M. G., LaBerge, G. L., Myers, P. E., and Cordua, W. S., 1987, Bedrock geology of Wisconsin, northwest sheet, Wisconsin Geological and Natural History Survey Map 88-7. Quigley, A., 2016, Setting of the volcanogenic massive sulfide deposits in the Penokean Volcanic belt, Great Lakes region, USA: Unpub. M.S. thesis, Colorado School of Mines, 95 p. Schulz, K. J., and Cannon, W. F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4-25. Sims, P. K., Van Schmus, W. R., Schulz, K. J., and Peterman, Z. E., 1989, Tectonostratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean orogen: Canadian Journal of Earth Sciences, v. 26, p. 21452158. Van Schmus, W. R., 1980, Chronology of igneous rocks associated with the Penokean orogeny in Wisconsin: Geological Society of America Special Paper, v. 182, p. 159-168. Van Wyck, N., and Johnson, C. M., 1997, Common Lead, Sm-Nd, and U-Pb constraints on petrogenesis, crustal architecture, and tectonic setting of the Penokean orogeny (Paleoproterozoic) in Wisconsin: Geological Society of America Bulletin, v. 109, p. 799-808. Weber, E. M., and Lodge, R. W. D., 2022, New U/Pb Geochronology from the Proterozoic Penokean Orogen, Wisconsin: Implications for VMS Metallogeny: Society of Economic Geologists Annual Meeting, Denver, CO, paper P5.10. Welsch, B., Faure, F., Famin, V., Barronet, A., Bachelery, P., 2013, Dendritic Crystallization: A Single Process for all of the Textures of Olivine in Basalts?, Journal of Petrology, v.543, p. 539-574. Zi, J-W., Sheppard, S., Muhling, J. R., and Rasmussen, B., 2021, Refining the Paleoproterozoic tectonothermal history of the Penokean Orogen: New U/Pb age constraints from the PembineWausau terrane, Wisconsin, USA: Geological Society of America Bulletin, v. 134, p. 776-790. 26 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Field Trip 2 – Wisconsin’s Paleozoic stratigraphy and tour of Crystal Cave Carsyn Ames, Esther Stewart, William “Bill” Batten, Eric Stewart, Ian Orland Wisconsin Geological and Natural History Survey, University of Wisconsin- Madison, 3817 Mineral Point Rd. Madison, WI 53705 Introduction The Cambrian-Ordovician strata exposed in western Wisconsin were deposited during the major Sauk and Tippecanoe marine transgressions onto the interior of the Laurentian continent (Sloss, 1963). These rocks compose the regional aquifer system, host disseminated sulfide mineralization that contribute to groundwater contamination, and are locally mined as proppant for fracking in the oil and gas industry. Additionally, variable hardness of these units in part controls the formation of ledges and hillslopes in the fluvially-dissected Driftless Area of southwestern Wisconsin. During this field trip, we will focus on Cambrian and lower Ordovician strata of the Sauk sequence (Figures 1 and 2). We start our day touring Crystal Cave, a cave system developed along joints within the Ordovician Prairie du Chien Group dolostone. For the rest of the day, we will visit outcrop exposures of the Cambrian Jordan Formation, Tunnel City Group, and Wonewoc Formation sandstones, and if time permits- the Eau Claire and Mount Simon Formations. We hope this field trip will provide an opportunity to discuss similarities and differences between units deposited on the western side of the Wisconsin Arch and those deposited on the eastern side, where field trip authors have focused much of their work. Additionally, we welcome and encourage discussion between participants that have knowledge of or experience working with these stratigraphic units. Figure 1. Correlation of map units showing relative ages of Cambrian-Ordovician units. COpg: Parfreys Glen Formation, Ce: Elk Mound Group, Ctl: Lone Rock Formation of the Tunnel City Group, Ctm: Mazomanie Formation of the Tunnel City Group, Ctc: Tunnel City Group, Ct: Trempealeau Group, Opc: Prairie du Chien Group including the Oneota and Shakopee Formations, Oa: Ancell Group, including the St. Peter and Glenwood Formations, Osp: Sinnipee Group, including the Platteville, Decorah, and Galena Formations. From Stewart (in revision). Ages from Gradstein et al. (2020). Cambrian-Ordovician strata in the southern Lake Superior Region were deposited on an essentially flat continental shelf in a shallow epeiric sea well within the Laurentian continent (Figure 3, Runkel et al., 2012, 2020). These strata overlie Precambrian bedrock of variable ages across the Great Unconformity, a surface characterized by locally significant topographic relief and weathering and exposed in outcrops around the Eau Claire area. The regional paleogeography that controlled sediment source to sink was defined by several structural highs, including the Transcontinental Arch, Wisconsin Dome, and Wisconsin Arch, and several basins, including the Hollandale Embayment, Illinois Basin, and Michigan Basin (Figures 3 and 4, A very brief geologic history of the Cambrian-Ordovician strata in Western Wisconsin Regional depositional model and setting 27 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 2. Generalized stratigraphic column of Wisconsin. From: Bedrock Stratigraphic Units in Wisconsin Bedrock Stratigraphic Units in Wisconsin [small] - WGNHS. 28 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 3. From Runkel 2020, Figure 4. This figure illustrates a depositional model developed for southeastern Minnesota. The Cambrian- Ordovician strata in Wisconsin are thought to have been deposited in a similar fashion. Runkel et al., 1998). The field stops we will visit in western Wisconsin lie west of the Wisconsin Arch and straddle the eastern edge of the Transcontinental Arch and the southwest flank of the Wisconsin Dome. These structural highs were periodically subaerially exposed and eroded during deposition of Paleozoic units. dominated Great American Carbonate Bank (Figures 3 and 5; Runkel et al., 2012). Sandy Cambrian sediments of the Mt. Simon, Wonewoc, and Jordan Formations were deposited in shoreface, aeolian, wave-, and tideinfluenced settings within the inner detrital belt. Mixed, fine-grained sandstone, siltstone, shale, and carbonate of the Eau Claire Formation, Trempealeau and Tunnel City Groups were winnowed and trapped within a transitional, relatively deeper water moat that separated the inner detrital belt from the Great American Carbonate Bank (Runkel et al., 2012). Dolomite of the Prairie du Chien Group was deposited in relatively shallow water, subtidal to peritidal settings on this carbonate bank. Interfingering sandstone, shale, and carbonate record marine transgressions and regressions that caused reciprocal expansion and contraction of the facies belts. During sea level rise, the carbonate bank advanced landward as siliciclastic-dominated nearshore environments were drowned. Cambrian-Ordovician siliciclastic and carbonate units were deposited in a nearshore, sandstone-dominated inner detrital belt that passed offshore into a relatively deeper water moat, which in turn transitioned into a carbonate- Figure 4. from Runkel and others 1998, regional map showing locations of the Wisconsin Dome, Wisconsin Arch, Transcontinental Arch and Hollandale Embayment. Other depositional basins are shown on map (Michigan and Illinois Basins), as well as regional extent of Paleozoic units. 29 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 5. from Runkel et. al., 2012 showing the depositional environments that produce interfingering of different Cambrian-Ordovician siliciclastic and carbonate units across the Midwest. Conversely, during sea level fall, sandy nearshore facies of the inner detrital belt expanded seaward, limiting carbonate deposition. Paleozoic sedimentary rocks were gently folded and faulted in the Paleozoic, probably related to far-field effects of continental margin orogenic events. Structures in Wisconsin rarely exceed 200 feet in structural relief. Recent mapping in Wisconsin and Minnesota suggests folds and faults are probably related to reactivation of much older Precambrian structures (Figure 8). Deformation probably occurred in at least two pulses: once during the Ordovician (Mossler, 2006; Steenberg and Retzler, 2016; Stewart E.K., 2021) and at least once later in the Paleozoic (Heyl and others, 1959; Carlson, 1961). The importance of these folds and faults for groundwater studies is a topic of active interest. In northern Illinois, sandstones in the core of the Sandwich Fault zone have an order of magnitude reduction in horizontal hydraulic conductivity compared to the surrounding rocks (Hadley and others, 2020). In eastern Wisconsin, the Beaver Dam anticline is associated with a statistically significant increase in detection of dissolved arsenic in groundwater wells (Stewart E.D. and others, 2021). The Wisconsin Arch and its influence on Cambrian-Ordovician strata Strata deposited in areas east (for example, Dodge, Fond du Lac, and Jefferson Counties, Wisconsin) and west (for example, the outcrops we will visit today) of the Wisconsin Arch (Figure 4) were deposited in different sub-basins and tapped different local sediment source areas. In addition, the Dodge, Fond du Lac, and Jefferson County map areas were situated in more proximal locations on the Wisconsin Arch relative to today’s field stop locations. Therefore, the eastern sections include more pronounced exposure surfaces, condensed, or eroded sections, and typically include thinner and less abundant fine-grained intervals. Figure 6 shows a generalized stratigraphic column for Jefferson County (east of Wisconsin Arch), with accompanying pXRF elemental data. Figure 7 shows a stratigraphic column from Trempealeau County in western Wisconsin, south of this trip’s field stops, and west of the Wisconsin Arch. Regional and county scale mapping The most recent regional map for West-Central Wisconsin was published in 1988 by Bruce Structural observations on the CambrianOrdovician strata in Wisconsin 30 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 6. From Stewart (in revision), Bedrock geology of Jefferson County. Jefferson County is east of the Wisconsin Arch. 31 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 7. Core log, Gamma Ray log, and pXRF logs from the Arcadia core, Trempealeau County. Modified slightly from Zambito et al. (2018). Trempealeau County is west of the Wisconsin Arch. Brown (WGNHS). This map (Figure 9) includes all of the Paleozoic units we will see today, as well as the older Proterozoic and Archean rocks that make up the bedrock to the east of Eau Claire. Many of the stops for this field trip were found using the Cambrian contacts from this map. Field Trip Stops Stop 1: Crystal Cave, Spring Valley, WI (Contributed by Ian Orland, WGNHS) UTM location 4964692.71N) for stop (559180.94E, and includes walking, ducking, and climbing 7 stories. Please exercise caution while inside the cave as surfaces may be uneven. Below is a brief We will be touring the cave with staff from Crystal Cave. The tour is moderately strenuous 32 �Proceedings of the 69th ILSG Annual Meeting - Part 2 synopsis of a recent collaboration between WGNHS and UW-Madison’s Geoscience Department on speleothems from southern Wisconsin: Caves are fascinating natural features, and can preserve geologic records of past environments. Relatively recent advances in the methods and precision of geochemical analyses have established cave formations (speleothems) as important scientific tools for understanding climate changes of the last 500,000 years. A number of groups have studied the geochemistry of speleothems in the Lake Superior region. In Wisconsin, much of this work has happened at Cave of the Mounds in Blue Mounds, WI, just outside of the terminal moraine of the Laurentide Ice Sheet and some 20 miles southwest of Madison. While that cave is not the destination for this field trip, this section is intended to highlight the types of information we can learn from caves like Crystal Cave. Both caves are privately-owned show caves that were opened for tours in the late 1930s/early 1940s. Crystal Cave is situated in Prairie du Chien Group dolomites of the Early Ordovician (~475 Ma), while Cave of the Mounds is in Sinnipee Group dolomites of the Middle Ordovician (~465 Ma). The formation ages of passages in each cave are poorly constrained. Stalagmites and stalactites from Cave of the Mounds, however, have recorded environmental signals for >250,000 years. Cave of the Mounds: permafrost record Researchers from UW-Madison collected the first seven stalagmite samples in 2015 for modern U-Th geochronological analysis at UM-Twin Cities. Initial results prompted further sampling and analyses; Batchelor et al. (2019) reports 141 U-Th dates from 19 cave carbonate (speleothem) samples ranging from 250–2 ka. The temporal distribution of these ages revealed hiatuses of stalagmite growth in the cave during both of the last glacial maxima, demonstrating the presence and duration of permafrost (Figure 10). Notably, Figure 8. Cross-section from Dodge County, eastern flank of the Wisconsin Arch, south-central Wisconsin. Note offset of Precambrian basement and Cambrian Elk Mound Group (Ce) through Ordovician Prairie du Chien Group (Opc) and subtle folding of younger units. From Stewart E.K. (2021). 33 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 9. Bedrock geologic of west-central Wisconsin from Brown, 1988. 34 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 10 (*from Batchelor et al., 2019). Speleothem U‐Th ages at Cave of the Mounds (COM) in context with regional and global paleoclimate records. (a) MIS boundaries (odd numbers=interglacial periods, even numbers=glacial periods). (b) Stacked records of δ18O (‰) from benthic marine foraminifera annotated with MIS substage names (Lisiecki & Raymo, 2005). (c) Summer insolation (21 June at 43°N). (d) Atmospheric CO2 (ppm) and (e) CH4 (ppb) concentrations. (f) U‐Th ages from COM speleothems with associated 2σ uncertainties (this study) and statistically significant growth hiatuses (gray and red vertical bars). (g) Paleo‐permafrost reconstructions based on geomorphic features in Wisconsin, including ice wedge casts and polygons (Clayton et al., 2001). (h) U‐Th dates of speleothems from caves in the Midwestern United States in order of decreasing latitude. References provided in the main text. MIS = Marine Isotope Stage changes in the δ18O signal during a time period when warm periods are recorded in polar ice cores and stronger monsoons are recorded in tropical stalagmites (Figure 11). the 18 ky duration of the growth hiatus at MIS 2 was much longer than the hiatus that overlaps MIS 6 (5 ky), consistent with more extensive continuous permafrost in the region during the last glacial period. A combination of microscopic imaging and analysis showed that the δ18O changes each happened in ~10 years, and comparison to a climate model demonstrated that the δ18O changes likely happened as a result of >10°C warming above the cave. These results speak to how quickly and dramatically those polar Cave of the Mounds: Decadal warming events during the last glacial period Earlier this year, Batchelor et al. (2023) published a record of the oxygen isotope ratios (δ18O) of calcite from a Cave of the Mounds stalagmite that grew during the last glacial period. Their interpretation focused on a number of rapid 35 �Proceedings of the 69th ILSG Annual Meeting - Part 2 warming events were propagated across the Laurentide Ice Sheet, which is important for better understanding the dynamics of rapid climate change. As you enjoy the tour of Crystal Cave, consider what geologic stories might be captured in its Figure 11 (*from Batchelor et al., 2023). Stalagmite CM-5 δ18O record in comparison to other regional δ18O records of the last-glacial period. a, Cave of the Mounds (COM; this study) δ 18O record (black line), with associated U-Th ages (black dots/2SD error). Note the error of our age model ranged from 520 to 2800 years and was on average 730 years. b, A stalagmite δ18O record from Buckeye Creek Cave, WV (red line) showing relatively lowmagnitude δ18O changes during the last glacial period. c, A compilation of Chinese speleothem δ 18O records (orange line), showing high-magnitude δ18O changes, which reflects the sensitivity of the East Asian monsoon system to high-latitude warmings (DO events) during the last glacial period. *Note the scale of the y-axis in panels A-C are the same to allow for one-to-one comparison. d, The North Greenland Ice Sheet Project (NGRIP) δ18O record (blue line), showing the timing of abrupt warming DO Events (labeled #s). 36 �Proceedings of the 69th ILSG Annual Meeting - Part 2 speleothems. If you have ideas or questions, feel free to reach out to Ian Orland (orland@wisc.edu)! Stop 2: Prairie Du Chien Group- Kraemer Quarry Entrance Outcrop- 850th Ave between Lincoln Rd and 870th Ave intersections. UTM location 4966004.74N) for stop (564719.58E, We do not have permission to enter the quarryDo not enter the quarry. There is an outcrop of the Prairie Du Chien Group just outside of the quarry gate that continues down the hill from the quarry entrance. This outcrop appears to be a very sandy portion of the Prairie Du Chien Group, possibly representing the lower most Stockton Hill Member of the Oneota Formation, or an interfingering of the Jordan Sandstone within the basal Prairie Du Chien Group. Figure 12. Massive beds, of sandy, carbonate cemented Prairie Du Chien Group. Just to the right of the quarry entrance are massive, 1-2m thick beds (Figure 12). To the left, and down the hill, the massive beds continue and just below them thinly bedded, lighter color units begin to appear (Figure 13). The portion of the outcrop that continues down the hill also contains what may be the Prairie Du Chien Gp./Jordan Fm. contact in the ditch just below the road grade between the outcrop and the road (Figure 14). While not recognized in the formal bedrock stratigraphic column for Wisconsin, the thinly bedded, lighter color units may also represent the Coon Valley Member of the Oneota Formation, often recognized and mapped in Minnesota (Steenberg, J.R., and Retzler, A.J., 2016). We will depart on 850th Ave. by continuing down the slope. To the left near the toe of the slope, there is a valley floor with a barn and small pasture. Figure 13. Massive beds of Prairie Du Chien GroupStockton Hill Member? Possibly atop thinly bedded interfingerings of Jordan sandstone. Looking across the valley floor, there is an outcrop of Jordan sandstone just across the creek (Figure 15). 37 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 15. View of valley floor, at toe of slope driving down 850th Ave., looking across pasture towards Jordan outcrop just across creek. 18). This outcrop (Figure 17) is likely close to the base of the Figure 14. Arrow pointing to possible Prairie Du Chien Gp/Jordan Fm. contact in ditch just below road grade. Jordan/St. Lawrence contact (labeled in Figure 16), and the floodplain that the Eau Galle River runs through most likely represents the top of the St. Lawrence Formation. Stop 3: Jordan Formation- Cth B and 770th (Spring Lake, Wisconsin) UTM location 4962594.39N) for stop The Jordan Formation of the Trempealeau Group has been highly studied in both Wisconsin and Minnesota (Mudrey, M.G. Jr. ed, 1997 and references therein). The distinction between and regional application of the quartzose and feldspathic sandstones in this formation have also been debated (Runkel 1994 and Byers and Dott, 1995). Overall, the Jordan Formation represents a coarsening upward sequence that is conformable (562605.40E, Lithofacies of the Jordan Formation are described in Runkel, 1994 and are as follows: 1) very fine-grained hummocky cross-stratified and burrowed sandstone, 2) fine-grained, trough cross-stratified and burrowed sandstone, 3) medium- to coarse-grained, large-scale crossstratified sandstone and 4) thinly interbedded sandstone, mudstone and shale. They note that lithofacies 4 may only be relevant to certain areas in Minnesota (Figure 18). Authors are open to discussion as to where this particular outcrop falls in Runkel’s 1994 classification schema (Figure Figure 16. View of the outcrop across Cth B with approximate contacts labeled. 38 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Iron staining Figure 17. A. View of outcrop. Note the cross bedding and iron staining above hammer. B. Possible iron concretions? PDC Jordan Silcrete at its basal contact with the St. Lawrence, and unconformable with the Prairie Du Chien Group contact at its top. The Coon Valley member of the Jordan is not formally recognized in the stratigraphic column of Wisconsin (WGNHS Figure 19. Photo of core from Jefferson County, Wisconsin (east of Wisconsin Arch) modified from Kusick, 2022 M.S. Thesis. This interval of core shows the contact between the Jordan Formation and the Prairie Du Chien Group. 2011), though it has been noted above the Jordan Fm. in southern parts of the state. A recent M.S. dissertation (Kusick, 2022) discussed, in detail, both the stratigraphy and depositional environments of the CambrianOrdovician units east of the Wisconsin Arch. Kusick (2022) describes the Jordan sandstone as being comprised of only 2 facies of cross stratified sandstone and shaly sandstone, and as being deposited in an upper to lower shoreface environment. These authors would also like to note that locally, the Jordan Formation east of the Wisconsin Arch includes silcrete and clay, and hosts disseminated sulfides (Figure 19). Stop 3a: Rock Elm Impact Structure- Rock Elm, WI- lunch at Nugget Lake County Park UTM location 4948450.76N) Figure 18. Figure 2 from Runkel, 1994 illustrating the different lithofacies of the Jordan Sandstone in Minnesota. 39 for stop (561573.58E, �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 20. Core photos showing jumbled and deformed Cambrian strata from the southern edge of the Rock Elm central uplift - from WGNHS archives. No set stop, informational only as we’ll be eating lunch at a park within the crater. detected in detrital zircon grains and is interpreted to be caused by the impact (Cavosie et al., 2015). The field trip will go through the Rock Elm impact structure (Figure 20), located in Pierce County around 35 miles WSW of Eau Claire (Figure 21). The Rock Elm impact structure is the largest deformation event recorded in the Paleozoic section of western Wisconsin. The structure contains a 6.5 km diameter ring boundary fault and a central uplift 1 km across (Cordua, 1985). Where control exists, the ring boundary fault is thought to have accommodated 45 meters of down-in-the-center displacement (French and others, 2004). Much of the interior of the ring boundary fault is filled with the relatively flat-lying Rock Elm shale and the overlying Washington Road sandstone, which have a combined thickness of approximately 48 meters. These units are unique to the area, and do not exist outside of the ring fault. These units are described based on numerous outcrops, many given in Cordua (1987) and Cunningham and others (2011). The central uplift contains outcrops of tilted Mt. Simon Formation (Figure 20), which suggests 250 to 300 meters of uplift within the core zone relative to rocks outside the impact structure (French et al., 2004). Reidite, a high pressure polymorph of zircon, has been Stop 4: Skolithos burrows in Tunnel City Group-330th Ave. between HWY 25 and Cth Y (Private Property!!!) UTM location 4961335.76N) for stop (586184.90E, We will park on a private drive and walk east along the road to this outcrop. This stop in the Tunnel City Group is an excellent example of Skolithos burrows (Figure 22) which are common in the Tunnel City Group. This outcrop is likely the Tomah Member of the Lone Rock Formation. Excellent examples of cross-stratification can be seen at this outcrop as well. 40 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 21. Map plate from the 2007 Wisconsin Geological and Natural History Survey Open File Report on the Rock Elm impact structure (Cordua and Evans, 2007). 41 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 22. Tunnel City Group outcrop with excellent examples of Skolithos burrows and possibly multiple types of cross-stratification. “Tunnel City Group (Cambrian) The Tunnel City Group is comprised of the Lone Rock and Mazomanie Formations. Similar to neighboring La Crosse County (Evans, 2003), the Mazomanie Formation was not recognized in Trempealeau County. Stop 5: St. Lawrence Formation/Tunnel City Group road cut- Cth C and Cth Y UTM location 4958918.89N) for stop (587368.83E, Lone Rock Formation. The Lone Rock Formation (Figure 23) and its members are identifiable in the map area. The members, from oldest to youngest, are Birkmose, Tomah, and Reno; these are not differentiated at the map scale. The Birkmose Member is a dolomitecemented, coarse-grained, glauconitic sandstone to sandy dolostone with flat-pebble conglomerates; the Tomah Member is a tan to white-colored, medium-grained, glauconitic quartz sandstone; and the Reno Member is a glauconitic medium- to coarse-grained quartz sandstone with flat-pebble conglomerates. Palaeophycus and Skolithos are common, as is hummocky cross-stratification and crossstratification bounded by horizontal bedding surfaces. The contact with the overlying St. Lawrence Formation is sharp and unconformable. The road cut is just west of the intersection of Cth C and Cth Y. We will park and walk to this outcrop. Cth C is a fairly busy road, please exercise caution when you decide to cross. Recent mapping in Trempealeau County, southeast of stops 5 and 6, has produced interesting work on both the geological relationships of the Cambrian- Ordovician rocks, and the quality of groundwater in the west-central part of the state. Zambito and others, 2018 published the following unit description for the Tunnel City Group: 42 �Proceedings of the 69th ILSG Annual Meeting - Part 2 phyllosilicate mineral. Another interesting part of this road cut is the bench approximately 35ft up. This bench feature likely represents the unconformable contact Zambito and others, 2018 alluded to with the overlying St. Lawrence Formation. The Mazomanie Formation is generally not observed in this part of the state and is more prevalent in the southern parts of the state where it interfingers the Lone Rock Formation (Mudrey, M.G. Jr. ed, 1997 and references therein). St. Lawrence Fm. above bench Reno Member Tomah Member A 2019 study by Zambito and others investigated the relationship between groundwater quality and the geochemistry of the Tunnel City-Wonewoc units in western Wisconsin. This study notes that sulfide bearing minerals are disseminated between the two units in west-central Wisconsin, and they call for more work to better understand the geochemical effects of oxidation of sulfide minerals during groundwater pumping in this part of the state. Our next stop will be at an outcrop of Wonewoc sandstone, and we will pass other outcrops of this unit on our drive. Figure 23. Road cut showing contacts between the St. Lawrence Fm., and Reno and Tomah Mbrs. of the Lone Rock Fm. This is the view from the north side of Cth C. The Lone Rock Formation is commonly exposed in shale pits, along roads leading to ridgetops, and at the top of sand mine high walls where the Wonewoc Formation is extracted and the Birkmose Member forms the caprock. The formation is approximately 150 feet thick in the map area. Elemental data for part of the Lone Rock Formation is shown in plate 2 [figure 7]. These data show the formation’s lithologic variability, in particular the distinct upper carbonate-cemented and lower sandstone dominated intervals in the Birkmose; the lower interval consists of reworked quartz grains from the underlying Wonewoc with interspersed, rare glauconite grains and phosphatic brachiopods.” These authors find this to be an excellent, and representative description of the unit for the westcentral region. – Zambito and others (2018) Overall, the Tunnel City Group both east and west of the Wisconsin Arch are quite similar. As examined at this stop, west of the Arch, the Tunnel City Gp. East of the Arch is also a quartz sandstone with glauconite and trace amounts of shale. Stop 6: Wonewoc road cut- Cth Y UTM location 4959793.83N) Figure 24 shows a small part of the westernmost portion of the outcrop. The very dark, greenish-black bed just below the more resistant dolomitic bed is rich in the mineral glauconite, which is an iron potassium for stop (593272.27E, The Wonewoc Formation is a fine to coarse sandstone unit with medium to thick beds, highangle trough cross-stratification and some 43 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 24. A. View of westernmost point of the outcrop. B. close up of the phosphatic rich, friable sandstone of the Tomah Member. orange, iron rich bed on the south side of Cth Y that doesn’t seem to appear in the north face of the outcrop. feldspar (Mudrey, M.G. Jr. ed, 1997). Brachiopod fragments, Skolithos burrows (which we observed at stop 4), and Climactichnites are somewhat common in the fossiliferous Ironton Member of this formation. The upward contact (Figure 25) with the Tunnel City Group is gradational and fines upward; the basal contact with the Eau Claire Formation has been debated as to whether it is gradational or not (Mudrey, M.G. Jr. ed, 1997 and Ostrom 1978). Note the In the 1990s, to better characterize aquifer and confining units, the Minnesota Geological Survey began focusing on the hydrostratigraphic characteristics of geologic units (1998 ILSG field guide). Hydrostratigraphic subdivisions include: 1) fine clastic; 2) coarse clastic; 3) carbonate; 4) clastic/carbonate mix. While this approach has not been implemented as part of bedrock mapping in Wisconsin, its importance has been recognized by Wisconsin hydrogeologists in lithologically complex units such as the Eau Claire Formation (Bradbury and Runkel, 2011). The authors are open to questions, and discussion of this method as it may pertain to future groundwater study needs across the Midwest. Additional stops if time permits: Devil’s Punchbowl – Eau Claire Formation: UTM location 4966909.29E) Figure 25. View of the Wonewoc outcrop on Cth Y. 44 for stop (582783.00N, �Proceedings of the 69th ILSG Annual Meeting - Part 2 Park in the parking lot, walk east towards the stairs, and take them down the trail into the Punchbowl. Acknowledgements A special thanks to Eric Stewart and Ian Orland for contributing content to this guide. A very special thanks to Bill Batten for helping scout field locations and always knowing where to find the best contacts. Additionally, this guide would not have been possible with consulting Dave LePain’s mapping notes and field guides from Pierce and St. Croix counties and the work of others who have previously published work on these Paleozoic units. This is a classic stop for field trips in this part of the state. Devil’s Punchbowl is managed by the Landmark Conservancy- please do not use rock hammers on the outcrops and be good stewards of the landscape. This outcrop shows the relationship between the Eau Claire Formation and the Wonewoc Formation. Expect to see a fine-grained sandstone with swaley cross beds in the Eau Claire Formation, and mediumto coarse-grained, cross-stratified sandstone in the Wonewoc Formation at this location (Mudrey, M.G. Jr. ed, 1997). References Batchelor, C. J., Marcott, S. A., Orland, I. J., He, F., and Edwards, R. L., 2023. Decadal warming events extended into central North America during the last glacial period. Nature Geoscience 16: pages 257261, Hwy 37/Hendricks Ave and Silver Springs Dr.- Mt. Simon Formation: UTM location 4958712.01E) for stop (615774.61N, Batchelor C. J., Orland I. J., Marcott S. A., Slaughter R., Edwards R. L., Zhang P., and Li X., 2019. Distinct permafrost conditions across the last two glacial periods in mid-latitude North America. Geophysical Research Letters 46: pages 1331813326, This is a typical Mt. Simon Formation exposure (Figure 26), coarse- to mediumgrained, cross-bedded, iron stained, sandstone, interbedded with shale and fine grained sandstone (Mudrey, M.G. Jr. ed, 1997). Bradbury, K. R., & Runkel, A. C., 2011. Recent advances in the hydrostratigraphy of Paleozoic bedrock in the Midwestern United States. GSA Today, v. 21, pages 10-12. Byers C.W. and Dott R.H. Jr., 1995 Sedimentology and depositional sequences of the Jordan Formation (Upper Cambrian), Northern Mississippi Valley, Journal of Sedimentology, v. B65, no.3, pages 289-305. Cavosie, A. J., Erickson, T. M., & Timms, N. E., 2015. Nanoscale records of ancient shock deformation: Reidite (ZrSiO4) in sandstone at the Ordovician Rock Elm impact crater. Geology, 43(4), pages 315-318. Cordua, W. S. 1985. Rock Elm structure, Pierce county, Wisconsin: a possible cryptoexplosion structure. Geology, 13(5), pages 372-374. Figure 26. View of the Mt. Simon Formation outcrop. This location is heavily iron stained and exhibits excellent examples of sedimentary structures like trough cross stratification and channel forms. Cordua, W. S., 1987. The Rock Elm Disturbance, Pierce County Wisconsin, in Balaban, N. (ed.), Field trip guidebook for the Upper Mississippi 45 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Valley, Minnesota, Iowa and Wisconsin, prepared for the 21st annual meeting of the Geological Society of America North-central section, Minnesota Geological Survey Guidebook Series #15, pages 123-152. Central Section, Geological Society of America, May1-2, 114 pages. Ostrom, M.E.,1978. Stratigraphic relations of Lower Paleozoic rocks of Wisconsin, Wisconsin Geological and Natural History Survey Field Trip Guidebook 3, pages 3-22 Cordua W.S. and Evans T.J., 2007. Geology of the Rock Elm Complex, Pierce County, Wisconsin, Wisconsin Geological and Natural History Survey Open File Report WOFR2007-02, Map, 1 plate. Runkel A.C., 1994. Deposition of the uppermost Cambrian (Croixian) Jordan Sandstone, and the nature of the Cambriand-Ordovician boundary in the Upper Mississippi Valley, Geological Society of America Bulletin, vol. 43: pages 60-71 Cunningham, J., Dolliver, H., and Cordua, W., 2011. Flaming meteors, dark caves and raging water: geological curiosities of western Wisconsin, in Miller, J.D, Hudack, G., Wittkop, C., and McLaughlin, P.I. (eds.), Archean to Anthropocene: Field Guides to the Geology of the Mid-continent of North America, Geological Society of America Guidebook Field guide 24, pages 411-424. Runkel A.C. McKay, R.M., and Palmer, A.R., 1998. High-resolution sequence stratigraphy of lower Paleozoic sheet sandstones in central North America: The role of special conditions of cratonic interiors in development of stratal architecture. GSA Bulletin, v.110 no.2., pages 188-210. doi:10.1130/B26117.1 French, B. M., Cordua, W. S., & Plescia, J. B., 2004. The Rock Elm meteorite impact structure, Wisconsin: Geology and shock-metamorphic effects in quartz. Geological Society of America Bulletin, 116(1-2), 200-218. Runkel, Anthony C., Robert M. McKay, Clinton A. Cowan, James F. Miller, and John F. Taylor, 2012, The Sauk megasequence in the cratonic interior of North America: Interplay between a fully developed inner detrital belt and the central great American carbonate bank, in J. R. Derby, R. D. Fritz, S. A. Longacre, W. A. Morgan, and C. A. Sternbach, eds., The great American carbonate bank: The geology and economic resources of the Cambrian – Ordovician Sauk megasequence of Laurentia: AAPG Memoir 98, p. 1001 – 1011. Carlson, J.E., 1961. Geology of the Montfort and Linden Quadrangles, Wisconsin, in Geology of parts of the Upper Mississippi Valley zinc-lead district: U.S. Geological Survey Bulletin 1123– B, pages 95–138, 2 pls., scale 1:24,000, Gradstein, F.M., Ogg, J.G., Schmitz, M.D. and Ogg, G.M. eds., 2020. Geologic time scale 2020. Elsevier. Runkel, A.C., 2020. Minnesota at a Glance Paleozoic History of Southeastern Minnesota-Ancient Tropical Seas. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, Heyl, A.V., Jr., Agnew, A.F., Lyons, E.J., Behre, C.H., Jr., and Flint, A.E., 1959, The geology of the Upper Mississippi Valley zinc-lead district: U.S. Geological Survey Professional Paper 309, 310 pages., 24 pls. Sloss, L.L., 1963. Sequences in the cratonic interior of North America. Geological Society of America Bulletin, 74(2), pages 93-114. Kusick, A. R., 2022. Stratigraphy, Sedimentology, and Deformational Significance of Cambrian and Early Ordovician Strata Along the Southeast Wisconsin Arch (M.S. dissertation, The University of Wisconsin-Milwaukee). Steenberg, J.R., and Retzler, A.J., 2016. Bedrock geology, plate 2 of Geologic atlas of Washington County: Minnesota Geological Survey County Atlas Series C–39, Part A, scale 1:100,000, Mossler, J.H., 2006, Bedrock Geology of the Prescott quadrangle, Washington and Dakota counties, Minnesota: Minnesota Geological Survey Miscellaneous Map Series M–167, scale 1:24,000, Stewart E.D., Stewart E.K., Bradbury, K.R., Fitzpatrick, W.A., 2021. Correlating Bedrock Folds to Higher Rates of Arsenic Detection in Groundwater, Southeast Wisconsin, USA, Groundwater, v59, no.6, pages 829-838. Mudrey, M.G. Jr. ed, 1997. Guide to field trips in Wisconsin and Adjacent areas of Minnesota. Prepared for the 31st Annual meeting of the North- 46 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Stewart, E.K., 2021. Bedrock geology of Dodge County, Wisconsin: Wisconsin Geological and Natural History Survey Map Series M–508, scale 1:100,000, Stewart (in revision). Bedrock Geologic map of Jefferson County, Wisconsin: WGNHS Map Series, 1 plate, 1:100,000-scale. Wisconsin Geological and Natural History Survey [WGNHS], 2011, Bedrock stratigraphic units in Wisconsin: Wisconsin Geological and Natural History Survey Educational Series 51, 2 p. Zambito J.J IV, Mauel, S. W., Haas, L.D., Batten, W.G., Chase, Streiff, C.M., P.M., Niemisto, E.M., Heyrman, E.J., 2018. Preliminary Bedrock Geology of Southern Trempealeau County, Wisconsin, Wisconsin Geological and Natural History Survey Open File Report WOFR2018-01, 2 plates scale 1:100,000, 27 pages Zambito J.J IV, Haas, L.D., Parsen, M.J., McLaughlin, P.I., 2019. Geochemistry and mineralogy of the Wonewoc-Tunnel City contact interval strata in western Wisconsin, Wisconsin Geological and Natural History Survey Open File Report WOFR2019-01: 28. 47 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Field Trip 3 – Precambrian Geology of the Eau Claire River Valley: Re-discovering the Eau Claire Volcanic Complex Robert W.D. Lodge, Evan M. Weber, Robert L. Hooper Department of Geology & Environmental Science, University of Wisconsin-Eau Claire, Eau Claire, Wisconsin 54701 of the “prove-it-first” law, or sulfide mining moratorium, in 1997 effectively shut down mineral exploration and mining activities in the region. More recently, the mineral exploration industry has been reinvigorated because of the 2002 discovery of the 18.2 Mt Back Forty deposit in Michigan, easing of the sulfide mining moratorium in 2017, and a recent national push for securing domestic critical mineral resources. However, this has also highlighted the lack of modern datasets on Wisconsin’s mineral deposits that could be used to further our knowledge of the mineral-forming systems in the belt. The Pembine-Wausau Terrane has received most of the historic and recent attention since it hosts approximately 150 million tonnes of known VMS mineralization. However, little attention has been given to the Penokean volcanic deposits that overprinted the Marshfield Terrane that are presented in this guidebook. These volcanic deposits host a VMS prospect (Butler Prospect) and therefore the geodynamic setting of these volcanic rocks clearly are favorable for submarine hydrothermal activity. DeMatties (2022) recognized the gap in knowledge for these Penokean volcanic deposits, known as the Eau Claire Volcanic Complex, within the Marshfield Terrane and their exploration potential. It is a little embarrassing how little we know about the Eau Claire region considering the mineral wealth of the rest of the orogen. Current research at the University of Wisconsin-Eau Claire is aimed at the addressing this issue. Introduction The erosional outliers of Precambrian bedrock in the Eau Claire River valley represent the southernmost extent of the Canadian Shield before it is completely covered by Paleozoic sedimentary strata. The rocks exposed here are part of the Paleoproterozoic Penokean Orogeny, a collisional orogen that resulted from the accretion of the Pembine-Wausau and Marshfield terranes onto the southern margin of the Superior Province. This region was last visited by members of the Institute of Lake Superior Geology in 1980 when a field trip through the region was conducted by Paul Myers and colleagues (Myers et al., 1980) when it was called the “Chippewa Amphibolite Complex”. Since then, the “Eau Claire River Complex” was defined and described in detail by Cummings (1984). There has been ‘new’ U/Pb data collected by the USGS (Sims et al. 1989) and others (Van Wyck et al, 1997; Klier, 2019; Weber and Lodge, 2022), regional syntheses of the Penokean volcanogenic massive sulfide (VMS) mineralization (DeMatties 1989; 1994; 2018; 2022), maps published by government surveys (Brown, 1988), and orogen-wide tectonic model (Shultz and Cannon, 2007) that is being revisited based on new U/Pb data (Zi et al., 2021). The rocks that will be visited on this trip are a critical part of evaluating the tectonic models for the Penokean Orogen and have not been examined using modern analytical techniques. The Penokean Orogen is perhaps best known for hosting numerous VMS deposits. In fact, one of the most complete descriptions of several deposits was published by the Institute of Lake Superior Geology (LeBarge, 1996). The passing The portion of the Eau Claire Volcanic Complex that is visited in this guidebook is not well exposed and its regional context is poorly constrained. Students from the University of 48 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Wisconsin-Eau Claire have been visiting Big Falls and Little Falls locations in this guidebook for decades to learn how to map and describe rocks in the field, measure structures and interpret geologic histories, and learn the basic mechanics of field work. Faculty, students, and alumni from Eau Claire consider these outcrops classic. This guidebook will (re-)introduce these rocks and present some of the ongoing research with the Eau Claire Volcanic Complex. The outcrops visited in this guidebook are accessible by foot, but many others were accessed by kayaking in the Eau Claire River. Ongoing research in this region hopes to expand the lithogeochemistry and zircon petrochronology database to better delineate the geodynamic evolution and crustal architecture of this region. Determining the presence or absence of Archean basement throughout the Marshfield terrane will help refine terrane boundaries and improve our understanding of the metallogeny of the region to assist in future mineral exploration efforts. Regional Geology The Paleoproterozoic Penokean Orogen (ca. 1.8 Ga) in the Lake Superior region (Figure 1) is a classic Precambrian orogenic belt comprised of Figure 1: Geologic map of the major tectonic assemblages and major structures of the Penokean Orogen. Notable abbreviations that are important for this guidebook are EPSZ, Eau Pleine shear zone; NFZ, Niagara fault zone. Figure from Shultz & Cannon (2007). 49 �Proceedings of the 69th ILSG Annual Meeting - Part 2 dominantly submarine volcanic rocks formed in a suprasubduction zone setting that are now structurally juxtaposed along the southern edge of the Archean Superior Province during the earliest phases of forming the Columbia, or Nuna, supercontinent (LaBerge and Myers, 1984; Sims et al., 1989; Schulz and Cannon, 2007). The orogen is host to at least 150 million metric tonnes (Mt) of VMS and associated mineralization (DeMatties, 1994, 2018) but remains one of the more poorly understood and underexplored mineral districts in North America. The Penokean Orogen has been divided into the Interior and Exterior domains. These domains are sutured by the Niagara Fault Zone (Figure 1). The Exterior domain consists of passive margin, rift, and forearc basin sediments and Archean crustal blocks from the Superior Province that were deformed in the folded and faulted foreland part of the orogen. The Interior Domain consists of two accreted terranes, the Pembine-Wausau and Marshfield terranes. These terranes are sutured by the Eau Pleine Shear Zone (Figure 1). The PembineWausau Terrane is a composite accreted oceanic arc overprinted by continental-margin magmatism and hosts numerous VMS deposits and occurrences (DeMatties, 1994; Shultz & Cannon, 2007) (Figure 2). The Marshfield Terrane is composed of Archean crustal fragments of unknown origin that was overprinted by Penokean-aged magmas during the Penokean orogen (Figure 2) and is described in more detail in the sections to follow. Figure 2 - Schematic tectonic evolution of the Penokean Orogen provided by Shultz and Cannon (2007) based on geophysical, sedimentological, and geochronological compilations. continental arc volcanism and back arc extension developed until about 1850 Ma until the collision with the Marshfield terrane began. During this ocean closure, a double subduction zone with concurrent northward and southward subduction resulted in arc magmatism on both the PembineWausau and Marshfield terranes. Sedimentation related to this convergence in a foreland basin setting continued until about 1835 Ma. The end of the orogen was constrained by undeformed post-tectonic plutons dated at 1830 Ma that stich shear zones. Shultz and Cannon (2007) synthesized the tectonic events that formed the Penokean Orogen (summarized in Figure 2) based on a detailed compilation of lithologic, structural, sedimentological, and geochronological datasets. This classic model proposed that an oceanic arc, now the Pembine-Wausau Terrane, collided with the southern margin of the Superior Province around 1880 Ma. Following a subduction flip from south-directed to north-directed subduction, 50 �Proceedings of the 69th ILSG Annual Meeting - Part 2 However, this classic tectonic model for the evolution of the Penokean Orogen has recently been re-evaluated considering new U/Pb data. The first contradictory data came when Quigley (2016) obtained a high-precision U/Pb zircon age of 1832.98 ± 0.52 Ma from a rhyolite at the Back Forty deposit via CA-ID-TIMS. This younger age was in stark contrast to the other VMS deposits that yielded U/Pb zircon ages of ca. 1875 Ma. Additional U/Pb zircon ages reported by Zi et al. (2021) from volcanic units (Beecher Formation) and plutonic rocks (Dunbar Gneiss, Newingham Tonalite) in the eastern part of the orogen supported the younger extensional tectonic event proposed by Quigley (2016). These new ages resulted in a revised Penokean tectonic model where long-lived northward subduction along a continental margin with repeated extensional and contractional regimes in response to retreat and advance of the subducting oceanic plate (Figure 3). Weber and Lodge (2022) obtained a U/Pb age of 1831.4 ± 2.0 Ma on the dacite unit hosting the Eisenbrey deposit in the western part of the orogen, suggesting that this second VMS forming event was widespread. A summary of the geochronology is presented in Figure 4. Figure 3 - Schematic illustration of the revised tectonic model of the Penokean Orogen. Figure is from Zi et al. (2021). Abbreviations: NF—Niagara fault zone; EPSZ—Eau Pleine shear zone. Marshfield Terrane This guidebook visits the only Penokean volcanic complex south of the Eau Pleine Shear Zone and is interpreted to part of the Marshfield Terrane. The Marshfield Terrane represents an Archean microcontinent of uncertain origins (Sims et al., 1989; Schulz and Cannon, 2007; Zi et al., 2021). Some of the earliest works on the terrane by Sims et al. (1989) noted eight Archean U/Pb ages from isolated outcrops along the Wisconsin, Black, and Chippewa Rivers; many of which were compiled from unpublished sources. Paleoproterozoic volcanic rocks in the Marshfield terrane were deposited about 1835-1865 Ma (Sims et al., 1989; Van Wyck, 1995; Klier, 2019; Weber and Lodge, 2022). These supracrustal rocks were referred to as the Eau Claire River Complex by Cummings (1984) or the Eau Claire Volcanic Complex by DeMatties (2018; 2022). They consist of an interlayered sequence of felsic to mafic volcanic rocks, dacite porphyry, and a variety of clastic and chemical sedimentary rocks (Sims et al., 1989). Some conglomerates contain granitic gneissic clasts that were interpreted to be Archean (Myers et al. 1980), but no definitive ages were determined on the clasts. Otherwise, our knowledge of the Archean Marshfield terrane and associated Paleoproterozoic volcanic rocks remains as sparse as the outcrop exposures. Eau Claire Volcanic Complex The Eau Claire Volcanic Complex is poorly documented and understudied mainly due to its inaccessible outcrops in remote parts of the Eau Claire River valley. Myers et al. (1980) described 51 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 4 - Time-space plot for the tectonic components of the Penokean Orogen. Plot is from Zi et al. (2021). See citation for references on data sources. supracrustal amphibolites, metarhyolites and metasediments in the Eau Claire River valley and classified them as part of the Chippewa Amphibolite Complex. This informal classification of the high metamorphic grade rocks in the Eau Claire-Chippewa River area was eventually grouped with the Marshfield Terrane by Sims et al. (1989) and Shultz and Cannon (2007). The first time that that the Eau Claire “Complex” was officially referred to was by Cummings (1984) when discussing the petrology and geochemistry of the gneisses in the Big FallsLittle Falls area (Stops 1 and 2 in this guidebook). After that, research in the Eau Claire Volcanic Complex essentially ceased. Sims et al. (1989) reported a U/Pb rutile age from Big Falls of ca. 1835 Ma. In fact, the words “Eau Claire” are not used in the Shultz and Cannon (2007) regional synthesis. DeMatties (2018; 2022) refers to the Eau Claire Complex when discussing the volcanic complexes in the Penokean, but largely cites the work of Myers et al. (1980). 52 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Preliminary Hf-isotope data and zircon trace elements reveal that the rocks analyzed in the Eau Claire Volcanic Complex are juvenile, mantlederived melts with no inheritance from older sources (See Weber et al. 2023, Part 1 of this volume). This suggests that these volcanic rocks are not forming on Archean basement, as one would expect if the Eau Claire Volcanic Complex was emplaced onto the Marshfield Terrane. Additionally, magnetic lineaments on aeromagnetic maps for the region appear to crosscut the interpreted position of the Eau Pleine Shear Zone and the overall fabric as outlined by magnetics appears constant (Figure 5). Ongoing research in the region seeks to better define the relationship of the Eau Claire Volcanic Complex to the Marshfield Terrane and the architecture of the basement in this area. Field Trip Stops The overall objective of this guidebook is to tour the accessible parts of the Eau Claire Volcanic Complex as exposed in the Eau Claire River valley and surrounding tributaries. The guidebook can be divided into two main regions: The Big Falls-Little Falls and North Fork regions. The Big Falls-Little Falls region represents the classic “Eau Claire Complex” originally described by Cummings (1984). The North Fork region is much more remote and rarely visited by geologists. In fact, it is not obvious that anyone has studied these rocks since they were first reported by Myers et al. (1980). These more remote parts of the complex are currently being studied (see Leahy and Lodge, Part 1 of this volume) to determine their regional context. Some of that data will be presented herein. The goal of this work is to determine if the Eau Claire Volcanic Complex is a volcanic center built upon Archean crust (continental arc) or is juvenile Figure 5 - Total field aeromagnetic map of the Eau Claire River region showing the location of Eau Pleine Shear Zone and field trip regions (Big Falls, North Fork). White dashed lines highlight a couple of magnetic lineaments that extend through the suturing shear zone. Magnetic maps from Daniels and Snyder (2002). 53 �Proceedings of the 69th ILSG Annual Meeting - Part 2 (oceanic arc) as this is important implications for the regional metallogeny and mineral systems. Most of the locations in this guidebook are on riverside outcrops. These areas are prone to sudden flooding and the upmost caution and careful planning should be used prior to visiting these locations. In addition, rocks here are uneven and potentially slippery. To access larger sections of outcrops, low water conditions may be required. In addition, all locations in this region may contain poisonous plants (e.g. nettle, poison ivy) and black-legged ticks that can transmit diseases. While this is unlikely to be a concern in early spring during the 2023 ILSG conference, future users of this manual should plan appropriately. Figure 6 - Generalized Precambrian geologic map of the Big Falls-Little Falls area of the Eau Claire River. Figure modified from Cummings (1984). Big Falls Region The banded amphibolite, gneisses, and intrusions in the Big Falls region of the Eau Claire River are some of most studied Precambrian exposures in this region and are visited multiple times a year by introductory and upper division geology classes at the University of WisconsinEau Claire. It is in this region that the term “Eau Claire River Complex” was first introduced by Cummings (1984) and this terminology has since been adopted by others (e.g. DeMatties, 2018) to describe the volcanic rocks present in the Marshfield Terrane. Figure 7 - Metamorphic conditions from geothermobarometic studies at Big Falls indicated by the yellow star. Data is from unpublished student project at the University of Wisconsin-Eau Claire. The region consists of mostly amphibolitic and felspathic gneisses that are intruded and brecciated by tonalite (Figure 6). Regional metamorphism in this region is at lower to upper amphibolite facies. A sample of amphibolitic gneiss from the Eau Claire Volcanic Complex using the edenite-richterite thermometry determined temperatures between 719-769 °C (Hannack and Radwany, 2018). Unpublished data from University of Wisconsin-Eau Claire class projects using garnet-biotite thermobarometry estimate peak metamorphic conditions at 765 °C and 11.5 kbars (Figure 7). A rutile U/Pb age of 1835 Ma from Sims et al. (1989) in the Eau Claire Volcanic Complex may indicate the timing of metamorphism. New research in this region provides our first glimpse into the trace element characteristics of these rocks (Figure 8). Rocks from both Big Falls and Little Falls have mafic protoliths with EMORB to oceanic arc like abundances of Th, Nb, and Yb (Pearce, 2008). On normalized trace element diagrams, samples have elevated LREE, low Th/La ratios, negative Nb and Ti anomalies. These trace element characteristics features are common in back-arc environments. Additionally, feldspathic units sampled have extremely 54 �Proceedings of the 69th ILSG Annual Meeting - Part 2 depleted trace element signatures and positive Eu anomalies, suggesting that they may be fractionated crystal cumulates. This broadly supports the interpretation of Cummings (1984) that the protolith of the Big Falls gneisses are a layered mafic intrusion. Stop 1: Amphibolite Gneisses and the “Great Unconformity” at Big Falls County Park Lat: 44.8215° Long: -91.2953° Figure 8 – Preliminary trace element geochemistry from the Big Falls-Little Falls area of the Eau Claire Volcanic Complex. Top: Trace element classification diagram from Pearce (1996) modified from Winchester and Floyd (1977). Middle: Mantle source discrimination diagram from Pearce (2008). Bottom: Primitive mantle-normalized trace element diagram using values from Sun and McDonough (1989). This location is accessible from the north entrance to Big Falls County Park off Eau Claire County Highway Q. There is a parking lot at this entrance with plenty of parking for park visitors. Follow the paved foot path eastward toward the river. Once on the riverbank, walk northward for 55 �Proceedings of the 69th ILSG Annual Meeting - Part 2 about 50 m to reach the outcrops at the falls. Note that outcrops on the south bank of the falls will have to be accessed via the south entrance to the park on County Highway K. During very low water conditions, it is possible to hop across the outcrops to access the south bank. This region has some steep-edged rock cliffs adjacent to the river and there are springs that keep some areas wet and slippery. Please watch your step. This stop highlights the geology along the north side of Big Falls County Park where the rocks are exposed along the Eau Claire River. In addition to the Precambrian rocks, this location also has a great exposure of the “Great Unconformity” with overlying Cambrian Mount Simon Formation. The Eau Claire River flows along the nonconformity between the Cambrian and Precambrian rocks, where the river has eroded the overlying Cambrian units away exposing the Precambrian basement rocks. At this stop, we highlight four locations that highlight different units seen here at Big Falls County Park (Figure 9). Photo 1: Photographs of the banded amphibolitic gneiss at Big Falls County Park. (A) Banded amphibole gneiss at Big Falls. B: Photomicrograph in plane-polarized light (25x) of large garnet porphyroblasts in quartzofeldspathic and hornblende matrix. Location 1: The Banded Amphibole Gneiss The banded amphibole gneiss (Photo 1A) is best described as a fine-grained banded gneiss with alternating hornblende-rich and plagioclaserich layers. The hornblende-rich layers range anywhere from less than 1 cm to ~15 cm and contain about 85% hornblende and 15% granulated plagioclase with sparse idioblastic garnet. The plagioclase-rich layers are consistently thicker and contain approximately 15% hornblende. The garnets though scarce occur as coarse grained porphyroblasts in both layers (Photo 1B) although these garnets often show retrograde alteration back to hornblende. The garnets are typically poikioblastic (Photo 2) with quartz, plagioclase and occasionally biotite inclusions. The hornblende occurs as euhedral to subhedral grains and the plagioclase as very-fine grained, dynamically recrystallized, matrix. The granulated plagioclase is typically labradorite to bytownite but anorthite (An92) occurs as cores in some of the idioblastic hornblende to create an Photo 2: Photomicrograph in plane polarized light of poikioblastic garnet in a hornblende and granulated plagioclase matrix (50X magnification). unusual bi-modal plagioclase population (Photo 3). Multiple shear zones and isoclinal folds are present throughout the outcrop, providing evidence for multiple deformation periods. Partially annealed shear zones can be traced across almost the entire unit that truncate and offset banding. There are also asymmetric 56 �Proceedings of the 69th ILSG Annual Meeting - Part 2 amphibole schist is enriched in MgO and FeO and depleted in CaO and Al2O3 in comparison to the banded amphibole gneiss (Table 1). Both major element (Table 1) and trace elements (Figure 8) characterize this as a MORB-like composition. Isoclinal folds of the foliation and the presence of ductile shear zones indicate deformation throughout this unit. The contact between the banded amphibole gneiss and the amphibole schist is buried by slumping blocks of the Cambrian Mount Simon Formation and vegetation. Photo 3: Photomicrograph in partially crossed polars of idioblastic hornblende with anorthite cores (An 92) in a matrix of granulated plagioclase (An72) in the banded amphibole gneiss at Big Falls (Magnification is 50x). amphibolite inclusions kinematics. that show variable This unit was sampled for a recent zircon petrochronology study and produced significant results that question what is currently understood about the southern portion of the Penokean Orogen. This unit yields a U/Pb age of 1874.7 ± 2.1 Ma which temporally correlates with other VMS-forming events across the PembineWausau terrane. Zircon trace element geochemistry of the sample indicate the sample formed in a hydrated but reduced melt in a backarc setting where decompression was occurring in a metasomatized mantle (Weber and Lodge, 2022). Hf-Lu data from the banded amphibolite gneiss indicates a lack of older basement inheritance. The zircon petrochronology from these rocks contradicts the interpretation that Eau Claire Volcanic Complex was emplaced into the Archean Marshfield Terrane. These results have motivated additional research in the Eau Claire Volcanic Complex. Photo 4: Outcrop photo of the feldspathic gneiss at Big Falls. Location 3: Transition Gneiss and Feldspathic Gneiss The amphibole schist gradually grades into the transition gneiss for a few meters as the unit contains fewer amphibole-rich layers and the plagioclase rich layers become more prominent (Photo 4). The feldspathic gneiss is primarily made of plagioclase, with 10-20% hornblende and lesser amounts of chlorite, epidote, and localized sulfidation with some pyrite mineralization. Location 2 Amphibole Schist Despite strong metamorphic recrystallization and structural fabric overprinting, there are some primary igneous textures that are preserved (Cummings, 1984). The compositional banding and layering throughout all units appear to be The further west along the Eau Claire River, the amphibole schist is exposed. This unit is best described as a dark green to black, fine grained thinly banded amphibole schist. Hornblende is the primary amphibole with lesser amounts of plagioclase. Based on whole rock XRF data, the 57 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 9: Geologic map showing outcrop locations at the Big Falls stop. Modified from Cummings (1984). Table 1. Whole rock major element geochemistry (via XRF) from Big Falls for banded amphibole gneiss (location 1), amphibolite (location 2), and altered rocks at the Precambrian-Cambrian contact (location 3). Average MORB composition (Winters, 2010) is included for comparison. Unit Unaltered Banded Gneiss 1 Banded Gneiss 2 Banded Gneiss 3 Altered (Depth) surface .5m 1.0m 1.5m Unaltered Amphibolite 1 Amphibolite 2 Average MORB (Winter 2010) Altered (Depth) surface .25m .5m .65m .75m 1.0m SiO2 TiO2 Al2O3 Fe2O3T CaO MgO MnO Na2O K2O P2O5 Totals 47.10 48.43 46.63 0.17 0.21 0.23 30.43 30.56 30.24 2.14 3.35 2.47 15.18 14.65 14.51 0.72 1.38 0.88 0.03 0.04 0.03 2.21 2.97 2.85 0.20 0.25 0.18 0.05 0.04 0.03 98.23 101.88 98.05 52.06 50.26 52.55 52.92 0.44 0.74 0.34 0.32 16.66 15.28 17.36 18.23 6.72 11.56 7.28 5.79 0.88 0.63 1.06 0.80 5.21 6.18 5.17 4.91 0.04 0.05 0.03 0.02 0.11 0.05 0.05 0.09 9.28 9.14 9.04 9.37 0.03 0.28 0.03 0.03 91.43 94.17 92.91 92.48 51.10 52.66 50.50 1.34 1.71 1.56 15.50 13.25 15.30 13.73 12.12 11.50 9.07 8.05 11.50 5.79 6.57 7.47 0.19 0.20 n/a 3.11 3.56 2.62 0.29 1.46 0.16 0.23 0.17 0.13 100.35 99.75 100.74 61.93 62.79 61.48 62.62 60.07 58.69 0.94 0.99 0.84 0.88 0.95 0.87 17.36 17.18 17.04 16.26 16.84 16.31 4.98 6.10 5.04 5.42 5.60 7.30 0.63 0.65 0.64 0.64 0.64 0.66 2.70 3.05 3.02 2.68 2.89 3.24 0.02 0.04 0.02 0.04 0.04 0.07 0.32 0.20 0.21 0.23 0.07 0.08 7.10 7.22 6.97 7.15 7.45 7.18 0.19 0.19 0.18 0.18 0.21 0.19 96.17 98.41 95.44 96.10 94.76 94.59 58 �Proceedings of the 69th ILSG Annual Meeting - Part 2 primary. Anorthositic autoliths are incorporated in a fine-banded, more mafic matrix near the transitional gneiss. The banding in the autolith is discordant to banding in the matrix and appear to be concentrated in bands but are not associated with boudinage fabrics. These observations were critical in interpreting the protolith of this region. Location 4: The Great Unconformity and the basal portion of the Mt. Simon Formation The large hillside on the north bank of the river and above the Precambrian outcrops is the Cambrian Mount Simon Formation. The base of the Mt. Simon Formation is a mix of coarsegrained quartz arenites and quartz pebble conglomerates. Photo 5: Nonconformity between the Cambrian Mount Simon Formation and the amphibolites at Big Falls County Park. The Great Unconformity creates a nonconformable contact between the Precambrian units at the previous three locations and the base of the Mt. Simon Formation. At Big Falls, a thin blue-green celadonite clay layer (Figure 10) has formed along the nonconformity as a result of Kmetasomatism from basinal brines. The Kmetasomatism at the unconformity is recognized throughout the midcontinent region and is related to MVT lead zinc deposits in the Tri-state region (See field trip 1 in this volume). Locally the celadonite acts as a fluid barrier for springs that flow along the unconformity. In other places, the contact appears to be relatively sharp with little alteration (Photo 5). Figure 10: A-CN-K diagram showing chemical change due to weathering and the alternative path of Kmetasomatism. The celadonite at Big Falls is not a weathering profile and requires adding substantial potassium and to produce the celadonite and authigenic K-spar seen along the unconformity (see Table 1 for chemistry). Stop 2 –Tonalite Breccia and Gneisses at Little Falls. Lat: 44.8103° Long: -91.2825° bridge at this location for the Eau Claire River flood stage measurement. Just north of the bridge there is a small parking area on the west side of the road. There are several small foot paths that lead down to the river’s edge. The outcrops are mostly exposed immediately around and under the bridge. The quality of exposure here changes all the time as flooding conditions sometimes This location is just on the north side of the County Highway K bridge that crosses the Eau Claire River and is 300 m north of the exit to the south entrance of Big Falls County Park. Google Maps calls this place the East Eau Claire Canoe Landing, but the USGS refers to this location as Little Falls (so does the faculty and students at the University of Wisconsin-Eau Claire) and uses the 59 �Proceedings of the 69th ILSG Annual Meeting - Part 2 buries parts of the outcrop with sand and downed vegetation. To access the south bank of the river, a short bush traverse (100-300 m) will be required from the southside of the bridge along the riverbank. If water level and time permit, this stop has four locations of interest that highlight the intrusive history in the region. The outcrops in this area expose an inclusion-rich intrusive contact between a foliated tonalite and lensoidal amphibolite and are cut by younger pegmatitic and mid-continent rift diabase dykes (Figure 11). This location is used to teach students at the University of Wisconsin-Eau Claire about interpreting relative geologic time and observing contact relationships. The absolute ages of these rocks are unknown as recent attempts to isolate zircons from the tonalite were unsuccessful. Photo 6: Gneissic tonalite breccia highlighting some of the banded gneiss xenoliths. Some of the smaller xenoliths here are elongated. significant assimilation of amphibolite. Biotite commonly produces a crude foliation that may have formed from hornblende during a later deformation. Location 1: Gneissic Tonalite Breccia It is clear the tonalite is metamorphosed and is an important part of determining the nature of the Eau Claire Volcanic Complex. However, efforts to constrain the timing of this intrusive event have yielded conflicting results. Van Schmus (1980) yielded a U/Pb age of 1842 ± 10 Ma utilizing zircon fractions (i.e. not modern single crystal methods). Sims et al. (1989) reported a U/Pb age of 1856 ± 5 Ma from a xenolith at Little Falls. Assuming that the amphibolite xenoliths at Little Falls are from the same amphibolite unit at Big Falls that was dated at 1875 Ma, then there is a clear conflict. The tonalite was sampled for that recent petrochronologic study (Weber and Lodge, 2022) but yielded very few zircons. Resolving the timing of emplacement and tectonic setting of this intrusion will help better understand the geodynamic setting of the Eau Claire Volcanic Complex. The gneissic tonalite breccia is the most prominent unit at Little Falls (Photo 6). Roughly 90% of xenoliths in the breccia are characterized as banded gneiss to banded amphibolite containing 50-80% hornblende and 20-40% plagioclase with lesser amounts of biotite. These xenoliths range in size anywhere from less than 1cm to greater than 20 cm, and they are hosted in a biotite tonalite intrusion which destroys the older banded gneiss. The banded gneiss xenoliths are also elongated and contain folds. Ultramafic xenoliths are scarce but also present. These xenoliths contain over 90% hornblende with lesser amounts of epidote-clinozoisite and plagioclase. Occurring in localized clusters, the fragmented ultramafic xenoliths indicate these were most likely part of a larger block but separated during the tonalite intrusion event (Myers et al., 1980). The tonalite is composed of 35-40% plagioclase (An50-55), about 30% hornblende, 25-30% quartz, 5-10% biotite, and accessory epidote. Myers et al. (1980) interpreted the fabric in the rock as flow-lamination, however it is parallel to regional magnetic lineaments suggesting it may be a structural fabric. Large variation in mafic mineral abundance indicates Location 2: Diabase Along the north side of the river and east of the bridge, lies one of many mid-Proterozoic diabase dykes associated with the mid-continent rift in the 60 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 11: Geologic map of the Little Falls area showing the locations of interest at this stop. Figure modified from Myers et al (1980). Eau Claire region. Like many other diabase dike outcrops in the area, this diabase exhibits both clean columnar jointing and well-defined chilled margins. This diabase is only a few meters in size and disappears beneath the surrounding overburden (Figure 11). The diabase has a medium-grained, equigranular texture and does not have any foliation or recrystallization textures and is clearly post-metamorphism. throughout the Eau Claire volcanic complex along the Eau Claire River where the Precambrian rocks are exposed. Their macro- and microscopic characteristics indicate they are Location 3: Pegmatite Dike On the west bank of the river lies a 2 m wide pegmatite dyke (Photo 7). Outlier boulders of this pegmatite can be found on the eastern bank of the river. The alkali-feldspar crystals in this outcrop reach sizes greater than 30 cm. Various sizes of quartz veins also crosscut this unit. This granite pegmatite dike is one of a handful seen Photo 7: 18m-wide pegmatite in the Eau Claire River downstream from Little Falls. 61 �Proceedings of the 69th ILSG Annual Meeting - Part 2 clearly younger than the Penokean deformation and metamorphism and could be related to Mazatzal or Yavapai orogenic events to the south like the Wolf River Batholith in northcentral Wisconsin. The mineralogy of the pegmatites is very complicated with many accessory carbonate, phosphate and oxide phases enriched in Nb, Y, F U, Th and REE. Zircon in the pegmatites indicate extreme fractionation (Figure 12). The zircons also show considerable xenotime (Y,P) substitution and considerable solid solutions with both coffinite (USiO4) and thorite (ThSiO4). The pegmatites, despite their pink color, are primarily composed of plagioclase with an overprint of potassic alteration. Most of the alkali-feldspar occurs along cleavages, crystal boundaries and fractures indicating it is a late phase in pegmatite formation. The plagioclase in the pegmatites is primarily albite but ranges from An0 to An30. The euhedral and inclusion free garnets (Photo 8) have a limited range of chemistry close to 50% almandine and 50% spessartine which is similar to magmatic garnet compositions in other garnet-quartz-albite pegmatites (Muller et al., 2018). Figure 12. Zr/Hf in pegmatites from the Eau Claire River Complex. The Eau Claire River pegmatites have many characteristics of pegmatites in the Nb/Y/F (NYF) family of rare element pegmatites and NYF pegmatites are always associated with metaluminous to alkaline (or peralkaline) granites (Cerny and Ercit, 2005) Location 4: Amphibolitic Gneiss The xenoliths within the tonalite are assumed to be derived from the nearby outcrops of the amphibolitic gneisses. Unlike the planar banding at Big Falls, the amphibolitic gneisses here is more lensoidal with cm- to dm-scale lens-shaped, hornblende-rich pods surrounded by more plagioclase-rich “matrix” and quartz-veining. Photo 8. Top: Photomicrograph (25X in plane polarized light) of garnet cluster in quartz and albite from Little Falls pegmatite dike on the west side of the river. Bottom: Almandine/spessartine garnet clusters at the same location at Little Falls are magmatic in origin. 62 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Detailed work has not been completed on these outcrops but are assumed to be petrogenetically related to the amphibolites at Big Falls. Future research will examine these exposures more closely. the trace element geochemistry and zircon petrochronology of the rocks in this region to make better links with the rest of the Eau Claire Volcanic Complex. Much of that data is still pending or preliminary, so results will be forthcoming soon. The goal for the field trip in this region is to show as many of the rocks as possible, regardless of how much we know about them. North and South Fork Region This is the part of the trip where information on these rocks is sparse and new data is only just becoming available. To our knowledge, the rocks in the North and South Fork areas of the Eau Claire River have not been studied in any detail since Myers et al. (1980). Much of the area is remote and sparsely developed and very few outcrops are easily accessible. Field work in the 2022 summer relied on one-way, day-long kayak trips along different segments of the North Fork. Aside from the occasional powerline, field work on these stretches of the Eau Claire River felt wild and remote. This field work aimed to characterize The region consists of amphibolites, feldspathic gneisses, and foliated granitoids and are cross-cut by younger, undeformed granitic pegmatites (Figure 13). A metarhyolite from this region yielded an age of 1858 ± 5 Ma (Sims et al, 1989) and was one of the key samples that linked Penokean volcanic processes to the Marshfield terrane. Myers et al. (1980) interpreted mappable contacts between intrusions and foliated Figure 13: Geologic map of the North and South Fork of the Eau Claire River in eastern Eau Claire County and western Clark County. Figure modified from Myers et al. (1980). 63 �Proceedings of the 69th ILSG Annual Meeting - Part 2 supracrustal rocks in this area to be sheared and nearly vertical and that they enclose lensoidal fault slices which have been juxtaposed mainly by strike-slip displacement. Outcrops of amphibolite in the southern part of the map (Figure 13) near the confluence of the North and South Forks of the Eau Claire River were interpreted to be part of the “Chippewa Amphibolite Complex” (Myers et al., 1980) which is broadly supported by regional aeromagnetic maps (Figure 5). Myers et al. (1980) interprets the volcanic rocks in this region to unconformably on amphibolites, but geochronologic data is lacking to make any absolute local or regional correlations. Regional metamorphic grade is estimated to be upper greenschist to lower amphibolite based on the presence of garnet, epidote, muscovite, and hornblende. Preliminary geochemical results from the North Fork region begin to reveal the setting of these volcanic and intrusive rocks. Volcanic protoliths are bimodal (Figure 14) with tholeiitic mafic rocks with oceanic affinities (Figure 15) and FI- to FII-type felsic rocks arc-like affinities (Figure 16). More work needs to be done before we can concretely interpret the setting of this region of the Eau Claire Volcanic Complex. Figure 15: Trace element geochemical characteristics of the mafic rocks from the North Fork region of the Eau Claire River. Top: Magmatic affinity diagram for sub-alkaline basalts from Ross & Bedard (2009). Middle: Mantle source discrimination diagram from Pearce (2008). Bottom, Primitive mantle-normalized trace element diagram using values from Sun and McDonough (1989). Figure 14: Trace element classification diagram from Pearce (1996), modified from Winchester & Floyd (1977). 64 �Proceedings of the 69th ILSG Annual Meeting - Part 2 These settings are not typical of continental settings, which continues to question the relationship between the Eau Claire Volcanic Complex and Marshfield Terrane. Stop 3 – Amphibolite and Intrusions at Knights Pool Lat: 44.7482° Long: -90.9669° Figure 16 – Trace element geochemical characteristics of felsic rocks from the North Fork region of the Eau Claire River. Top: Nb/Y discrimination diagram for granites from Pearce (1984). Middle: F-type felsic discrimination diagram from Hart et al. (2004). Bottom: Primitive mantle-normalized diagram using values from Sun & McDonough (1989). The directions to get to this stop are a little more elaborate since it is in a more remote location. From the community of Augusta, take State Highway 27 north for 4.4 miles to County Road GG. Turn east on to County Road GG and drive 4.7 miles to the intersection of Channey 65 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Road just after the bridge over the Eau Claire River. Turn east on Channey Road (note that this is an unpaved road) and drive for 4.6 miles until the road crosses the North Fork of the Eau Claire River. This location is called Knights Pool and is labelled by signage. The accessible outcrops in this region are immediately beneath the bridge and are accessible by foot trails. There are larger outcrops north of the bridge that are accessible by a small, 150 m bush traverse along a sparsely used trail. At both locations, rocks are immediately adjacent to a shallow but fast-moving river and caution should be used. Complex exposed in this section of the river upand downstream of this location. However, this is the only easily accessible section by foot. At Knights Pool, there are mostly strongly foliated and deformed amphibolites that are intruded by a biotite granodiorite (Figure 17). Amphibolite: The amphibolite at Knights Pool is characterized by a fine to very fine grained mafic lineated amphibolite with stretched quartzfilled amygdules, relict pillow structures, and wispy textures suggesting a mafic flow (Photo 9). Thin sections of the amphibolite clearly show the lineations present in the amphibolite here (Photo 10). Several stages of deformation occurred starting with the amphibolite being isoclinally folded, then intruded by aplite veins, and intruded by large granodiorite body (Myers et al., 1980). The shearing in of the granodiorite body created a mylonite gneiss along the contact with the amphibolite. Knights Pool is located at the bridge on Channey Road as it crosses the North Fork of the Eau Claire River along the southern edge of the North Fork Eau Claire River State Natural Area. There is more of the Eau Claire Volcanic Photo 9: Outcrop of the amphibolite showing both the isoclinal folds and strained amygdules present in the unit. Trace element geochemistry of the amphibolites at Knights Pool are notably LREEdepleted with strong negative Nb anomalies (Figure 15). This suggests it was derived from strongly depleted but metasomatized mantle. This type of environment, presumably a mature backarc, rarely exists in a continental setting. The granitoids in the map area have classic enriched LREE and Th with depleted Nb and HREE signatures suggesting they are related to a Figure 17: Geologic map of the Knights Pool area. Map is modified from Myers et al. (1980). 66 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Photo 10: Photomicrographs in plane-polarized light of the amphibolite at Knights Pool amphibolite under thin section at 25x magnification (A) and 100x magnification (B). Hornblende forms a clear lineation. Photo 11: Photomicrographs of the biotite granodiorite at Knights Pool in (A) plane-polarized light, and (B) cross-polarized light. Both images are 25x magnification. the village of Rock Dam. Turn northward on Butler Road and use the parking lot on Hay Creek Lake just south of the bridge over Hay Creek. The outcrop of phyllite are under the bridge and near the Rock Dam spillway. The nonconformity and metarhyolite can be better accessed from within the Rock Dam Campground near campsite 90. different tectonic event when the crust was thicker, and garnet was stable to deplete HREE. Granodiorite Intrusion: The medium to coarse grained biotite granodiorite intruded the amphibolite creating a mylonitic fabric along the contact. The intrusion, shearing along the contact, caused the folding of the aplite veins seen in the amphibolite. Quartz is strongly recrystallized and biotite concentrations define a weak foliation. Thin section photos highlight how the quartz is being recrystallized, as well as show the alignment of biotite aggregates (Photo 11). This location reveals another nonconformity between the Precambrian and Cambrian Mount Simon Formation. The metarhyolites and phyllites in this area are strongly mylonitized and primary structures are difficult to interpret. The metarhyolite at this location described by Myers et al. (1980) is the only reference to a rhyolitic unit in the Eau Claire Volcanic Complex. Since Sims et al. (1989) dated a metarhyolite at 1858 ± 5 Ma in the Eau Claire River and cited Myers et al. (1980), it presumably came from this location. If that is the case, then the rocks at Rock Dam are of regional significance because it is one of Stop 4 – Metarhyolite and Phyllite at Rock Dam Lat: 44.7338° Long: -90.8469° From the previous stop, continue eastward on Channey Road until it ends at County Highway H. Turn southward and drive 1.5 miles to Rock Dam Road. Turn eastward and take this road into 67 �Proceedings of the 69th ILSG Annual Meeting - Part 2 estimated mineral percentages and matrix of this rock is composed of a very fine-grained alkalifeldspar (57%), quartz (35%), muscovite (3%), magnetite (2%), and biotite (1%). The quartz eyes along with the absence of feldspar porphyroclasts suggests that the quartz either originated as phenocrysts or clasts in a tuff (Myers et al. (1980). Foliation trends east-west and is near vertical. Phyllite: Closer to the base of Rock Dam lies a muscovite-rich phyllite composed of alkalifeldspar, quartz, and muscovite (Photo 12). This outcrop lacks both the quartz eyes and biotite possibly indicating a separate protolith than the metarhyolite (Myers et al. 1980). Photo 12: Outcrop photo of strongly foliated phyllite at Rock Dam near spillway. Cambrian strata can be seen in background on opposite bank of river. Mt. Simon Formation: The basal part of the Mt. Simon Formation, a conglomerate layer containing pebbles of vein quartz and rhyolite, is exposed at this location (Photo 13). The contact between the Mt. Simon and the Precambrian metarhyolite shows about 5 m of relief. Many of the locations exposing the Great Unconformity in the Eau Claire region show deep weathering of the underlying Precambrian rocks, however there is not much weathering of the Precambrian metarhyolite here. the very few dated Penokean supracrustal rocks within the Marshfield terrane. Mylonitized Metarhyolite: Myers et al. (1980) admittingly conceded that determining the protolith of this outcrop can be challenging considering the similarities between sheared porphyritic rhyolites and leucogranites. The mylonite here is pale pink metamorphosed porphyritic rhyolite containing quartz eyes that can be described as phenocrysts or clasts. These quartz eyes are roughly 1-2.5 mm in size and under thin section show a subrectangular to lenticular shape (Myers et al. (1980). The 68 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Photo 13: Basal pebble conglomerate in Cambrian strata overlying metarhyolite at Rock Dam. Stop 5 – Metavolcanic Rocks at Mead Lake Lat: 44.7885° Long: -90.7742° From the previous stop, continue northward on Butler Road for 0.8 miles to the intersection with Willard Road. Drive eastward on Willard Road for 2.0 miles and turn north onto County Road M. Drive northward on County Road M for 1.6 miles and turn eastward onto Rocky Run Road. Drive 1.2 miles on Rocky Run Road and turn northward on Bruce Mountain Road that will turn into South Lake Road. South Mead Lake Park will be 1.4 miles down this road. Park there, and the outcrops are on the riverbank west of the Mead Lake Dam spillway. The bedrock exposed at this location is primarily a foliated, fine-grained chloritic metavolcanic rock (Photo 14). There has been no known study of this rock, and our data is still pending. Nonetheless, it is apparent that the metamorphic grade seems to be decreasing in this part of the Eau Claire Volcanic Complex. This is in stark contrast to the rocks in the Chippewa River valley (Fieldtrip 1, this volume) and Big Falls region (Stops 1-2). Future work in the region will utilize every outcrop, even small ones like this location, to better describe and define the tectonics and metallogeny of the Eau Claire Volcanic Complex. Photo 14: Outcrop photo of metavolcanic rocks at Mead Lake Dam. 69 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Wisconsin: An overview: Economic Geology, v. 89, p. 1122-1151. Acknowledgements Despite decades of regular visits from groups from the University of Wisconsin-Eau Claire, the most extensive detailed maps and rock descriptions were provided by Paul Myers and collaborators in the 1980 ILSG guidebook. Outside of Big Falls County Park, many of those locations have not been visited since then. A lot of the geologic descriptions from those lesserknown areas have been updated from Myers et al. (1980) and figures have been digitized while adding new data and insights where available. DeMatties, T. A., 2018, Effects of paleoweathering and supergene activity on volcanogenic massive sulfide (VMS) mineralization in the Penokean Volcanic Belt, northern Wisconsin, Michigan and east-central Minnesota, USA: Implications for future exploration: Ore Geology Reviews, v. 95, p. 216-237. DeMatties, T. A., 2022, Exploration-resource assessment of productive felsic volcanic centers in the Paleoproterozoic Penokean Volcanic Belt of northern Wisconsin, Michigan and east-central Minnesota, USA: Ore Geology Reviews, v. 141, article 104489. In addition, the authors of this guidebook would like to thank the countless undergraduate and graduate students that have worked on these outcrops and have continued to inspire new work in the region. Specific acknowledgement is deserving to Matt Leahy and his undergraduate research project in the North Fork region in providing some insight into that part of the complex. Hannack, G., and Radwany, M., 2018, HornblendePlagioclase thermometry of the Eau Claire River Complex, western Wisconsin: Proceedings of the Institute on Lake Superior Geology 64th Annual Meeting, Iron Mountain, Michigan, p. 47-48. Hart, T. R., Gibson, H. L., and Lesher, C. M., 2004, Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic massive Cu-Zn-Pb sulfide deposits: Economic Geology, v. 99, p. 1003-1013. References Klier, J. J., 2019, The Marshfield Terrane: Refedinition of origin through zircon geochronology and geochemistry: Unpub. M.S. thesis, Ball State University, 115 p. Brown, B. A., 1988, Bedrock geology of Wisconsin, west-central sheet, Wisconsin Geological and Natural History Survey Map 87–11b. Cummings, M. L., 1984, The Eau Claire River complex: A metamorphosed Precambrian mafic intrusion in western Wisconsin: Geological Society of America Bulletin, v. 95, p. 75-86. LaBerge, G. L., 1996, Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume, Proceedings of the 42nd Annual Meeting of the Institute on Lake Superior Geology, Cable, Wisconsin. Cerny, P. and Ercit, T. S., 2005, The classification of granitic pegmatites revisited. Canadian Mineralogist, v.43, p. 2005-2026. LaBerge, G. L., and Myers, P. E., 1984, Two early Proterozoic successions in central Wisconsin and their tectonic significance: Geological Society of America Bulletin, v. 95, p. 246-253. Daniels, D. L., and Snyder, S. L., 2002, Wisconsin aeromagnetic and gravity maps and data, U.S. Geological Survey Open-File Report 02-493. Myers, P. E., Cummings, M. L., and Wurdinger, S. R., 1980, Precambrian geology of the Chippewa Valley, Wisconsin, Institute of Lake Superior Geology 26th Annual Meeting, Eau Claire, Wisconsin, Field Trip Guidebook 1, 123 p. DeMatties, T. A., 1989, A proposed geologic framework for massive sulfide deposits in the Wisconsin Penokean volcanic belt: Economic Geology, v. 84, p. 946-952. Muller, A., Spratt, J., Thomas, R., Williamson, B.J., and Seltmann, R., 2018, Canadian Mineralogist, v. 56, p. 657-687. DeMatties, T. A., 1994, Early Proterozoic volcanogenic massive sulfide deposits in 70 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Pearce, J. A., 1996, A users guide to basalt discrimination diagrams: Trace Element Geochemistry of Volcanic Rocks: Applications for Massive Sulphide Exploration. Geological Association of Canada, Short Course Notes, v. 12, p. 79-133. pre- and early Proterozoic rocks in Wisconsin: Unpub. Ph.D. thesis, University of Wisconsin Madison, 295 p. Van Wyck, N., and Johnson, C. M., 1997, Common Lead, Sm-Nd, and U-Pb constraints on petrogenesis, crustal architecture, and tectonic setting of the Penokean orogeny (Paleoproterozoic) in Wisconsin: Geological Society of America Bulletin, v. 109, p. 799-808. Pearce, J. A., 2008, Geochemical fingerprinting of oceanic basalts with applications to ophiolite classification and the search for Archean oceanic crust: Lithos, v. 100, p. 1-4. Weber, E. M., and Lodge, R. W. D., 2022, New U/Pb Geochronology from the Proterozoic Penokean Orogen, Wisconsin: Implications for VMS Metallogeny: Society of Economic Geologists Annual Meeting, Denver, CO, paper P5.10. Pearce, J. A., Harris, N. B. W., and Tindle, A. G., 1984, Trace element discrimination diagrams for the tectonic interpretation of granitic rocks: Journal of Petrology, v. 25, p. 956-983. Quigley, A., 2016, Setting of the volcanogenic massive sulfide deposits in the Penokean Volcanic belt, Great Lakes region, USA: Unpub. M.S. thesis, Colorado School of Mines, 95 p. Winchester, J. A., and Floyd, P. A., 1977, Geochemical discrimination of different magma series and their differentiation products using immobile elements: Chemical Geology, v. 20, p. 325-343. Ross, P.-S., and Bédard, J. H., 2009, Magmatic affinity of modern and ancient subalkaline volcanic rocks determined from trace-element discriminant diagrams.: Canadian Journal of Earth Sciences, v. 46, p. 823-839. Winter, J. D., 2010, An Introduction to Igneous and Metamorphic Petrology, Prentice Hall, 697p. Zi, J-W., Sheppard, S., Muhling, J. R., and Rasmussen, B., 2021, Refining the Paleoproterozoic tectonothermal history of the Penokean Orogen: New U/Pb age constraints from the PembineWausau terrane, Wisconsin, USA: Geological Society of America Bulletin, v. 134, p. 776-790. Schulz, K. J., and Cannon, W. F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4-25. Sims, P. K., Van Schmus, W. R., Schulz, K. J., and Peterman, Z. E., 1989, Tectonostratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean orogen: Canadian Journal of Earth Sciences, v. 26, p. 21452158. Sun, S., and McDonough, W. F., 1989, Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes, in Saunders, A. D., and Norry, M. J., eds., Magmatism in the Ocean Basins, Geological Society Special Publication, v. 42, p. 313-345. Van Schmus, W. R., 1980, Chronology of igneous rocks associated with the Penokean orogeny in Wisconsin: Geological Society of America Special Paper, v. 182, p. 159-168. Van Wyck, N., 1995, Oxygen and carbon isotopic constraints on the development of eclogites, Holsnpy, Norway, and, Major and trace element, common Pb, Sm-Nd, and zircon geochronology constraints on petrogenesis and tectonic setting of 71 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Field Trip 4 – Quaternary Geology and Geomorphology of the Eau Claire Region Douglas J. Faulkner Department of Geography and Anthropology, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 J. Elmo Rawling, III Wisconsin Geological and Natural History Survey, Madison, WI 53705 Phillip H. Larson Earth Science Programs, EARTH Systems Laboratory, Minnesota State University Mankato, Mankato, MN 56001 72 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Introduction Eau Claire lies close to the outermost edge of the former Chippewa Lobe of the Laurentide Ice Sheet as it existed during late Wisconsinan time (MIS-2) (Fig. 1). To the south are older glacial deposits and then the Driftless Area, which apparently was never glaciated. This all-day field trip will concentrate on three aspects of the region’s landscape development from the late Wisconsinan to the late Holocene: glacial, fluvial and aeolian. Glacial Landscapes Northern Wisconsin was glaciated multiple times in the Quaternary. The oldest glacial deposits were derived from the northwest and were likely deposited prior to 780,000 ka. These include the Pierce and Marathon Formations (Rawling et al., in review; Syverson et al., 2011). These deposits are poorly preserved where they occur at the surface (Rawling et al., in review) and although their occurrence is documented in the subsurface (Attig 1985 and 1993; Woodruff et al., 2004), their regional distribution is poorly documented. During the most recent glaciations, ice flowed from the northeast through the Superior Basin until it was thick enough to spill over the regional bedrock divide (Attig and Rawling, 2018). Ice formed during an earlier glaciation deposited glacial and meltwater sediment of the River Falls Formation (Syverson, 2007). River Falls tills and outwash are preserved on uplands in the Eau Claire area, and landforms associated with this advance have been eroded and are not preserved. The best preservation of landforms is associated with the late Wisconsinan ice (ca. 25–11.5 yr B.P.), which formed the Chippewa Lobe that reached its maximum extent at the Chippewa Moraine. This ice was subject to stagnation whenever the ice profile in the Superior Basin lowered, resulting in an ice margin landscape consisting of broad (10s of kilometers) zones of stagnant ice features such as disintegration ridges, ice-walled lakes, and kettles. Figure 1. Top: Map of Wisconsin showing areas covered by lobes of the southern Laurentide Ice Sheet during the late Wisconsinan (MIS-2) Glaciation; inset map shows distribution of ice in the Great Lakes region. The red circle shows the location of Eau Claire near the southern edge of Chippewa Lobe. Bottom: Schematic illustration showing how moraines form over time. Supraglacial sediment accumulates at a stable ice margin (time one). As glacier retreats, buried ice is preserved under supraglacial sediment and minor moraines form if margin temporarily stabilizes (time two). The distribution of moraines after ice has melted (time three; modified from Attig and Rawling, 2018). 73 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Fluvial Landscapes The process of knickzone migration and incision up the LCR was episodic and unexpectedly prolonged. The episodic history of knickzone migration and incision is clearly indicated by the number and spatial distribution of terraces found in the LCR valley below the Wissota. Instead of two terrace levels resulting from the two episodes of abrupt base-level fall, there are as many as seven (Fig. 2). Each of these levels represents a period when the river was migrating laterally and forming a floodplain, followed by an episode of renewed incision that left the floodplain as a terrace. The prolonged history of knickzone migration and incision along the LCR is revealed by the optically stimulated luminescence (OSL) ages of terrace alluvium from several sites in the LCR valley (Fig. 3). These OSL ages indicate that knickzone migration took thousands of years longer than studies of modern alluvial streams affected by minor base-level falls suggest it should’ve taken (Begin 1986, 1988; Begin et al., 1981) The Chippewa River is the second largest stream in Wisconsin, draining a watershed of approximately 25,000 km2 to the upper Mississippi River (UMR). During the Late Wisconsinan, it was the primary stream draining meltwater from the Chippewa Lobe. Overloaded with glacigenic sediment, the lower Chippewa River (or LCR, which refers to the river beyond the Chippewa Moraine) filled its bedrock valley with sandy outwash to depths exceeding 50 meters. Then, sometime between 18-16 ka, the UMR incised 15 m, and at ~13.4 ka, it incised an additional 40 m (Knox 2007; Loope 2012; Gran et al. 2013). Each of these incision episodes abruptly lowered the base level of the LCR, creating knickzones that migrated up the LCR and its tributaries. The incision resulting from knickzone migration created the Wissota terrace, a prominent landform in the LCR valley that marks the maximum height of LCR aggradation during the Late Wisconsinan (Andrews 1965). Figure 2. Terraces of the LCR valley. The names of terraces below the Wissota terrace are based on their height above the modern floodplain and distance below the Wissota. From lowest to highest these are T-1, T-2, T-3, T-4, T-5, and T-6. One terrace that does not fit into the T1-T6 schema is found in the relatively narrow bedrock valley downstream from the Eau Galle-Chippewa River confluence. Named the Maxville Terrace, this terrace slopes from the level of the Wissota at its upvalley end to a level equivalent to T-4 in the UMR valley. (Figure from Faulkner et al. (2016).) 74 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 3. A proposed model of the evolution of the longitudinal profile of the LCR in response to UMR incision labeled with OSL ages of terrace alluvium. The profiles were constructed by connecting scattered terrace remnants, except for the Wissota and Maxville terraces, which are relatively continuous features. (Modified from Faulkner et al. (2016).) Autogenic variations in the amount of sediment supplied to the river likely explain why the migration of knickzones and incision up the LCR was episodic and prolonged. Incision resulting from knickzone migration would’ve created a relatively deep narrow channel with steep, unstable banks. Bank collapse, promoted by lateral stream erosion, would have greatly increased the supply of sediment to the stream. With more sediment to transport, the stream would no longer have had excess power, causing knickpoint migration and incision to slow or cease altogether. This, in turn, would’ve allowed lateral stream migration and floodplain formation to occur. Over time, lateral erosion and bank failure would’ve caused the banks to move away from the stream and become less steep, leading to a reduction in the amount of sediment supplied from them to the LCR. With a declining sediment supply, the stream would’ve once again had excess power, resulting in renewed knickzone migration and incision, at least until the process repeated itself farther upstream. The supply of sediment to the LCR from its tributaries was also subject to autogenic variations. LCR incision resulting would have lowered base level for its tributary streams, creating knickzones that then migrated up them. The subsequent tributary incision would have caused a dramatic increase in the supply of sediment to the LCR. Sediment supplied from tributaries would have remained high until their incised channel banks began to stabilize. But until that happened, high amounts of tributary sediment would have affected knickpoint migration and incision on the LCR, slowing it down and possibly causing it to stop. It is likely that autogenic variations in sediment from the largest tributaries, the Red Cedar River and the Eau Claire River, had the biggest impact on the LCR. Aeolian Landscapes There is abundant evidence throughout the Eau Claire region that wind has been a significant geomorphic agent during the late Quaternary. The most widespread evidence of aeolian activity is provided by deposits of loess, which were mainly sourced from the outwash plains of meltwater streams, including that of the Chippewa River (Schaetzl et al., 2014; Schaetzl et al., 2018; Fig. 4). Loess deposition during the Late Wisconsinan in the Eau Claire region began no later than 24 ka and continued until as recently as 10 ka (Schaetzl et al., 2014). Over this interval, the dominant processes of loess transport and deposition 75 �Proceedings of the 69th ILSG Annual Meeting - Part 2 2014). Later, strong northwest winds entrained sands from outwash and weathered sandstone. As they traveled over existing loess, saltating sands remobilized it and kept it in suspension until topographic barriers blocked further sand movement, which allowed loess to accumulate in their lee. Today, large swaths of the region are loess-free, with the thickest loess found on the southeast sides of prominent sandstone inselbergs and ridges (Schaetzl et al., 2018). A variety of sandy aeolian landforms, such as parabolic dunes, sand sheets, sand ramps, and sand stringers, also attest to the geomorphic significance of wind in the Eau Claire region. While these landforms are generally subtle and apparent only on LiDAR-derived DEMs, they are widespread (Fig. 5). They also have a generally consistent orientation, which indicates that they were primarily formed by west-northwesterly winds. In addition, a dozen OSL ages from different landforms reveal that sandy aeolian landforms in the region were being deposited between 13 and 9 ka (Schaetzl et al., 2018; Millett, 2019; Mataitis, 2020; Shandonay et al., 2022). Figure 4: Extent and thickness of loess within in western Wisconsin, as derived from Natural Resources Conservation Service county soil surveys (from Schaetzl et al., 2018.) apparently changed (Schaetzl et al., 2018). Existing evidence indicates that, early on, loess was primarily deflated from the outwash plains of the Chippewa River and its meltwater tributaries and deposited downwind of them (Schaetzl et al., Figure 5. Parabolic dunes (Millett, 2019) and sand stringers (Schaetzl et al., 2018; Mataitis, 2020) identified from lidar-derived DEMs, aerial photographs, and soil survey data in and near the LCR valley. 76 �Proceedings of the 69th ILSG Annual Meeting - Part 2 buried and melted. The dominant landforms in the area are ice-walled lake plains (Figs. 6 and 7; Clayton et al., 2001), such as the one the Obey Center is built upon. These form when lakes develop in the stagnant ice landscape preserved in permafrost conditions. It is likely that permafrost conditions were in northern Wisconsin until ~13.5 ka (Attig and Rawling, 2018; Batchelor, 2019). They are composed of laminated finegrained sediment that is typically fine sand and silt. These landscapes contain organic material further south that have aided in interpreting the timing of the Lake Michigan Lobe (Curry et al., 2018); however, organic material is typically not preserved in northern Wisconsin. Field Trip Stops UTM coordinates are in zone 15, WGS84 datum Stop 1: Copper Falls Glacial Deposits UTM coordinates 623464E, 5008658N Till of the Copper Falls Formation is reddish brown, sandy (~30 – 80% sand; Syverson, 2007; Syverson et al, 2011), and sourced from the northeast. Copper Falls till is distinguished from older River Falls till primarily by the landscape they underlie. Copper Falls sediment is found in relatively unmodified landscapes formed during the late Wisconsinan. Glacial landforms (moraines, eskers, drumlins, ice-walled lake plains…) are well persevered. River Falls sediment is associated with a highly eroded landscapes and likely formed before the late Wisconsinan Glaciation. This roadcut exposes typical unsorted glacial deposits of the Copper Falls Formation deposited as stagnant ice melted. Stop 2. David R. Obey Ice Age Interpretive Center UTM coordinates 624514E, 5009009N The interpretive center is located in the Chippewa Moraine State Recreation Area and within the Chippewa Moraine. The moraine formed when late Wisconsin ice was at its maximum position (Syverson, 2007). The landscape here is generally described as hummocky, and formed as stagnant ice was Figure 6. Schematic illustration showing the formation of ice walled lake plains (from Clayton et al., 2001). (A) Supraglacial sediment forms at the surface of stagnant ice and lakes occupies low areas where sorted sediment accumulates. (B) Hummocky topography with ice-walled lake plains remain after the ice has melted. Figure 7. Lidar-derived DEM showing the hummocky topography of the Chippewa Moraine near Stops 1 and 2. Ice walled lake plains are abundant within the moraine, which contrasts with the flat landscape formed by meltwater streams (outwash). 77 �Proceedings of the 69th ILSG Annual Meeting - Part 2 suggest that it happened during the late Holocene (Fig. 3). Stop 3. Colluvium Exposure UTM coordinates 625123E, 5001761 N The falls consist of a series of small knickpoints formed in early Proterozoic bedrock consisting of banded amphibolites with granitic intrusions (Myers and Maercklein, 1978). The angular form of the knickpoints, along with angular boulders of the same lithology scattered along the channel bed, suggest that the river is incising here primarily by mean of hydraulic plucking. Hydraulic plucking occurs when flows are deep and fast, leading to a zone of flow separation and low pressure on the downstream face of knickpoints. If the bedrock of a knickpoint is sufficiently jointed and weathered, the resulting drag force will pull blocks away from the knickpoint face. Polished rock surfaces with rare grooves and potholes indicate that abrasion by bedload sediment is also playing a role here in channel incision, although its effects appear secondary to that of plucking. The landscape beyond the LIS margin was greatly affected by periglacial processes. One of the most profound effects was the stripping of hillslopes by the mass-wasting process of solifluction (Clayton et al., 2001). Evidence of solifluction is provided by relict deposits of colluvium (colluvial aprons) that mantle bedrock slopes in areas of former permafrost. Here we see a prime example of such a colluvial deposit, which consists of an unsorted mixture of sandstone clasts (pebble to boulder in size) supported in a matrix of silty sand. The sand and the sandstone clasts were likely derived from the underlying bedrock (sandstone of the Cambrian Eau Claire Formation) by intense freeze-thaw weathering during the Late Wisconsinan. The silty material probably is loess that winds deflated and transported to the sight from nearby outwash plains. While the Chippewa has clearly incised into the Jim Falls bedrock, incision overall has been minimal. The lack of incision is likely due, in part, to the weathering and erosion-resistant nature of the bedrock. An additional factor is the relatively short amount of time that the river has been incising at this location. Given the model of long-profile evolution in Figure 3, incision didn’t start here until sometime after 4.7 ka. Stop 4. Jim Falls UTM coordinates 635875E, 4990538 N) The Chippewa at Jim Falls is an example of a superimposed river (Fig. 8). Here, the Chippewa River incised through a cover of outwash and till and encountered a topographic high in the buried a bedrock landscape. It did not, however, incise to its present level in one episode of downcutting, as evidenced by two terraces that are apparent at and near this site. The highest terrace grades to the Wissota terrace, the maximum level of aggradation of the lower Chippewa River. This indicates that the river continued to flow at this level after the glacial margin had retreated to the north of this site. Remnants of a terrace approximately 3 meters below the Wissota (best seen downstream from the east end of the pedestrian bridge), which is cut into till, indicate a period of channel stability before the river incised further to the buried bedrock. When deeper incision occurred is unknown, although OSL ages of terrace alluvium farther down valley The site today clearly is highly modified by a dam, appropriately named the Jim Falls Dam. The original Jim Falls Dam was built in 1923 to utilize the hydraulic head provided by the falls to generate electricity. It was designed so that the falls were bypassed and left dry except during high flows, when excess water was released through a spillway that was located at the falls upstream end. The dam was redeveloped in 1988 and now has the highest generation capacity of any hydropower dam in Wisconsin (~60 MW). This redevelopment included moving the main spillway from the head of the falls to a location adjacent to a new main powerhouse. It also included constructing a smaller spillway and 78 �Proceedings of the 69th ILSG Annual Meeting - Part 2 auxiliary powerhouse where the main spillway had been. This was done so that a minimum flow of 240 cfs could be released down the bypassed reach year-round, except for the period April 1- May 31, when flow through the reach is increased to 850 cfs to enhance fish spawning habitat. Figure 8. Lidar-derived DEM of Jim Falls and surrounding area. Stop 5. Wildenberg Quarry Evidence for permafrost during the Late Wisconsinan is widespread in Wisconsin, with a hypothesized permafrost interval in central Wisconsin from ca. 33 to 15 ka (Batchelor et al., 2019) and as late as 13.5 ka in northern Wisconsin (Attig and Rawling, 2018). This interval is, however, poorly constrained in the Eau Claire region due to a lack of 14C datable materials in features diagnostic of permafrost. OSL dating now makes it possible to date proxy geomorphic features, like ice-wedge pseudomorphs, to help constrain this (Schaetzl et al., 2021). UTM coordinates 616648 E, 4970372 N Note: This is private property. No access is allowed without owner’s permission. Glacial sediment of the River Falls Formation includes till that is lithologically like the Copper Falls Formation and melt-water stream sediment (Syverson, 2007; Severson et al., 2011). These can be distinguished from the Copper Falls formation because they occur at the top of highly eroded landscapes and the soils in them are more developed. There are no primary glacial landforms associated with the River Falls Formation, likely due to intense modification by periglacial processes in permafrost conditions during the late Wisconsinan. The ice-wedge pseudomorphs in this quarry are sand wedges (Fig. 9). Sand wedges such as these form in periglacial settings when thermal contraction of frozen ground in winter forms 79 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 9. Sand wedges exposed in the headwall of the Wildenberg quarry in July 2018. (Photograph by Randy Schaetzl.) cracks in the soil. If this happens in a cold, dry, wind-swept environment with sand available for transport, sand will blow into and fill the cracks. Over time, repeated cracking and filling will form vertical structures that generally taper with depth. OSL dating of sand wedges in this quarry and in another quarry located 60 km to the south indicate that thermal contraction cracks existed and were filling with sand from no later than 19.3 ka until 14.7 ka. Schaetzl et al. (2021) interpret these ages as documenting when permafrost in the region most likely ended. Interestingly, the OSL ages from this quarry (15.1 and 14.7 ka) are younger than those from the quarry 60 km to the south (19.3, 19.1, and 18.3 ka). These may represent a time-transgressive spatial relationship in that permafrost possibly degraded earlier at the more southerly location and remained longer at the more northerly one, although this is purely speculative given the large errors on the OSL ages (1.4 to 2.2 ka at 1). That said, the larger and more complex morphologies of the sand-wedges found at this quarry do suggest the possibility of more intense sand-wedge development due to more prolonged permafrost conditions. City incorporated it into its bicycle-pedestrian trail system. In addition to being historically significant, the High Bridge (Fig. 10) affords an excellent view of many of the terrace levels found in the LCR valley. The High Bridge itself is at the level of T6. To the east, trees and houses can be seen at the top of the Wissota terrace scarp, which is 6-7 meters above T-6. The Wissota can also be seen to the west where the pedestrian-bicycle trail cuts into its scarp. Looking downstream, lower terraces are difficult to discern from this vantage point, although the residential and business districts located near the river provide clues. These built-up areas are all above the 100-year flood level. That is, they all are on terraces. In Eau Claire, there is little active floodplain. This suggests that incision below the lowest terrace level occurred here recently (within the last 2.3 ka according to the model of long-profile evolution in Fig. 3). Upstream from the High Bridge is the Dells Dam. This dam, which was built in 1924 for the purpose of generating electricity, is situated at an unusually good site for a dam on the LCR – a bedrock gorge. This gorge was formed when the river incised into a cover of glacial outwash that buried a low ridge of Cambrian sandstone (Mt. Simon Formation) connecting Mt. Simon (the conical bedrock hill located approximately 500 m northeast of the dam) to the bedrock uplands located west-northwest of it. In other words, the river here did not incise into its pre-glacial bedrock valley, which is located east of Mt. Stop 6. High Bridge UTM coordinates 617832E, 4964561N This stop is on the so-called High Bridge in the city of Eau Claire. Standing 26 m above the Chippewa River, the High Bridge was built in 1880 as a railroad bridge and was an innovative bridge for its time. It was abandoned in 1992, and the City acquired ownership in 2007. In 2015, the 80 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Simon and the High Bridge. This is an example of an epigenetic gorge (Ouimet et al., 2008), and the river likely carved it sometime after 7.4 ka (see Fig. 3). Incision is actually still occurring here, as evidenced by a 1.5-m bedrock knickpoint located 180 m downstream of the bridge. Figure 10. Lidar-derived DEM showing the stream terraces found in Eau Claire in the vicinity of the High Bridge. upvalley and on the valley’s other side. The last episode of incision, below T-1, occurred within the last 2.3 ka based on an OSL date of T-1 fill from a site also located 6.5 km upvalley (Fig. 3). In the downstream direction, the valley below the Wissota is predominantly floodplain with only rare lower terrace remnants. In addition, the river’s planform switches from a single-thread meandering shape to a multi-thread anabranching one that extends downvalley for a distance of 8.5 km. At that point, it returns to a single-channel meandering form. It is uncertain why this anabranching reach exists, although its similarity to the sedimentation zones of wandering gravel bed streams in British Columbia described by Desloges and Church (1987) suggests a cause. Given its location immediately downstream of Stop 7. Sand Hill Cemetery UTM coordinates 599549E, 4958254N This stop is at the edge of the Wissota terrace tread and top of the Wissota terrace scarp. To the north, a braided channel is apparent in the subtle rolling topography of the Wissota tread, providing evidence of the Chippewa’s glacial past (Fig. 11). To the south, the Wissota scarp descends nearly 30 m to the lowest terrace in the valley (T-1). Looking upvalley, this terrace can be identified by the farmland situated on it. Lowlying land that is wooded is either floodplain or paleochannels cutting across the T-1 surface. Incision here below the Wissota level occurred between 10 and 9 ka, based on four OSL dates of Wissota fill obtained from sites located 6.5 km 81 �Proceedings of the 69th ILSG Annual Meeting - Part 2 the reach that was recently incised below T-1, it could be the result of sedimentation resulting from that incision event (Fig. 12). A pronounced convexity in the long profile of the modern river along the anabranching reach supports this hypothesis (Faulkner et al., 2016). Figure 11. Lidar-derived DEM showing the fluvial and aeolian landforms in the vicinity of the Sand Hill Cemetery (Stop 7), which are discussed in the text. Figure 12. Cartoon showing the setting of the anabranching reach (and hypothesized sedimentation zone) downstream of the reach that is incised below the T-1 terrace (adapted from Adams et al., 2016). 82 �Proceedings of the 69th ILSG Annual Meeting - Part 2 A short distance (~200 m) west of the Sand Hill Cemetery, an elongate wooded hill—informally named the Steffes-Zanoni site—can be seen rising above the low-relief landscape of the Wissota terrace. (This hill can also be seen in Fig. 11.) Its summit is approximately 8 m higher than the surrounding Wissota terrace surface, and its elongate form runs parallels to the Wissota terrace scarp. The landform is composed of sand that is mineralogically indistinguishable from the terrace sediments beneath it, although it is relatively finer-grained and more well-sorted, indicating that it is aeolian sand (Millett, 2019). It is also morphologically complex, consisting of multiple parabolic forms coalesced together with smaller parabolic forms on top of them. These forms--smaller parabolic dunes superimposed on larger older ones—indicate repeated aeolian activity at this site. The OSL ages of samples obtained from depths of 1.7 m and 2.5 m near the hill’s northwest end suggest a depositional age of the landform’s upper part to be ~0.5 ka. Larson et al. (2008) identified a dune similar to this one in the city of Eau Claire and called it a cliff-top parabolic dune (because of its form and its proximal relationship to the Wissota terrace scarp). Since then, many cliff-top parabolic dunes have been noted in the Eau Claire region (note the large number of parabolic dunes situated along the Wissota scarp in Fig. 5). Why these dunes exist will be discussed at our final stop. with parabolic forms oriented perpendicular to them—are hypothesized to have had a similar genetic origin (Larson et al., 2008; Millett, 2019). In their proposed model, a river cuts into the base of a high fill-terrace scarp. This creates an unstable cutbank and promotes mass wasting that removes vegetation from the scarp face. Wind that then blows against such a scarp gets compressed, which causes its velocity to go up. The increase in velocity enhances the wind’s ability to entrain exposed sandy sediment and transport it up the scarp face. When this happens, the sandy sediment ultimately settles out at the top of the scarp as wind velocity is reduced there. This leaves behind “cliff-top dunes.” (Fig. 14). Given this model of dune genesis, one should expect cliff-top dunes to have different orientations and depositional ages compared to other parabolic dunes not in cliff-top positions. This is indeed the case. Non-cliff-top dunes generally have a northwest-southeast orientation and depositional ages older than 10 ka. Cliff-top dunes display a variety of orientations (perpendicular to their scarps) and are generally much younger. OSL ages from a cliff-top dune in the city of Eau Claire indicate a period of aeolian deposition at ca. 6.0 ka, while two from the Steffes-Zanoni site (Stop 7) indicate that deposition in the uppermost dune sediments occurred at ca. 0.5 ka. At the Kiwanis site, eight OSL ages point to two depositional episodes: at ca. 0.9 ka and 0.5 ka. Stop 8. Town of Union Conservancy The model of Larson et al. (2008) of cliff-top dune genesis suggests that these should be forming wherever the Chippewa River is eroding laterally into Wissota terrace fill. This, however, is not the case; today, all cliff-top dunes in the LCR valley are stabilized by vegetation and no longer moving. Thus, the genesis of these landforms is doubtless more complex than the model of Larson et al. (2008) suggests, with climatic variability likely playing a key role in the process of aeolian sedimentation and dune formation at cliff-top locations (Millett, 2019). For example, during humid climate intervals UTM coordinates 607504E, 4959523N Note: This site involves walking on unpaved trails for an approximate distance of one mile. There are some short sections of the trail that are moderately steep. Parabolic dunes in cliff-top positions are especially prominent at this location, which is informally called the Kiwanis Site (Fig. 13). Like those seen at the previous stop, these dunes are situated atop the Wissota terrace scarp. These dunes and others that are similarly situated in the LCR valley—above high Wissota terrace scarps 83 �Proceedings of the 69th ILSG Annual Meeting - Part 2 (such as at the present), high cutbanks carved by the river into Wissota fill are colonized readily by vegetation, which inhibits the entrainment and transport of sand up terrace scarps. In contrast, during arid intervals, vegetation cover is greatly reduced, especially on steep well-drained terrace scarps, allowing wind to entrain and transport sand up them. OSL ages from the Kiwanis Site and the Steffes-Zanoni Site support the significance of climate variability in the formation of cliff-top dune. At Kiwanis, these ages indicate two pulses of aeolian deposition – the first at ca. 0.9 ka and a second at ca. 0.5 ka, with the latter coinciding with ages from the Steffes-Zanoni Site. If correct, these pulses happened during the Medieval Climatic Anomaly when well documented dry periods affected the mid-continent of North America (reviewed in Millett, 2019). forms enclosed by a subtle linear ridge appear to be anthropogenic features. It is likely, based onthe morphology and distribution of these features, that their creation was tied to the genesis of the prominent aeolian dunes found nearby and, potentially, culturally linked to a wellestablished late Woodland period of mound building in the upper Mississippi River valley – a period which would have coincided with the formation of the Kiwanis and Steffes-Zanoni dunes. Based on the configuration of these features, it is hypothesized that this site represents the “Thunderer,” an effigy of a bird-like deity, or sky being. The spotted Thunderer, with “spots” represented by the location of the mounds within the linear structure, is associated with the West and the bringer of storms. If true, Native peoples may have watched the dunes form during a period of more aridity during episodes of higher winds/storms, resulting in this site being spiritually significant at that time and now an important site of cultural heritage (R. Schirmer, personal communication). Lastly, closer examination of the Kiwanis Site (Fig. 13) reveals landforms in close proximity to the dunes that do not look to be of natural origin. Several hemispherical (or conical) mound-like Figure 13. The Kiwanis Site. Cliff-top parabolic dunes are situated directly above the Chippewa River on top of the Wissota terrace scarp. Also note the hemispherical mounds and linear ridge of hypothesized anthropogenic origin. 84 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Figure 14. Model of cliff-top dune genesis induced by lateral river erosion (adapted from Larson et al., 2008). Begin, Z.B. 1986. Determination of “diffusion” erosion coefficients for some tributaries of Oaklimiter Creek, North-Central Mississippi, in Hadley, R.F., ed., Drainage Basin Sediment Delivery. IAHS Publication, p. 447–462. References Adams, H.R., Vincent, A.M., and Faulkner, D.J. 2016. Patterns of downstream fining on the lower Chippewa River, Wisconsin: Abstracts, Annual Meeting - American Association of Geographers, San Francisco, CA Begin, Z.B. 1988. Application of a diffusion-erosion model to alluvial channels which degrade due to base-level lowering: Earth Surface Processes and Landforms, v. 13, p. 487–500. Andrews, G.W. 1965. Late Quaternary geologic history of the Lower Chippewa Valley, Wisconsin. Geological Society of America Bulletin: v. 76, p. 113–124. Begin, Z.B., Meyer, D.F., and Schumm, S.A. 1981. Development of longitudinal profiles of alluvial channels in response to base-level lowering: Earth Surface Processes and Landforms, v. 6, p. 49–68. Attig, J.W. 1985. Pleistocene Geology of Vilas County, Wisconsin: Wisconsin Geological and Natural History Survey, Information Circular 51, 32 p. Clayton, L., Attig, J.W., and Mickelson, D.M. 2001. Effects of late Pleistocene permafrost on the landscape of Wisconsin: Boreas, v. 30, p. 173–188. Attig, J.W. 1993. Pleistocene Geology of Taylor County, Wisconsin: Wisconsin Geological and natural History Survey, Bulletin 90, 25 p. Curry, B.B., Lowell, T.V., Wang, H., and Anderson, A.C. 2018. Revised time-distance diagram for the Lake Michigan Lobe, Michigan Subepisode, Wisconsin Episode, Illinois, USA, in Kehew, A., and Curry, B.B., eds., Quaternary Glaciation of the Great Lakes Region: Process, Landforms, Sediments, and Chronology: Geological Society of America Special Paper 530, p. 1–12. Attig, J.W., and Rawling, J.E., III. 2018. Influence of persistent buried ice on late glacial landscape development in part of Wisconsin’s Northern Highlands, in Kehew, A., and Curry, B.B., eds., Quaternary Glaciation of the Great Lakes Region: Process, Landforms, Sediments, and Chronology: Geological Society of America Special Paper 530, p. 1–12. Desloges, J.R., and Church, M. 1987. Channel and floodplain facies in a wandering gravel-bed river, in Ethridge, F.G., Flores, R.M., Harvey, M.D., Weaver, J.N., eds., Recent Developments in Fluvial Sedimentology: Special Publication. Society of Economic Paleontologists and Mineralogists, p. 99–109. Batchelor, C.J., Orland, I.J., Marcott, S.A., Slaughter, R., Edwards, R.L., Zhang, P., Li, X., and Cheng, H. 2019. Distinct permafrost conditions across the last two glacial periods in midlatitude North America: Geophysical Research Letters, v. 46, no. 22, p. 13318-13326. 85 �Proceedings of the 69th ILSG Annual Meeting - Part 2 Faulkner, D.J., Larson, P.H., Jol, H.M., Running, G.L., Loope, H.M., and Goble, R.J. 2016. Autogenic incision and terrace formation resulting from abrupt late-glacial base-level fall, lower Chippewa River, Wisconsin, USA: Geomorphology, v. 266, p. 75–95. Schaetzl, R.J., Forman, S.L., and Attig, J.W. 2014. Optical ages on loess derived from outwash surfaces constrain the advance of the Laurentide Ice Sheet out of the Lake Superior Basin, USA: Quaternary Research, v. 81, p. 318–329. Schaetzl, R.J., Larson, P.H., Faulkner, D.J., Running, G.L., Jol, H.M., and Rittenour, T.M. 2018. Eolian sand and loess deposits indicate west-northwest paleowinds during the Late Pleistocene in western Wisconsin, USA: Quaternary Research, v. 89, p. 769–785. Gran, K.B., Finnegan, N., Johnson, A.L., Belmont, P., Wittkop, C., and Rittenour, T. 2013. Landscape evolution, valley excavation, and terrace development following abrupt postglacial baselevel fall: Geological Society of America Bulletin, v. 125, p. 1851–1864. Schaetzl, R.J., Running, G.L., Larson, P., Rittenour, T., Yansa, C., and Faulkner, D. 2022. Luminescence dating of sand wedges constrains the Late Wisconsin (MIS 2) permafrost interval in the upper Midwest, USA: Boreas, v. 51, p. 385– 401. Knox, J.C., 2007. The Mississippi River System, in Gupta, A., ed., Large Rivers: Geomorphology and Management. John Wiley & Sons, Chichester, England ; Hoboken, NJ, p. 145–182. Larson, P.H. McDonald, J., Baker, A., Dryer, W.P., Running, G.L., Faulkner, D.J. and Jol, H.M. 2008. Geomorphology of cliff-top parabolic dunes within the lower Chippewa River valley, upper Putnam Park, Eau Claire, Wisconsin: Abstracts, Annual Meeting - Association of American Geographers, Boston, MA. Schirmer, R. 2023. Personal communication regarding archeology in the upper Mississippi valley and the Kiwanis Site. 3/8/2023. Shandonay, K.L., Bowen, M.W., Larson, P.H., Running, G.L., Rittenour, T., and Mataitis, R. 2022. Morphology and stratigraphy of aeolian sand stringers in southeast Minnesota and western Wisconsin, USA: Earth Surface Processes and Landforms, v. 47, p. 2863–2876. Loope, H.M., Mason, J.A., Knox, J.C., Goble, R.J., Hanson, P.R., Young, A.R., and Curry, B.B. 2012. Late Wisconsinan aggradation and incision history of the upper Mississippi River, USA: Abstracts with Programs - Geological Society of America, v. 44, p. 455. Syverson, K.M. 2007. Pleistocene Geology of Chippewa County, Wisconsin: Wisconsin Geological and natural History Survey, Bulletin 103, 53 p. Mataitis, R. 2020. “Geomorphology and chronology of sand stringer deposition beyond the ice margin: Southeastern Minnesota and western Wisconsin, USA.” M.S. Thesis. Minnesota State University, Mankato. Syverson, K.M., Clayton, L., Attig, J.W., and Mickelson, D.M. 2011. Lexicon of Pleistocene Stratigraphic Units of Wisconsin: Wisconsin Geological and Natural History Survey Technical Report 1, 180 p. Millett, J. 2019. “Cliff-top dunes in the lower Chippewa River valley of west-central Wisconsin.” M.S. Thesis. Minnesota State University, Mankato. Woodruff, L.G., Attig, J.W., and Cannon, W.F. 2004. Geochemistry of glacial sediments in the area of the Bend massive sulfide deposit, north-central Wisconsin: Journal of Geochemical Exploration, v. 82, p. 97–109. Myers, P.E., and Maercklein, D.A. 1978. Amphibolites and Granites at Jim Falls: Wisconsin Geology of Wisconsin – Outcrop Descriptions, Geological and Natural History Survey, 7 p. Rawling III, J.E., Carson, E.C., Attig, J.W., Mickelson, D.M., Mode, W.N., Johnson, M.D., Syverson, KM. (in preparation). The Quaternary Geology of Wisconsin. Wisconsin Geological and Natural History Survey. Map Sale 1:500,000. 86 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology: Proceedings, 2023 Description An account of the resource Institute on Lake Superior Geology. Eau Claire, Wisconsin. April 24-25, 2023. Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2023-04-24 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English Type The nature or genre of the resource Text https://digitalcollections.lakeheadu.ca/files/original/f3eca25e11be27330162009a0a576dd3.pdf c652fed41c0ca84f2ff1ac5aa5644ac5 PDF Text Text The Volcanoes of the Island of Hawaii Field Trip Guide Institute on Lake Superior Geology Special Publication 3 Allan MacTavish, M.Sc., P.Geo. George Hudak III, Ph.D., P.Geo. December 2023 �Table of Contents 1. Introduction .................................................................................................................................. 2 1.1. Field Stop and Information Sources: ......................................................................................... 2 1.2. Field Trip Overview: .................................................................................................................. 3 1.3. 2022 Mauna Loa Eruption:........................................................................................................ 4 1.4. Acknowledgements:.................................................................................................................. 4 1.5. Photo Credits for Field Guide Cover: ........................................................................................ 4 1.6. Important Notes on Spelling, Location, Distance, and Units of Volume: ................................. 4 2. 2.1. 3. Geology of the Hawaiʻian Islands .................................................................................................. 9 Volcano Descriptions .............................................................................................................. 15 Field Trip Stops ............................................................................................................................ 40 3.1. Day 1: Kailua-Kona to Hawai‘i Volcanoes National Park ........................................................ 40 3.2. Day 2: Highway 11 and South Point ....................................................................................... 45 3.3. Day 3 (Part 1): Mauna Loa Road and Mauna Loa Strip .......................................................... 49 3.3. Day 3 (Part 2): Mauna Iki Trail/Kaʻū Desert Trail and the Southwest Rift Zone..................... 53 3.4. Day 4: Kīlauea Caldera, Kīlauea Iki, Hilina Pali ....................................................................... 63 3.4. Day 4 (Part 1): Kīlauea Caldera............................................................................................... 64 3.4. Day 4 (Part 2): Koaʻi Fault Zone and Hilina Pali ...................................................................... 79 3.5. Day 5 – Chain of Craters Road, Napaū/Mauna Ulu Trail, Hōlei Pali ........................................ 84 3.5. Day 5 (Part 2) – Chain of Craters Road.................................................................................. 102 3.6. Day 5 (Part 1): Helicopter Flight over Kīlauea (Morning) ..................................................... 107 3.6. Day 6 (Part 2): Kīlauea Lower East Rift Zone (afternoon) .................................................... 110 3.7. Day 7 (Mauna Loa) and Day 8 (Mauna Kea).......................................................................... 119 3.7. Day 7 (Part 1) – Hilo Area ...................................................................................................... 120 3.7. Day 7 (Part 2): Saddle Road (Highway 220) and Mauna Loa Observatory Road ................. 123 3.8. Day 8 – Mauna Kea Summit Road ......................................................................................... 129 3.9. Day 9: Mamalahoa Highway (Hawai‘i Belt Road; Hāmākua Coast) ..................................... 133 3.10. Day 10 – Kohala Volcano – Waimea to Hāwī ........................................................................ 138 3.11. Day 11 – Waimea to Kailua-Kona .......................................................................................... 146 4.1. Glossary References .................................................................................................................. 168 5. Field Guide References ............................................................................................................. 172 1 �1. Introduction The original version of this field guide was written to accompany the successful Institute on Lake Superior Geology (ILSG)-sponsored ‘Volcanoes of the Island of Hawaiʻi’ Field Trip’ that took place between February 11 and 21, 2020. The aims of the field trip were to observe the characteristics of modern volcanoes in an intra-plate, non-rift environment (see Figure 1) and to provide contrast and comparison with modern and ancient rift environments (i.e. North American Mid-Continent Rift (MCR); Mid-Altlantic Rift in Iceland). The 9 main Hawai‘ian Islands are shown in Figure 2. This field trip is confined to the Island of Hawai’i, which is the easternmost and largest island of the chain and is host to 5 extinct, dormant, or active volcanoes (see Figure 3). Short volcano descriptions are given within this introductory section immediately below (USGS HVO website); whereas longer descriptions, including 2 associated submarine volcanoes, will be presented in the ‘Volcano Descriptions’ sub-section of the ‘Geology of the Hawaiʻian Islands’ section, also below: 1) Kīlauea (4009ft, 1222m), is the most active subaerial (exposed above sea level) volcano on earth and was continuously active between 1983 and 2018, mainly in the vicinity of the Pu‘u Ō‘ō vent; however, the 35-year eruption ended after a caldera subsidence event and a voluminous 3month eruption from the Lower East Rift Zone occurred between April 30 and August 9, 2018. After the 2020 field trip there have been three successive eruptions at the volcano’s summit, all consisting of a lava lake that began infilling the 2018 caldera subsidence crater centred on Halemaʻumaʻu Crater. The first eruption began on December 20, 2020 and ceased on May 13, 2021; the second commenced on September 29, 2021 and ended on December 9, 2022; and the third began less than a month later on January 5, 2023 and is ongoing. 2) Mauna Loa (13,681ft, 4170m) is active and the largest, most massive volcano on earth; after last erupting in 1984 Mauna Loa began erupting from fissures within its summit caldera on November 27, 2022. The vigorous eruption migrated north to several fissures located on the upper Northeast Rift Zone on the 28th where lava continued to flow until the eruption was declared over on December 10, 2022 (USGS HVO Website). 3) Mauna Kea (13,803ft, 4207m) is considered by some to be active, but is presently within the waning stages of activity (i.e., dormant); Mauna Kea last erupted at ~3300 BP (Before Present); 4) Hualālai (8278ft, 2523m) is in its waning stages of activity and last erupted in between 1801 and 1802; and 5) Kohala (5480 ft, 1670m) is considered extinct; it is deeply eroded with its last eruption occurring at ~100,000 BP. This version of the guide is no longer designed to be used by a specific field trip group (i.e. the 2020 group) and has been modified to be used by anyone, or any group, who so wishes to examine the volcanology of the island. Since the primary author (AM) is very visual-centric this guide contains a large number of photographs and maps. This is primarily to help those using the guide to more readily find and identify field trip outcrops and other described features. Also, the guide has, in some ways, become a travelogue of the 2020 field trip with many of the included photos taken during that trip and an earlier field field trip reconnaissance completed by the primary author (AM) in August 2019. 1.1. Field Stop and Information Sources: Several excellent, pre-existing geological field guides, written by experienced and knowledgeable authors, and publications available from the US National Park Service (NPS) were critical to the preparation of this guide. Those used (in no particular order) were: Hazlett (2014); Hazlett and Hyndman (2007 edition of 1996 book); Easton and Easton (1995); Robinson (2010, 2012); Merguerian and Okulewicz (2007); Tilling et. al. (2010); and the NPS Mauna Ulu Eruption and Kilauea Iki Trail guides. 2 �Full source citations are listed in the reference section of this guide. Some field stops were generated, and others discarded, after a field reconnaissance using a preliminary version of the guide was completed by AM during the summer of 2019. Also, several field stops were added after the 2020 field trip using descriptions and locations discovered and then generated by AM during the field trip. Both the satellite and street views of Google Earth Professional were of inestimable use for planning, locating, and examining field stops proposed for the original version of the guide. AM cannot stress how important it was to have the ability, within a few minutes to examine, download a high resolution image, and obtain a UTM location of a proposed stop while sitting in front of a computer located over 4100mi (6600km) away in northwestern Ontario. Google Earth is an incredible resource, particularly Google Earth Street View. 1.2. Field Trip Overview: This 11-day field trip consists of 96 stops and 49 sub-stops. The suggested daily field trip routes are shown in Figure 4; however, those using this guide do not have to stick to the routes or visit all of the stops shown. The daily routes were designed to decrease daily driving distances and radiate, or extend linearly from the accommodations used during the 2020 field trip. One recommended variation of the routes given would be to modify the days such that the shortest hikes are done first and the longest hikes are done last. This allows participants to a build-up a conditioning level over the course of the field trip making each successive hike easier to complete. Experience during the 2019 summer reconnaissance by AM clearly showed that reordering the hikes from shortest to longest with 1 or 2 non-hiking days between hikes worked well during the 2020 trip, particularly for the older participants. The helicopter tour completed on the morning of Day 6 of the 2020 field trip provided an excellent overview of the structure and eruptive activity of Kīlauea, helped add perspective and scale to the surface visits, and hopefully tied the many Kīlauea stops into a cohesive whole. Those using this guide do not have to take a helicopter tour, or if taking a tour, even to use the route shown in Figure 4. The authors strongly recommend taking a helicopter tour if your trip budget allows. It is surprisingly affordable and well worth the cost, particularly if you can fill the helicopter with your group. There are several tour companies operating out of the International Airports in Hilo and Kailua-Kona that offer packaged aerial tours over Kīlauea. The 2020 field trip chartered 2 Bell 407 aircraft from Paradise Helicopters in Hilo who, since we chartered the helicopters rather than booking an existing fixed tour, allowed us to tailor the specific route flown within the chartered time-frame (paid by the hour rather than by a per/person cost. Over the last 15 years AM has taken many helicopter flights over the islands and volcanoes of Hawai’i and every flight was well worth the cost. Participants of the 2020 field trip mainly stayed in budget accommodation in order to keep field trip costs within a reasonable range. These accommodations included the Kīlauea Military Camp (dormitory accommodations, 8 nights), the Kamuela Inn in Waimea (2 nights), and the Marriott Courtyard King Kamehameha’s Kona Beach Hotel (Kailua-Kona, 1 night). Of course, users of this guide can stay whereever they wish, again depending upon their trip budget. For the most part, meals were not provided during the 2020 trip. The Kīlauea Military Camp has restaurant facilities for breakfast and dinner meals. The food is relatively basic, but quite good and bag lunches are available for purchase with notice given the night before. This guide initially contained daily road logs; however, those logs were cumbersome and took up a lot of space. After the 2019 reconnaissance trip it was decided to drop the road logs and state all stop locations in the appropriate UTM GPS co-ordinates. Most people now own, and can effectively use, handheld GPS units (particularly geologists) and many rental vehicles now have built-in GPS units. 3 �1.3. 2022 Mauna Loa Eruption: With impeccable timing the most recent eruption of Mauna Loa (November 27 to December 10, 2022) began just as final edits were being completed on this guide. This obviously threw a major spanner (wrench to American readers) into the works, particularly the field trip stops planned for the afternoon of Day 7. This required a revision of the Mauna Loa information, including the addition of a short history of the new eruption (with maps and photographs) and the inclusion of a new stop where the road was truncated by the new 2022 flow field (USGS HVO website). None of the originally planned stops were destroyed by the advancing flows; however, the upper 2 stops became inaccessible once the upper half of the road was cut off by 2022 Fissure 3 and 4 flows. Some of the stops planned for the lower half of the road are probably now unnecessary (and probably less interesting) with the presence of a nearby new flow field, but have been left in the guide as contrast to the flows from the new eruption. Therefore, the afternoon of Day 7 was modified to fit the current conditions. The authors cannot in good conscience add other new stops to the guide without visiting the potential sites beforehand, either in person or using Google Earth, other than the point where the easternmost flow of the eruption (a single Fissure 4 flow) cut the Observatory Road. Google Earth will not complete a ground resurvey the Mauna Loa Observatory Road until access is restored to the observatory, which could be several years. However, we do recommend driving the Mauna Loa Observatory Road as far as is allowed, or is possible, to examine the new flow field at your leisure. The existing stops past where the road was truncated have been left in the guide so that they can be visited once the road is reestablished. 1.4. Acknowledgements: There are 2 people who need to be gratefully acknowledged for their contribution to the success of the 2020 field trip and ultimately, this final, albeit revised version of the field guide: • • Peter Hinz, H.B.Sc., P.Geo. (Ontario Geological Survey, retired) who, along with AM (the field trip leader) was co-organizer of the trip. The trip would not have happened, or proceeded in such a relatively flawless manner, without his considerable input, his wonder and fascination with the rocks, and his constant good humour. Dr. Prajukti (Juk) Bhattacharyya (Professor of Geology, University of Wisconsin, Whitewater). Juk had registered for the originally planned August 2019 field trip, but could not participate after it was postponed to February 2020. She graciously agreed to be AM’s extremely overqualified “field assistant” for the 2-week field reconnaissance completed in July and August 2019. The resulting trip would not have been as successful without her help. 1.5. Photo Credits for Field Guide Cover: Clockwise from the upper left: Pu‘u Pua‘i cinder cone and Kīlauea Iki crater, A.D. MacTavish (2008); lava falls entering the Pacific Ocean, south flank Kīlauea, A.D. MacTavish (2011); ‘a‘ā flow crossing Makamae Street, Leilani Estates, May 6, 2018, Lower East Rift Zone, USGS HVO website; steam escaping from mostly buried lava tube within the active Puʻu Ōʻō flow field, A.D. MacTavish (2008). 1.6. Important Notes on Spelling, Location, Distance, and Units of Volume: • • • Since the primary author (AM) is Canadian all spelling within the body of the guide is British/Canadian standard. The exception is the ‘Glossary of Volcanic Terms’ compiled by the guide’s co-author, George Hudak in 2020. He used US standard spelling practices. The GPS datum for all locations stated is UTM Zone Q5, WGS84. Distance and volume units are stated initially in imperial units (US standard) followed by the corresponding metric units (world standard) in brackets; i.e., 100ft (33m). 4 �Iceland MCR Figure 1: Location of the Hawai‘ian Islands with respect to world tectonic plates, the MCR, and Iceland. Modified after Tilling et al., 2010. Kauaʻi Ni‘ihau Oahu Ka‘ula Moloka’i Lānaʻi Maui Kahoʻolawe Hawai‘i Figure 2: Topographic relief and bathymetric map of the nine main Hawai‘ian Islands. Modified from Easton et al., 2003. The inset shows the Kea and Loa volcanic trends first postulated by Dana in the 1840’s (Figure modified from Gazdar, 2003). 5 �6 Waimea 5 Hilo Kailua-Kona 4 3 LO‘IHI 2 1 Figure 3: Volcanic Hazard Map of the Island of Hawai‘i. Each volcano associated with the island is named and numbered (in red) according to age from youngest, Loʻihi (1) to the oldest, Kohala (6). Volcanic hazard severity is shown by the smaller black numbers with the hazard key at the top right of the figure. The estimated location of Loʻihi is shown by the #1. Figure modified after USGS Volcanic Hazard Map from Lava Flows (2010) downloaded from temblor.net website. 6 �Figure 4: The 11 planned daily field trip routes for the Island of Hawai‘i presented within this guide. Modified from Hazlett and Hyndman, 2007 (p.50). 7 �Table 1: Suggested Field Trip Itinerary Day Suggested Accommodation and Location Activity Summary Planned Stops 0 Kailua-Kona; there are a variety of hotels, Night before start of the field trip condos, or B&Bs available None 1 Kīlauea and Volcano; Kīlauea Military Camp (KMC); Volcano House Hotel; or B&Bs in the Village of Volcano Drive to Kīlauea; geology of Mauna Loa southern flank; Kealakekua Bay and pali; Pu‘uhonua o Hōnaunau (Place of Refuge) D1-1 to D1-7 2 Same as Above Punalu‘u Black Sand Beach; Ka Lea (South Point); Pāhala Ash; and Papakōlea green sand beaches D2-8 to D2-9b 3 Same as Above Part 1: Mauna Loa Road and Scenic Lookout; Part 2: Kīlauea Southwest Rift Zone; Ka‘ū Desert trail; Keanakāko‘i Ash; fossilized footprints; Mauna Iki shield Part 1: D3-10 to D3-14 Part 2: D3-15 to D3-25 4 Same as Above Part 1: Kīlauea Caldera; Kīlauea Iki trail; Devastation Trail; Keanakākoʻi Crater; Part 2: Koaʻi Fault Zone; Hilina Pali Road Part 1: D4-26 to D4-32 Part 2: D4-33 to D4-38 5 Same as Above Part 1: Chain of Craters Road; Na Paū Trail; Puʻu Huluhulu; perched lava lake; Mauna Ulu summit and lava channels Part 2: Hōlei Pali; Pu‘u Loa trail and Pu‘u Loa Petroglyphs; Hōlei Sea Arches; Roads End; and Pu’u O’o flow field Part 1: D5-39 to D5-41 Part 2: D5-42 to D5-49 6 Same as Above Part 1: Guided helicopter tour of Kīlauea summit caldera, Southwest and East Rift zones, Mauna Ulu, Pu’u O’o Lower East Rift Zone (LERZ); southern coastline Part 2: June 27th Flow (Pāhoa); New Kaimu Black Sand Beach, MacKenzie State Park; 2018 LERZ eruption flow fields; Rifts 8, 9, 12; 20,21 (Leilani Estates) Part 1: D6-50 Part 2: D6-51 to D6-58 7 Same as Above; alternate hotels and B&Bs in Waimea Part 1: Hilo area; Coconut Island Park; Tsunami Park; Part 1: D7-59 to Rainbow Falls; Kaūmanu Cave D7-61 Part 2: Saddle Road; Puʻu Huluhulu; Mauna Loa Observatory Part 2: D7-62 to Road; 2022 Mauna Loa flows; multicoloured flows, Mauna D7-68 Loa Observatory, & lava flow diversion barriers if Mauna Load Road has re-opened 8 Same as Above; or alternatively hotels and B&Bs in Waimea Mauna Kea Summit Road; Puʻu Kalapeamoa cinder cone; Mauna Kea Visitors Centre; Ellison B. Onizuka Astronomical Complex; Lake Waiau Trail; Mauna Kea Observatories; and summit D8-69 to D8-75 9 Waimea; Kamuela Inn or other area hotels and B&Bs Mamaloa Highway from Hilo to Waipio Valley and Waimea; NE coast Mauna Kea geology; Hawaiʻi Tropical Botanical Gardens; ʻAkaka Falls; Lauapāhoehoe flows; Waipiʻo Valley; Kohala Saddle D9-76 to D9-81 10 Waimea; Kamuela Inn or other area hotels and B&Bs Kohala Volcano; Puʻu Kawaiwai cinder cone; benmoreite D10-82 to D10-91 flows; Pololu Valley; residual boulders; Moʻokini Luakini and King Kamehameha birthplace; Lapakahi State Park; Mugearite flows; pseudodykes 11 Kailua-Kona, Hawai’i; there are many hotels, condos and B&B’s available Last field trip day; Mauna Kea West Rift Zone and cinder cone D11-92 to D11-96 field; Hapuna Beach basaltic ankaramite lave; Hualālai and 1959 Mauna Loa flows contact; Hualālai trachyte flows; Pu’u Waʻawaʻa trachytic cinder cone; Kaʻūpūlehu flow Scenic Lookout; Hualālai Northwest Rift Zone 8 �2. Geology of the Hawaiʻian Islands The Hawaiʻian Islands are the best-known examples of oceanic intraplate volcanoes (those which lie within tectonic plates). Another oceanic example would be the Canary Islands. Not all intraplate volcanoes are oceanic, many are continental such as the volcanoes of the Yellowstone volcanic system in North America and the Quaternary volcanic fields of the Eifel Region in Germany (Schmincke, 2004). About 25% of the world’s volcanoes are of the intraplate variety (Lockwood and Hazlett, 2010) and are thought to be related to ‘hot spots’, also referred to as ‘melting sources’ (Lockwood and Hazlett, 2010) and ‘mantle plumes’ (Morgan, 1971). According to Foulger and Anderson (2006) the Hawaiʻian Islands are part of the ~70 million-year-- old, ~3230mi (5200km) long Hawaiʻian-Emperor Chain comprising more than 100 individual seamounts, atolls, islands, and volcanoes. From the Island of Hawaiʻi the chain extends west-northwestward for ~1680 mi (2700km) to Yuryaku Seamount (near Midway Atoll) as the Hawaiʻian Ridge, where it abruptly bends north-northwestward and extends for another ~1550mi (2500km) to the western end of the Aleutian Trench as the Emperor Seamount Chain (Lockwood and Hazlett, 2010). The oldest island in the Hawaiʻan chain is ~25-million-year-old Kure Atoll and the youngest is the still forming Island of Hawaiʻi (commonly referred to as the Big Island). The distinctive northwest-southeast alignment of the Hawaiʻian chain was known to the early Hawaiians. Their legends clearly reveal that they recognized that the islands were progressively younger from the northwest to the southeast. The first geologic study of the Hawaiian Islands (1840-1841) was directed by James Dwight Dana (Dana, 1890) who deduced that the islands young to the southeast from the differences in their degree of erosion. He also suggested that other island chains in the Pacific showed a similar general decrease in age from northwest to southeast. The 9 main islands of the Hawaiian chain are composed of 15 volcanoes apparently comprising two strands of volcanoes located along distinct but parallel curving pathways. Multiple volcanoes line up to form each strand. Dana (1890) coined the terms Loa and Kea series for the two prominent trends. The Kea trend includes Kīlauea, Mauna Kea, Kohala, East Maui (Haleakalā), and West Maui volcanoes. The Loa trend includes Kamaʻehuakanaloa (Lō‘ihi), Mauna Loa, Hualālai, Mahukona (a submerged volcano and the original volcano comprising the Big Island), Kaho`olawe, Lana`i, West Moloka`i, and possibly Kauaʻi (Figure 2). The pair of volcano trends may exist all the way along the Hawaiian and Emperor chains, though this is less clear amongst the older islands and seamounts (Clague and Dalrymple, 1987). Relative age of an island or atoll can be determined based on its state of growth or erosion (Mattox, 1994). The Hawaiian archipelago rides on the northwesterly moving Pacific Plate. The oldest islands of the archipelago are located far to the northwest of the main Hawaiian Islands. The youngest member of the chain, Kamaʻehuakanaloa (Lō‘ihi), is presently forming as a submarine seamount ~ 19mi (30km) south of the southern coastline of the Island of Hawai‘i. The Big Island is approaching mid-life, while Lō‘ihi is still submerged. By contrast, Kure and Midway atolls are in the final stages of the life cycle of an island. The formation of fringing reefs combined with gradual sinking and erosion of an island causes it to eventually disappear from the surface of the ocean. During this process, reefs grow vertically (upward) and begin to surround the island, eventually becoming separated from the island by a lagoon. If sinking continues, the island disappears and only a circular reef remains - an atoll. Eventually, once the reef sinks below the surface, the original island becomes a submerged, flat-topped guyot. Most of the ancient volcanoes comprising the northwestern Hawaiʻian and Emperor chains are now guyots. 9 �Physically, Hawaiʻian volcanoes exhibit distinct developmental eruptive stages and, from birth to the post-shield/declining stage, their life-span is usually <1,000,000 years. The available references rarely agree on the number of stages and the various lists contain between 4 and 10 stages and substages. The most common listings comprise between 6 and 8 distinct stages. The stages below are a fusion of data from Moore and Clague (1992), Clague and Sherrod (2014), Mattox (1994), and Seach (2022) and consists of 7 stages with the shield stage split into 3 sub-stages: 1. Initial Deep Submarine Pre-Shield Stage: This initial stage of Hawaiʻian volcano growth is characterized by infrequent, small-volume, effusive alkalic eruptions with pillowed lava flows forming an unstable edifice with up to 45o slopes and a summit caldera (see Figure 5-1). Rocktypes comprise basanite, alkalic basalt, and transitional basalt. The steep slopes are due to the alkalic composition of the flows which is more viscous than shield stage tholeiitic basalt. Inherently unstable, steep-sided seamounts composed of immense piles of uncemented, watermelon-shaped lava pillows tend to collapse many times before reaching the ocean surface. The only Hawaiʻian volcano presently at this stage is Lō’ihi which has recently been renamed Kamaʻehuakanaloa . 2. Shield Stage: This is the most voluminous stage of Hawaiian volcanism where more than 95% of the volume of the tholeiitic basalt of a Hawaiian volcano is erupted over a period that may last up to 2 million years. The oceanic crust of the Pacific Plate, unaccustomed to the enormous weight of the volcanoes building atop it, subsides greatly during this stage, as much as 3mm per year. Mauna Loa and Kīlauea are both within this stage of development. Mattox (1994) splits the shield building stage into Stage 2 Submarine, Stage 3 Sea Level (termed emergent above), and Stage 4 Subaerial stages. Clague and Dixon (2000) place Kīlauea into a shield explosive phase (low-elevation subaerial shield) where the volcano exhibits effusive, strombolian, and Hawaiian eruption styles interspersed with phreatomagmatic explosive episodes. Possible examples of this were the phreatomagmatic explosions that accompanied the recent 2018 Kīlauea summit caldera collapse. a. Submarine Shield Sub-stage: Early shield-building eruptions are entirely underwater during this stage (see Figure 5-1) and occurs after the switch from alkalic to tholeiitic volcanism. After the switch in chemistry the rate of growth of the volcano exceeds the rate of subsidence due to the increase in eruptive volume. The bulk of the material produced is tholeiitic in composition with voluminous eruptions of pillowed flows. Volcanic edifice slopes during this stage are in the 10 to 20o range There is no explosive eruptive activity due to water depth. b. Emergent or Sea-level Shield Sub-stage (none at this present time): Eventually the volcano approaches, and then emerges, above sea level with a combination of effusive, Hawaiian, and strombolian volcanism, interspersed with explosive phreatomagmatic eruptions due to the decrease of confining water pressure. At this point an island begins to form (see Figure 5-2). Phreatomagmatic pyroclastic debris (known as hyaloclastite) covers the submarine sub-stage pillowed flows and steepens the flanks of the emerging volcanic island to between 10 and 15o (Figure 5-2). Wikipedia refers to this as the explosive phase which lasts to when the volcano has sufficient mass and height (1000m or 3000ft) and the interaction between sea water and lava finally fades. c. Subaerial Shield Sub-stage: This is the dominant above sea level growth stage and results in the formation of a permanent island by very frequent (almost continuous), voluminous, central- and rift effusive Hawaiian eruptions and the production of ʻaʻā and pāhoehoe flows (see Figure 5-3). Calderas and pit craters form repeatedly, continued submarine and sea level eruptions expand the volcano outwards and with time only a 10 �3. 4. 5. 6. 7. small percentage of tephra erupts from the volcano. The slopes of the volcano during the subaerial shield stage vary between 3 and 10o and are responsible for the characteristic shape of a shield volcano. The cover of fluid, very thin, but dense flows produced during this stage sits precariously upon the weak support of the underlying pillowed flows and hyaloclastites. This weak, unstable base often results in large blocks of the island sliding into the sea and causing enormous, extremely destructive tsunamis with wave heights that can exceed 350m. Post-shield, Capping, or Declining Stage: During this stage eruption volumes decrease and the summit magma chamber solidifies. Transition to post-tholeiitic volcanism is gradual and can last up to 100,000 years with post-shield lavas comprising only a small part of the total erupted volume. The eruption rate usually decreases to zero over a span of between 500,000 and 1 million years. The short-lived, periodic eruptions of this stage tend to cap the volcano with a steep, hummocky carapace of short-length alkaline lava flows and clusters of steep-sided cinder cones (see Figure 5-4, Declining Stage). These lavas typically consist of alkalic basalt, hawaiite, and trachyte and commonly fill and overflow the shield-stage caldera. Explosive eruptions become more common because alkaline flows are more viscous in nature. Erosional Stage: The post-shield stage is followed by a stage where erosion and subsidence are dominant over lava production. During this stage deep canyons and sea cliffs may form along the flanks of the volcano (see Figure 5-5; e.g., Kohala volcano) and the island begins to shrink in size. As the volcanic islands erode and subside, fringing coral reefs grow. Post-Erosional or Rejuvenated (Renewed) Stage: Volcanism of this stage is characterized by strongly alkaline, strombolian to effusive, sometimes phreatomagmatic, volcanism (see Figure 56). Rock-types produced include melilitite, nephelinite, basanite, and alkalic basalt. This stage is characterized by low eruption rates with sporadic activity that may occur over several million years and comprise much less than 1% to the cumulative eruptive volume of a volcano. Most Hawaiian volcanoes do not pass through this stage; however, those that do exhibit periods of erosion that may precede or be interspersed with the eruptions. Lavas commonly erupt through reefs that form offshore as erosion progresses or near the shoreline to produce volcanic maars (e.g., Diamond Head on Oʻahu). There are no Hawaiʻian volcanoes presently within this stage; however, Haleakalā is thought by some to represent an early form of the stage. This stage may be related to remelting of still hot rocks at depth beneath the volcano due to decompression caused by uplift accompanying erosion of the volcanic edifice. Atoll Stage: This stage occurs after erosional processes and subsidence due to the increasing weight of the cooling seafloor eventually lowers the surface of a subaerial volcano to sea level forming flat islands surrounded by coral reefs (see Figure 5-7). Midway and Kure Islands are examples of Hawaiʻian atolls. As the island continues to subside the reefs grow further and further away from the volcanic edifice forming broad lagoon-like planes. Late Seamount or Guyot Stage: Erosional processes and subsidence eventually overtake reef building and the island sinks below the ocean’s surface to form an elevated flat-topped submarine seamount surrounded by, and capped by, dead, submerged coral reef remnants (see Figure 5-8). This type of seamount is referred to as a guyot. 11 �Walker (1990) has described four styles of Hawaiʻian volcanic activity which are summarized below: 1. Hawaiʻian: This is the most common style and is characterized by fountains of gas-rich foamy lava (pumice) spurting from a fissure and being torn into tatters as it flies through the air. Fountain height depends upon lava volumetric discharge rate and gas content and ranges from <5m to >500m (Head and Wilson, 1989). Most of the lava erupted by this style forms pyroclastic spatter deposits such as spatter cones, rings, or ramparts of various heights, diameters, or lengths. If the fountain is concentrated enough some of the lava will flow away. Most of the modern eruptions from Kīlauea are of this style. 2. Strombolian: This style results from both higher gas content and viscosity when compared to Hawaiʻian-style and results in a higher eruptive column and a more highly fragmented lava producing scoria or cinders that cool significantly before landing. This activity forms cinder cones ranging from 50 to 200m in height, with 100 to 200m diameter craters, and extensive cinder deposits located around or downwind of the cones. This style characterizes the declining as well as rejuvenated volcanism stages. Examples include the spectacular-coloured cinder cones within the erosional summit crater of Haleakalā and the slopes and summit of Mauna Kea. 3. Surtseyan: Surtseyan eruptions occur in shallow water (emergent sub-stage) or along sea coasts where large amounts of water interact violently with ascending lava within the vent. This results in a characteristic large, ascending, steam cloud (eruption column) producing highly fragmented lava in the range of sandy and glassy ash. This material accumulates around the vent forming an ash ring which, with time, form indurated tuff rings such as Diamond Head (on Oʻahu) and Kapoho Crater (on the Big Island). 4. Phreatic: Phreatic steam-flash eruptions are very violent and sometimes occur at the summit of Kīlauea when lava drains back into the magma conduit system and interacts with hot groundwater trapped within the conduits. This eruption-style was responsible for the explosive eruptions and spectacular, debris-laden eruption columns that were produced from Kīlauea’s summit caldera between May and August 2018 after the 10yr old lava lake resident within Halemaʻumaʻu crater drained away. After the magma drained away the subsidence into the void left by its disappearance resulted in a summit caldera collapse. When eruption columns collapse, they can generate pyroclastic base surges. It is thought that a base surge of this type killed part of the Hawaiʻian war-party that made the fossilized footprints in freshly fallen Keanakāko‘i Ash erupted from Kīlauea in 1790 (see field trip Stop D3-19, below). Petrochemically Hawaiʻian volcanoes exhibit 4 well-documented eruptive stages (Clague and Dalrymple, 1987; Clague and Dixon, 2000) consisting of: 1. 2. 3. 4. An alkalic submarine pre-shield stage; A main tholeiitic stage that dominates submarine, emergent, and subaerial shield growth; An alkalic post-shield stage; and A strongly alkalic post-erosional or rejuvenated stage. The Island of Hawaiʻi hosts 2 active (Mauna Loa, Kīlauea), 2 dormant (Mauna Kea, Hualālai), and 1 extinct volcano (Kohala). The submerged remains of Mahukona, the precursor volcano to Kohala, are located a short distance northwest of the Big Island (Moore and Clague, 1992). The order of volcano growth that makes up the island and its submarine base are: Mahukona; Kohala; Mauna Kea; Hualālai; Mauna Loa, Kīlauea, and Lō’ihi (see Figure 6). Mahukona started the formation of the Island of Hawaiʻi over 500,000 years ago and after extinction slid into the sea about 300,000 years ago (Moore and Clague, 1992). The youngest active volcano in the chain is the Lō’ihi seamount and is located ~20mi (30km) south of Kilauea at a little over 1000m water depth. The stages of growth, with approximate ages, of the Island of Hawaiʻi are graphically depicted in Figure 6, below. 12 �Figure 5. The 8 volcanic growth stages of the Hawaiʻian islands are shown in this composite diagram modified by Walker (1990) after Stearns (1946), Macdonald et al. (1983), and Peterson and Moore (1987). Not all of the stages represented in the written descriptions above are shown in this figure with the Shield stage beginning during the Submarine stage and including the Emergent stage. The Declining stage in this diagram is equivalent to the Capping or Post-shield stage described in the written text. 13 �Figure 6. This figure graphically shows the growth of the island of Hawaiʻi over the past half million years. The growth is shown at 100,000-year (100ka) intervals with shoreline and volcano boundaries (heavy shaded lines), vigorous subaerial volcanic centers (solid stars), waning subaerial volcanic centers (open stars), and dormant or feebly active subaerial volcanic centers (open circles). Volcano shortform notations are: Kohala (KO); Mahukona (M); Hualālai (H); Mauna Kea (MK); Mauna Loa (ML); and Kilauea (KI). The diagram is modified after Moore and Clague (1992). 14 �2.1. Volcano Descriptions The individual volcano descriptions for the Island of Hawaiʻi are presented in order of decreasing age and, along with the 5 subaerial volcanoes on the island, will also include the submerged, extinct Mahukona volcano, which is thought to be the precursor volcano for the island, and the youngest active volcano, Kamaʻehuakanaloa (Lō’ihi), which is located south of the island. 2.1.1. Mahukona: The extinct submarine volcano Mahukona was discovered in 1987 and is located ~30mi (50km) west of Kohala volcano (see Figures 6 and 7). It was first identified and named by Garcia et al (1990) who later showed (Garcia et al, 2012) that it is a chemically distinct and separate volcano and not part of another volcanic edifice such as Kohala or Hualālai. Mahukona’s existence was predicted by Dana (1890) based on an interpreted gap in the Loa trend of his then hypothesised, subparallel Loa and Kea volcanic trends (see Figure 8) that comprise the Hawaiʻian volcanic chain. Mahukona is interpreted by some as the precursor volcano to the Island of Hawaiʻi and is thought to have begun forming between 1.5 and 1.0 million years ago (Clague and Sherrod, 2014) and to have ceased erupting ~350,000 years ago (Garcia et al, 2012). Earlier researchers (Clague and Moore, 1991; Moore and Clague, 1992) interpreted that the volcano became subaerial ~800,000 years ago and ceased shield-building ~463,000 years ago; whereas, later researchers (Garcia et. al., 2012) believe that it never emerged above sea level (ASL) although they agree on the timing of the growth stages. Mohukona is small compared to most other Hawaiʻian volcanoes with a volume of ~1440mi3 (6000km3), a height above the sea floor of ~9500ft, a cone diameter of ~ 2.5mi (4km), and a summit that is ~3600ft (1100m) below sea level (BSL) (Robinson and Eakins, 2006). Garcia et al (1990 and 2012) have shown that the surface lavas of the cone are weakly alkalic. Figure 7: This bathymetric map shows the approximate location of extinct submarine volcano Mahukona (identified by the yellow arrow) located ~50km west of the extinct, subaerial Kohala volcano. This figure was downloaded from lovebigisland.com who modified it from a publicly available map on the USGS website. The map also shows, in red, the locations of flows erupted on the island between 1800 and 2018. 15 �Figure 8: Map of the Hawaiʻian Islands showing the Loa and Kea trends originally hypothesized by Dana (1890) and now almost universally accepted. The location of Mahukona is highlighted by the box in the centre of the figure, the bold text and the green triangle. Figure from Garcia et al. (1990). 2.1.2. Kohala: Kohala Mountain is described by Robinson (2010) as a 20mi (32km) long, extinct volcano that forms the large, ridge-shaped northern peninsula of the Island of Hawaiʻi. It is the oldest volcano on the island at ~1,000,000 years of age, emerged ASL ~500,000 years ago, entered the post-shield stage and began erupting alkalic lavas ~280,000 years ago, and last erupted ~60,000 years ago (Moore and Clague, 1992). Robinson (2010) further states that Kohala is presently transitioning between the post-shield and erosional stages, and, at its greatest areal extent, was thought to be at least twice its present size with a length of >50km. The youngest eruptive activity on Kohala was contemporaneous with the shieldbuilding stage on Mauna Kea (Easton and Easton, 1995). Kohala is presently 5479ft (1670m) high, has an area of 235mi2 (609km2), a volume of ~3400mi3 (14,172km3), and was at least 5300ft (1615m) higher when it last erupted (Robinson, 2010; Hazlett and Hyndman, 2007). The volcano has a dominant northwest-trending rift zone and a shorter southeasttrending rift zone. It is mainly veneered by alkalic cinder cones and lava flows of the Hāwī Formation which is underlain by the older, late shield stage Pololū Formation (~400,000 years old). The oldest exposed Kohala lavas are the also the oldest on the island and have been dated at ~460,000yrs. The southern flank of Kohala is buried beneath Mauna Kea and Mauna Loa flows (Robinson, 2010). The geology of Kohala is shown in Figure 9. The western flank of Kohala is thought to partially overlie the eastern flank of Mahukona. Erosion has had a dramatic effect on Kohala, particularly its northeastern coastline, where canyons as deep as 2460ft (750m), including the Waipiʻo Valley, have been cut into the mountain. The seven prominent deep canyons incised into the northeast coast cut deeply through the capping alkalic (hawaiitic to trachytic) lavas into the underlying shield-stage tholeiitic lavas and have subsequently been partially infilled by alluvial sediments (Walker, 1990). This section of coast is also characterized by a 16 �relatively straight, fault-controlled line of ~1970ft (600m) high cliffs that form the headwall of the mainly submarine Pololū debris avalanche (Moore et. al., 1989). Hazlett and Hyndman (2007) state that this slide occurred somewhere between 400,000 and 150,00 years ago and resulted in much of the northeastern flank of Kohala sliding into the ocean and travelling for about 80mi (130km) along the seafloor. Robinson (2010) estimates that the slide was ~12.4mi (~20km) wide at the shoreline and extended back to Kohala’s summit. NW Rift Zone 20km SE Rift Zone Figure 9: The Geology of Kohala volcano. Dashed black lines show the locations of the northwest (NW) and the southeast (SE) rift zones. Modified from Aciego et. al. (2010). The numbered circles are where samples were taken for dating purposes, but those results are not included within this guide. 2.1.3. Mauna Kea: The following description is modified from Robinson (2010); Hazlett and Hyndman (2007); and the USGS HVO website. Dormant Mauna Kea volcano is the tallest mountain on earth, if measured from its base at 19,678ft (5998m) below sea level (BSL) to its highest point at 13,796ft (4205m) ASL at the top of the Pu‘u Wekiu cinder cone. Measured in this way Mauna Kea is 33,474ft (10,203m) high. The volume of the volcano is 10,075mi3 (42,000 km3), which is about 55% less than the 22,800mi3 (95,000 km3) of Mauna Loa. Mauna Kea, like dormant Hualālai and extinct Kohala, has evolved past the shield-building stage into the hawaiitic substage of advanced post-shield stage, as indicated by (from Hazlett and Hyndman, 2007 and the HVO website): • • • • Very low eruption rates compared to Kīlauea and Mauna Loa; The absence of a summit caldera and elongated fissure vents that radiate from its summit; Steeper and more irregular topography; i.e., the upper flanks of the volcano are twice as steep as those of Mauna Loa; and The different chemical compositions of the lava, which are now alkalic. In part, these differences are due to a low magma supply rate that produces occasional eruptions from isolated, small batches of magma that rise periodically into the volcano and then solidify without producing continuously active summit reservoirs. The lavas produced are more viscous, with a higher 17 �volatile content, which produces thick flows that steepen the sides of the volcano and explosive eruptions that build large cinder cones. The generalized geology of Mauna Kea is in Figures 10 and 11. Mauna Kea is presently dormant and last erupted ~4500 years ago. It is likely to erupt again since the volcano’s quiescent periods are long compared to: the more active late-stage Hualālai, which last erupted in the late 1700’s and early 1800’s; the much more active Mauna Loa which erupts every few years to tens of years; and the extremely active Kilauea which erupts every few years. The oldest dated rocks exposed on the flanks of Mauna Kea are 237,000±31,000 years BP. The volcano is estimated be ~1,000,000 years old, with its shield stage lasting ~800,000 years, and it is estimated to have begun the post-shield stage somewhere between 200,000 and 250,000 years ago. Mauna Kea and Mauna Loa are often snow-covered between November and March and on 3 occasions during the Pleistocene Mauna Kea hosted permanent ice-caps and summit glaciers. The ice cap on Mauna Kea reached to below the 11,000ft ASL level (see Figure 12). The preserved glacial moraines and glacial outwash formed on 2 occasions from between 70,000 years ago and 40,000 to 13,000 years ago. Mauna Loa also probably hosted and ice-cap, but any evidence has long been buried by younger lava flows. The Mauna Kea summit road was closed in early August 2019 due to political and indigenous Hawaiʻian protestors against the building of another telescope at the summit astronomical observatory. Negotiations with various state, federal, and observatory officials lead to the protestors temporarily removing the blockade just before the February 2020 Field Trip and allowed the field trip to drive the summit access road and reach the summit of the volcano. As of December, 2022 access to the summit of the mountain is again allowed. Figure 10: This geological map of Mauna Kea shows the generalized surface distribution of the Hamakua Volcanics. The younger Lauapāhoehoe Volcanics are inferred to overlie a vast area of Hamakua Volcanics on the upper flanks and summit. Map downloaded from the HVO website. 18 �Figure 11: This geology map of Mauna Kea shows the generalized surface distribution of the lava flows, cinder cones, and glacial deposits of the Lauapāhoehoe Volcanics. Map downloaded from the HVO website. Figure 12: Map of Mauna Kea showing the extent of the summit icecap during the Pleistocene. From Merguerian and Okulewicz (2007, p.75) who took the figure from Macdonald et al (1983, Figure 13.3, p.257). 19 �2.1.4. Hualālai: Hualālai is an 8300ft (2530m) high, post-shield stage volcano with a volume of 2,975 mi3 (12,400 km3) and an area of 290mi2 (751 km2) (Robinson, 2010). The mountain has well-defined northwest (NW) and southeast (SE) rift zones and a poorly-developed north rift zone (see Figure 13). Hualālai is thought to have emerged above sea level on the southwest flank of Mauna Kea ~300,000 years ago. Shield building tapered off about 130,000 years ago (Moore and Clague, 1992). The oldest exposed surface rocks on Hualālai are dated at ~128,000 years old with geological mapping indicating that 95% of its surface is covered by flows that are <10,000 years old. Of those flows most are <5000 years old. This suggests that eruptions can at times be quite common and that the mountain could potentially pose a considerable hazard to the large number of people presently living on or around it (i.e., Kailua-Kona and surroundings). Estimations on the growth of Hualālai are uncertain due to the burial of parts of the volcano by Mauna Loa lavas and because of a gravitational failure (slump) of the southwest flank of the mountain that produced the North Kona Landslide before 130,000 years ago (Moore and Clague, 1992). This slide produced a >40km wide up to 4km high scarp that extended into the shield to the northwest rift zone. (Moore and Clague (1992). Hualālai is the third youngest volcano on the island and has erupted from at least 7 different vents during the last 2100 years (Walker, 1990). Sometime between 1200 and 1400 AD (the timing is uncertain) the large Wahapele eruption was active for a few weeks or months, possibly as long as several years. Flows from the Wahapele eruption reached the sea about 10mi (16km) south of KailuaKona (USGS HVO website) in the Keauhou Bay area. The mountain’s last eruptions were in the late 1700’s (the dates are uncertain) and between 1800 and 1801, where flows issued from 6 vents. Two flows erupted during 1800 and 1801 reached the sea (USGS HVO website). Flows active in 1801 form the southernmost expression of this eruption and underlie Keahole-Kona International Airport located 7mi (11km) north of the City of Kailua-Kona (USGS HVO website). The flows comprising the 1800 to 1801 eruption are viewed at Stop D11-96. The 1800 Hualālai flows contain large numbers of dunitic and gabbroic xenoliths/blocks that are thought to be sourced from intrusions that core the volcano (Walker, 1990, Kirby and Green, 1980; and Jackson et al., 1981) Hualālai presently erupts viscous alkaline lavas, including alkalic olivine basalt and trachyte, with the entire subaerial surface of the volcano covered with flows of alkalic composition. In a similar manner as Mauna Kea, the viscosity of these alkaline lavas, particularly trachyte, has steepened the volcano’s flanks and covered its summit and northwest and southeast rift zones with cinder cones and pit craters. Trachyte forms the large Puʻu Waʻawaʻa cone and the 330 to 660ft (100 to 200m) thick flow that emerged from it (Walker, 1990). The viscous, alkalic eruptions during the Wahapele period included a powerfully explosive phase that spread pyroclastic material over a large area. This again underscores the potential danger that Hualālai volcano poses to those who live on, or near, the mountain. 20 �Kiholo Bay 1800 Flows 2022 NERZ Flows NW Rift Zone KeaholeKona Int’l Airport 1800 Flows Summit SE Rift Zone KailuaKona Keauhou Bay Figure 13: Modified USGS relief map showing Hualālai volcano and the lava flows extruded from a series of vents over that last 1000 years. The Wahapele flows (pink) are thought to have been erupted sometime between 1200 and 1400AD, but did not come from a vent located on either of the rift zones. The salmon-coloured flows were erupted during the late 1700’s and 1800 to 1801. The NW and SE rift zones are defined by black dashed lines. The inset map shows the surface expression of the 5 volcanoes comprising the island and the lava flows erupted from Mauna Loa and Kīlauea since 1800AD, including the rough location of the 2022 Mauna Loa flows (green). Figure was downloaded from the USGS HVO website and then modified. 2.1.5. Mauna Loa: The following description is modified from information provided by Robinson, 2010; Hazlett and Hyndman, 2007; and the USGS HVO website. Mauna Loa (Hawaiʻian for ‘Long Mountain’) is the largest and most massive volcano on earth, but not the highest, which is Mauna Kea (see Day 8, below). It dominates just over half of the Big Island; has an area of 2035mi2 (5271km2); reaches to a height of 13,678ft (4169m) ASL; has a volume of 19,000mi3 (79,195km3); and a weight of 207 x 106 tons (188 x 106 tonnes). This immense weight greatly depresses the sea floor around the island and results in a base to summit elevation of ~31,000ft (9449m). Mauna Loa first erupted on the seafloor along the flanks of either Hualālai or Mauna Kea between ~1,000,000 and 600,000 years ago and emerged above sea level ~300,000 years ago. The oldest exposed flows on the volcano are between 100,000 and 200,000 years old with ~98% of those exposed lavas <10,000 years old. The volcano is less active than Kīlauea, but it characteristically produces greater lava volumes over shorter periods of time because it is fed from a much larger magma chamber than the one beneath Kīlauea. The Mauna Loa summit hosts the elongated, northeast-southwest-orientated Mokuʻāweoweo caldera, with dimensions of 2.8 by 1.6mi (4.5 by 2.5km), including the summit collapse pits. The caldera floor, is 21 �~180 m (590 ft) below the volcano’s summit, which is located on the western rim of the caldera. Three rift zones radiate down the flanks of the volcano from the caldera. The southwest rift zone enters the ocean west of Ka Lae (South Point); the northeast rift zone arcs down the slope of the mountain ending in rain forest 15mi (24km) south of Hilo; and the third, diffuse northwest rift zone traverses the flank of the volcano and extends to the foot of Mauna Kea located 20mi (32km) north of Mauna Loa. Geological mapping and radio-carbon dating of flows produced during the last 4000 years shows that vent locations have cycled twice between summit-dominant and rift-dominant vents. Rift-dominant eruptions have dominated the last 700 to 800 years, whereas the last summit-dominant stage occurred between ~200AD and 1200AD, a period of almost 1000years. HVO geologists suggest that the decline of summit-dominant eruptions and the increase in rift-dominant activity was related to the summit collapse that led to the formation of Moku‘āweoweo Caldera. Once the caldera formed the lava flows erupted within it were trapped and were then usually unable to overflow the caldera rim. There are several possible causes for the transition from summit-dominant to rift zone-dominant eruptions such as: significant changes in the magma supply or reservoir plumbing system of the volcano leading to the formation of the summit caldera; the advent of explosive activity; and/or flank instability. HVO scientists refer to five broad areas on the volcano where eruptions occur: the summit area, which is that part of the volcano above 12,000 ft (3,660m) ASL and includes Moku‘āweoweo Caldera and the uppermost parts of the northeast and southwest rift zones; the northeast and southwest rift zones below the summit area; and the southeast, north, and west flanks (considered as one). At least 33 radial vents have been mapped in the north and west sectors signifying that lava can erupt from these sectors in addition to the rift zones and summit area. The volcano is in the closing part of its shield stage and has erupted 34 times since 1843, making it one of the earth’s most active volcanoes (see Figure 14). Mauna Loa’s large, voluminous, basalt flow eruptions have reached the ocean 8 times since 1868. Previous to November 27, 2022 the next to last eruption began on March 24, 1984 and continued until April 15, 1984 (23 days). During that short period of time lava flows from the eruption approached to within 4mi (6.4km) of the City of Hilo. Hazlett and Hyndman (2007) state that Mauna Loa may be the most threatening Hawaiʻian volcano, mainly due to the volume and length of erupted flows which have threatened Hilo on 7 occasions since its founding. Lava flow hazard zones are shown in the lower right of Figure 14. 2.1.5.1. 2022 Mauna Loa Eruption: The most recent eruption of Mauna Loa began at 1130PM on Sunday November 27, 2022 within Mokuʻāweoweo Caldera where lava initially issued from fissures quickly covered much of the caldera floor. By the morning of the 28th lava was issuing from several active fissures to the southwest of the caldera and within the upper northeast rift zone (NERZ). By mid-day of the 28th activity to the southwest and within the caldera had ceased and lava fountains up to 200ft (60m) high were observed issuing from Fissures 3 and 4 (F3 and F4, see Figure 15) located downslope from the caldera within the upper part of the NERZ. F3, at an elevation of ~11,500ft (3510m) ASL, quickly became the dominant vent source. Flows from F3 first cut the Mauna Loa Weather Observatory Road, located ~3.9mi (6.3km) downslope from the vent, in 2 places on the 28th. F4, located 1mi (1.6km) northeast of F3, continued to extrude flows until December 2nd and associated flows again cut the Mauna Loa Observatory Road on December 1st. By December 5th the access road had been further cut multiple times over a wide area by flows from F3, after it had become the only active vent. By December 5, 2022 flows had advanced a further 6.2mi (10.0km) downslope to reach the relatively flat plateau (the Saddle) that occupies the area between Mauna Loa, Mauna Kea, and Hualālai at ~ 6500 (1829m) ASL. The flatter slopes of the saddle caused the flows to slow, spread out, inflate, and split into several separate sub-flows, some of which were channelized, and small overflows from main channels were common. Lava fountains from the vent 22 �attained heights of >330ft or 100m (higher than at the beginning of the eruption) and were feeding a >10.5mi (16.70km) long lava flow. Eruptive activity at the F3 vent began to decrease over the night of December 7 and 8. By the morning of the 8th flow volume was much reduced causing the flow front on the saddle to stall ~1.7mi (2.8km) south of, and before reaching, the Daniel K. Inouye Highway (Saddle Road). The height of the F3 spatter cone, when eruptive volume began to decrease on the 8th, was 98ft (29.9m). Eruptive activity continued to wane with lava volumes decreasing such that by the morning of the 10th only a few weak flows were active. These flows were fed by a lava lake within the vent rather than the fountains which typified the eruption to this point. By the morning of the 11th all activity on the flow field appeared to have ceased; however, the main flow still glowed and inched forward on occasion as it settled. The F3 vent was still incandescent at night. The eruption was deemed over late on December 10, 2022 (USGS HVO website) after being active for 12 days. The location of flows and fissures associated with the 2022 eruption are summarized on Figure 15. Waimea Hilo KailuaKona Figure 14: Map showing the extent of historic Mauna Loa lava flows; hazard zones are designated by the USGS. The 2022 flows are located about where the 1843 flows are shown. Map downloaded from the HVO website. 23 �Figure 15: Map showing the lava flows erupted from Mauna Loa’s summit caldera and the Upper Northeast Rift Zone during the November 27 and December 10, 2022 eruption. Map downloaded from the USGS HVO website. 24 �2.1.6. Kīlauea (description after Hazlett, 2014; the USGS HVO Website, 2022): Kīlauea is the most active volcano on earth and has the appearance of a bulge on the southeastern flank of Mauna Loa. For a long time it was thought to be a satellite of Mauna Loa; however, research shows that Kīlauea has a distinct and deep magma-plumbing system that extends into the earth for >60km. Eruptive activity for Kīlauea since 1790 is shown in Figure 16. The first alkali-basalt lava flows of Kīlauea’s submarine pre-shield stage erupted onto the seafloor on the southeastern flank of Mauna Loa between 210,000 and 280,000 years ago. It transitioned to the submarine shield-building stage about 155,000 years ago and emerged above sea level between 50,000 and 100,000 years ago. The oldest exposed surface lavas are the Hilina basalt formation which is exposed along various Hilina fault scarps on Kīlauea's central south flank. These flows are the oldest found above sea level and erupted between 50,000 and 70,000 years ago. >90% of the remaining surface flows are <1000 years old, and 70% of those are <600 years old. The summit of the volcano is located at 4080ft (1240m) ASL. Research and mapping clearly show that Kīlauea exhibits cycles of explosive and non-explosive (effusive) eruptions that individually last for prolonged periods of time. This pattern of activity has persisted for at least the last 2500 years and possibly longer, but since the surface of Kīlauea is very young it is difficult to accurately determine the eruption record earlier than that time. The known eruption record shows that effusive (non-explosive eruptions) were the norm up to ~2200 years ago. At this time the Powers Caldera, which is the precursor to the present caldera, formed by a collapse of the crater floor to a depth of at least 2030ft (620m) where magma and external water interacted to trigger powerful phreatomagmatic eruptions. Tephra from the many explosive (pyroclastic) eruptions that occurred over the next 1,200 years produced the Uwekahuna tephra. The most powerful known explosive eruption from Kīlauea occurred between 850 and 950CE and sent golf ball-sized rocks as far away as the southern coast of the island, a distance of 11mi (18km). Effusive activity began again ~1000 years ago and completely filled the summit caldera to where it overflowed to form the Observatory Shield. Eruptions were also frequent along the east and southwest rift zones. Observatory Shield construction ended ~1400CE when activity migrated to the east and over the next 60years produced the longest-lasting flow ever witnessed in Hawaiʻi. This Ailāʻau flow covered much of Kīlauea from the summit to the coast on the north side of the East Rift Zone. The Ailāʻau eruption ended ~1470CE and the collapse after the withdrawal of lava from the summit formed the present-day Kīlauea Caldera. The Keanakāko‘i explosive eruption period began ~1500CE when the caldera floor dropped to a depth of ~1970ft (600m) and had a diameter of 2.2mi (3.5km) by 1.9mi (3 km). The Keanakāko‘i period ended in ~1800CE after at least 4 strong explosive eruptions over a 300yr period ejected ash over a broad area east of the volcano and deposited the 35ft (11m) thick Keanakāko‘i tephra (ash bed). In 1790, near the end of the period, a series of strong explosive eruptions produced several pyroclastic base surges, which are a type of turbulent, very hot (>100oC), fast-moving, low-density pyroclastic density currents that can sweep over ridges, hills, and other topographic boundaries and are almost impossible to escape, particularly on foot. These surges seared down the west side of the summit area killing several hundred, possibly several thousand, indigenous Hawaiʻians (see The Footprints Trail, Field Trip Day 2). This is the deadliest known eruption of a volcano on U.S. soil. The control on this explosive-effusive cycle may be magma supply. High volumes of magma will allow the caldera to fill, as well as feed large amounts of magma to summit lava flows and rift zone vents. When the magma supply drops the caldera will collapse. If the floor of the crater drops sufficiently to approach or cross the water table then that water interacts with the magma in the vent to produce phreatomagmatic (magma-steam) explosions. When the magma supply again increases to allow 25 �effusive eruptions to dominate then the cycle begins again. Research strongly suggests that a caldera is necessary for prolonged periods of explosive summit eruptions and it is estimated that a deep caldera has existed at the summit for ~60% of the last 2,500 years. Before the end of April 2018, the summit caldera (see Figure 18A, B, C, and D) was 541ft (165m) deep with an outermost diameter of 3.7mi (5.95km) and an elongate, north-northeast-trending main depression measuring 3.1mi by 1.9mi (4.99km by 3.06km). As mentioned above this caldera largely formed between 400 and 500 years ago with lesser collapses leading up to ~1790CE when pyroclastic eruptions deposited the thick, complex Keanakāko’i Ash over a wide area. This ash is best observed along the Footprints Trail (Day 2). Also, before the end of April 2018, the Kīlauea summit caldera hosted the 0.62mi (1km) diameter Halema‘uma‘u Crater (see Figures 17 and 19) which represented the top of a low-lying lava shield within the greater caldera. Halema‘uma‘u Crater began forming in July 1894 when the original lava shield collapsed. The crater reached its pre-2018 size and shape in 1924 when a long-lived lava lake (1905 to 1924) contained within an earlier, smaller version of the crater drained away. Over a 9-day period between April 29, 1924 and May 7, 1924, a series of collapses and steam-blast eruptions roughly doubled the size of the crater to its pre-May 2018 dimensions. The most recent pre-2018 effusive summit eruption, began on March 19, 2008 with an explosion blew a narrow subvertical vent into the bottom of Halema‘uma‘u Crater near its southern margin. This vent initially emitted sulphur-dioxiderich (SO2) gasses and eventually hosted a lava lake that occasionally overflowed onto the crater floor. A protracted Central East Rift Zone (CERZ) eruption (see Figure 20A, B, C, D) began on January 3, 1983 as a series of localized fissure eruptions and by June 1983 had focussed on the Pu‘u Ō‘ō vent. The eruption focussed there for ~3 years until it shifted 1.8mi (3km) east down-rift to the Kupaianaha vent. The next 5½ years were a period of almost continuous and destructive eruption from the Kupaianaha vent that destroyed many homes and the communities of Kapa‘ahu and Kalapana. In February 1992 eruptive activity shifted west up-rift via a series of fissure eruptions that culminated in a fissure eruption on the west flank of Pu‘u Ō‘ō. The eruption focus stayed near, Pu‘u Ō‘ō for the next 26 years (see Figure 21). The only community threatened during this time was the town of Pahoa, located ~12mi (20km) east of Pu‘u Ō‘ō on Highway 130, between July and December 2014 (see Figure 22). Eruption from Pu‘u Ō‘ō and vicinity was continuous until April 30, 2018 when, after a series of earthquakes, the magma moved from Pu‘u Ō‘ō and the eruptive vent collapsed. The magma was observed by seismic monitoring to move >20km east to the Lower East Rift Zone (LERZ). Vigorous eruption began on May 3rd via a series of 24 fissures along a 14km segment of the LERZ and continued until August 9th. Also, in early May the 10yr old summit lava lake within Halema‘uma‘u Crater began to drain away and by May 10th had disappeared from view. Once the supporting magma beneath Halema‘uma‘u disappeared the area around the crater began to dramatically subside. This caldera collapse was accompanied by numerous steam-generated pyroclastic eruptions and Halema‘uma‘u Crater was eventually replaced by a much larger and deeper crater 1.7mi (3km) in length, 0.93mi (1.5km) in width, and >1640ft (500m) in depth. A small lake was present at the bottom of the crater in February 2020 (see Figure D6-3, right). Since the February 2020 field trip and the final edit of this Guide there were 3 summit eruptions within the new Halema‘uma‘u Crater. The deep, cone-shape crater that formed in mid-2018 has been partially infilled by lava lakes formed during the 3 eruptions. The first eruption, between December 20, 2020 and May 26, 2021, partially infilled the crater by 732ft (223m) bringing the base of the crater up to 2431ft (743m) ASL. The second eruption began September 29, 2021 and ended on December 9, 2022. The third eruption commenced on January 5, 2023 and was ongoing by the completion of this guide in February 2023. 26 �Figure 16: Map of Kīlauea volcano showing subaerial extent of historic lava flows extruded between 1790 and 2018; hazard zones are designated by the USGS. Map downloaded from the HVO website. Figure 17: Incandescent volcanic ash and lava fragments are blasted from the Halemaʻumaʻu Crater vent at Kīlauea’s summit during an explosive eruption on October 12, 2008 (left). The volcanic-gas plume emitted from that vent a month later in November 2008 on the right. Photographs by Janet L. Babb from Tilling et. al. (2011), USGS General Information Product 135 that were downloaded from the USGS HVO website. 27 �A. B. C. D . Figures 18A, B, C, and D: Kīlauea Caldera as it appeared before May, 2018: A. Halema‘uma‘u Crater viewed from the now closed Jagger Museum observation deck with the vase plume from the lava lake clearly visible; B. Northwest caldera rim from the caldera floor below the Volcano House Hotel; C. Northeast caldera rim from the caldera floor; and D. Eastern Caldera rim from northern caldera rim. Photos by A.D. MacTavish (2009). A. Figures 19: These 2 photos are close-ups of Halema‘uma‘u Crater and the lava lake that was resident within the crater between 2008 to 2018: A. Aerial view of Halema‘uma‘u Crater with lava lake/gas vent (2010); B. Halema‘uma‘u lava lake at night (2012). Photos downloaded from the HVO Website. 28 �A. A. B. C. D. Figures 20A, B, C, and D: Kīlauea’s Central East Rift Zone eruption began in January 1983 and ended in May 2018. Photos of the earlier stages of this eruption can be seen here: A. Lava fountain at Pu‘u Ō‘ō (1983); B. A‘ā flows from Pu‘u Ō‘ō vent passing through the Royal Gardens Subdivision located about 4km southeast of the vent (1983); C. Kupaianaha vent and perched lava pond with the Pu‘u ‘Ō‘ō cone in background (1986); and D. Flows from the Kupaianaha Vent destroying a house in Kalapana (1990). All photos downloaded from the HVO Website. Figures 21: These 2 photos show eruptive activity associated with Pu‘u Ō‘ō: The left photo shows a perched lava channel with Pu‘u Ō‘ō in the background (2007); the right photos show the Kamomoa lava fountains with Pu‘u Ō‘ō in the background (2011). Photos downloaded from the HVO Website. 29 �Figure 22: These photos show the Pu‘u Ō‘ō ‘June 27th Flow’ threatening the town of Pahoa in 2014: Lava flows from Pu‘u Ō‘ō are approaching Pahoa (left); a pāhoehoe lava flow advancing west of Pahoa between the town and the Waste Transfer Station on Apa‘a Street (right). Photos downloaded from the USGS HVO Website. A. 2.1.6.1. 2018 Kīlauea LERZ Eruption and Halema‘uma‘u Summit Caldera Collapse The description within this sub-section was modified and summarized from publicly available data on the USGS HVO website. During late April 2018 the focus of the 35-year Kīlauea eruption shifted from Pu‘u Ō‘ō on the Central East Rift Zone (CERZ) and the volcano’s summit (Halemaʻumaʻu lava lake) to the Lower East Rift Zone (LERZ) after the collapse of the long-term Pu‘u Ō‘ō crater on April 30, 2018. On May 2nd this collapse was followed by a series of strong earthquakes and the opening of the first ground cracks along the LERZ and the dropping of the level of the 10yr old Halemaʻumaʻu summit lava lake. The LERZ fissure eruption began on May 3rd with one vent (Fissure 1) opening in the area of Mohala and Leilani Streets in the Leilani Estates subdivision. By the next day there were 6 open fissures, with lava issuing from Fissure 2 and a magnitude 6.9 earthquake on south flank of Kīlauea. On May 8th there was a pause in eruptive activity after 15 new LERZ fissures opened. On May 10th the Halemaʻumaʻu lava lake had disappeared from view and on May 11th Hawaiʻi Volcanoes National Park was closed to the public. After 4 days of inactivity Fissure 16 opened on May 12th and by May 14th Fissures 17, 18, and 19 were open with flows issuing from Fissures 16 and 17. On May 15th a 12,000ft (3660m) ash plume issued from Halemaʻumaʻu after a rock fall and subsequent explosions. On the same day Fissure 20 opened in the Lanipuna Gardens Subdivision, located <0.6mi (1km) SE of Leilani Estates, and a slow, narrow flow from Fissure 17 was creeping toward the ocean. Explosive events at the Kīlauea summit commenced on May 16th with ash clouds rising up to 30,000ft ASL. The Hawaiʻi Volcano Observatory (HVO) building was evacuated and subsequently permanently closed, and cracks were observed on Highway 11, a short distance northeast of the Park Entrance (between mile markers 28 and 29). Explosive events of various sizes continued at the summit until early August with caldera subsidence beneath Halemaʻumaʻu beginning on May 25th. By May 19th fountaining from Fissures 16 through 20 had merged into the single, continuous fissure referred to as Fissure 20. Lava from this fissure entered the ocean near the MacKenzie State Park Recreation Area the same day (this ocean entry lasted about 10 days). By May 28th there were 24 LERZ fissures with at lava erupting from least 10 of the fissures at the same time. A large, over 100ft (30m) high, spatter rampart had been built around Fissure 7 by lava fountains 30 �reaching up to 200ft (45 to 60m) high, that fed a perched, 20 to 40ft (6 to 12m) thick, pāhoehoe flow. On the same day a magnitude 4.1 earthquake occurred on the Koa`e fault zone south of the caldera. Caldera down-drop accelerated with the onset of near-daily summit collapse events with each event releasing energy equivalent to a Magnitude 5.0 earthquake. By May 31st impressive fountaining from Fissure 8 had formed a broad, levéed, channelized flow. Fissure 8 quickly became the most active and longest acting fissure of the 2018 eruption. On June 2nd the channelized Fissure 8 flow crossed Highway 137, near the junction with Highway 132, advanced into Kapoho Crater, and entered and completely filled up Green Lake within the crater. This voluminous and fast-moving flow entered Kapoho Bay on the Pacific Ocean late the next day after travelling over 8mi (13km) from Fissure 8. The flow immediately began building a lava delta into Kapoho Bay on a 500ft (150m) wide flow front and, by the following morning, the flow had completely infilled the bay. The width of this flow front had expanded to 3.7mi (6km) by July12th. By July 18th an increase in lava supply from Fissure 8 produced several overflows that destroyed more homes. Explosions were evident near the main ocean entry, which had shifted to near Ahalanui Beach Park. The margin of the flow at the ocean entry continued to extend southwards and advanced to within <575ft (175m) of the Isaac Hale Park boat ramp. The eruption of lava from Fissure 8 continued vigorously until August 4th when the eruption rates began to decrease. Summit deflation stopped after a single collapse event earlier in the day. By August 7 the only eruptive activity at Fissure 8 consisted of a small active lava lake within the cone. The lake’s surface was located between 15 and 30ft (5 to 10m) below the spillway entrance of the cone. Small, active ooze-outs near the coast on the Kapoho Bay and Ahalanui lava lobes were greatly diminished and active lava remained close to the Pohoiki boat ramp at Isaac Hale Park, but had not advanced significantly toward it. Deformation at the summit had virtually stopped. From August 9th to 13th the activity and lava output from Fissure 8 remained low with no signs of reactivation or new subsurface intrusion. Up-rift Fissures 9, 10, and 24 and down-rift Fissures 3, 7, 13, and 23 continued to steam. The crusted Fissure 8 lava pond was deep within the cone and on the 10th was about 130ft (40m) below the rim of the cinder cone. On the 11th there were 2 lava ponds – one active and the other stagnant and crusted over. On the 12th the only molten lava visible was oozing into the ocean between the Kapoho Bay and Ahalanui areas; the summit remained quiet except for a few small rock falls. As of August 16th, Kīlauea volcano had remained quiet for over a week with no collapse events at the summit and, other than a crusted-over lava pond deep within the Fissure 8 cone and a few scattered ocean entries, there was no lava flowing in the Lower East Rift Zone. On September 22, 2018 Hawaiʻi Volcanoes National Park partially reopened and the eruption appeared to be over. Lava from the 2018 LERZ eruption covered >5914 acres and destroyed at least 533 homes. The lava delta built by the Fissure 8 Flow within Kopoho Bay covered an area of over 380 acres and extended out over 0.5mi (800m) from the former shoreline. Figures 23 through 25 and Figure 30 highlight the 2018 Kīlauea summit caldera collapse beneath Halemaʻumaʻu Crater. Figure 26 shows the 2023 lava lake that now occupies, and has infilled, much of the 2018 crater as the result of the 3 summit eruptions that have occurred between 2019 and February 2023. Figures 27 and 28 highlight various aspects of the 2018 Lower East Rift Zone eruption. 31 �Figure 23: Airborne radar maps (upper 2 panels) of Halema‘uma‘u Crater and Kīlauea Caldera taken in June 2009 and August 2018 show the changes in the caldera floor before and after the withdrawal of the lava lake that occupied the crater between 2008 and 2018. The lower panel shows a vertical cross-section through the crater showing the over 500m subsidence of the floor of the crater between May and August 2018 (USGS HVO Website). HVO Volcano House Hotel Halema‘uma‘u Crater Crater Rim Drive Figure 24: Satellite image of Kīlauea Caldera and Halema‘uma‘u Crater in January 2003. Photo from USGS HVO Website. 32 �Volcano House Hotel HVO Old Halema‘uma‘u Crater Outline Crater Rim Drive Figure 25: Satellite image of Kīlauea Caldera and Halema‘uma‘u Crater after the collapse of Halema‘uma‘u taken on August 11, 2018. Photo from USGS HVO Website. Volcano House Hotel Kīlauea Iki Crater Crater Rim Drive Figure 26: Aerial photograph of a greatly modified Kīlauea caldera showing the changes, due to the collapse of the caldera floor beneath Halema‘uma‘u Crater, that took place between May and August 2018. The new crater is ~2.5km by 2km, and 350 to 400m deep. Photo by A.D. MacTavish, August 4, 2019. 33 �Figure 27: The post-2018 Halema‘uma‘u Crater at 645AM January 6, 2023 showing the active lava lake from the Kīlauea summit eruption that began on January 5, 2023. Sunrise-lit Mauna Loa is in the background to the west. Most of the crater formed in 2018 is now infilled with lava. Photo taken from the USGS HVO website. Figure 28: Map showing the flows (represented by the salmon colour) produced by the 2018 Lower East Rift Zone eruption. Figure taken from the USGS HVO website. 34 �A. B. C. D. E. F. Figure 29: 2018 Lower East Rift Zone eruption photos (all photos downloaded from the USGS HVO website): A. Leilani Estates, fountaining from a new fissure, May 4, 2018; B. ‘A‘ā flow crossing Makamae Street, Leilani Estates, May 6, 2018; C. Fissure 20 channelized flow, May 19, 2018; D. Channelized flows entering the Pacific Ocean, evening May 23, 2018; E. Breached spatter cone and fountaining at Fissure 8, June 5, 2018; F. Flows entering Kapoho Bay, June 5, 2018; the bay has been completely infilled and the lava delta has extended ~300m outward from the former entrance to the bay. 35 �A. B. C. D. E. Figure 30: Kīlauea 2018 Summit Events: A. Summit eruption cloud, May 15, 2018; B. ‘Summit eruption plume, May 23, 2018; C. Halemaʻumaʻu Crater subsidence, June 5, 2018; D. The new Halemaʻumaʻu Crater, March 6, 2019; E. New Halemaʻumaʻu Crater taken from a helicopter overflight on August 4, 2019; the view is to the south. Photos A, B, C, and D from the USGS HVO website; Photo E by A.D. MacTavish (2019). 36 �2.1.7. Loʻihi (Kamaʻehuakanaloa): As mentioned at the beginning of Section 2, above, Lō’ihi is the youngest active volcano associated with the Island of Hawaiʻi and has recently been renamed. According to the USGS HVO website the name Lō’ihi was introduced in 1955 to describe the elongated shape of the seamount. More recently, Hawaiʻian scholars have found that stories of “Kama‘ehu” , the red island child of Haumea (earth) and ‘Kanaloa’ (sea) that rises from the deep in the ocean floor, may also be a reference to the submarine volcano, hence the proposal and acceptance of the name ‘Kamaʻehuakanaloa’, which means ‘red island child of the earth and sea’. The name ‘Kamaʻehuakanaloa’ was adopted by the Hawaiʻi Board of Geographic Names in 2021 (USGS HVO website) and is not yet in common use, or even known about other than by researchers, and will not be used further in this guide. Lō’ihi, which means ‘long’ in Hawaiʻian, is considerably easier to remember, to spell, and to pronounce (by the author’s at least). Lō‘ihi is presently forming as a submarine seamount located ~ 19mi (31km) south of the southern coastline of the Island of Hawai’i (see Figure 30). It has a volume of 407mi3 (1,700km3) with its summit at ~10,100ft (3078m) above the abyssal ocean floor at a water depth of 3235ft (986m). The seamount comprising Lō’ihi was long thought to be a young submarine volcano, but that was not confirmed until a series of sub-sea earthquake swarms detected by the HVO’s seismic network in 1971, 1972, and 1975 quickly lead to the recovery of fresh, glassy lava samples and the identification of active hydrothermal vents and deposits near the summit (Moore et al., 1979, Frey and Clague, 1983, Garcia et al., 1989; Clague et al., 2019). According the USGS HVO website the summit of Lō’ihi is nearly flat and marked by a caldera-like depression that is ~1.7mi (2.8km) wide and ~2.3mi (3.7km) long. The southern part of the caldera is host to three collapse pits or craters. The most recent, Pele's Pit, formed during an intense 1996 seismic swarm that was subsequent to an eruption from a shallow magma chamber (Clague et al., 2019). This new crater is about ~1,970 ft (600 m) in diameter with its base ~985ft (300m) below the previous surface. Figure 31 shows a Clague et al. (2019) interpretation of the summit caldera complex and its various pit craters. The volcano has grown from eruptions along distinct northwest and southeast rift zones that extend out from the caldera. Clague et al (2019) state that the asymmetric, east-west-dipping slopes of the rift zones and the summit platform suggest that the flanks of Lō’ihi have been modified by landslides. They also state that the west flank has been modified by 2 landslides and the eastern flank has been modified by one much larger slide or several merged slides (Malahoff, 1987). Lō’ihi is nearing the end of its deep submarine, pre-shield stage and is starting to switch from the alkaline-dominant volcanism characteristic of the pre-shield stage to the tholeiitic-dominant volcanism characteristic of the beginnings of the submarine shield sub-stage (Clague and Dixon, 2000). Lō’ihi could emerge above sea level in as little as ~30,000 years or as much as 200,000 years, depending upon eruption rate (USGS HVO website). 37 �Figure 30: This is a regional map of the Lō‘ihi Seamount, the surrounding seafloor, and the sub-aerial south flank of the island of Hawaii. The summit calderas of Kīlauea and Mauna Loa are labeled, as are the Punalu’u and Papa’u slumps. The red line surrounding Lō‘ihi is the extent of known lava flows. The letter ‘L’ indicates the locations of flank landslides. The color ranges from blue for deep and shallowing through of green and yellow shades into orange for shallow. Figure from Clague et al., 2019. 38 �Figure 31: Lōʻihi summit bathymetry with interpretive overlay of caldera- and pit crater-bounding scarps. Hatched lines indicate down-thrown side of the caldera- or pit crater-bounding scarps or ring faults R1 (oldest) to R9 (youngest). The exact sequence of formation for some collapse events could not be determined. P-A to P-D are pit craters. EP is East Pit, WP is West Pit, PP is Pele’s Pit, C1 to C3 indicate cones, S1 to S5 indicate the remnants of lava shields, B indicates 1996 basaltic breccia, and V indicates 5-11m thick volcaniclastic sediment with a basal age date of ∼5900 years. Arrow labeled “flow” indicates direction of channelized flow from S1 to the east. Figure from Clague et al. (2019). 39 �3. Field Trip Stops 3.1. Day 1: Kailua-Kona to Hawai‘i Volcanoes National Park The 7 stops planned for Day 1 are located between the city of Kailua-Kona and the junction between Highway 11 and the South Point Road. The stops are listed below and are shown in Figure D1-1: 1. Kealakekua Bay, Kapu o Keōua Pali (fault scarp), and the 1779 Captain Cook landing location monument; 2. Puʻuhonua o Hōnaunau (Place of Refuge) National Historic Park; 3. The 1950 Mauna Loa Kaʻapuna Flow and Pali Kaholo; 4. 1926 Mauna Loa Ho‘ōpūloa flow at the site of the destroyed village of Ho‘ōpūloa; 5. Old Mauna Loa ʻaʻā flow infilling an older flow channel and complex lava draping. 6. 1907 flows, view of Ka Lea slide scarp, and coastal littoral cones; massive olivine-porphyritic flow core; 7. 1868 Mauna Loa ʻaʻā lava channel; columnar jointing; and olivine-rich pāhoehoe basalt flows with lava stalactites, dripstone, a pseudodyke, and horizontal tree moulds. Figure D1-1: Map showing the Day 1 field stops associated with Mamaloa Highway 11 between Kailua-Kona and the South Point (Ka Lae) Road. Map modified from Hazlett and Hyndman (2007, p.102). 40 �Stop D1-1. Kealakekua Bay and Kapu o Keōua Pali (Fault Scarp). Data Source: Robinson (2010). • UTM 193505E, 2156045N; several parking areas near the end of, or alongside, the road. From the waterfront village of Nāpoʻopoʻo an impressive view of the Kapu o Keōua Pali is possible. This pali, or fault scarp, is thought to have originated during the enormous Alikā landslide that occurred between 100,000 to 150,000 years ago. This landslide may have generated an immense tsunami that scoured the island of Kahoʻolawe to a height of 800ft (243m) and washed blocks of coral as high as 1000ft (305m) up the slopes of the Island of Lānaʻi. Looking closely at the pali (see Figure D1-2) will reveal the presence of several old landslide scars. Captain James Cook landed across the bay in 1779 near the white pylon monument (in the distance south of the pali) and was killed near this spot later in the year during a battle with Hawaiʻians natives. Landslide Scars Captain Cook Monument Figure D1-2: Pali Kapu o Keōua (fault scarp) with several old landslide scars (shown left) and the location of the Captain Cook Monument (shown right). Photo Credits: A.D. MacTavish (left, 2019; right 2020). Stop D1-2. Puʻuhonua o Hōnaunau (Place of Refuge) National Historic Park. Data Source: Robinson (2010). • UTM 194405E, 2150110N; park in parking lot. The Puʻuhonua o Hōnaunau is an ancient Hawaiʻian religious site built on a 750 to 1500yr old pāhoehoe flow delta and is the best-preserved site of its kind in the islands (see Figure D1-3, left). It was originally built about 1650 and was restored in 1968. A large drystone wall over 1000ft (305m) long, 15ft (4.6m) thick, and 10ft (3m) high (see Figure D1-3, right) was built about 1550AD and leads up to the main temple. The wall separated the temple site from the royal grounds on the east side of the wall. The city was a sacred sanctuary that provided temporary shelter to defeated warriors and kapu (taboo) breakers. If kapu breakers were able to reach the site after swimming through the shark infested ocean they were cleansed during a ceremony of absolution and their lives were spared. During times of war this site was also a place of temporary shelter for women, children, the aged, and defeated warriors. 41 �Every hour on the half hour NPS Park Ranger’s present talks in a small amphitheatre near the park entrance building. The talks add detail about the site and its cultural and religious significance. Figure D1-3: The upper left photo is a silhouette of the Place of Refuge temple. The upper right photo shows the 10ft (3m) high drystone wall built with blocks of lava and no mortar. Some of blocks are greater than 2m in length. Photo Credit: A.D. MacTavish (2012). Stop D1-3. Kaʻapuna Flow. Data sources: Hazlett and Hyndman (2007); Robinson (2010). • UTM 197773E, 2132571N, parking area along shoulder of the Highway. At this stop is the very rough Kaʻapuna ‘a‘ā flow which is the widest (~1100ft or 640m) and most easily recognizable of the multiple flow tongues formed during the spectacular 1950 Mauna Loa eruption. The 1950 eruption began when a 15mi (24km) long fissure opened along the southwest rift zone, which is located towards the sea from Mokuʻāweoweo, Mauna Loa’s summit caldera. Approximately 491,789,433 cubic yards (491.8x106 yd3) or 376,000,000 cubic metres (376x106 m3) of dark, blocky ‘a‘ā erupted from the fissure over a period of a few days. The lava flowed down-slope to the sea at a speed of ~6mph (10km/hour) and buried a community in the process. Accretionary lava balls can be seen at the north edge of parking area. The steep slope below and to the west of the highway is Pali Kaholo which is the top of a large slump that cut into the lower western flank of Mauna Loa. Stop D1-4. 1926 Mauna Loa (or Ho‘ōpūloa flow). Data sources: Hazlett and Hyndman (2007); Robinson (2010). • UTM 194595E, 2124810N; pull-off area at north side of road. This stop is the approximate location of Ho‘ōpūloa Village which was destroyed by the underlying Mauna Loa flow in 1926. There has been considerable rebuilding over the old village site. This channelized ‘a‘ā flow (see Figure D1-4) is the result of an eruption along the southwest rift of Mauna Loa that began on April 14, 1926 at an elevation of 7478ft (2316m). The eruption lasted 14 days, extruded a volume of at least 130.8x106 yd3 (100 x 106 m3), and destroyed the Ho‘ōpūloa Village. 42 �The Miloli‘i road taken to get to this site traverses down the pali from Highway 11 to the coast in multiple switchbacks comprising >100 turns over a 5mi (8km) distance; the experience of this road is similar in some ways to the much longer, busier, and much more exhausting Hana Road located on the northeast coast of the Island of Maui (which is well worth a careful drive, in good weather). Figure D1-4: Channelized 1926 ‘a‘ā flow at the former site of Hoʻōpūloa Village. Photo Credit: Google Earth. Stop D1-5. Lava Channel, Drape Structures. Data sources: Easton and Easton (1995); MacTavish (2019). • UTM 209680E, 2111865N; park 160m to the east-southeast near the entrance to King Kamehameha Boulevard (the safer option) or on the southern road shoulder at the field stop. The roadcut on the north side of the road hosts an infilled, 46 to 54ft (14 to 16m) wide ʻaʻā flow channel occupying a depression within underlying pāhoehoe flows (see Figure D1-5, left). These olivine-phyric flows form complex drape structures over the levée margins of channel on both sides of the highway (see Figure D1-5, right). The flows infilling the channel were subsequently overlain by a thin columnarjointed flow. Figure D1-5: On the left is an ʻaʻā flow infilling a channel in underlying pāhoehoe flows and overlain by columnarjointed flow on the north side of the highway. The right photo shows complex drape structures forming a levée to the underlying flow channel infilled by the later ʻaʻā flow. Photo Credits: A.D. MacTavish (left, 2020; right, 2019). 43 �Stop D1-6. 1907 Mauna Loa flows; Ka Lea slide scarp; littoral cones; massive olivine-porphyritic flow core. Data sources: Easton and Easton (1995); MacTavish (2019). • UTM 211515E, 2111125E; scenic overlook (Mile Marker 75). A 15-day Mauna Loa eruption began on January 9, 1907 at an elevation of 6200ft (1890m) ASL. The vent, located 8mi (13km) inland, erupted 981x106 yd3 (750x106 m3) of lava over this period. The roadcut at this stop exposes the core of a massive ‘a‘ā flow with oxidized top and bottom breccias, and abundant green olivine crystals up to 0.16in (4mm) in diameter. Numerous Mauna Loa Southwest Rift Zone cones and flows are located upslope from this location (not readily visible). The Pali o Kūlani fault and slump scarp, located west of Ka Lea (South Point), is visible east and south of the overlook (see Figure D1-6). This pali formed when a large submarine landslide broke away from Mauna Loa’s submerged western slope. In the distance, to the right of the pali at the shoreline, are numerous littoral cinder cones comprising the Pu‘u Ho‘u Cone Complex. The largest, most prominent cone is wave-eroded Pu‘u Ho‘u, which is located to the left of the smaller cones (Figure D1-6). Puʻu Hoʻu Littoral Cone Figure D1-6: 1907 flow (dark-coloured lava) with the Puʻu Hoʻu littoral cone field at the coastline in the distance. Photo Credit: A.D. MacTavish (2019). Stop D1-7. Olivine-rich 1868 Mauna Loa pāhoehoe basalt flows; lava tubes (pyroducts); and horizontal tree moulds. Data sources: Easton and Easton (1995); Robinson (2010); Hazlett and Hyndman (2007). • UTM 216865E, 2109545N; parking on the right (south) side of highway. On the north side of the highway the eastern portion of the 1868 Mauna Loa flow is characterized by thin, very vesicular pāhoehoe flows that range from olivine-poor at their base upward through olivinerich to picritic (very olivine rich with >13 weight % MgO) at their tops (see Figure D1-7, left). The cooling units consist of 2 layers, separated by a central cavity (small lava tubes/pyroducts?), contain dripstone (see Figure D1-7, right) and occasionally small lava stalactites. Lava drapes are common and there are some small, horizontal tree moulds. On the south side of the highway there is a pseudo-dyke produced by channel overflow from a crack in the 1868 levée. Some volcanologists now using of the term pyroduct instead of lava tube; however, the term lava tube is much more intuitive for non-specialists to understand without a lot of explanation and will continue to be used within this guide. 44 �Figure D1-7: The left photo shows the thin 1868 Mauna Loa pāhoehoe olivine basalt flows. Note the wall built on top of the outcrop from blocks of the thin flows. The right photo shows drip-stone and incipient lava stalactites within the central cavity (small lava tube/pyroduct) of a single cooling unit. Photo Credits: A.D. MacTavish (2019). 3.2. Day 2: Highway 11 and South Point Note that as you drive southwest along Highway 11 you will be driving back and forth across the contact between the usually well-vegetated and blocky Mauna Loa ‘a‘ā flows and the primarily less vegetated, Kīlauea pāhoehoe flows. For much of the drive the Mauna Loa flows will be to the right and the Kīlauea flows will be to the left. You will also drive past several thin, difficult to see deposits of Pāhala Ash; however, the best exposures are most easily observed at South Point 8. Punalu‘u Black Sand Beach Park and Hawaiʻian green sea turtles. 9. South Point, Pāhala Ash (Stop D2-9a), and green sand beaches (Stop D2-9b). 8 9b 9a Figure D2-1: Map showing the Day 2 field stops associated with Mamaloa Highway 11, south and west of the town of Pahala and the South Point (Ka Lae) Road. Map modified from Hazlett and Hyndman (2007, p.96). 45 �Stop D2-8. Punalu‘u Black Sand Beach. Data sources: Hazlett and Hyndman (2007); Robinson (2010); MacTavish (2020). • UTM 236403E, 2117534N; parking lot. The Punalu‘u Beach Park protects a black sand beach (see Figure D2-2, left) that is thought to be the site of the first landing in the islands by Polynesians. The sand on the beach was derived from both Mauna Loa and Kīlauea flows and consists of black particles of obsidian (volcanic glass) formed when the lava flows entered the sea, chilled very rapidly, and broke into glassy, sand-sized grains. The beach was larger in the past; however, much of the original sand was removed by a series of tsunami’s that pounded this coastline in 1868, 1960, and 1975. Hawaiʻian green sea turtles (honu) are often present and sleeping in the beach sand (see Figure D2-2). AM has visited this beach 5 times over a twelve-year period and there has always been at least 1 honu sunning itself. Please do not touch the turtles because the bacteria on our skin is potentially deadly to them. Turtles Figure D2-2: The left photo shows the black sand beach at Punalu‘u Beach Park. The right photo is a close-up of 2 Hawaiian green sea turtles (honu) sunning themselves on the beach. Photo credits: A.D. MacTavish (2008). Stop D2-9a, Ka Lea (South Point) and the Pahala Ash. Data sources: Easton (1978, 1987); Easton and Easton (1995). • UTM 217575E, 2093130N; parking lot. Ka Lea is the southernmost point in the United States and is the site of one of the island’s oldest settlements, with artifacts dating back to circa 300AD. Also, at this stop the ~31,000yr old Mauna Loa pāhoehoe flows are overlain by 3.3 to 4.9ft (1 to 1.5m) of >22,000yr old, fine-grained, yellow-brown, palagonatized, wind-reworked Pāhala Ash (see Figure D2-3). The Pāhala Ash is an enigmatic, widespread deposit used as a stratigraphic marker despite the uncertainty of its source or age. It is present on Kīlauea, Mauna Loa, Mauna Kea, and Kohala volcanoes. The oldest Pāhala Ash on Mauna Loa and Kīlauea is ~31,000 years old with the youngest ash covered by flows that are between 10,000 and 200 years old. This ash formation may represent a long period where Kīlauea was primarily explosive, possibly due to a period of higher rainfall during the last glacial maximum. This caused more phreatomagmatic (water and magma interaction) summit activity, increased the size of the caldera, and resulted in both explosive and magmatic activity (Easton, 1978). 46 �Figure D2-3: These photos show the 1 to 1.5m thick deposits of well-bedded, but poorly consolidated Pāhala Ash at Ka Lea (South Point). Photo credits: A.D. MacTavish (2019). The sea cliff (see Figure D2-4) that forms South Point’s western coastline and eventually heads inland to the north is known as Pali o Kūlani. The cliff is a slump scarp where Mauna Loa’s Southwest Rift Zone dropped due to a large slide that broke from the submerged Mauna Loa slope to the south and west. The Puʻu Ha‘u Littoral Cone can be seen ~4.35mi (7km) to the northwest (observed first from Stop D1-7) from the top of the western sea cliff (noted by arrow in the upper left of Figure D2-4). The cone marks the location where a tongue of lava from the 1868 Mauna Loa ‘a‘ā flow, that erupted from the inland extension of the scarp, entered the ocean. Wave erosion has cut the cone in half. Puʻu Ha‘u Littoral Cone Figure D2-4: The western sea cliff at Ka Lea (South Point) with the Puʻu Ha‘u cone in the distance to the northwest at top left. Photo credit: A.D. MacTavish (2019). 47 �Stop D2-9b. Papakōlea Green Sand Beach. Data sources: Easton (1978, 1987); Easton and Easton (1995). • UTM 218448E, 2094263N; parking Area. • UTM 221295E, 2095850N; intermediate green sand beach located south of 4x4 road. • UTM 221295E, 2095850N; Papakōlea Green Sand Beach. South Point is famous for its green olivine sand beaches. The source of the olivine comprising the beach is the weathered and eroded, olivine-rich, Puʻu Mahana Littoral Cone. The at least 2 to 3-hour return hike from the parking lot northeast to the Papakōlea Green Sand Beach is approximately 2mi (3.2km), as the nene flies, and at least 3mi (5km) along a complex series of 4x4 roads incised into the Pāhala Ash. Some field guides, and the present authors, recommend that if you have any trouble walking that you do not attempt the full hike. There are other, smaller, pocket green sand beaches along the shoreline seaward of the route about 2/3 of the distance to the main beach (see Figure D2-5), that can be visited if there is not enough time (or energy) to complete the full walk (bring lots of water and snacks and liberally apply sunscreen). This is not an easy walk, even though it is relatively flat. Access to the main beach is difficult so good hiking boots or hiking shoes are a must. For a fee very rough transportation in the back of 4x4 trucks is available from a group of indigenous Hawaiʻians at the parking lot at the beginning of the access trail. This transportation is not recommended if you have a bad back, knees or neck or suffer from vertigo or motion sickness. Figure D2-5: The left photo shows a small, pocket green sand beach located along the coastline between South Point and the Papakōlea Green Sand Beach. The right photo shows the small, green, olivine sand grains mixed with small amounts of carbonate sand and some magnetite that comprise the pocket beach in the left photo. Photo credits: A.D. MacTavish (2019). Robinson (2010) states that the olivine comprising the beach sand was derived by wave erosion of the Puʻu Mahana Littoral Cone (or tuff ring volcano; see Figure D2-6) which formed approximately 28,000 years ago. A tuff ring forms from interaction of magma with shallow groundwater or seawater. Wave action has eroded the seaward side of the tuff ring, formed a small bay, and removed the light grains of ash while leaving the denser and heavier olivine grains behind. The sides of the cliffs above the beach are quite steep and access to the beach is difficult. Do not attempt to descend to the beach if you have difficulty walking or climbing 48 �Figure D2-6: Papakōlea Green Sand Beach from the western crater rim. Please note how small the people on the beach appear, which gives you an idea of the height of the poorly consolidated cliffs. The ash layers comprising the walls of the cone are readily visible in the upper centre of the photo. Photo credit: A.D MacTavish, 2020. 3.3. Day 3 (Part 1): Mauna Loa Road and Mauna Loa Strip The Mauna Loa Road starts at Highway 11 and ascends up the eastern flank of Mauna Loa for 12mi (19.3km) to an elevation of 6725ft (2050m). This road traverses the ‘Mauna Loa Strip’ which is the portion of Hawai‘i Volcanoes National Park linking the summits of Kīlauea and Mauna Loa (Moku‘āweoweo Caldera). The Mauna Loa Strip from the western rim of Kīlauea caldera comprises a narrow, dark green, forested belt enclosed by the grassy, light-coloured grassland that ascends Mauna Loa’s eastern flank. This strip of subtropical Hawaiʻian upland forest is one of the world’s most rare and fragile ecosystems. This forest type was once much more widespread; however, it has been almost completely eliminated due to overgrazing, over-logging, eruptions, and displacement by introduced species (Hazlett 2014). Since this is a unique, ecologically fragile area with a high fire danger it is recommended that all visitors be especially careful. PLEASE DO NOT SMOKE FOR ANY REASON. Mauna Loa Road Field Trip Stops (see Figure D3-1): 10. Large tree moulds. 11. Kīpuka Puaulu; ecologically diversified old land surrounded by younger lavas. 12. Ke‘āmuku Flow; thin lobe of a larger composite flow derived from several eruptions that have been grouped together. 13. Ke‘āmuku Flow channel; a spectacular and well-developed flow channel. 14. Road’s End Scenic Overlook; panoramic view of Kīlauea Caldera, the Ka‘ū Desert, and the upper Southwest and East rift zones (on a clear day); this is also the start of the Mauna Loa Trail. Backtrack from Stop D3-14 to Highway 11 and drive southwest to the Ka‘ū Desert Trail (Footprints/Mauna Iki Trail) for Part 2 of Day 3. 49 �Map from National Geographic Hawai‘i Volcanoes National Park Illustrated Trails Map (2010) 14 13 12 11 10 Figure D3-1: Field trip stops on the Mauna Loa Road and within the Mauna Loa Strip. Map taken from the National Geographic Hawai‘i Volcanoes National Park Illustrated Trails Map (2010). Stop D3-10: Lava Tree Moulds Area. Data sources: Hazlett and Hyndman (2007); Hazlett (2014); Easton and Easton (1995); MacTavish (2019). • UTM 260480, 2150490; parking area and turnaround located at the end of a 0.4mi (640m) long side road to the right of the Mauna Loa Road. The lava tree moulds here are very large, well-preserved, and encased by 700 to 800yr old Kīlauea pāhoehoe lava and according to Hazlett and Hyndman (2007) are: ‘Among the largest and deepest in Hawai‘i and preserve the shapes of mature acacia koa tree trunks.’ These moulds can exceed ~5ft (1.5m) in diameter with depths of up to 10ft (3m). The largest mould is located to the right of the road about 200ft (60m) west of the parking area (at approximately UTM 260412E, 2150479N) as you are exiting the road loop back to the Mauna Loa Road. Some of the moulds have trees growing out of them or provide a place for roots to grow (see Figure D3-2). In all cases at this location there is a well-developed weathering-resistant, up 15cm thick, radially jointed chill margin surrounding each mould (also see Figure D3-2). Figure D3-3 graphically illustrates how tree moulds and lava trees form. 50 �Figure D3-2: Large fenced lava tree mould with tree root and thick chilled mould rim on the left and the preserved bark pattern in mould wall on the right. Photo credits: A.D. MacTavish (2019). Figure D3-3: Lava trees and tree moulds form when a forest is invaded by a lava flow and the lava surrounds the trees. The lava chills against the tree trunks, the ground, and the top of the flow and forms a solid crust around the trees. As the lava supply diminishes the remaining liquid drains away. ‘Tree moulds’ are hollow impressions of trees left in the lava that are enveloped, but not instantaneously incinerated by the lava and where the surface of the flow does not drop. If the mould is preserved as a shell rising above the surface of the flow after the flow surface drops it is termed a ‘lava tree’. (Hazlett, 2014; Easton and Easton, 1995). Stop D3-11. Kīpuka Puaulu and self-guided trail. Data sources: Hazlett and Hyndman (2007); Easton and Easton (1995); National Parks Service (NPS) Kipukapuaulu Trail Guide. • UTM 258180E, 2150865N; parking area. Kīpuka are areas of old, often forested land, that are usually surrounded by unvegetated younger terrain, often flows. There are innumerable kīpuka on the island and they provide isolated habitats for many, often rare, plants, animals, and birds that can be found nowhere else on earth. The fragile, ecologically diverse Kīpuka Puaulu is located at the base of the long slopes where Kīlauea and Mauna Loa meet. Here the kīpuka is underlain by up to 20ft (6m) of 2200yr old volcanic ash that accumulated as fallout strata and windblown ash on top of older flows. The kīpuka is enclosed on 3 sides by a <500yr old Mauna Loa flow and, at the trail entrance, by a 700-800yr old Kīlauea flow. If time allows the trail may be hiked by those who are interested. Walking the 1mi (1.6km) self-guided nature trail (please keep the gate closed) provides striking contrast in vegetation that corresponds to different ages, compositions, and weathered surfaces along the edge of the kīpuka. 51 �Stop D3-12. Thin southeast offshoot lobe of the composite Ke‘āmuku Flow. Data Sources: Hazlett (2014); MacTavish (2019). • UTM 254110E, 2153350N; parking area. • UTM 254065E, 2153405; centre of flow lobe ~75m northwest along road from cattleguard. This narrow (~330ft or ~100m wide), offshoot ʻaʻā flow lobe from the main composite Ke‘āmuku Flow, located to the north, is a weakly-developed levéed channel containing a broken, roughly spherical accretionary lava ball located ~35m (115ft) south (downslope) from the road (see Figure D3-4). This lava ball was formed from pieces of solidified lava which were pulled into the stream of the lava channel and rolled along with the current such that the ball eventually accumulated a series of coats of lava in a similar manner to a rolled snowball increasing in size due to added snow. Accretionary Lava Ball Figure D3-4: Broken accretionary lava ball within offshoot of the Keʻāmoku Flow. Photo source: Google Earth. Stop D3-13. Ke‘āmuku Flow. Data sources: Lockwood (1979); Hazlett (2014); MacTavish (2019). • UTM 251868E, 2155185N; centre of flow. • UTM 251940E, 2155225N; eastern parking area located just before the flow opposite the 5630ft (1730m) ASL sign. • UTM 251697E, 2155250N; western parking area (located ~330ft or ~100m past the flow). Flows from several eruptions have been grouped together by the USGS and referred to as the Ke‘āmuku Flow. Here this specific ʻaʻā flow exhibits a spectacular, well-developed, levéed lava channel (see Figure D3-5) that Lockwood (1979) estimates is between 400 and 500yrs old and was erupted from vents on Mauna Loa’s northeast rift zone. There are numerous, partially buried accretionary lava balls within the channel. 52 �Figure D3-5: Well-developed, levéed, lava channel within the Keʻāmoku Flow. Photo source: A.D. MacTavish (2019). Stop D3-14. Road’s End Scenic Overlook. Data Sources: Hazlett (2014); Easton and Easton (1995). • UTM 249630, 2157090; parking area, restrooms, and picnic tables. From the road, and the overlook shelter, panoramic views of Kīlauea Caldera and the upper Southwest and East Rift zones are available from an elevation of 6725ft (2046m), if cloud cover permits. This is also the trailhead for the 30.5km (19.0mi) Mauna Loa Trail which leads upslope to Moku‘āweoweo caldera and the summit of Mauna Loa at 13,677m (an approximate 2-day hike). 3.3. Day 3 (Part 2): Mauna Iki Trail/Kaʻū Desert Trail and the Southwest Rift Zone Field trip stops along the western portion of the Ka‘ū Desert Trail/Mauna Iki Trail (see Figure D3-6): 15. Ka‘ū Desert Trailhead with a view of Ka‘ōiki Pali to the west. The trail starts on the lower Ke‘āmoku Flow. 16. Several Large accretionary lava balls similar to those observed earlier at Stops D3-12 and D3-13. 17. Trail drops down from Ke‘āmoku ʻaʻā flows onto Kīlauea pāhoehoe flows. Sand dunes, palagonitized Pele’s hair, and some ash layers are visible along the trail. 18. Keanakāko‘i Ash with well-developed graded, cross-bedded and laminated ash layers. 19. Fossil footprints in 1790 pisolitic (accretionary lapilli-bearing) Keanakāko‘i Ash. Recent pyroclastic activity from Kīlauea summit (2018) has obscured many of the footprints. However, wind will probably uncover other, presently buried footprints in the future. 20. Accretionary lapilli layer. 21. Pāhoehoe toes formed from small lava breakouts from the base of a small tumulus. 22. Pre-Mauna Iki lava channel with lava level marks visible on the northern levée of the channel. 23. Trail crosses 700yr old pāhoehoe flows covered with drifts of Keanakāko‘i Ash and recent, circa 2018, Pele’s hair with a large tumulus rising near the edge of the Mauna Iki shield. 24. The southern edge of the large tumulus hosts the broken and congealed remnants of several small lava falls. 25. Mauna Iki summit which was active in 1919 and 1920. This shield is relatively low topographically. 53 �15 16 17,18 19 20 21 22 23,24 25 Figure D3-6: Field trip stops along the Kaʻū Desert Trail between Highway 11 and the summit of Mauna Iki. Map taken from National Geographic Hawai‘i Volcanoes National Park Illustrated Trails Map (2010). Stop D3-15, Ka‘ū Desert Trailhead. Data Sources: Lockwood (1986); Hazlett (2014); MacTavish (2019). • UTM 251315E, 2143330N; turnout (parking area) on left. Approximately 1.0km (0.5mi) west of the parking area is Ka‘ōiki Pali which is a 100m (325ft) high fault scarp related to the Ka‘ōiki fault system. This fault system is a region of recurrent seismic activity along the southeast flank of Mauna Loa (Hazlett, 2014) The Ka‘ū Desert Trail begins on the lower part of the Ke‘āmoku Flow visited earlier in the day at Stops D3-12 and D3-13. The ʻaʻā flow at this location has been dated at ~500yrs old (Hazlett, 2014); many well-developed accretionary lava balls are present on the surface of the flow. This date was obtained via a 1986 personal communication between Hazlett (2014) with J.P. Lockwood. 54 �Stop D3-16a. Large accretionary lava balls. Data source: MacTavish (2019). • UTM 251375E, 2143250N. This area has several good examples of large, well-developed, accretionary lava balls (see Figure D3-7) that are similar to those observed earlier at Stops D3-12 and D3-13. Some of the best examples are located to the left (east) of the trail near the 2 trail signs at this location. Figure D3-7: Large well-developed accretionary lava balls on the Lower Ke‘āmoku Flow. Photo credit: A.D. MacTavish (2019). Stop D3-16b, Large, >2m accretionary lava ball. Data source: MacTavish (2019). • UTM 251430E, 2143150N At this field stop, on the left side of the trail, is a >6.5ft (~2m) diameter accretionary lava ball (see Figure D3-8, left). This lava ball has several cavities that show the interior structure of the ball (see Figure D3-8, right). Figure D3-8: The left photo shows a large >2m diameter accretionary lava ball on the north edge the trail (Dr. Juk Bhattacharyya as scale). The right photo shows the interior of the lava ball exposed by a large cavity (lens cap for scale). Photo credits: A.D. MacTavish (2019). 55 �Stop D3-17. Edge of Ke‘āmuku Flow. Data sources: Easton and Easton (1995); Hazlett (2014). • UTM 251705E, 2142620N The trail descends between 20 and 30ft (6 to 9m) from the top of the Ke‘āmuku Flow onto the surface of an 800 to 900yr old (Swanson, 2000), ropey Kīlauea pāhoehoe flow (see Figure D3-9). The surface of the flow is locally covered by patches of a dark sandy ash which is the upper part of the Keanakāko‘i Ash erupted from the Kīlauea summit in 1790 (Easton and Easton 1995; Hazlett 2014). Figure D3-9: In the left photo the Ka‘ū Desert Trail descends from the Ke‘āmoku Flow onto and older, underlying, ropey Kīlauea pāhoehoe flow, shown in the right photo. Photo credits: A.D. MacTavish (2019). Stop D3-18. Keanakāko‘i Ash. Data sources: Swanson and Christianson (1973); MacTavish (2019). • UTM 251632E, 2142286N; ~30m east (left) of where the paved trail ends. This location and its vicinity has several good exposures of finely bedded and laminated, graded, and cross-bedded greyish brown and light grey to yellowish-brown Keanakāko‘i Ash that erupted from the summit of Kīlauea in 1790. This ash is easily eroded and reworked by the wind and varies from weakly lithified to completely unlithified. Partially lithified examples are often preserved in cracks, crevasses, and wind-protected locations such as the lee sides of basalt outcrops and small ridges, or where stabilized by the roots of trees or grass (see Figure D3-10). Where more than a few centimetres are preserved it is possible to see that the greyish-coloured ash is unlithified, whereas, the yellowishcoloured ash is partially lithified and more weathering-resistant. The cross-bedded nature of some of the preserved ash supports the Swanson and Christianson (1973) suggestion that the ash, at least in part, was formed from a series of pyroclastic base surge events. 56 �Figure D3-10: The upper photo shows graded, cross-bedded, finely bedded to laminated Keanakāko‘i Ash of various colours. The buff to tan-coloured layers are the ones that preserve the 1790 fossil footprints hopefully exposed near Stop D3-19. The lower left photo shows ash plastered against an older flow. The lower right photo shows ash preserved in cracks and below overhangs in the older flow. Photo credits: A.D. MacTavish (2019). Stop D3-19. Fossil footprints in Keanakāko‘i Ash. Data sources: Easton and Easton (1995); Hazlett (2014); Swanson and Christianson (1973). • UTM 251562E, 2142195N; NPS Hut. The 230yr old fossil footprints (see Figure D3-11) are preserved within yellowish-grey, partially lithified beds of Keanakāko‘i Ash erupted from Halema’uma’u crater within the Kīlauea summit caldera in 1790. Many of the unprotected footprints within the yellowish-grey to yellowish-brown beds of Keanakāko‘i Ash have been covered by both windblown 1790’s and 2018 ash and are often very difficult to find and to see. An NPS shelter built on site exhibits moulds of some of the better fossil footprints, but does not cover any of the actual footprints, which are all exposed to outside weather. 57 �Figure D3-11: These photos show some of the footprints preserved in the yellowish-brown variety of Keanakāko‘i Ash. Most of these footprints had been covered by windblown 1790 and 2018 vintage ash during AM’s visits to the site during the summer of 2019 and winter of 2020 and did not provide any illustrative photographs. Photo credits: Donald A. Swanson, U.S. Geological Survey, Hawaiian Volcano Observatory website. 3.3.1. History of the Footprints Area: The following description of the events in 1790 that produced the fossil footprints was taken verbatim from Easton and Easton (1995, p.37 and 38; also see Figure D3-12) with the spelling of Hawaiʻian words as in the original document: “Footprints are preserved in indurated ash from the 1790 eruption from Halemaumau. In addition to the footprints at the shelter, other footprints are visible along the trail. The following account of the 1790 eruption is condensed from Swanson and Christianson (1973). An army led by King Keoua camped on the northern rim of Kilauea Caldera. That night Kilauea erupted violently. The next day King Keoua was afraid to travel, and Kilauea again erupted explosively that night. The same pattern held for the next day. On the third day, Keoua split his army into three groups of about 80 men (and their families) each, and resumed their march to Kau. The groups left at intervals of about 2 to 4 hours apart. Soon after the second group left, a violent phreatic or phreatomagmatic eruption occurred at Kilauea, and a hot base surge composed mainly of superheated steam spread SW of Kilauea Crater, enveloping the second group of King Keoua’s party, suffocating the army [see Figure D3-18]. The lethal front also overwhelmed the first group, but had dissipated somewhat, and caused only a few injuries or deaths. The third group was in a protected area and quickly joined the first group (after discovering the deaths of the second group) and quickly left the scene. Other base surges probably accompanied the explosions witnessed by Keoua’s army on the previous 3 nights, but the encampment was on the high upwind side of the caldera, and the high caldera walls would have served to protect the encampment. The death of part of King Keoua’s army has historical significance, since the loss of warriors may have aided Kamehameha in his unification of the Island of Hawaii, and later the archipelago. 58 �The footprints are preserved in soft ash 7 to 9 km SW of the 1790 eruption site and occur in two ash layers that contain numerous pisoliths (accretionary lapilli) and are separated by 90cm of dune sand. The lower footprint layer contains few footprints, most heading away from Kilauea. The upper footprint layer contains more footprints, most heading to Kilauea Crater. Swanson and Christianson (1973) speculate that the lower footprints were made when the army fled the eruption site, and the upper footprints when the army returned days or weeks later.” Figure D3-12: Sketch map showing the location of King Keōua’s army, the area of the eruption, and the footprint locality. Map after Swanson and Christianson (1973) and taken from Easton and Easton (1995). Stop D3-20. Accretionary Lapilli/Lapillistone. Data Source: A.D. MacTavish. • UTM 251565E, 2142071N; located ~650ft (~200m) south of the Footprints shelter (NPS hut). Visible here is a thin, 10-15cm thick, slightly erosion-resistant, weakly-indurated (lithified), brownishgrey lapillistone layer consisting of 3 to 4mm diameter, light brown, accretionary lapilli (pisoliths) concentrated within a light brown ash matrix (see D3-13). The lapilli within the layer are marginally matrix-supported. Figure D3-13: Accretionary lapillistone layer. Left photo shows the surface of the slightly weather-resistant layer. The right photo shows the accretionary lapilli and ash comprising the layer. Photo credits: A.D. MacTavish (2020). 59 �Stop D3-21: Budding of pāhoehoe toes from small tumulus. Data Source: A.D. MacTavish (2020). • UTM 251852E, 2141953N. This stop shows several small pāhoehoe toes formed from small lava breakouts from the base of a small tumulus (see Figure D3-14). These small breakouts only travelled a couple of metres from the tumulus before congealing. Figure D3-14: Pāhoehoe toes budding from small tumulus. Photo credit: A.D. MacTavish (2020). Stop D3-22: Lava levels on margins of lava channel. Data Source: A.D. MacTavish (2020). • UTM 252041E, 2141725N; Northern levée of the channel. The shallow lava channel at this stop has prominent levées with well-preserved, horizontal lava levels that mark the level of lava within the channel when flowing lava was present and lava volume from the source was decreasing (see Figure D3-15). This channel is pre-Mauna Iki Shield. Figure D3-15: The left photo shows a shallow lava channel and its northern levée. The right photo shows the horizontal lava levels progressively marking the top level of the lava within the channel as the volume of lava decreased. Photo credits: A.D. MacTavish (2020). 60 �Stop D3-23. Large fractured tumulus on the margins of the Mauna Iki Shield. Data source: Hazlett (2014). • UTM 252285E, 2141518N; located ~215ft (~65m) west of the trail and easily identifiable by its height and size. The feature here is the spectacular, very large, fractured, 10 to 12m high tumulus that looms west of the trail near the northern edge of the Mauna Iki shield (see Figure D3-16). Mauna Iki sits on Kīlauea’s Southwestern Rift Zone about 8.8km southwest of the summit caldera. Tumuli form when pāhoehoe flows develop a crust on their surface due to cooling in contact with air. A subsequent influx of lava beneath this crust will lift or inflate it. This inflation is not uniform, with some portions of the flow inflating more than others. Tumuli are essentially focused inflation features, whereas areas of arrested inflation (depressions) are referred to as inflation pits (Hazlett 2014). This tumulus is very large compared with most other tumuli observed on this field trip. Figure D3-16: Large, fractured tumulus with February 2020 tour participants for scale. Photo credit: A.D. MacTavish (2020). Stop D3-24. Three small lava falls formed from late lava breakouts (?) from tumulus. Data source: A.D. MacTavish (2020). • UTM 252287E, 2141497N; located ~80m west of the trail. On the south side of the large tumulus observed at Stop D3-23 are 3, possibly syn-tumulus, lava breakouts (buds) that flowed down the side of the tumulus, possibly at the end of its formation. The buds oozed over the broken lip of an earlier pāhoehoe lava tube to form several small lava falls. These small falls are now mostly broken (see Figure D3-17). Participants in the February 2020 Field Tour designated these features as ‘post-tumulus budding lava ooze blobs’, which is an entertaining, somewhat non-geological, description which describes the sense of fun and wonder embodied by the participants of the field trip, if nothing else. 61 �Figure D3-17: ‘Post tumulus budding lava ooze blobs’ (not a technical term, but entertaining nonetheless) located on south side of the large, fractured tumulus described at Stop D3-23, with a trekking pole for scale. Photo credit: A.D. MacTavish (2020). Stop D3-25. Summit of Mauna Iki Lava Shield. Data source: Hazlett (2014); Rowland and Munro (1996). • UTM 252610E, 2141010N This stop is at the summit of the small Mauna Iki lava shield which formed during an 8-month eruption along Kīlauea’s Southwest Rift Zone in 1919 and 1920. During most of the eruption the Mauna Iki summit contained an active lava lake which often overflowed and poured through lava channels and lava tubes down the flanks of the shield. At the same time there was also a lava lake within Halema‘uma‘u crater at the summit of Kīlauea. It was noted that periods of high-stand within the Halema‘uma‘u lava lake coincided with vigorous overflow from Mauna Iki. This strongly suggests that a direct connection existed between the 2 vents even though they were geographically separated, northeast to southwest by 5.45mi (8.8km). The shield grew laterally more than vertically which accounts for its relatively low topographic height above the surrounding terrain. The total volume of lava erupted from Mauna Iki was about 1.23 billion ft3 (35 million m3) (Rowland and Munro, 1996). Safety Note: Please be very careful and remain on the trail when at the summit of Mauna Iki. The summit area is relatively dangerous due to the instability of the crater rim, lava tube skylights, and embankment edges. 62 �3.4. Day 4: Kīlauea Caldera, Kīlauea Iki, Hilina Pali 3.4.1. The Kīlauea East Rift Zone: Figure D4-1 (see below) is a generalized depiction of the structure of Kīlauea Volcano from Hazlett (2014). The East Rift Zone trends east-northeast, except for the relatively short distance between the summit caldera and Pauahi Crater, where the trend is southeast. For some unknown reason the bend in the rift to the east-northeast occurs near Mauna Ulu. There is a second, limited fissure system extending from Halema‘uma‘u through Kīlauea Iki Crater that may represent an ancient trace of the East Rift Zone, most of which has shifted southward as the volcano has grown. The East Rift Zone (Figure D4-1) is informally subdivided into: the Upper East Rift Zone (UERZ); the Central East Rift Zone (CERZ); and the Lower East Rift Zone (LERZ). Figure D4-1: Generalized structure of Kīlauea Volcano showing the caldera, the East Rift Zone, Southwest Rift Zone, Kaōiki Fault Zone, Koaʻe Fault System, and Hilina Fault System. Map from Hazlett (2014, p.78). 63 �3.4. Day 4 (Part 1): Kīlauea Caldera Much of Crater Rim Drive and many Kīlauea Caldera area trails were closed during the 2018 eruption and the caldera collapse of Halema‘uma‘u Crater. This collapse happened after the withdrawal of the lava lake that had been active within the crater between 2008 and 2018. Most trails were closed during the February 2020 field trip; however, all are now open. Crater Rim Drive is permanently closed from the Jagger Museum/HVO to the junction with the Pu‘u Pua‘i/Devastation Trail access road. The National Park Service (NPS) plans on re-routing Crater Rim Drive south of the destroyed central portion of the road. The Uēkahuna Bluff viewpoint located north of the closed Jagger Museum/HVO has reopened. The Field Trip Stops for Day 4 (Part 1) are (see Figure D4-2): 26. 27. 28. 29. 30. Steaming Bluffs Overlook; Volcano House Observation Deck; Kīlauea Visitor Center and park headquarters; Kīlauea Iki Scenic Viewpoint; Kīlauea Iki trailhead; Kīlauea Iki Trail (partially closed in 2020); the portion of this trail leading northwest from Substop D4-30-4 to Waldron Ledge was closed during the 2020 field trip due to the instability of the northeastern caldera wall, but it is now open all the way to Volcano House; 31. Pu‘u Pua‘i and Devastation Trail; walk to cinder cone located at the south rim of Kīlauea Iki; and 32. Keanakāko‘i Crater; road access past Crater Rim Drive/Chain of Craters Road junction is blocked to vehicular traffic due to the partial destruction of Crater Rim Drive by caldera subsidence. Access by foot is allowed. 28 26 27 30-1 to 30-14 29 31 32 Figure D4-2: Map of Kīlauea Caldera and Kīlauea Iki Crater showing Day 4 (Part 1) field trip stop locations. Map taken from Hazlett (2014, p.28). 64 �Stop D4-26. Steaming Bluffs (Wahinekapu) and Sulphur Banks. Data source: Hazlett (2014). • UTM 262875E, 2149795N, parking area. The Steaming Bluff Viewpoint is located ~600ft (180m) south of the parking lot along a flat, wide, and well-marked trail. The Sulphur Banks solfatara are located to the northeast of the parking area and are associated with ring faults occurring along the northern edge of the caldera (low cliff-face observable ~1000-1300ft or 300400m north). The solfatara can be reached by taking the Sulphur Banks (Ha‘akulamanu) Trail starting on the north side of Crater Rim Drive opposite the parking lot. If you look around before moving out of the parking lot you will see numerous steam clouds issuing from steam vents located along various structures flanking the slightly down-dropped caldera block at this location. To access the Steaming Bluffs Viewpoint, take the viewpoint trail south from the parking area. There are also several steam vents below the viewpoint that issue from one of the faults bounding the main part of the caldera. The viewpoint provides a good view of the Kīlauea Caldera and the new crater formed after the collapse of the caldera floor beneath Halema‘uma‘u crater in mid-2018. Figure D4-3: The left photo shows steam issuing from below the Steaming Bluffs Viewpoint. The right photo provides a view of the modified Halemaʻumaʻu Crater within Kīlauea Caldera as seen from the Steaming Bluffs Viewpoint. Photo sources: A.D. MacTavish (2020). Stop D4-27. Volcano House Hotel, Kīlauea Caldera Viewpoint. Data source: MacTavish (2019). • UTM 262875E, 2149795N; parking lot. The viewpoint area on the caldera side of the hotel provides another good place to observe Kīlauea Caldera. Features visible from the viewpoint: • • • Almost directly ahead (at 1100 o’clock) is the large, crater formed during the 2018 caldera collapse events that swallowed the original Halema‘uma‘u Crater (see Figure D4-4, left); To the right is the vertical cliff-face marking one of the northwestern rim faults of the caldera (see Figure D4-4, right). This not the northern caldera rim. The northern rim is located a further 2000-2300ft (600 to 700m) to the north and can be easily seen from the Steaming Bluffs; On a clear day, in the distance to the west and past the caldera rim, can be seen the mass of Mauna Loa Volcano. Please note the gentle slopes involved. We cannot see the summit from this location since the eastern flank of the volcano bulges somewhat and blocks the view; and 65 �• Visually following the cliff rim further to the left the buildings comprising the Hawaiian Volcano Observatory (HVO) and the Jagger Museum will eventually come into view (just before the cliff face drops down a level). Both are now indefinitely, possibly permanently, closed due to earthquake damage from the violent subsidence of the caldera floor in mid-2018. HVO New collapse crater HVO Figure D4-4: The left photo shows the new collapse crater (caldera) that engulfed Halema‘uma‘u Crater. The right photo allows a good view of the northern rim of Kīlauea Caldera. All photos taken from the Volcano House viewing area. Photo sources: A.D. MacTavish (left photo, 2020, right photo; 2012). Stop D4-28. Kīlauea Visitor Center. • UTM 262995E, 2149900N, parking Lot; this stop will be made on Day 1 or Day 2. Maps, books, trail information, weather, T-shirts, washrooms, etc. can be obtained at this visitor’s centre. There are usually some park rangers around that will be happy to answer any questions you may have. Stop D4-29. Kīlauea Iki Viewpoint Trailhead. Data Source: Hazlett (2014). • UTM 264475E, 2148490N; parking lot. Kīlauea Iki crater (see Figures D4-5 and D4-6) is the site of the mid-15th century collapse of the ‘Ailā‘au lava shield. In 1832 and 1868 eruptions began with crater floor collapse followed by partial lava infill. Prior to 1959 the crater was ~600ft (180m) deep and almost completely forested; The 1959 eruption began at 808PM on November 14th after a 3-month swarm of earthquakes and summit inflation. The eruption lasted for 36 days, filled the crater to the 400ft (120m) level with a lava lake; produced 17 episodes of lava fountaining (Figure D4-5, right), some reaching heights of 1900ft (580m); and built the Pu‘u Pua‘i (gushing hill) cinder cone (left centre in both photos in Figure D4-5). The northern flank of the cone has partially collapsed since the eruption; however, parts of the steep, unstable face of Pu‘u Pua‘i occasionally slid into the lake during the eruption and were rafted across the lake where today they form topographic highs on the floor of the crater. The volcanic haze (‘vase’) plume from the lava lake that resided in Halema‘uma‘u Crater from early 2008 until mid-2018 is visible behind the cinder cone in the background of the left photo in Figure D4-5. The present crater floor, which is the solidified top of the 1959 lava lake, still steams in places, providing evidence of continuing heat release due to cooling of the crystallized lava at depth. 66 �Figure D4-5: The left photo shows Kīlauea Iki Crater in December 2012 with Puʻu Puaʻi cinder cone in the left middle distance. The right photo is a wonderful picture of lava fountaining and the active lava lake within the crater during the 1959 eruption. Both photos were taken from about the same location. Photo sources: The left photo is by A.D. MacTavish (2012); the right photo is by the USGS (1959) and was downloaded from the USGS HVO website. Figure D4-6: Map of Kīlauea Iki Crater and immediate vicinity. Taken from Hazlett (2014, p.66). 67 �Stop D4-30. Kīlauea Iki Trail. Data Sources: Hazlett (2014); NPS Kīlauea Iki Trail Guide. • UTM 262995E, 2149900N; parking lot at trailhead and overlook. The Kīlauea Iki Trail leaves from the northern end of the Overlook parking lot and heads in an anticlockwise direction around the crater’s north rim, winds down to Byron Ledge (western rim) and down to the crater floor where it proceeds east along the long axis of the crater floor, climbs the eastern crater rim to the Nāhuku (Thurston Lava Tube) parking lot, and then proceeds northwest along the crater rim back to the Overlook parking lot. This 4mi (6.4km), moderate-difficulty hike makes a 440ft (122m) elevation change from 3874ft (1180m) at the eastern rim down to ~3474ft (1060m) on the crater floor. The trail ends at the southern end of the overlook parking lot. Small brown NPS trail markers with yellow numbers mark the trail stops and the 14 D4-30 sub-stops within this field guide match those numbers (see Figure D4-7). After the 2018 main caldera subsidence events the northern half of the trail was closed due to instability of the Kīlauea Iki Crater walls; however, it was re-opened just prior to the 2020 field trip and remains open. While on this trail please abide by the following safety rules: • • • • • • • Stay on the trail; Avoid unstable cliff edges; Keep away from ground cracks; Wear sturdy walking shoes or hiking boots (flip-flops and sandals are dangerous on this trail); Wear sunscreen; Bring a raincoat, particularly if you are hiking the trail in the afternoon since, in this part of the island, it rains most afternoons; AM has been rained upon both times he has hiked the trail; and Carry plenty of drinking water (it can get very hot within the crater). Weather conditions can change very quickly so please take protective gear for both rain and sun. 30-6 30-8 30-4 30-3 30-2 Overlook Parking Lot Stop D4-29 30-1 30-7 30-10 30-11 30-12 30-9 30-13 30-14 Figure D4-7: Field trip Day 4 sub-stops within Kīlauea Iki Crater. Map modified from Hazlett (2014, p.66). 68 �Stop D4-30-1. Kīlauea Iki Crater, NPS Marker 1. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 264250E, 2148540N. This location is identified with a brown NPS sign with the yellow #1 at the beginning of the opening. It provides a good view of the Pu‘u Pua‘i cinder cone. The main 1959 eruptive vent (see Figure D4-8) is the linear depression/cavity located at the base of the cone. During the eruption this fissure was up to 800m in length. The photo below is the only one available that properly shows the fissure, even though it was taken from the top of the crater wall located across the crater to the south of this stop location. North face of Puʻu Puaʻi Cone Hiker for scale Eruptive vent fissure Figure D4-8: The 1959 Kīlauea Iki eruptive vent seen from the top of south wall of the crater on the eastern flank of the cinder cone (the only available photo properly showing the fissure). Photo source: A.D. MacTavish (2019). Stop D4-30-2. Concrete Trolley Platform, NPS Marker 2. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263820E, 2148505N. After the eruption ended in late 1959 USGS scientists used this site to descend into the crater using a trolley system to study and sample the cooling lava lake. The NPS Kīlauea Iki Trail Guide states: “An old jeep powered the trolley system. Workers suspended a steel cable from a tripod on the crater rim to an A-frame on the crater floor. Rope wrapped around a spool on the rear axle of a Jeep moved the trolley along the cable transporting heavy equipment into and out of the crater.” As you walk to the next stop you will cross a deep crack formed during the collapse of the crater ~500 years ago. Please stay on the trail to view this crack. Stop D4-30-3. Lava fountain spatter, NPS Marker 3. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263665E, 2148465N During 1 of the 17 fountaining episodes that characterized the 1959 eruption the fountain was deflected to the north by slabs of rock partially blocking the vent. During a 20-minute bombardment the surrounding forest and this spot were completely denuded of vegetation by spatter with blobs up to 3.28ft (1m) in diameter. The surface of the trail here is lumpier than the rock that was previously walked over due to the lumpy spatter surface. This location also provides an excellent view of the cinder cone. 69 �Stop D4-30-4. Byron Ledge and Pu‘u Pua‘i; NPS Marker 4. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263515E, 2148565N This stop is <30ft (10m) past the junction with the Crater Rim trail and provides another view of the cinder cone (which will be off to the left) as well as Uwēaloha (Byron Ledge, see Figure D4-9) which is a tree-covered ridge that separates Kīlauea Iki Crater from the eastern floor of Kīlauea Caldera. In the distance, to the west, you should be able to see the buildings of the HVO and Jagger Museum perched on the western rim of the caldera and, on a clear day, the massive bulk of Mauna Loa rising up in the distance Byron Ledge Figure D4-9: Byron Ledge in the middle distance in December 2008. Photo credit: A.D. MacTavish (2008). Stop D4-30-6. Byron Ledge trail junction, NPS Marker 6 (NPS Marker 5 was skipped). Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263142E, 2148392N. From this location on Byron Ledge the unobstructed view is to the east along the axis of the crater and the floor of the lava lake. There is also a good view of the western flank of the Pu‘u Pua‘i cinder cone and north flank slump scars. These slump scars formed when over-steepened slabs of congealed spatter occasionally broke loose during the eruption, slid down the side of the cone, and exposed the hot interior of the cone. Due to irregular steps placed on the trail by the NPS the trail from this point to the crater floor is steep and uneven and makes several switchbacks across the slope as it descends. Please proceed slowly and be careful of your footing. The trail becomes slippery when wet. 70 �Stop D4-30-7. ‘Lava subsidence terrace’; NPS Marker 7. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263207E, 2148344N. Here the base of Byron Ledge is also near the base of a ‘lava subsidence terrace’ (or ‘volcanic bathtub ring’) that marks the high mark of the lava lake at 50ft (15m) above the present surface (see Figure D410). During the eruption the lake occasionally filled higher than the vent causing the fountains to stop erupting. Lava often drained back into the vent dragging pieces of the lake’s crust with it. This drainback was often up to 4 times faster than during eruption and formed a noisy lava whirlpool. Figure D4-10: Lava subsidence terrace (volcanic bathtub ring) near the western crater floor at the base of the Byron Ledge. Photo credit: A.D. MacTavish (2020). Stop D4-30-8. Base of cinder cone slump, NPS Marker 8. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263317E, 2148247N. Here, the rock (see Figure D4-11) looks like jumbled ‘a‘ā flow; however, it formed when welded spatter, formed during the lava fountaining episodes, broke apart as it slid down the side of the cinder cone during a collapse (slump) on the cone’s north flank, which looms over the trail immediately to the south. Figure D4-11: Base of cinder cone slump on the floor of the crater marked by the darker lava. Photo credit: A.D. MacTavish (2020). Photo taken during a rain shower (forming water drops on the camera lens) which are common during winter afternoons. 71 �Stop D4-30-9. Western lip of the main 1959 eruptive vent; NPS Marker 9. NPS Kīlauea Iki Trail Guide. • Approximate UTM 263485E, 2148193N. Here the trail slopes down to the western edge of the main eruptive vent which erupted 17 times during the 26-day eruption. The NPS field guide states that: “Each eruptive episode played out differently. Some went on for days, while others only lasted for hours. Molten rock sometimes poured from the vent in a rolling boil. At other times lava burst skyward to form towering fountains in a matter of seconds. Every episode ended with lava draining back into the vent”. Stop D4-30-10. Buckled lava lake crustal plates, NPS Marker 10. Data: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263535E, 2148257N. Here, the surface of the crater shows evidence of the ~50ft (15m) collapse of the lava lake’s crust after the eruption ceased and the level of the lava lake dropped. As the lake drained back into the vent the lava level dropped and the rigid crust of the lake buckled and cracked creating the uneven rocky ridges observed here (see Figure D4-12). The floor of the crater continues to subside at approximately 2cm/year due to ongoing cooling and contraction of the crystallized interior of the >60yr old lava lake. Figure D4-12: Buckled and uneven crustal plates formed during subsidence of the lava lake after the eruption ended. Photo credit: A.D. MacTavish (2019). Please note that the rain obscuring the crater wall in the distance near the east end of the crater is a common occurrence during the afternoon. Stop D4-30-11. Raised terraces, NPS Marker 11. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263777E, 2148155N Flanking the trail are several raised, 13 to 16ft (4 to 5m) high ‘terraces’ formed when blocks of the cinder cone slumped into the lava lake and were slowly rafted away as floating cinder islands by the molten lava gushing from the active vent (see Figure D4-13). Every time the lake level rose the terraces were covered by lava; however, when the lava drained back into the vent the blocks were again exposed above the surrounding lake surface. Steam is often seen rising from the terraces and cracks in the crater floor due to the heating of rainwater by the still cooling interior of the lava lake. Safety Note: Be very careful approaching any of the escaping steam since it is extremely hot and dangerous. 72 �Hikers for scale Raised lava terraces Figure D4-13: Raised terraces formed from slumped blocks of the cinder cone that were rafted away by gushing lava. Photo taken from the south rim of the crater. Photo credit: A.D. MacTavish (2011). Stop D4-30-12. Crustal overturn plates, NPS Marker 12. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 263993E, 2148115N. During the eruption a dark solid crust rapidly formed on the surface of the lava lake. This crust readily broke into plates between 10 and 20ft (3 to 6m) across. The cracks between the plates were filled by less dense lava rising from beneath the crust and oozing over and eventually covering the rigid plates and swallowing them back into the lake. This process is referred to as ‘crustal overturning’ and moved across the entire lake in a matter of minutes and continued for about a week after the eruption ceased. The existing surface at this stop was formed by the final crustal overturn (see Figure D4-14). Figure D4-14: The final crustal overturn plates forming the surface of the 1959 lava lake after the eruption was over are shown in these 2 photos. The left photo is a view to the west and shows the cinder cone and Byron ledge in the distance. The right photo looks east with the 2020 field trip participants for scale and the eastern crater wall in the distance. Photo credits: A.D. MacTavish (2019, left; 2020, right). 73 �Stop D4-30-13. Lava lake drill holes; NPS Marker 13. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 264080E, 2148095N. The lava lake was first drilled in early 1960, 4 months after the eruption ended on December 20, 1959. That hole terminated when it hit molten lava at a depth of 9ft (2.7m). Later drilling showed that the crust, as one would expect, grew thicker with time (see Figure D4-15). The last hole, drilled in 1988, intersected only traces of residual intercumulus melt occurring between 240 and 330ft (73 to 100m) depth. The interior of the lake is now solid, but is still hot. The 1988 hole also showed that the pre-1959 crater floor had dropped during the eruption with the present base of the lake at about 440ft (135m) depth. The visible collars of the drill holes are located between 60 and 175ft (20 to 50m) north of NPS Marker 13. Figure D4-15: On the left is a cross-section of the lava lake showing the crystallization stages as determined from the regular drilling of the lake between 1959 and 1988. Figure taken from the NPS Kilauea Iki Trail Guide. The photo on the right shows the collars of 3 drill holes of differing vintages. Photo credit: A.D. MacTavish (2019). Stop D4-30-14. Eastern bathtub ring and plants revegetating lava; NPS Marker 14. Data source: NPS Kīlauea Iki Trail Guide. • Approximate UTM 264548E, 2147955N. The eastern bathtub ring of the crater (see Figure D4-16) shows the highest level up the crater walls that Drill hole collars the lava lake attained during the eruption and dramatically shows the approximately 50ft (15m) drop of the surface of the lava lake. The ring in this location is better preserved than that present at the western end of the crater, as was previously observed at Stop D4-26-7 (above). Please note the ‘ōhi‘a trees and other plants that are revegetating the lava. These plants have taken root in cracks where moisture and nutrients have collected. Given enough time the whole of the crater floor will be forested as it was before the 1959 eruption commenced. Also, keep an eye on the sky so that you may be lucky enough to see an ‘io (an endangered Hawai‘ian hawk) soaring on the updrafts above the forested walls of the crater. The trail ahead switchbacks up the eastern wall of the crater from this point. Once at the top of the crater rim the trail connects with the Thurston Lava Tube (Nāhuku) parking lot. The lava tube was closed during the 2020 field trip, but reopened shortly thereafter. At the northern end of the lava tube parking lot another trail leads to the northwest along the crater rim and ends at the Kilauea Iki trailhead parking lot where the trail began. 74 �‘Bathtub Ring’ showing 15m drop ʻōhi‘a tree Figure D4-16: Eastern ‘bathtub ring’ and small ‘ōhi‘a trees. Photo credit: A.D. MacTavish (2019). Stop D4-31. Pu‘u Pua‘i Overlook; Devastation Trail Trailhead. Data source: NPS website; Hazlett (2014); Hazlett and Hyndman (2007). • UTM 263755E, 2147885N; parking area. The Pu‘u Pua‘i Overlook provides a good view of Kīlauea Iki crater from the southern rim. The old road at the west end of the overlook is buried under the Puʻu Puaʻi (Gushing Hill) cinder cone. The Devastation Trail leaves from the western side of the parking lot for both the overlook and the trail. The trail is an easy 30 minute, approximately 1/2mi (800m) return walk, crossing the eastern flank of Pu‘u Pua‘i above the southern Kīlauea Iki crater rim. It initially moves northwest towards Pu‘u Pua‘i and then heads southwest along the eastern edge of the cone and then eventually south through a forest devastation zone caused by the rain of spatter and cinder during the 1959 eruption. Reticulite, cinder, spatter, and ash were blown from the fountains to the southwest by tradewinds into the forest, which was stripped of leaves or buried for about 2.5mi (4km) downwind. The pumice cinders that fell to the surface of the cone close to Kīlauea Ikiʻs fountaining were hot enough to weld themselves together (see Figure D4-17, left). Further downwind, the falling material cooled sufficiently to form a blanket of cinders (see Figure D4-17, right). The zone of devastation is shown graphically in Figure D4-18. Skeletal tree trunks (Figure D4-17, right) and small tree moulds can be observed in the welded spatter. Few of the skeletal tree trunks remain; however, it can be readily observed in Figure D4-17 (right) that these remnants provide a somewhat stable location for seeds to collect that eventually allowed grasses and then new trees to grow. 75 �Figure D4-17: Welded spatter, eastern flank of the Pu‘u Pua‘i cinder cone (left) and partially revegetated devastation zone (right). Photo credits: A.D. MacTavish (left, 2012; right, 2020) Figure D4-18: Distribution of the tephra blanket that formed the Pu‘u Pua‘i cinder and spatter cone, the downwind forest devastation zone, and lava from Kīlauea Iki during the 1959 eruption. Figure taken from Hazlett (2014, p.71). 76 �Stop D4-32. Keanakāko‘i Crater and Vicinity. Data Source: Hazlett (2014); NPS website. • Approximate UTM 263775E, 2147530N; junction between Crater Rim Drive and Pu‘u Pua‘i Overlook Access Road. The portion of Crater Rim Drive west of the junction with the Pu ‘u Pua‘i Overlook access road was closed when visited by the 2020 field trip group due to 2018 earthquake damage, and at the time of final writing (February 2023) was still closed. The Park plans to reroute the road, but construction has not yet begun. The 4 sub-stops that comprise Stop D4-32 (shown in Figure D4-19) are accessed on foot along Crater Rim Drive west from where the Pu ‘u Pua‘i Overlook access road joins Crater Rim Drive. 32-3 32-2 32-4 32-1 Figure D4-19: Location of field stops in the Keanakākoʻi Crater area. Figure taken from Hazlett (2014, p.73). Stop D4-32-1. Keanakāko‘i Crater Overlook (south side of Crater Rim Drive). Data sources: Hazlett (2014); Hazlett and Hyndman (2007); Robinson (2012). • Approximate UTM 262072E, 2146835N. The 115ft (35m) deep, 1500ft (460m) wide Keanakāko‘i Crater (see Figure D4-20) is a collapse pit that is the westernmost crater of a crater chain that defines Kīlauea Volcano’s Upper East Rift Zone. It is thought that most of the craters comprising the chain are <1000 years old and that many formed during the great 1790 eruption. Pit craters like this form by collapse after the underlying magma drains away, which is essentially a smaller version of the caldera collapse observed at Halema‘uma‘u in 2018. If not for lava infill from the eruptions of 1877 and 1974 the crater floor would be funnel-shaped and at least 394ft (120m) deep. Prior to the 1877 eruption the early Hawai‘ians quarried dense, fine-grained basalt from the crater to make stone tools. The quarry site was buried by lava during the 1877 eruption and was further covered by tephra from the 1959 Kīlauea Iki eruption. On a clear day, good views of the Mauna Loa summit, to the west, are afforded from this area. 77 �Figure D4-20: Keanakākoʻi Crater from the viewing area. Photo source: Corine Michel Giron on Google Earth. Stop D4-32-2. Spatter rampart across from Keanakāko‘i Crater Overlook. Data sources: Hazlett (2014); Hazlett and Hyndman (2007); Robinson (2010, 2012). • Approximate UTM 262130E, 2146903N. North of the Keanakākoʻi Crater overlook parking area on the north side of Crater Rim Drive is a small spatter rampart from which issued a thin pāhoehoe flow that in July 1974 crossed the road and spilled over the Keanakāko‘i crater rim in a cascade onto the floor of the crater. Stop D4-32-3. Drainage channel. Data sources: Hazlett (2014); Hazlett and Hyndman (2007); Robinson (2010, 2012). • Approximate UTM 262060E, 2146935N; the viewing area is accessed via a short trail on the north side of Crater Rim Drive, starting a short distance west of the northern parking area. This location allows inspection of another spatter rampart and eruptive fissure formed during the 3-day July 1974 eruption (see Figure D4-21). Lava from the fissure flowed down a drainage channel to the west where, as it rounded a bend, it washed up against the channel wall as it flowed towards, and into, the pre-2018 caldera floor. Stop D4-32-4. Keanakākoʻi Ash deposit. Data sources: Hazlett (2014); MacTavish (2020); White and Houghton (2000). • Approximate UTM 261997E, 2146795N. This roadcut stop is ~100m southwest of the Keanakākoʻi Crater overlook on the northwest side of the road. It hosts well-bedded, variably sorted, graded, weakly indurated, locally cross-bedded pyroclastic rocks of Keanakākoʻi Ash that accumulated over ~130 years during the 17th and 18th centuries. The unit at this site consists of layers of tuff, lapilli tuff, and localized lapillistone containing variable amounts of ash- and lapilli-size fragments (see Figure D4-22, left). The rocks locally contain accretionary lapilli and angular to subangular volcanic bombs (Figure D4-22, right). Accretionary lapilli are spherical aggregates (commonly with a concentric structure) formed by the accretion of moist ash in eruption clouds. They are equivalent to volcanic hailstones. The one observed cross-bedded interval is probably a base surge deposit. The bombs observed are pieces of pre-existing rock blown out of the vent during the eruption. 78 �Spatter Rampart 2018 Halemaʻumaʻu Crater Eruptive Fissure Figure D4-21 Spatter rampart and eruptive fissure formed during a 3-day, July 1974 eruption as viewed from field trip stop D4-28-3. The new 2018 Halemaʻumaʻu Crater is in the background. Photo source: A.D. MacTavish, 2020. Volcanic Bomb Accretionary Lapilli Lapillistone Figure D4-22 Well-bedded Keanakākoʻi ash deposits consisting of ash, lapilli, accretionary lapilli, and volcanic bombs. The left photo shows variably-bedded, variably sorted pyroclastic layers with an angular bomb at left centre (pencil magnet for scale in lower centre). The right photo shows a layer containing accretionary lapilli (volcanic hailstones). Photo sources: A.D. MacTavish, 2020. 3.4. Day 4 (Part 2): Koaʻi Fault Zone and Hilina Pali According to Hazlett (2014) the Hilina Pali Road allows access to the finest examples of faulting on the Hawai‘ian islands and allows examination of the fresh, well-exposed, easily approached escarpments comprising the Koa‘e Fault Zone. Robinson (2010) describes the fault zone as a series of nearly parallel, NE-SW-trending normal faults with a surface length of 9mi (14.5km) and a width of approximately 1.6mi (2.6km). At the end of the road is Hilina Pali which is part of an extensive network of faults that downdrop the southern flank of Kīlauea Volcano (Hazlett, 2014). The road was closed during the 2020 field trip due to road-surface damage by fault down-drop along the Koa‘e Fault Zone during earthquakes associated with the 2018 eruptive events. The damage has since been repaired and the road is now open to vehicular traffic. 79 �The Day 4 (Part 2) Field Trip Stops are (see Figure D4-23): 33. 34. 35. 36. 37. 38. Koa‘e Fault Zone; can easily view fault scarps on left during drive Kulanaokuaiki Pali (Koa‘e Fault Zone); Kulanaokuaiki Campground; rest stop; eastern end of Mauna Iki Trail. Kulanaokuaiki Pali; brown ropey basalts at top of scarp. Tumuli and brown ropey basalts. Hilina Pali (also a rest stop). 33 35 34 36 37 38 Figure D4-23: Locations of Day 4 (Part 2) field stops. Map taken from National Geographic Hawaiʻi Volcanoes National Park Illustrated Trails Map (2010). 80 �Stop D4-33. The Koa‘e Fault System. Data source: Hazlett (2014). • UTM 264210E, 2142940N; small parking area on left with enough room for 1, possibly 2 vehicles. A short distance west of Mauna Ulu the Upper East Rift Zone bends sharply to the northwest from a general east-northeast trend (Hazlett, 2014). The Koa‘e Fault System is a 9mi (14.5km) long, ~1.2mi (2km) wide segment (extension?) of the East Rift Zone and forms a series of small grabens (see Figures D4-24 and D4-1). What is unusual with this fault system is that elsewhere on the south flank of Kīlauea the faults are predominantly south-dipping, whereas within the Koa‘e Fault System both north-dipping and south-dipping faults are common, resulting in the formation of grabens. Easily visible to the south at this stop is an excellent example of a north-facing fault bounding the southside of one of the grabens. This particular fault trends subparallel the road for the next 0.75mi (1.2km). Fault Trace Fault Trace Field Stop 33 Figure D4-24: Koaʻe Fault system on either side of Stop D4-33. Google Earth satellite image. Stop D4-34. Entrance to Kolanaokuiki Campground. • UTM 261190E, 2140295N; park on south side of the road opposite to the campground entrance where the ground is solid enough to take the weight of most vehicles. The ~30m high north-facing scarp of Kulanaokuaiki Pali is very visible, easily reachable, and is located ~165ft (50m) south of the campground entrance. Stop D4-35. Kolanaokuiki Campground. • UTM 261125E, 2140430N; parking lot. This campground provides a potential rest area with washrooms and picnic tables, if required. Stop D4-36. Kolanaokuiki Pali. Data source: Hazlett (2014). • UTM 260966E, 2140228N; the only parking is on the top of the slope on the south (right) side of the road ~330ft (100m) south of the top of the Pali at UTM 261042E, 2140173N. • The parking lot also services the Mauna Iki Trailhead which is located 165ft (50m) westnorthwest of the parking lot at UTM 260995E, 2140190N, about halfway to the pali. At this stop the road curves around and up a short slope near the western terminus of the north-dipping Kolanaokuiki Pali normal fault. 81 �This fault marks the southern margin of the Koa‘e Fault System in this area, where in December 1965 the road was vertically offset 8.25ft (2.5m) during a major faulting episode (see Figure D4-25, left). Hazlett (2014) states that: ‘…along the base of the north-facing scarp the crust of the down-dropped block alternates between monoclinal up-warps and fissured down-warps, referred to as ‘rollovers’ (see Figure D4-24, right). The base of some of the monoclinal up-warps show vertical fracturing or thrust buckles, with the up-warps shoved onto the down-dropped blocks. Rollovers are indicators of listric (curved) fault planes and in this case the controlling fault plane may curve northward at depth’. Figure D4-25: The left photo shows Kolanaokuiki Pali, looking east-northeast along one of the north-facing escarpments of the Koaʻe Fault system. The diagram on the right explains rollovers, monoclines, and listric normal faulting. Figure taken from Hazlett (2014, p.119). Photo source: Google Earth. Stop D4-37. Tumuli and ropey pahoehoe flows. Data source: Hazlett (2014); Hazlett and Hyndman (2007); Robinson (2012). • UTM 259720E, 2137885N; carefully park along the right (north) side of the road where the road shoulder is slightly wider than elsewhere; be very careful of other road traffic. The pāhoehoe flows here were erupted in the 13th or 14th Century from the summit shield of Kīlauea. The area is characterized by good examples of ropey and entrail pāhoehoe flows and well-developed tumuli (see Figure D4-26, left). The reddish weathering of the flows indicates age despite the lack of vegetation. Tumuli is the plural form of tumulus which is described by various authors as (Figure D4-26, right): • • ‘a steep mound from a few feet to tens of feet across in a pāhoehoe flow that may form from molten lava heaving up plates of chilled crust or from subsiding flow crust draping over bedrock obstacles’ (Hazlett and Hyndman, 2007). ‘elliptical, domed structures that form on the surfaces of pāhoehoe flows extruded on flat or gentle slopes. Tumuli form when the upward pressure of slow-moving lava swells or pushes the overlying solidified crust. Sometimes the lava can drain away and leave a hollow shell’ (Robinson, 2012). 82 �Figure D4-26: Good examples of ropey flows are in the photo foreground with several tumuli in the background. Photo source: Google Earth. The diagram on the right (from Hazlett 2014, p.125) shows tumulus formation by focussed hydrostatic pressure with molten pāhoehoe beneath a cooling, thickening crust. Stop D4-38. Hilina Pali Overlook. Data source: Hazlett (2014). • UTM 257575E, 2135150N; overlook parking lot. Walk ~50ft (15m) down the trail from the overlook shelter to a triangulation station and small memorial. Hazlett (2014) states that: “Hilina Pali is a part of the extensive network of faults that drop the south flank of Kīlauea seaward in stepwise fashion. The Hilina Fault itself, with a throw of about 1150ft (350m), dips 50-60o to the south, possibly flattening out at depth within the Kīlauea volcanic pile, though this interpretation is controversial. Easton and Garcia (1980) estimated that the fault system has been active for at least 20,000 years”. Hazlett (2014) also states: “Look to the east of this location (see Figure D4-27) to see the dark, freshlooking lava flows from the Mauna Ulu eruption cascade over Poliokeawe Pali. On the downthrown fault block, 1600ft (500m) below, meandering stream channels, lava fans, and talus piles are exposed. The lava fans were formed as lava from the Kālu‘e eruptions piled up at the base of the pali.’ Figure D4-27: View looking east from the Hilina Pali Overlook. Photo source: Google Earth. 83 �3.5. Day 5 – Chain of Craters Road, Napaū/Mauna Ulu Trail, Hōlei Pali The Chain of Craters Road extends for 18.3mi (30km) from Crater Rim Drive to the Puʻu Ōʻō-Kupaianaha lava field. The road follows the northern part of the Upper East Rift Zone, descends the southern flank of Kīlauea Volcano, winds down Hōlei Pali, crosses the southern coastal plain, and then works east along the coast to where it is truncated by the Puʻu Ōʻō-Kupaianaha lava field about ~0.6mi (1km) east of the Hōlei Sea Arch. Numerous volcanic features are present along the route and Field Trip Day 5 will observe or visit many of them. The field stops on the Chain of Craters Road are shown in Figure D5-1 and consist of 10 stops on, or adjacent to, the road as well as various sub-stops associated with the Napaū and the Puʻu Loa Petroglyph trails. The planned stops are: 39. 40. 41. 42. 43. 44. 45. 46. 47. 48. 49. Luamanu Pit Crater and Lava Trees; Pauahi Crater; Mauna Ulu/Napaū (trail); Mau Loa o Mauna Ulu (Alternate;) Muliwai o Pele (Lava River); Kealakomo Overlook; Pāhoehoe transitioning to ‘a‘ā; Alanui Kahiko; Pu‘u Loa Petroglyphs (trail) (Stops D5-47-a and b); Coastal Sea Arches; and Puʻu Ōʻō-Kupaianaha lava field (mid-1980’s). 39 40 41 42 43 44 45 46 47a,b 48 49 9 Figure D5-1: Day 5 of the field trip consists of 10 stops along, or near, the Chain of Craters Road. Figure modified from Hazlett and Hyndman (2007, p.80). 84 �3.5.1. The Formation of Mauna Ulu: This sub-section is summarized from Hazlett (2014, p.88-89). The 5-year Mauna Ulu eruption began on May 24, 1969 when a lava curtain erupted along the length of a new, east-northeast-trending fissure extending from south of Pauahi Crater (near the present highway and the start of the Nāpau Trail) through ‘Ᾱloi Crater to the north of ‘Alae Crater. The fountaining soon became focused in the area between ‘Ᾱloi and ‘Alae craters where Mauna Ulu presently stands. Over the next 5 years episodes of sustained fountaining from this vent, sometimes reaching heights of 1770ft (540m), and associated overflows were interspersed with lava lake activity. By the end of 1970 both ‘Ᾱloi and ‘Alae craters had been completely infilled with molten lava. Flows piling up around the vent formed a low shield that was the beginning of the Mauna Ulu edifice. A lava tube from the top of the shield fed the ‘Alae lava lake which, in turn, fed long-lasting flows that extended down the southern flank of Kīlauea. These flows reached and entered the ocean in June 1969 and September 1970. Mauna Ulu’s central crater extended northeastward in 1971 by merging with a series of small secondary pits on the flank of the shield, forming a trench. Lava continued to be fed to the ‘Alae lava lake and from there down Kīlauea’s southern flank. Lava again entered the sea between March 8 and May 25, 1971. In late May 1971 the lava lake within Mauna Ulu began to subside; lava stopped flowing from ‘Alae in July; and lava completely disappeared from view in Mauna Ulu in October 1971. It was initially thought that the eruption had ended; however, lava returned to the summit crater in February 1973 and the ‘Alae lava tube reactivated soon afterward. After a month the level of the Mauna Ulu lava lake dropped a few metres and new vents opened at Mauna Ulu and ‘Alae. A levéed lava lake formed at ‘Alae and fed numerous overflows which gradually formed the low ‘Alae satellite shield. The ‘Alae flows were active for more than a year with some of the far-travelling flows pouring into and completely filling the western end of Makaopuhi Crater. Other flows eventually formed lava tubes that reached the ocean between August and October 1972 and from February to May 1, 1973. An M6.2 earthquake on April 26, 1973, centred north of Hilo, may have triggered a drastic change at Mauna Ulu, where, on May 5, 1973 the lava completely drained from both Mauna Ulu and ‘Alae. A few hours later a fissure eruption began at nearby Hi‘iaka and Pauahi craters that lasted less than a day. The Mauna Ulu summit lava lake began to refill 2 days later with lava returning to ‘Alae at the end of May. The last lava to be observed at ‘Alae was on June 7th. Sluggish activity continued at the summit crater until September when a gradual increase began and by early November there was vigorous fountaining accompanied by overflows. On November 10, 1973 the lava lake suddenly drained away at the beginning of another short rift eruption at Pauahi Crater and did not return until a month later. Strong activity within the lava lake was again present in January 1974 and a steep-sided spatter cone grew within the lava lake. On January 24th a series of vigorous fountaining episodes began with fountains attaining heights of 130ft (40m) and overflows rapidly travelled several miles from the vent. This irregular pattern of fountaining and overflow lasted for 5 months and by June had built the Mauna Ulu shield an additional 100ft (30m) to a height of 390ft (120m). Activity in the lava lake became sluggish in June 1974. Harmonic tremors and deflation were recorded near the summit of Kīlauea on the morning of July 19th, the Mauna Ulu lava lake level dropped, and a day long fountaining eruption began at Keanakāko‘i Crater within Kīlauea’s summit caldera . This marked the end of the Mauna Ulu eruption. Figure D5-2 shows the lava flows and the year they were erupted from Mauna Ulu and vicinity between May 1969 and June 1974. Figure D5-3 shows the Mauna Ulu area pre-eruption and post-eruption. 85 �Figure D5-2: Geology associated with the 5-year Mauna Ulu eruption. Figure source: Easton and Easton (1995); the circled numbers are stops from the NVO Chain of Craters Road guide; the letters are stops from the NPS Napaū/Naulu Trail Guide; and the geology is from Holcombe (1976 and 1987). 86 �Figure D5-3: The Mauna Ulu area pre-eruption (top half of sketch) in January 1969 and post-eruption (bottom half of sketch) in August 1974. Map from Hazlett (2014, p.94). Stop D5-39. Lua Manu Pit Crater, Lava Trees. Data Sources: Hazlett (2014); Hazlett and Hyndman (2007), Robinson 2012. • UTM 263360E, 2146520N; overlook parking area. The small Lua Manu Pit Crater is thought to be the uppermost East Rift Zone crater along the Chain of Craters Road. The 328ft (100m) wide crater was once tree-filled; however, on July 19, 1974 a fissure eruption a short distance north and east of the pit partially infilled the crater to a depth of ~50ft (15m). After the eruption ceased much of the lava drained back into a fissure in the crater’s east wall. The highlava level is easily seen as prominent ‘bathtub ring’ on the crater walls. Spatter ramparts marking the locations of the eruptive fissures are visible on both sides of the Chain of Craters Road. Many fragile lava trees and tree moulds are present on the flows north of the crater. Many of the lava trees have collapsed since the eruption; however, drain-back features and remnant charcoal are preserved on the edges and within some, respectively. Lava quickly chills against the moist, tough ōhi‘a hardwood trees as it moves around them. There are numerous preserved, unburned ʻōhi‛a trunks that protrude from the lava trees or have toppled onto the surface of the flow. Safety Warning: Do not attempt to climb or lean against the lava trees since they are fragile and topple easily. Also, the pāhoehoe flow surfaces are brittle and are underlain by large voids, so be very careful where you walk. 87 �Stop D5-40. Pauahi Crater Overlook. Data sources: Hazlett (2014); Robinson (2012). • UTM 266230E, 2143410N; large parking area located on the left (east-side) of the highway and reached by 2 short access roads at the north and south ends of the parking lot. Pauahi Crater is the largest of the easily reached East Rift Zone pit craters. It is a composite double-pit crater about 1650ft (500m) long and 360ft (110m) deep. A small unnamed pit crater, a short distance east of the main pit can be seen in in the upper left corner of Figure D5-4. Pauahi Crater was the site of 3 eruptions in the 1970’s: 1. May 5, 1973: A small amount of lava was extruded onto the crater floor, but ceased when the eruption moved north to nearby Hiiaka Crater. 2. November 10, 1973: A fissure opened on the pit floor near the present overlook a few hours after an active lava lake at Mauna Ulu suddenly drained. Lava fountaining was initially confined to the floor of Pauahi but activity soon moved up the eastern and western crater walls and into the adjacent forest to quickly form a 1.8mi (3km) long west-southwest-striking, en-echelon fissure system. This fissure vented lava from near Pu‘u Huluhulu to the east to a short distance west of Pauahi Crater. The bulk of this eruption lasted about 10 hours with lava fountains feeding 2 lava lakes (one in each of the 2 parts of the crater). The individual lakes eventually combined into 1 large lava lake exhibiting huge whirlpools that drained lava out of the lake almost as fast as it was erupted from the fissures. This activity is easily visible from the visitor’s platform as dark lines in the far northeast crater wall and just north of the platform. Broken lava trees are found in the lava flow near the entrance to the parking lot showing the high point of the lava flow. Minor eruptive activity continued in the crater until December 9, 1973. 3. November 16, 1979: This less than 24-hour eruption was preceded by an 11-hour earthquake swarm where the earthquakes migrated upward through the crust from ~1.9mi (3km) to ~0.6mi (1km) depth. It is estimated that 915,000yd3 (700,000m3) were erupted during an initial fountaining stage followed by flows issuing from fissures west of the crater wall cutting the Chain of Craters Road. Most of the floor of Pauahi is covered by a thin layer of 1979 lava. Figure D5-4: Pauahi Pit Crater as seen from the viewing platform. Photo credit: A.D. MacTavish (2012). 88 �Figure D5-5 shows the volcanic features in the vicinity of Pauahi Crater and Mauna Ulu as well as the field trip sub-stops along the Napaū Trail. The 13 sub-stops comprising Stop D5-41 are: 1. Napaū trailhead and trail sign. 2. Eastern exposed end of the initial (1969) Mauna Ulu fissure and its associated spatter rampart partially covered by later Mauna Ulu ‘a‘ā flows. 3. Spatter rampart and lava drain-back pits and hornito (?) field. 4. Broken and toppled lava trees. 5. Small-scale pahoehoe lava channel. 6. Old lava rampart with Mauna Ulu lava ‘bathtub ring’. 7. Large lava tree (possibly a hornito?). 8. Ecological Stop. 9. Lower overlook near summit of Pu‘u Huluhulu. 10. Upper Pu‘u Huluhulu Overlook. 11. Perched lava Pond. 12. Well-developed lava channel. 13. Mauna Ulu Summit 8 11/79 Flow 40 10 9 7 11/79 Flow 1974 pāhoehoe flow 6 11/73 pāhoehoe flow 11 5 12 11/79 Flow 13 1974 ‘a‘ā flow 4 1 41 3 2 1974 ‘a‘ā flow Map from Hazlett (2014; p.84) Figure D5-5: Day 5 field trip stops, sub-stops, and volcanic features in the vicinity of Pauahi Crater, Napaū Trail, Puʻu Huluhulu, and Mauna Ulu. 89 �Stop D5-41. Napaū Trail. • UTM 267210E, 2142705N; large parking lot. Trail access from the present Chain of Craters Road is via a remnant of the original Chain of Craters Road that was buried by a Mauna Ulu flows in 1973. Please note: The brown NPS markers with yellow letters scattered along the Napaū Trail do not match the numbers of the field stops within this portion of the guide. The descriptions within the NPS guide tend to be general rather than specific. Where a field stop has an NPS Marker that number is noted. Stop D5-41-1 (NPS1). Data sources: Hazlett (2014); NPS Mauna Ulu Eruption Guide; Hazlett (2014). • UTM 267326E, 2142700N, Napaū Trailhead; NPS Marker 1. Walk 330ft (100m) east-southeast along the old highway from the parking lot to the Napaū Trailhead exhibit board located ~15ft (5m) north of the old highway. This trail has 2 segments: 1. A loop that leads east, initially along the old highway, then south to a fissure and spatter rampart, and then west allowing a close-up view of the western end of the fissure that was the beginning of the 1969 Mauna Ulu eruption; and 2. A 2mi (3.25km) round-trip trail that leads northeast to the Pu‘u Huluhulu summit, which is a 400 to 600yr old spatter cone located ~1650ft (500m) north-northwest of the Mauna Ulu lava shield. The trail follows 1973 Pauahi flows and flanking 1974 Mauna Ulu flows for most of its length, passes an old undated spatter rampart, and then through a lava tree forest. The Pu‘u Huluhulu cone is truncated by a 165ft (50m) deep collapse crater, the top of which provides a panoramic view that includes (on a clear day): Mauna Loa; Mauna Kea; Kīlauea summit; the ‘Ailā‘au lava shield; Puhimau and Pauahi Craters; Mauna Ulu; ‘Alae and Kanenuiohamo lava shields; Makaopuhi Crater; and numerous recent flows, fissures, and many vent edifices . Stop D5-41-2a. Mauna Ulu Spatter Rampart and Eruptive Fissure. Data source: NPS Mauna Ulu Eruption Guide. • UTM 267430E, 2142590N Walk east-southeast from the trailhead to the end of the paved road. From here a trail leads south (right) to a fissure that opened after a swarm of earthquakes on May 24, 1969 and marked the beginning of the 5-year Mauna Ulu eruption. The fissure first opened at ‘Alae Crater, near where Mauna Ulu now stands, passed through the now infilled Ᾱlo‘i crater, and propagated west-southwest like an opening zipper for over 1mi (1.6km) to near the location of the present Chain of Craters Road. The eruption at this location lasted less than a single day but during that time it spewed a 100ft (30m) high curtain of lava along the entire length of the fissure. Most of the lava moved south and downslope towards the sea, but what fell on the upslope side formed the present spatter rampart. When the eruption ended at 1100PM that evening most of the nearby lava drained back into the fissure and congealed in place as it poured over the rim. Activity continued at the main Mauna Ulu vent at the eastern end of the fissure between ‘Alae and Ᾱlo‘i craters and by December, 1969 had sustained 12 fountaining episodes where fountains sometimes reached heights of 1770ft (540m). Part of the fissure east of this point is covered by a 1974 Mauna Ulu ‘a‘ā flow (see Figure D5-6). 90 �Mauna Ulu Shield Eruptive Fissure 1969 Spatter Rampart 1974 Mauna Ulu ‘A‘ā Flow Top of 1969 Spatter Rampart Figure D5-6: The left photo shows the May 24, 1969 spatter rampart truncated by a 1974 Mauna Ulu ʻaʻā flow with the Mauna Ulu shield in the distance. The right photo shows the 1969 spatter rampart and eruptive fissure, north of where it is truncated by the 1974 Mauna Ulu ʻaʻā flow. Photo credits: A.D. MacTavish (2019). As mentioned above, to the left (east-northeast) of the trail the 1969-vintage fissure and spatter rampart are covered by a very irregular, blocky 1974 Mauna Ulu ‘a‘ā flow that covered the fissure and the rampart. This ‘a‘ā flow commonly contains small, green olivine grains. To get to this stop you walked over relatively smooth, 1973-vintage ropey pāhoehoe lava flows which formed as fluid lava flowing along a relatively flat or gentle slope. This type of flow advances as a series of small lobes and toes that continually expand and break out from the cooling crust along its leading edges. The surface textures of pāhoehoe vary widely; however, the most common is ropey where the numerous folds, wrinkles, and ropes form when the thin partially solidified crust of the flow is slowed or halted. The lava below the crust continues to move forward and drags the malleable crust along. Notice the extreme difference between the 2 types of flow even though they have essentially the same composition. An ‘a‘ā flow forms due to factors such as: lower temperature; gas loss; onset of crystallization; a loss of elasticity, such that it fractures instead of stretches; being forced to move faster, such as being pushed from behind by an upslope surge; or if it has to move down a steeper slope. Stop D5-41-2b. Spatter mound field. Data source: A.D. MacTavish (2020); AGI Glossary of Geology, 4th Edition (1997). • UTM 267370E, 2142546N. Directly adjacent to the spatter rampart and eruptive fissure described immediately above at Stop D541-2a is a small spatter mound field. This field was initially incorrectly identified by AM as a hornito field. The AGI Glossary of Geology, 4th Edition (1997) defines a hornito as ‘a small mound of spatter built on the back of a lava flow (generally pahoehoe), formed by the gradual accumulation of clots of lava ejected through an opening in the roof of an underlying lava tube’. After close examination the hornito interpretation was discarded in favour of spatter mounds that did not build vertically enough to become hornitos. At this location several good examples, up to ~8.2ft (2.5m) in height, are readily observable. The shape of these formations, particularly the one shown in the right photo of Figure D5-7, are reminiscent of mushrooms, or flowerpot structures formed by tidal action in the Bay of Fundy in New Brunswick, where the flare at the top of the flowerpot shows the maximum water level at high tide. In a similar way the flaring at the top of these features defines the maximum height of the host lava flow, upon which the spatter mound was building, before the volume of lava within the tube decreased and the upper surface of the lava flow dropped as eruption volume waned. 91 �Figure D5-7: The left photo shows the spatter mound field at Stop D5-37-2b. On the right is a 2.5m tall, mushroom-shaped mound showing the characteristic shape formed when the host lava flow surface deflates as the volume of lava in the underlying flow decreases. Steve Fox for scale. Photo credits: A.D. MacTavish (2020). Stop D5-41-3. Drain-back pits. Data source: NPS Mauna Ulu Eruption Guide. • UTM 26720E, 2142530N. Walk west along the southern edge of the fissure until you reach a series of small pits that occur along the fissure for about 200ft (60m). These are where lava drained back into the fissure and then solidified when lava fountaining ended. On the inner walls of these pits iron-rich minerals within the lava quickly oxidized in the residual stream of the lava and turned bright red and yellow. The fissure can be traced intermittently west of these pits for about another 825ft (250m). Good photos of the drain-back pits can be obtained from the top of the adjacent spatter rampart located a short distance north of the pits (see Figure D5-8, left). Also at this location is a good example of a welded spatter mound sitting on top of the spatter rampart (see Figure D5-8, right). Walking west from Stop D5-41-3 there may be solidified ejecta (or tephra) of several forms along the side of the trail such as: lapilli-sized pumice (cinders with a profusion of gas bubbles) are the most common; Pele’s Tears (black, teardrop- or sphere-shaped droplets of obsidian) are less common; reticulite, which is a delicate lava foam version of pumice is uncommon; and thin, very fragile and delicate threads of, often golden-coloured Pele’s Hair are rare. There are also some good examples of lava trees (and tree moulds) that can usually be identified as narrow structures projecting above the top of the lava flows. From this stop cross to the northern side of the spatter rampart and follow the trail to the east along the fissure and through the forest back to the northern trailhead, then walk northeast along the trail to the next stop (D5-41-4). 92 �Figure D5-8: A drain-back pit within the May 24, 1969 fissure is shown in the left photo with Dr. Juk Bhattacharyya for scale. The welded spatter mound on the top of the spatter rampart adjacent to the north of the drain-back pits is shown on the right. Photo credits: A.D. MacTavish (2019). Stop D5-41-4. Lava trees and Mauna Ulu and ‘Ᾱlo‘i Shields in the distance. Data source: NPS Mauna Ulu Eruption Guide. • UTM 267358E, 2142763N; Lava Trees located 100ft (30m) west-northwest (left) of the trail. There are numerous, usually toppled or broken lava trees northwest (left) of the trail at this location. The lava trees in this area are sometimes associated with pieces of the original ōhi‘a trees that formed the mould. Looking east from this point, across the 1973 Pauahi flows and the ‘a‘ā flow in the middle distance, you can readily see the summits of Pu‘u Huluhulu (tree-covered) and Mauna Ulu, and, to the right of Mauna Ulu the dimpled lava mound of the ‘Ᾱlo‘i Shield. The ‘Ᾱlo‘i Shield, which grew above the pre-1970 ‘Ᾱlo‘i Crater, and the ‘Alae Crater to the south were completely infilled as lava overflowed the growing Mauna Ulu shield. As lava flowed over the rim of the crater it formed several approximately 80ft (24m) high lava falls. The resulting lava lakes overflowed both craters with each new lava surge from Mauna Ulu. This process added layers and elevation to the surrounding terrain and eventually produce a low, dimpled, readily visible lava mound that grew above the older crater. Figure D5-9 (taken from the NPS Mauna Ulu Eruption Guide) diagrammatically illustrates the formation of the Mauna Ulu and the ‘Ᾱlo‘i Shields. 93 �Figure D5-9: This illustration, taken from the NPS Mauna Ulu Eruption Guide, diagrammatically illustrates the formation of the Mauna Ulu and the ‘Ᾱlo‘i Shields. Stop D5-41-5. Small-scale lava channel, lava trees, and tree moulds. Data sources: MacTavish (2019); NPS Mauna Ulu Eruption Guide. • UTM 267554E, 2143030N The trail here follows a small-scale lava channel preserved on the surface of a pāhoehoe flow from the 1973 Pauahi eruption (see Figure D5-10). Look for evidence of flowage around the lava trees with a lava crust on the north side of the trees and the flowage of lava ropes around the trees. In some instances, tree moulds and lava trees can be used as flow direction indicators with an obvious asymmetry where the upstream edge of the tree is rounded and the downstream portion is roughly pointed (see Figure D5-11). 94 �Figure D5-10: Small-scale pāhoehoe lava channel flowing down the slope. The trail follows the channel in this location. Photo credit: A.D. MacTavish (2019). Upstream Side Downstream Side Upstream Side Downstream Side Figure D5-11: The right photo dramatically illustrates a lava tree/tree mould in the process of formation with a rounded upstream portion and a pointed downstream portion. These features are preserved in the tree mould in the left photo at Stop D5-41-5. Photo credits: A.D. MacTavish (left, 2019); NPS Mauna Ulu Eruption Guide (right). 95 �Stop D5-41-6 (NPS-8). 500yr old spatter rampart. Data source: NPS Mauna Ulu Eruption Guide. • UTM 267632E, 2143270N. In the distance in the northeast (well left of the trail) along the forest edge are the spatter ramparts from the 1973 eruptive vents that formed the lava flow presently under foot. The low, vegetated hill to the right of the trail southeast of the NPS #8 marker is an approximately 500yr old spatter rampart that was encroached upon by 1973 Pauahi Crater eruption flows. These 1973 flows built up and inflated against this side of the rampart and were deflected to the southwest along the present path of the trail. When the <1-day long eruption ended the lava drained away and left a black, bathtub ring-like, high lava mark against the rampart (3 to 5ft or 1 to 1.5m, above the present flows). Stop D5-41-7. Large well-developed lava trees. Data sources: NPS Mauna Ulu Eruption Guide; MacTavish (2019). • UTM approximately 267869E, 2143482N. About 25m south of the trail are several very well-developed lava trees (see Figure D5-12). A single larger formation (Figure D5-13), initially thought by AM to be a hornito, was determined by closer inspection to be a large, very well-developed, composite lava tree composed of multiple, closely-spaced lava tree impressions. The interior of this formation exhibits lava drips. Figure D5-12: The left photo shows a large lava tree and the right photo shows a tree mould along the edge of another lava tree. Photo credits: A.D. MacTavish (2019). 96 �Figure D5-13: On the left is a large, complex, well-developed, composite lava tree looking a lot like a hornito, that is composed of multiple, closely-spaced lava trees. The interior of this formation (right photo) exhibits lava drips. Small, silver jackknife for scale. Photo credits: A.D. MacTavish (2019). Stop D5-41-8 (NPS16). Plants and animals of recent lava flows in Hawai‘i. Data sources: Tom Callus, Hawaii Tribune Herald (March 25, 2019); Howarth (1979); NPS Mauna Ulu Eruption Guide; nature.Berkeley.edu. • UTM 268115E, 2143565N; NPS Marker 16. Several animals arrive to inhabit Hawai‘ian lava flows within a few months of their formation. The first two are a wingless, soundless cricket (Caconemobius fori, or ‘ūhini nēnē pele in Hawai‘ian; see Figure D514, left) and a large wolf spider (Lycosa sp.) that feeds on the crickets (Figure D5-14, right). Hawai‘an lava crickets are the first to arrive on new lava flows, are found nowhere else in the world, and eat decaying plants swept in by the wind. They were not documented in scientific literature until 1978, four years after scientists from the Bishop Museum in Honolulu discovered them on new Kīlauea lava fields. They abandon the flows once they are covered by vegetation and move on to younger flows, or they will die out. Soon after the crickets and spiders arrive come the algae, ‘ōhi‘a trees, lichens, and mosses – in that order. Within 15 years small shrubs are growing in cracks and the original plants are joined by pūkiawe, ‘a‘ali‘i, kūpaoa, and ‘ōhelo. Some species, like pāwale, only grow on the active volcanoes of the island of Hawai‘i and they disappear as other plants crowd them out. Figure D5-14: On the left is Caconemobius fori, a wingless, soundless cricket which is the first animal to colonize fresh Hawai‘ian lava flows. On the right is Lycosa sp., a large wolf spider that eats the crickets and is the second animal to colonize fresh Hawai‘ian lava flows. Photo credits: nature.berkeley.edu. 97 �Stop D5-41-9. Lower Pu‘u Huluhulu Overlook. Data source: NPS Mauna Ulu Eruption Guide. • UTM 268364E, 2143353N. This spot provides a good unobstructed view of Mauna Ulu and a perched lava pond located on the northern flank of the shield mid-way between Pu‘u Huluhulu and Mauna Ulu and partway up the northern slope of the shield (see Figure D5-15). Perched Lava Pond Mauna Ulu Summit Figure D5-15: Photo of Mauna Ulu and the perched lava pond from the lower Pu‘u Huluhulu viewpoint. Photo credit: A.D. MacTavish (2019). Stop D5-41-10 (NPS 14). Pu‘u Huluhulu Overlook. Data source: NPS Mauna Ulu Eruption Guide. • UTM 268382E, 2143376N; NPS Marker 14. Before the Mauna Ulu eruption the overlook provided a view of the forest that surrounded Pu‘u Huluhulu on all sides with the southern horizon view unobstructed down to the Pacific Ocean (see Figure D5-16, right). The overlook was constructed in 1934 at an elevation that was about 300ft (94m) above the surrounding land surface. When the eruption ended in 1974 the unobstructed view was blocked by the Mauna Ulu shield (Figure D5-16, left). During the eruption the overlook was the primary viewing platform for HVO scientists. The Mauna Ulu summit crater is >100ft (30m) deep. On a botanical note: The Pu‘u Huluhulu Crater protects rare native plants, like the ‘ōhā (see Figure D517) which is rare outside the crater, from feral pigs that roam the nearby forests. The other end of the platform provides a view of the summits of Mauna Loa and Mauna Kea in the far distance and in the middle distance is the caldera at the summit of Kīlauea. The brass plaque located at the northwest corner of the overlook was installed in 1934 and points to the volcanic features visible at that time such as ‘Alae and ‘Ᾱloi craters, which no longer exist, but does not point to Mauna Ulu, which did not then exist. 98 �Perched Lava Pond Tourist blocking the view Mauna Ulu Pacific Ocean Figure D5-16: The left photo shows a modern Google Earth image of the view south of the viewpoint at the summit of Pu‘u Huluhulu. The right photo is a 1969 image that provides a similar view as the modern photo on the left. Please note the steam rising from fissures near left centre and no Mauna Ulu shield, or tourists, blocking the view. Right photo credit: NPS Mauna Ulu Eruption Guide. Figure D5-17: Photo of a rare ʻōhā wai nui or Clermontia Hawaiiensis which grow within the Pu‘u Huluhulu crater and are therefore protected from the feral pigs roaming the surrounding forests. Photo credit: Google Earth. Stop D5-41-11 (NPS 13). Perched lava pond. Data source: NPS Mauna Ulu Eruption Guide; Easton and Easton (1995). • UTM 268499E, 2143189N; breach in southwestern perched lava pond wall. After climbing back down the trail from the top of Pu‘u Huluhulu to the trail junction turn left and walk about 720ft (220m) to the east along the Makaopuhi Trail. This will take you out into the flow field to examine the well-developed perched lava pond located mid-way up the northern flank of Mauna Ulu (Figure D5-15). The pond formed when lava pooled behind self-constructed levées that contained the lava surface and became perched when those levées (see Figure D5-18, left) kept the surface of the lava higher than the surrounding terrain. Breaches in the levée are visible locally (see Figure D5-18, right). 99 �Figure D5-18: The left photo shows the Southern levée of the Mauna Ulu perched lava pond with Tom Erikson for scale. The right photo shows one of many visible breaches in the levée wall (this is the southwestern breach) that allowed lava to escape the perched lava pond and flow downslope (Peter Hinz and Lindsay Smith for scale). Photo credits: A.D. MacTavish (2020). Stop D5-41-12. Lava channel, western flank of Mauna Ulu. Data Source: MacTavish (2020). • UTM 268555E, 2143020N. Numerous lava rivers once cascaded down the flanks of Mauna Ulu. The empty remnants of those lava channels now form a pattern somewhat reminiscent of curved spokes on a bicycle wheel. The example at this sub-stop (see Figure D5-19) is one of the better examples (Peter Hinz for scale) and exhibits quite high, well-defined marginal levées with well-defined drain-back lines, which are readily visible in the lower right corner of Figure D5-19. Figure D5-19: This photo shows one of the numerous lava channels that radiate out from the summit of Mauna Ulu (Peter Hinz for scale). Photo credit: A.D. MacTavish (2020). 100 �Stop D5-41-13. Mauna Ulu Summit. Data sources: NPS Mauna Ulu Eruption Guide; MacTavish (2020). • UTM 268747E, 2142870N; southern rim of crater. Walk upslope along the south side of the lava channel viewed at Stop D5-41-12 for about 500ft (150m) until reaching the summit crater of Mauna Ulu. The summit crater (see Figure D5-20, left) is over 100ft (33m) deep and often exhibits rising steam due to the still hot rock present in the interior of the shield. The rocks exposed at the summit, particularly arcing around the northwestern side, show considerable evidence of reddish to reddish-brown hydrothermal alteration due to the passage of superheated steam through the rock along fractures. This can be seen particularly well in the southwest adjacent to several small collapse pits. The pit shown in the right-hand photo of Figure D5-20 has a faint, but readily visible wisp of steam rising from it. The HVO constantly monitors Mauna Ulu. The instrument package is visible near the eastern rim of the crater. Warning: The entire crater rim is very unstable and prone to collapse. Do not approach too closely. Figure D5-20: The left photo shows the western rim of the central Mauna Ulu crater with 2020 field trip participants for scale. The right photo shows a collapse pit with wisps of visible steam rising from it. Note the clearly visible reddish-brown alteration of the rock surrounding the pit. Photo credits: A.D. MacTavish (2020). Figure D5-21: The reddish hydrothermal alteration and lizard skin-like weathering pattern characteristic of the summit of Mauna Ulu is easily visible in this photo. Trekking poles for scale. Photo credit: A.D. MacTavish (2020). 101 �3.5. Day 5 (Part 2) – Chain of Craters Road Stop D5-42. Mau Loa o Mauna Ulu; alternate stop. Data source: Hazlett (2014). • UTM 268450E, 2139575N; parking area. This stop is on the western border of the 1969 to 1974 Mauna Ulu flow field and is also the trailhead for the Keauhou Trail. This flow field is composed of approximately 440 million cubic yards (340 million m 3) of lava covering 10.5mi2 (45km2) of the southern flank of Kīlauea to depths ranging from 3.3ft (1m) to >330ft (100m). The flow field buried 12.5mi (20km) of the original Chain of Craters Road. The longest flows reached the ocean, a distance of 7.5mi (12km) from Mauna Ulu. Most of the surrounding sea of pāhoehoe at this location erupted from the Mauna Ulu vent in 1974. The flows in this area are quite thin and average only a few metres in thickness. Safety and Endangered Species Note: Nēnē, the endangered Hawai‘ian geese, frequently graze along the shoulders of Chain of Craters Road between this parking lot and Muliwai a Pele and also between Maulu and Hōlei arches at the end of the road below on the coastal plain. Take care accordingly since most of the nēnē killed in the park has been along these 2 stretches of highway. Stop D5-43. Muliwai a Pele. Data source: Hazlett (2014). • UTM 269630E, 2138540N; parking area. A short 160ft (50m) walk south of the parking area leads to a viewing platform that overlooks a welldeveloped ‘a‘ā lava channel with accretionary lava balls deposited on its levée (see Figure D5-22, left). It is possible that several portions of this channel were in the preliminary stages of transitioning into a lava tube. This channel formed in 1974 during one of the many Mauna Ulu lava overflows and travelled about 5mi (8km) from the vent to the base of Poliokeawe Pali located ~2135ft (650m) to the south. Mauna Ulu is easily visible to the north on a clear day (see Figure D5-22, right). The walk from the parking lot is over one of many pāhoehoe overflows from the channel. Close inspection of the edge of this flow away from the channel shows that the original underlying flows were all ‘a‘ā. The parking lot roadcut shows that the channel levées are built of several pāhoehoe overflows. Flows that start from the vent as ‘a‘ā, with ‘a‘ā flowing in a channel, are commonly overlain by later pāhoehoe within the channel. Figure D5-22: The left photo shows a channelized ‘a‘ā flow with accretionary lava balls on the top of welldeveloped channel levées. The photo view is north-northeast. The right photo shows Mauna Ulu located ~2.75mi (4.4km) north of this stop. Photo credits: A.D. MacTavish (left 2019; right 2008). 102 �Stop D5-44. Pāhoehoe transitioning to ‘a‘ā flows. Data source: MacTavish (2019). • UTM 272305E, 2137347N; narrow roadside parking area on right of the highway. Here can be seen what may be a pāhoehoe flow transitioning into an ‘a‘ā flow north of the highway. Another explanation is that an underlying ‘a‘ā flow is being covered by thin pāhoehoe flows. This stop also affords a good view of the coastal plain (see Figure D5-23). Look below and to the right to see the dark ‘a‘ā flow in the distance. This may be the flow that was viewed at the last stop (D5-39). Ᾱpua Point Figure D5-23: The view looking southwest from the Kealakoma viewpoint. The distant, dark ‘a‘ā flow is possibly the flow we viewed at the last stop. To the right of the flow, you can see Āpua Point. Photo credit: A.D. MacTavish (2019). Stop D5-45. Kealakomo Overlook. Data source: Hazlett (2014). • UTM 272825E, 2137340N; parking area. This overlook is in a kīpuka surrounded by Mauna Ulu flows and looks south over Hōlei Pali with a good view of the pali slope and the coastal plain. Mauna Ulu-era lava flows are very easy to distinguish from the older, originally grassy terrain with the dark gray areas representing ‘a‘ā flows and the light gray areas representing pāhoehoe flows. The prominent point of land to the south-southwest is Āpua Point and is the site of a village destroyed during a tsunami and coastal subsidence associated with the 1868 earthquake. Also destroyed during these episodes were the coastal villages of Keauhu Landing, at Keauhu Point located to the westsouthwest, and Kealakomo located to the south-southeast. Petroglyphs and house sites at Kealakomo were again destroyed by Mauna Ulu flows in 1971. Additional archeological sites at Āpua Point were drowned by subsidence during the November 29, 1975 earthquake. Stop D5-46. Alanui Kahiko Pullout; Hōlei Pali flows. Data source: Hazlett (2014). • UTM 273780E, 2135810N; pull-out on left side of highway. This narrow pullout affords a good view of the Hōlei Pali, from below, with 1972 Ᾱlae Shield/Mauna Ulu pāhoehoe and ‘a‘ā flows cascading down the steep slope of the pali (see Figure D5-24, left). It is easy to see in this location that most of the ‘a‘ā flows overlie the pāhoehoe flows. About 100ft (30m) southeast of the upslope half of the parking area are 2 small remnants of the pre-Mauna Ulu Chain of Craters Road that occur as uncovered windows within the pāhoehoe flows (see Figure D5-24, right). 103 �Figure D5-24: The photo on the left shows Hōlei Pali with 1972 (Ālae Shield) Mauna Ulu pāhoehoe and ‘a‘ā flows cascading down the steep slope. The photo on the right shows 2 windows through Mauna Ulu pāhoehoe flows revealing a remnant of pavement from the pre-1972 Chain of Craters Road. Photo credits: A.D. MacTavish (2008). Stop D5-47a. Pu‘u Loa Petroglyphs Trailhead. Data source: Hazlett (2014); MacTavish (2019). • UTM 276170E, 2134160N; parking area. The 1.4mi (2.25km) long trail exiting this parking area leads to the Pu‘u Loa archeological site where early Hawai‘ians carved at least 23,000 petroglyphs as figures, shapes, and forms into the pāhoehoe surface. There are no formal field trip stops on the trail until you reach the archeological site; however, several large tumuli and pressure ridges (which we have seen on other trails previously) occur along the path that may provide some informal stops (see Figure D5-25). Figure D5-25: This photo shows an excellent example of the turtle shell-like surface of a wind-polished pāhoehoe flow adjacent to the Pu‘u Loa Petroglyphs Trail near the trailhead. Photo credit: A.D. MacTavish (2019). 104 �Stop D5-47b. Pu‘u Loa Petroglyphs Site. Data sources: Robinson (2012); Hazlett (2014). • UTM 276175E, 2134180N; entrance to boardwalk at petroglyphs site. This site hosts the largest collection of petroglyphs (rock carvings) in Hawai‘i with most inscribed centuries before the arrival of westerners. Petroglyphs can be seen on both sides of the wooden walkway. Hawai‘ian fathers would places pieces of their children’s umbilical cords within small holes often surrounded by concentric circles (see left photo in Figure D5-26). There are ~16,000 of these holes at the site. The stick-like figure, also seen in Figure D5-26, was carved before the arrival of Westerners. After that arrival carved figures became more detailed and less stick-like (see right photo in Figure D526). Do not leave the walkway once you arrive at the site and please do not attempt to make any tracings of the petroglyphs. Figure D5-26: The photo on the left shows a stick human figure that was carved before the arrival of Europeans in the islands. The human figures in the photo on the right were carved after the arrival of Europeans and are less stick-like and more human-shaped. Photo Credits: A.D. MacTavish (2019). Stop D5-48. Hōlei coastal sea arches. Data source: MacTavish (2019). • UTM 279425E, 2134790N, parking area. Walk south (toward the ocean) from the parking lot to the viewing area located at the coastal cliff. The cliffs here are ~100ft (30m) high with vertical to locally undercut walls and sea arches are common (see Figure D5-27). A stone guard-wall is only present at the end of the trail. Arches are visible in both directions from the viewpoint. The east-facing photo in Figure D5-27 (right) was taken in late December 2008 when flows originating at the Pu‘u Ō‘ō vent were entering the ocean via a long-active lava tube. The lava entry point is the grey plume at the horizon in the distance. Warning: Do not venture along the cliffs past where the guard walls end since these cliffs can be very unstable. Also, beware of large waves which can break over the top of the cliffs. The 2020 field trip group were inundated by one such wave. Luckily no cameras we destroyed by the sea water. Over the years AM has had 2 cameras destroyed by seawater from breaking Hawaiʻian waves and almost lost a third at this viewpoint. 105 �Figure D5-27: The left photo shows the Hōlei Sea Arch located west of the viewing platform. The right photo shows another sea arch located east of the platform. In the distance is a volcanic haze (‘vaze’) plume produced in late December 2008 by lava entering the sea from a lava tube originating at Puʻu Ōʻō. Photo credits: A.D. MacTavish (2008). Stop D5-49. End of Road (Alternate stop due to long walk). Data source: MacTavish (2008). • UTM 280390E, 2135255N; end of road past gate. The road was gated past the buildings and restrooms in 2020. The end of the road pre-2014, was a 2950ft (900m) walk to the east past the gate; however, in 2014 a non-paved extension of the Chain of Craters Road was completed that connected with the road on the far side of the flow field. Vehicle access past the gate is not allowed, unless under special permission. This road extension was again cut by Pu‘u ‘Ō‘ō flows in 2016. Figure D5-28 shows a very appropriate road sign almost covered by 1980’svintage basalt flows. Figure D5-28: This photo shows one of the incongruities of Hawaiʻi – a Road Closed sign in the middle of an almost unending field of Kīlauea basalt. The photo was taken in 2008 several hundred metres east of Road’s End. The vase plume from the 2008 lava sea entry is visible in the centre background. Photo credit: A.D. MacTavish (2008). 106 �3.6. Day 5 (Part 1): Helicopter Flight over Kīlauea (Morning) During the morning of Day 6 the 2020 field trip participants partook in a 75-minute helicopter flight over Kīlauea and its East Rift Zone. Dr. Jack Lockwood (retired HVO volcanologist) and AM acted as aerial tour guides. Some of the features observed are shown in Figures D6-2, -3, and -4. The very approximate flight path is shown on Figure D6-1. The features observed during the flight were: A. The Kīlauea Caldera, including the new, larger, and deeper Halemaʻumaʻu crater; B. The Kīlauea Southwest Rift Zone and the Koaʻe Fault system; C. The Mauna Ulu Shield (1969 to 1974 upper East Rift Zone eruption) and its associated flow field and various pit craters; D. The Pu‘u ‘Ō‘ō Shield (primary vent for much of the 1983 to April 2018 eruption) on the Central East Rift Zone; E. The southern coastal plain covered by 1983 to 2018 Pu‘u ‘Ō‘ō flows; F. Kīlauea surface flows near the base of Hōlei Pali; flows from the 1983 to 2018 eruption and a landing on the flows (Stop D6-50); and G. The Lower East Rift Zone and features of the May to August 2018 eruption. G A C B F D E Stop D6-46, Helicopter Landing Site Figure D6-1: Map showing approximate 2020 helicopter flightpath over various portions of Kīlauea Volcano starting from Hilo Airport. Map is modified from USGS General Information Product 117 (2010), p.3. 107 �Figure D6-2: The Puʻu Ōʻō Cone located in Kīlauea’s Central East Rift Zone. This cone, which was in the process of transitioning into a shield, was the focus of most of the magmatism and lava flow activity during Kīlauea’s 1983 to 2018 eruption. Photo credits: A.D. MacTavish (2019, left; 2020 right). Figure D6-3: The new crater located within Kīlauea’s summit caldera that formed after the lava lake within the original Halemaʻumaʻu Crater drained away in May 2018. A water lake was forming at the bottom of the pit in the right photo during the 2020 overflight. Photo credits: A.D. MacTavish (2019, left; 2020, right). Figure D6-4: The left photo is an aerial view of the Mauna Ulu Shield taken from the east. Note the lava channel on the east-facing slope. The right photo is of the Fissure 8 area of the Lower East Rift Zone, formed during the 2018 eruption and taken from the east in August 2019. Note the steam still issuing from the rift zone and the various cinder cones located along it. Photo credits: A.D. MacTavish (2019). 108 �Stop D6-50. Helicopter landing site. Data source: MacTavish (2019, 2020). • UTM 289790E, 2140860N (approximate). This site (see Figure D6-3 for location) is owned by volcanologist Dr. Jack Lockwood (HVO retired) and is located on the coastal plain ~1.5mi (2.3km) west of the destroyed town of Kalapana and ~2.2mi (3.5km) west-southwest of the present Village of Kaimu. The pre-1983 forest floor at this location was completely covered in 1990 and 1992 by the extensive Pu‘u ‘Ō‘ō flow field to a depth of at least 80ft (25m) by basalt flows (J. Lockwood, personal communication, 2019). This stop (see Figures D6-5 and D66) allowed the 2020 field trip participants to view many pāhoehoe flow forms with a well-respected and exceptionally experienced volcanologist available to point out and explain the features observed. Figure D6-5: Helicopters (left photo) parked on the uneven, 1990 and 1992-vintage, Pu‘u ‘Ō‘ō pāhoehoe flows overlying property owned by retired HVO volcanologist Dr. Jack Lockwood (shown in the right photo). Photo credits: A.D. MacTavish (2020). Figure D6-6: The left photo shows and earlier pāhoehoe flow (1990 flow?) with a crack partially infilled by a later elongate rope of pahoehoe (Tom Erikson for scale). The right photo shows the complex irregularity of the flows. Photo credits: A.D. MacTavish (2020). 109 �3.6. Day 6 (Part 2): Kīlauea Lower East Rift Zone (afternoon) The field trip stops visited on the Lower East Rift Zone during the afternoon of Day 6 (see Figure D6-7): 51. 52. 53. 54. 55. 56. 57. 58. Pāhoa Flows (2014 June 27th Flow that issued from the Pu‘u ‘Ō‘ō shield); ‘A’ā flow, 1955 eruption; 2 large spatter and cinder cones to west (alternate stop); 1989 Pu‘u Ō‘ō-Kupaianaha pāhoehoe flows (49a) and the New Kaimu Black Sand Beach (49b); Palagonatized ash beds erupted approximately 1745AD; MacKenzie State Park; ‘a’ā flows and rest stop; Eruption-truncated Leilani Street within Leilani Estates subdivision; Fissure 9, located a short distance east of Moku Street, Leilani Estates; and 13-3574 Makamae St., Leilani Estates; home is located a short distance south of Fissure #8 which formed during 2018 Lower East Rift Zone eruption; provides access to Fissure 8 cone. Map modified from Hazlett and Hyndman (1996, p.88) 51 56 52 57 58 55 54 53 Figure D6-7: Location of field stops during the afternoon of Day 6. The map shown in Figure D6-8, below, shows the extent of the flows (salmon coloured region on the map) erupted from the Lower East Rift Zone between May 3 and August 14, 2018. The map also shows a closer view of 3 field trip stops made at the western end of the area affected by the eruption. 110 �56 57 58 Figure D6-8: Map showing the locations of field stops, fissures, and flows that issued from the fissures during the May 3 to August 9, 2018 eruption from Kīlauea’s Lower East Rift Zone. The salmon-coloured region represents the 2018 flows. Map modified from the USGS HVO website. Stop D6-51. June 27th Flow, Pāhoa, Lower East Rift Zone. Data Sources: USGS HVO Website; MacTavish (2019). • UTM 294350E, 2156355N; Pāhoa Transfer Station (Recycling Depot) Parking Area. Between 2014 and 2016 a series of flows issued from the north flank of Pu‘u Ō‘ō that eventually approached the town of Pāhoa, which straddles Highway 130, ~12mi (19.5km) northeast of the Pu‘u Ō‘ō vent. Two lobes of the ‘June 27th Flow’ threatened Pāhoa in 2014. The northern tongue of this particular flow-lobe (see Figure D6-9) threatened the town’s Transfer Station (recycling depot) and came within a few metres of destroying the station. The southern part of the Transfer Station fence and road were damaged. The fence acted as a partial barrier to the flow, which inflated against it until small breakouts at the base of the flow punched their way through the fence into the station after melting the fence steel (see Figures D6-10 and D6-11). The flow destroyed several outbuildings at a farm directly across Apa’a St. to the east from the station and part of a Japanese cemetery located midway between the station and the town. The southern tongue of the flow extended to within 625ft (190m) of the Town of Pāhoa before it permanently stalled. 111 �th June 27 Flow Apaa St. Fence Pāhoa Cemetery Partially destroyed farm Transfer Station Figure D6-9: Aerial view, looking northeast toward the Town of Pahoa. The location of the June 27 th flow is readily visible and the various features of interest are labelled. Photo credit: USGS HVO Website. Figure D6-10: The left photo shows a lobe of the June 27th Flow encroaching upon the Pahoa Transfer Station (photo taken on November 13, 2014). The right photo shows the same flow inflated against the barrier of the Transfer Station fence and the small flow lobes that oozed through the fence and onto the road of the station (photo taken on November 16, 2014). Photo credits: USGS HVO Website. 112 �Figure D6-11: These close-up photos were taken on August 4, 2019. The left photo shows the flow against the Pahoa Transfer Station fence after the station roadway has been repaired. The photo on the right shows the partially destroyed fence with the inflated flow abutting against it. Photo credits: A.D. MacTavish (2019). Stop D6-52: The 1955 ‘a‘ā flow, Lower East Rift Zone (Alternate Stop). Data source: Hyndman and Hazlett (1996). • UTM 294308E, 2148815N; pull-over on right side shoulder of Highway 130 a short distance past the outcrop. This stop is located at the southern edge of a huge ‘a‘ā flow erupted from the Lower East Rift Zone in 1955. These flows erupted from a pair of large spatter and cinder cones located on the horizon west of the highway. The flows contain large numbers of glassy green olivine crystals and are well exposed in the rock cuts located immediately northeast of the parking area. Because of the wetter climate in this area these flows are well-vegetated compared to what is visible farther to the west. Stop D6-53a. New Kaimu Black Sand Beach. Data sources: MacTavish (2019, 2020). • UTM 293050E, 2141970N. Parking area at Kalapana Village Café and Kaimu Korner Store. The beach and Stop D6-53b are located at UTM 293270E, 2141490N at the end of a 1800ft (550m) long jeep trail leading to the south-southeast from the parking lot. Vehicle traffic along this trail is restricted. The road to Kaimu and the edge of the vegetation to the north roughly mark the pre-1989 shoreline in this area. The former, famous Kaimu Black Sand beach used to lie immediately opposite this parking lot (see Figure D6-12). The Pu‘u ‘Ō‘ō-Kalapana flow field in this area developed between 1989 ad 1993. After completely destroying and covering the former site of the Kalapana village the flows were channelized by the low-tide shelf, and the former coastline beach, leaving the road and buildings at this site intact. • Trail to Stop D6-49b The 4x4 trail to the new Kaimu black sand beach passes through an area that is under active, guided, and tended revegetation by native Hawai‘ians. Identified species that have been planted and tended here include, but are not restricted to, coconut palm, papaya, and breadfruit (see Figure D6-13). Also, there are numerous pieces of recent indigenous art and shrines flanking the trail (Figure D6-13). Please be respectful of the indigenous culture since it is immediately apparent from the efforts made that this place is important to the native Hawai‘ians. 113 �Village of Kaimu Hwy 130 Approximate location of Pre-1989 Shoreline and the old Kaimu Black Sand Beach Hwy 137 New Kaimu Black Sand Beach Figure D6-12: Google Earth satellite image of the New Kaimu Black Sand Beach area. Figure D6-13: Planted and tended plants along the trail to the New Kaimu Black Sand Beach (left photo). Indigenous art along the trail to beach (right photo) Photo credits: A.D. MacTavish (2019). Stop D6-53b. New Kaimu Black Sand Beach. Data source: MacTavish (2019). • UTM 293260E, 2141485N; end of trail above new beach. The natural construction of the New Kaimu Black Sand Beach began immediately after the flows reached the ocean and formed fine hyaloclastites upon contact with sea water that were further broken down to sand-sized particles due to intense wave erosion at the shoreline. The beach is narrow and at high-tide. When there is a high on-shore wind, the beach is underwater and essentially inaccessible and invisible. The walk to the beach is well worth the effort, even if the beach is under water, just to view the revegetation efforts and the examples of indigenous art on the some of the adjacent flows. 114 �Stop D6-54. Palagonitized ash beds at Waste Transfer Station. Data sources: Easton and Easton (1995); Hazlett and Hyndman (2007). • UTM 294640E, 2142975N; parking on right (south) highway shoulder. At this location beds of palagonitized and limonitized ash and lapilli (see FigureD6-14), most probably erupted during littoral steam explosions along the ancient coastline, underlie a weathered ‘a‘ā flow immediately overlain by soil. The overlying flow has been dated using 14C methods (Carbon 14) at 2360±90 years Before Present (BP). The exposures of the ash are on the west side of the Waste Transfer Station near the base of an east-facing scarp and can be readily observed from the highway if the gates of the station are closed. Overlying ‘a‘ā flow Ash and lapilli beds Figure D6-14: Ash and lapilli beds overlain by 2360yr old ‘a‘ā flow. Photo credit: A.D. MacTavish (2019). Stop D6-55. MacKenzie State Park, Rest Area (Alternate). • UTM 304225E, 2150530N; park entrance. The park is a good afternoon rest stop, if not too busy or overcrowded. The park provides an excellent place to watch surf pound the resistant ‘a‘ā flows exposed in a shoreline cliff into rubble. Stop D6-56. Leilani Street truncated by 2018 Eruption. Data source: MacTavish (2019). • UTM 299335E, 2153310N; eastern end of Leilani Street, where truncated by 2018 flows; park along the side of the street, where it is wide enough. From this point you can readily see the spatter rampart formed along Fissure 24 and the southwestern slope of the Fissure 8 Cone (see Figure D6-15, left) formed during the 2018 Lower East Rift Zone eruption (LERZ). Figure D6-15 (right) provides a closer view of the Fissure 24 spatter rampart. 115 �Figure D6-15: These photos show Leilani Street, within Leilani Estates, truncated by flows and edifices from the 2018 LERZ eruption. In the left photo the Fissure 12 spatter rampart is in the distance beyond the road and the southwestern side of the Fissure 8 Cone is at the right, partially obscured by dead trees. The photo on the right provides a closer view of the irregular Fissure 12 spatter rampart. Photo credits: A.D. MacTavish (2019). Stop D6-57. Fissure 9 spatter rampart (2018 eruption). Data source: MacTavish (2019). • UTM 298740E, 2152523N; stop and park along the west side of Moku Street. At this location in 2020 the 2018 LERZ Fissure 9 was still issuing sulphurous gas and steam, as it was when the Figure D6-16 photos were taken on August 4, 2019. The opening of this fissure destroyed the house that was at the end of the truncated driveway at the lower left of the photo in Figure D6-16. Also, the neighbour on the right was lucky to escape destruction as the spatter rampart was built around the fissure. A close-up of steaming Fissure 9 and its associated hydrothermally steam-altered spatter rampart is shown in Figure D6-16 (right). The red pile to the right of the fissure are the remains of the roof of the house that once stood where the fissure is now located. Figure D6-16: These 2 photos show Fissure 9 and it’s accompanying spatter rampart located a short distance east of Moku Street, Leilani Estates. The house on the right in the left photo had a narrow escape from the destruction visited upon the house located where the fissure now sits. A close-up of Fissure 9 is shown on the right with the hydrothermally-altered spatter rampart readily visible. The red pile located a short distance right (south) of the fissure are the remnants of the roof of the home that once stood where Fissure 9 now gapes. Photo credits: A.D. MacTavish (2019). 116 �Stop D6-58. Lower East Rift Zone, Fissure 8 Area (2018 eruption). Data source: MacTavish (2019). • UTM 299760E, 2152625N, stop is at 13-3574 Makamae St., Leilani Estates. The winter home here is owned by Mark Bishop, from Minnesota, and he graciously allowed the 2020 field trip property access (one time access only; will not apply to anyone else using this guide), which provided an excellent view of the Fissure 8 cone and the edge of the 2018 LERZ eruption flow field. He also obtained permission from other landowners and led the group to the top of the Fissure 8 Cone All land underlying the flow field is privately owned and access permission is required from those landowners. Please do not walk onto the flow field without permission. Figure D6-17 shows the May 29, 2018 Fissure 8 fountaining episode, as viewed by helicopter from the north. The upper left edge of the photo is near the edge of Mr. Bishop’s property. The left photo in Figure D6-18 shows a close-up of Fissure 8 fountaining on June 2, 2018 and the right photo of the same figure shows the breached Fissure 8 cone on September 2, 2018 after the eruption had ended. Figure D6-17: Fountaining at Fissure 8 on May 29, 2018. Photo credit: USGS HVO website. Figure D6-18: The left photo shows a Fissure 8 fountaining episode on June 2, 2018 that fed a perched and channelized flow. The photo on the right shows the inactive Fissure 8 vent and breached spatter and cinder cone on September 2, 2018. Photo credits: USGS HVO website. 117 �Fissure 8 perched Lava channel Spatter rampart along Fissure 24 Fissure 21 Fissure 8 breached cone Mark Bishop house – Stop D6-58 Fissure 2 Fissure 7 Makamae Street Fissure 9 Figure D6-19: Aerial view of the Fissure 8 Cone area and it’s associated perched perched lava channel from the 2018 LERZ eruption (August 4, 2019 photo). Also shown are the house at Stop D6-58 and Fissures 2, 7, 9, 21, and 24. Photo credit: A.D. MacTavish (2019). Figure D6- 19 (above), taken from a helicopter late in the morning of August 4, 2019, shows the Fissure 8 cone and resulting perched lava channel as well as adjacent fissures that were active for a short time at some point during the eruption by issuing spatter, gas, or short-lived flows. Fissure 8 was the longest acting fissure and erupted by far the largest volume of lava during the 3-month eruption. The 2 photos in Figure D6-20 (below) were taken on August 4, 2019, almost a full year after the 2018 Lower East Rift Zone eruption ended, from the edge of the 2018 flow-field just past the present end of Makamae St. and adjacent to the home owned by Mark Bishop. Figure D6-20: The left photo shows 3 well-developed lava trees formed on the Fissure 8 flow field. The photo on the right shows the Fissure 8 cinder and spatter cone seen from Makamae Street. The. Photo credits: A.D. MacTavish (2019). 118 �3.7. Day 7 (Mauna Loa) and Day 8 (Mauna Kea) The various field trip stops for Day 7 and Day 8 are as follows (see Figure D7-1): Hilo Area: 59. Hilo Waterfront (Coconut Island Park) 60. Rainbow Falls 61. Kaūmana Cave Mauna Loa: 62. Pu‘u Huluhulu (summit of Saddle); 63. Rough lava channel with high levées 64. Road junction with Mauna Kea view 65. Multicoloured flows 66. 2022 Mauna Loa Fissure 4 Flow 67. Mauna Loa Observatory 68. Lava Flow diversion barriers Mauna Kea Summit Road: 69. Breached Pu‘u Kalepeamoa cinder cone 70. Mauna Kea Visitors Centre at Halepōhaku 71. Ellison B. Onizuka Astronomical Complex 72. Lake Waiau Trailhead 73. Top of pass on Lake Waiau Trail 74. Lake Waiau 75. Mauna Kea Observatory Complex, summit cinder cones, glacial features Map modified after Hazlett and Hyndman (1996, p.117) 72-74 75 70,71 69 60 59 Hilo 62 61 63 67 64 66 65 68 Figure D7-01: Location map of field trip stops on Day 7 and Day 8. 119 �3.7. Day 7 (Part 1) – Hilo Area Day 7 consists of 4 stops in the Hilo area (Part 1) and a drive up the Saddle Road and the Mauna Loa Weather Observatory Road, with several stops on the way up the mountain (Part 2). There is little walking on this portion of the field trip. The stops in the Hilo Area are listed below and shown on Figure D7-2, below: 59. Hilo Waterfront a. Coconut Island Park b. Tsunami Park (Alternate Stop if cannot get into Coconut Island Park) 60. Rainbow Falls, Wailuku Valley 61. Kaūmana Cave (Lava Tube) Figure D7-2: Hilo Area Day 7 field stops. Map modified after Hazlett and Hyndman (2007, p.57). Stop D7-59a: Hilo Waterfront. Coconut Island Park; alternate Stop if a clear day to view the volcano summits. Data sources: Hazlett and Hyndman (2007); Easton and Easton (1995). • UTM 283145E, 2182600N; parking area. This location provides a good view of Hilo Bay and, on a clear day looming in the background, the summits of Mauna Kea and Mauna Loa, and Hilo to the southwest along the bay coastline. Hazlett and Hyndman (2007) state that Hilo Bay is a notorious tsunami trap, most of which originate from Pacific Rim megathrust earthquakes. The funnel shape of the bay, its location on the northeast side of the island, and the steep offshore slope off the harbour causes tsunami waves to build quickly. Once inside the bay the waves reflect off the coastlines which causes positive wave interference where the wave crests combine to form waves with extraordinarily high crests. This positive reinforcement has produced numerous tsunami waves that have killed many people and have caused a lot of damage to the city. Tsunamis have killed more people in Hawaiʻi than all other natural disasters put together (www.darktourism.com). The 2 most damaging tsunamis in modern history took place in 1946 and 1960. 120 �Stop D7-59b. Tsunami Park (Alternate Stop). Data source: Hazlett and Hyndman (2007). • UTM 281880E, 2182227N, parking area. The April 1, 1946 and May 22, 1960 tsunamis killed a combined total of 141 people in the Hilo area. After the 1946 tsunami residents of Hilo rebuilt in the devastated area like they had always done in the past. However, after the 1960 tsunami rubble was cleared away, it seems that a lesson was learned and residents now prefer to build their homes on higher ground, away from the waterfront. The area cleared in 1960 remains as a wide grassy strip containing only a few structures that the locals now call Tsunami Park. Still, those parts of Hilo close to the sea remain vulnerable. Evacuation routes are mapped out and available online and are printed in every telephone directory and on signs throughout the city. An early-warning system is in place which includes sirens that sound if a tsunami is detected out to sea. Stop D7-60. Rainbow Falls. Data source: Hazlett and Hyndman (2007). • UTM 279015E, 2181735N, parking area. Rainbow Falls (see Figures D7-3 and D7-4) is on the Wailuku River at the western edge of Hilo. Hazlett and Hyndman (2007) state that the plunge pool at the base of the falls undercuts the thick ledge of an ‘a‘ā basalt flow. In the proper light the curving base of this flow can be seen to follow the outline of an older river bed that was infilled by an earlier eruption. The river occupies a low area between the volcanoes and cannot erode deeper because Mauna Loa flows commonly flow along it and displace the stream. This gorge contains excellent evidence of this happening repeatedly. How the channels formed is illustrated in the upper portion of Figure D7-3. An excellent palisade of columns within two thick Mauna Kea ‘a‘ā flows is visible in the walls of the gorge below the falls (above the infilled channel). How the palisades formed is illustrated in the lower portion of Figure D7-3. Figure D7-3: The upper panel illustrates the repeated infilling and cutting of the gorge by the river with no overall deepening of the resulting gorge. The lower panel illustrates the formation of joints produced when the flow infilling the lava channel cools to produce the palisades. Diagrams from Hazlett and Hyndman (2007, p.58). 121 �Figure D7-4: Rainbow Falls. Photo credit: A.D. MacTavish (2020). Stop D7-61. Kaūmana Cave (Lava Tube). Data source: Hazlett and Hyndman (2007). • UTM 276605E, 2178200N, a large parking lot on the west side of the highway. Be very careful crossing the highway to access the cave entrance due to vehicle traffic along the busy highway. At this small private park, a collapsed skylight opens into a large, easy-access lava tube with a 20 to 25ft (6 to 8m) high ceiling (see Figure D7-5). It is possible to walk or crawl downslope within the tube for almost 3000ft (915m). This tube was the main lava conduit during the 1880 to 1881 Mauna Loa eruption. The tube allowed lava to reach to within 2km of Hilo Harbour, which is 5km east of this point. The walls of the tube expose the internal structure of a pāhoehoe flow. The layering visible in the upper wall formed due to multiple lava spillovers over the levées of a channelized flow before it roofed over to form the lava tube. Fast flowing lava filled the tube when it first formed; however, later on when the lava level dropped the lava sometimes slopped from side to side building the smooth shelves visible near the tube entrance. When the lava finally drained away an empty channel was left between the shelves. Occasionally blocks of the roof fell into the flowing lava becoming partially embedded within, and coated by, lava. Visible on the upper walls and ceiling are numerous small stalactites and lava drip tracks. The cave is also home to a host of underground animals, the largest being insects and spiders. Figure D7-5: This figure is a map of the lava tube. Diagram was taken from Hazlett and Hyndman (2007, p.60). 122 �3.7. Day 7 (Part 2): Saddle Road (Highway 220) and Mauna Loa Observatory Road The Saddle Road (Highway 220) crosses the mainly unpopulated central plateau of the island and passes through the gap between Mauna Loa and Mauna Kea volcanoes known as the Humu‘ula Saddle. The top of the saddle is at an elevation of 6500ft (1980m) ASL. The Saddle Road was built in 1942 shortly after the Japanese attack on Pearl Harbour to access the U.S. Army’s Pohakuloa Training Area and Bradshaw Army Airfield. In 1945 the road was transferred to the Territory of Hawai‘i and designated Route 20; however, no maintenance funds were available for a road that was never designed for civilian travel. In 1959 the new State of Hawai‘i gave the road to the County of Hawai‘i, but, again, there were no funds for road maintenance. Finally, in 2004 Federal funds allowed for road re-alignment and upgrading to the present modern, paved, and well-maintained Highway 220. The Saddle Road starts at the Hilo Waterfront, as Waiānuenue Ave., and ends at Highway 190, a short distance southwest of the Town of Waimea. For much of the distance between Hilo and Pu‘u Huluhulu the road follows the 1935 and 1936 Mauna Loa flow which mostly covers the Humu‘ula Saddle plain. On approach Pu‘u Huluhulu the flow changes from ‘a‘ā to pāhoehoe. Stop D7-62: Pu‘u Huluhulu (Hairy Hill). Data sources: Hazlett and Hyndman (2007); Robinson (2010). • UTM 241390E, 2178890N; parking lot (if open). In August 2019 access to the cone was not possible due to roadblocks erected by political and religious protestors. The site was opened in February 2020 and was apparently still open as of December 2022. Pu‘u Huluhulu is a small, wooded, alkalic Mauna Kea cinder cone (>10,000 years old) located adjacent to the highway about 100m south of the junction with the Mauna Kea access road and 400m west the junction with the Hilo-Kona Road (see Figure D7-5). The cone is part of the alkalic Laupāhoehoe formation, is now a kīpuka surrounded by younger Mauna Loa flows, and is at the transition between the montane and sub-alpine vegetation zones. Near the northwest base of the cone is a partially buried stone wall constructed in 1935 in an attempt to deflect the 1935 Mauna Loa flow (see Figures D7-5 and D7-6). Hazlett and Hyndman (2007) state that: ‘[A]fter the lava pooled in the saddle and a hard crust formed on the surface, molten lava pouring down the long slope of Mauna Loa continued to feed the flow beneath the crust, lifting and splitting the crust as though it were rising bread dough [inflation]. You can judge the amount of the rise as you walk along the wall. In a few places the lava buried the wall, but in most places the flow was too thin and the crust too stiff to shift across the top of the wall’. The cone’s interior structure is exposed by an old quarry in its western flank (Figure D7-5). Clearly visible are basaltic dykes that intruded into the cone. The lava confined beneath the surface of the Mauna Loa flows was under pressure and was able to push sheets of molten lava into the unconsolidated debris of the cinder cone. A road leads from the quarry to the summit of the cone. From the summit the stone wall to the west is readily visible, as are numerous other alkalic Mauna Kea cinder cones formed during late-stage activity on the South Rift Zone. A group of cones located near the towers of the National Radio Astronomy Observatory, a few miles north of Pu‘u Huluhulu, are between 20,000 and 40,000 years old. Younger cinder cones, ~4500 years old, are located further upslope. The ash and cinder from these vents covered a wide area to a depth of ~1ft (30cm). This ash can be observed in a gully that is crossed by the Highway. This ash (now a soil) contains bits of charcoal from wood that is thought to have been burned during the eruption. 123 �Mauna Kea Road 1935-36 Mauna Loa Flows Pu‘u Huluhulu Hwy 220 Quarry Stone Wall Hilo-Kona Road Figure D7-5: Google Earth satellite image of the Pu‘u Huluhulu area showing various points of interest. Figure D7-6: The partially buried stone wall located west of Pu‘u Huluhulu. The wall was built in an attempt to stop or deflect a Mauna Loa flow during the 1935 to 1936 eruption. Photo credit: Hazlett and Hyndman (2007, p.120). 124 �3.7.1. Mauna Loa Observatory Road: The November 28 to December 10, 2022 Mauna Loa summit and Northeast Rift Zone eruption truncated road access and the final 2 of the planned stops for the field trip Day 7 and are no longer accessible. We cannot yet properly describe the flows and other volcanogenic features formed by the new eruption so we recommend that users of this guide examine the accessible portions of the flows at their leisure. Before the eruption the 17.4mi (28km) long, steep, and narrow Mauna Loa Observatory Road wound up the northern flank of the mountain to the Mauna Loa Weather Observatory operated by the National Oceanic and Atmospheric Administration (NOAA) situated at 11,000ft (3353m) ASL. The road passed through multiple climactic zones ranging from moist subtropical to alpine and provided a feel for the enormous size of the volcano. The road also passed over numerous, well-exposed lava flows of widely differing ages that due to sparse vegetation and slow weathering are almost indistinguishable from one another. All the originally planned stops have been left in place with the addition of one field stop where the flow now blocks the road. The road will eventually be rebuilt and the final 2 stops will again be accessible. Stop D7-63: Rough lava channel (Alternate). Data source: Hazlett and Hyndman (2007). • UTM 240950E, 2175580N, parking area is widened road within channel. The road passes through an extremely rough 1935-1936 ‘a‘ā lava channel with high levées. The huge blocks scattered downslope on the flow from the parking area originated when the levée walls of this flow channel collapsed into the lava stream and were carried downslope. The lava channel is very difficult to identify while driving so keep a close look at the GPS to be able to identify the right location. Stop D7-64. Road Junction at ~13.8km. Data source: Hazlett and Hyndman (2007). • UTM 242570E, 2167240N, park to the left near the microwave dishes. When the sky is clear this road junction provides a superb view of the Saddle Road and Mauna Kea in the distance (see Figure D7-7). This is also a good location for a short ½ to 1 hour break (possibly lunch) to partially acclimatize to the high altitude before climbing any higher. Figure D7-7: Spectacular view of Mauna Kea located to the north of Stop D7-60. Photo credit: Google Earth. 125 �Stop D7-65. Multicoloured flows. Data source: MacTavish (2019). • Approximate UTM 241400E, 2166690N; park to side of road where it is wide enough. Upslope, to the south, are multiple pāhoehoe and ‘a‘ā flows, the oldest with some vegetation. The flows exhibit multiple colours due to age and weathering (see Figure D7-8) with the oldest rocks comprising the red, partially vegetated pāhoehoe (left centre of the photo), the youngest are the dark brown ‘a‘ā, and in between are grey pāhoehoe flows (right centre of the photo). Figure D7-8: Multi-coloured flows of differing ages on the north flank of Mauna Loa at Stop D7-61. Photo credit: A.D. MacTavish (2019). Stop D7-66. 2022 Mauna Loa Fissure 4 Flow Blocking Mauna Loa Observatory Road. Data Sources USGS HVO website; Google Earth. • UTM 239365E, 2166345N (approximate location derived from Google Earth). Until road is reestablished travel by road past this point will be impossible. As of June 6, 2023 the lates Google Earth image shows that the road is still blocked. The ~1000ft (300m) wide 2022 Mauna Loa flow truncating the road here issued upslope from Fissure 4 (F4) on the northeast rift zone. It did not advance much past this point (600m) because lava ceased issuing from F4 shortly after the road was cut. It may be possible to walk around the north end of this flow and then walk further to the west to where the F3 flow field crosscuts the road; however, the older flow located west of the new flow is an ʻaʻā flow and walking across it would be extremely dangerous. Figure D7-9 shows an aerial view of the F3 flow field crosscutting the road in multiple places. Safety Note: If the new flow is pāhoehoe please be very careful walking across it since these flows are brittle and there will be a large number of hidden voids. If this is an ʻaʻā flow then do not attempt to cross it under any circumstances. ʻAʻā flows, particularly young ones, are very rough, very unstable, and all surfaces consist of razor-sharp edges that easily slice flesh and destroy footwear, even sturdy footwear. They are extremely dangerous to walk upon. The authors do not know how far the walk is around the toe of this flow. Only attempt it if the distance is relatively short and the underlying older flows are not ʻaʻā flows. The location of this and other flows, the F3 and F4 vents, and the Mauna Loa Observatory Road are shown in Figure 13, above. 126 �Inactive F4 Flow Active channelized F3 Flow Stop D7-66 (approximate) Inactive F3 Flows Truncated Mauna Loa Observatory Road Figure D7-9: Aerial view of the active F3 flow field crosscutting the Mauna Loa Observatory Road on December 5, 2022. The inactive F4 flow field is identifiable to the east at the top pf the photo. This view is looking roughly eastsoutheast on the north flank of Mauna Loa toward Stop D7-66. Photo credit: USGS HVO website (2022). Stop D7-67. Mauna Loa Weather Observatory (~11,000ft, 3355m ASL). Data source: Hazlett and Hyndman (2007). • UTM 229765E, 2162390N, parking area just before the gate to the observatory. Travel by vehicle, except for authorized NPS vehicles, past this point is restricted. Only foot travel is allowed. The road past this point ends at the summit of Mauna Loa. On clear days this location provides a spectacular, panoramic view of 3 of the other volcanoes on the island with Mauna Kea to the north, Kohala to the north-northwest, and Hualālai to the west. Barely visible in the distance in this photo, between Kohala and Hualālai, and above a narrow line of cloud is the summit of Haleakalā volcano located on the Island of Maui (see Figure D7-10). The Mauna Loa Observatory is a facility for studying the earth’s atmosphere, particularly atmospheric CO2 and ozone loss and is operated by the American National Oceanic and Atmospheric Administration. 127 �Hualālai Haleakalā (on Maui) Kohala Mauna Kea Summit Figure D7-10: Panoramic view from Stop D7-62 and the Mauna Loa Weather Observatory parking area. Photo credit: Google Earth. Stop D7-68. Lava flow diversion barriers. Data source: Hazlett and Hyndman (2007). • UTM 229550E, 2162165N; 160m south of the Mauna Loa Summit Trail a short distance past the gated road leading to the Mauna Loa Observatory. To get to this location either go through the observatory grounds (permission will be required) or walk along the summit road for a distance of 850ft (260m) until a window of the underlying pāhoehoe is exposed, then walk slightly east of south for 525ft (160m) until you see the end of the western diversion barrier. This diversion barrier was designed by Dr. Jack Lockwood of the HVO, retired (personal communication, 2019). These diversion barriers (see Figure D7-11) protect the observatory from Mauna Loa lava flows and were the first such structures constructed in the USA. Note how the 2 barriers form an acute angle into the direction of downslope flow. This shape and orientation are designed specifically to deflect, rather than stop, any flows striking the barriers. Barriers erected across the direction of flow will eventually be engulfed and overridden after the flows inflate at the barrier, as was seen at the wall observed at field trip Stop D7-58. If it is raining, please do not attempt to walk to the diversion barriers from the road since walking on the ‘a‘ā flows located between the road and the western diversion barrier is very dangerous. Observatory Road 128 �Mauna Loa Summit Trail Parking Area Western Diversion Barrier Eastern Diversion Barrier Figure D7-11: Google Earth satellite image of the area of the Mauna Loa Weather Observatory with the location of the lava flow diversion barriers shown by the labels and arrows. 3.8. Day 8 – Mauna Kea Summit Road Stop D8-69: Pu‘u Kalepeamoa Crater. Data source: Hazlett and Hyndman (2007). • UTM 242600E, 2186430N; gated entrance to tower access road. Having just passed through the breached crater wall, this stop is on the northern edge of the horseshoeshaped Pu‘u Kalepeamoa Crater The ridge west of the road is the crater rim where trade winds piled cinder high to one side. The cinder of this cone contains many fragments of older rock, including gabbro and green dunite. Stop D8-70: Mauna Kea Visitors Centre at Hale Pōhaku. Data sources: Hazlett and Hyndman (2007); Robinson (2010). • UTM 242640E, 2186710N; Visitors Centre parking lot. This is a rest stop because there are no public facilities at the summit. The store within the centre should be open and will be able to tell you whether travel is permitted higher, particularly if there is snow cover from a recent snowfall. The center sits on fresh-looking ash and cinder. All visitors planning on driving to the top of the mountain are required to stop here at an elevation of 9200ft (2804m) for at least an hour to acclimatize to the altitude before moving higher on the mountain. At the summit the atmospheric pressure is 40% of that at sea level and acute altitude sickness is common. Symptoms are: headaches, drowsiness, nausea, shortness of breath, and poor judgement (see Figure D8-5, left photo, taken at the summit of Mauna Kea, for a possible illustration of poor judgement). The optional 1 hour stay at this altitude will help reduce the symptoms. The high elevation 129 �also requires sunscreen and sunglasses. Appropriate clothing should be used as protection against the higher levels of UV radiation. It will be cold (usually below freezing) and windy at the summit and warm clothing will be necessary and should consist of gloves or mitts, warm jacket, hats with ear protection, and sunglasses. There is a short trail that heads to the summit of the nearby cinder cone west of the Visitor’s Center. This cone contains numerous cored spindle bombs formed around large, mainly gabbroic xenoliths. Stop D8-71. Ellison B. Onizuka center for International Astronomy; Astronomer’s Mid-Level Facility. • UTM 242630E, 2186970N, parking lot. The Ellison B. Onizuka Center for Astronomy is where astronomers live and work rather than having to physically be at the telescopes at the summit. This cuts down on the number of adverse altitude effects that would be suffered if astronomers had to physically be at the telescopes on the summit. The map in Figure D8-4, below, shows the geological features at the summit of Mauna Kea and the location of the astronomical observatories (black dots). Figure D8-4: Map of the summit of Mauna Kea showing geological features and the location of the various telescopes. From Hazlett and Hyndman (2007, p.125). Stop D8-72. Lake Wai‘au Trailhead. Data sources: Hazlett and Hyndman (2007); Meguerian and Okulewicz (2007). • UTM 241495E, 2192385N; parking lot (mile 12.5) on right (east) side of road. • UTM 241471E, 2192349N; Lake Wai‘au Trailhead; west side of the highway a short distance south of the parking area. The light-coloured patches visible on the flanks of the Pu‘u Wai’au cone, located due west of the parking area, are due to hydrothermal alteration produced when steam and hot water percolated through the cone near the end of its eruption. The clay within the alteration products decreased the permeability of the cinder and resulted in increased runoff from rain and melting snow producing more erosional gullies than is evident on the flanks of unaltered cones. The trail leads upslope past the steep lobate edge of a flow erupted from Pu‘u Hau Kea ~40,000 years ago. The fracture patterns along the edge of the flow suggest the lava cooled while banked against ice. 130 �Stop D8-73. Top of pass, 230ft (70m) northwest of trail junction. Data sources: Hazlett and Hyndman (2007); Meguerian and Okulewicz (2007). • UTM 240690E, 2192472N. This point is the top of the pass between Pu‘u Wai‘au to the south and west and the taller Pu‘u Hau Kea to the north (see Figure D8-5). Lake Wai‘au should be just visible several hundred metres downslope to the right within a blasthole located at the north end of the crater floor. Stop D8-74. Lake Wai‘au. Data source: Hazlett and Hyndman (2007). • UTM 240690E, 2192472N. Tiny Lake Wai’au is one of the few natural bodies of water found within the State of Hawai‘i and is the highest lake in the state at 13,160ft (4011m). It persists rather than draining away due to impermeable clay weathered from ~3300-year-old Mauna Kea ash and the clay-rich hydrothermally altered cinder that comprises the cone. The lake occasionally overflows through a notch in the northwest rim of the crater. On the floor of the crater south of the lake is the remains of a rock glacier composed of a mixture of rock and ice that once flowed toward the lake. It is now a mass of hummocky light grey debris (Figure D8-5). The rough lava embankment along the north side of the lake is part of the 40,000yr old flow that was walked past downslope toward the road. The cavernous voids, mosaic fractures, and lava pillows suggests that this flow stopped against ice. There are many inclusions of coarsely granular gabbro and green dunite within the flow. Pu‘u Hau Kea Lake overflow point Lake Wai‘au Dry overflow streambed Lake Waiʻau Parking Lot Mauna Kea Access Rd Pu‘u Wai‘au Trail Rock glacier remnant Altered Cinder Figure D8-5: Google Earth Satellite image of the Lake Wai’au area and the various geological features. 131 �Stop D8-75a. Mauna Kea Summit Trailhead. Data source: Hazlett and Hyndman (2007). • UTM 241195E, 2193670N, Trail and Summit Parking Lot. The 1970ft (600m) long Mauna Kea Summit Trail begins on the other side of the road from the north end of the parking area located a short distance southwest from the Gemini Telescope. Stop D8-75b. Mauna Kea Summit. Data source: Hazlett and Hyndman (2007). • UTM 241474E, 2193526N (top of Pu‘u Wēkiu Cinder Cone). Mauna Kea’s summit (see Figure D8-6, left) is at the top of Pu‘u Wēkiu Cinder Cone at 13,796ft (4205m). On a clear day Mauna Loa is south; Hualālai is southwest; Kohala is north; and in the distance past Kohala is Haleakalā, on Maui. To the north and northwest are the domes of the summit telescope complex. The westernmost 4 telescopes are shown in Figure D8-6 (right). Figure D8-7 (left) shows a happy guy, possibly feeling the effects of altitude sickness (although this may his normal). Nonetheless, he was entertaining and did act relatively normal, except for the lack of clothing at an ambient temperature of <0oC (<32oF). His warmly dressed girlfriend consented to take a photo of the seven summiteers from the 2020 field trip (D8-7, right). The field trip participants missing from the photograph did not make the climb due the altitude (that is their story, and they are sticking with it). Figure D8-6: On the left is the snow-covered cinder cone comprising Mauna Kea’s summit. 4 of the mountain’s astronomical observatories, as seen from the summit, are on the right. Photo credits: A.D. MacTavish (2020). Figure D8-7: The left photo may illustrate the effects of high-altitude judgement loss on an unidentified gentleman at Mauna Kea’s summit (<0oC). The gentle slopes of Mauna Loa are in the right background. The right photo shows the seven 2020 Field Trip summiteers. Photo credits: A.D. MacTavish (2020). 132 �3.9. Day 9: Mamaloa Highway (Hawai‘i Belt Road; Hāmākua Coast) Between the northern end of Hilo Bay and the town of Honoka‘a, Highway 19 crosses the slopes of Mauna Kea’s extinct shield stage. These rocks are mostly Hāmākua Formation basalt flows overlain by up to 15ft (4.5m) of Laupāhoehoe Formation ash deposits erupted from vents near the summit of the volcano. This area once supported vast sugar cane fields which thrived on the chemically-weathered, red, volcanic ash soil and the high rainfall experienced on the windward side of the island. Geologically young gulches, some quite large and deep, have eroded down through the ash into the shield-stage flows. Three major gulches are crossed by this stretch of highway. Between Honoka‘a and Waimea the highway passes into a scenically beautiful saddle located between Mauna Kea and Kohala where there are remarkable changes in vegetation as you pass from the wet east side to the dry west side of the island. The Highway follows along the mostly alluvium-buried contact between the 2 volcanoes. Dozens of eroded and vegetated Laupāhoehoe alkalic cinder cones, erupted over the last 65,000 years, can be seen scattered over the slopes of Mauna Kea The forested Kohala East Rift Zone lies along the horizon to the north; Mauna Kea dominates the south. Day 9 Field trip stops on the Hāmākua Coast, Mauna Kea (see Figure D9-1) are: 76. Hawai‘i Tropical Botanical Gardens; accessed via the Old Mamaloa Highway, which diverts right from Highway 19 at the village of Papaikou; the diversion rejoins Highway 19 at the town of Papeeko via a left turn on Kuliamano Road and a right tun onto the highway; 77. ‘Akaka (442ft or 135m) and Hakūnā Falls (400ft or 122m); 78. Laupāhoehoe basalts; 79. Waipi’o Valley Overlook; 80. Waipi’o Valley Road and valley floor; 81. Scenic saddle between Kohala and Mauna Kea; dozens of alkalic Laupāhoehoe cinder cones. 79 80 81 Waimea 78 77 76 Hilo Figure D9-1: Day 9 field trip stops on Mauna Kea’s Hāmākua Coast. Figure modified after Hazlett and Hyndman (2007, p.114). 133 �Stop D9-76. Hawaii Tropical Botanical Garden. Data source: Hawai‘i Tropical Botanical Garden website (2019). • UTM 280430E, 2191905N; large parking lot on right (east) side of the Old Mamaloa Highway with access to the gardens; visitors must pay at the building opposite the parking lot for entry. The Hawaii Tropical Botanical Garden website states that the garden ‘is a museum of living plants that attracts photographers, gardeners, botanists, scientists, and nature lovers from around the world’. The garden contains over 2,000 species of tropical plants (see Figures D9-2 and D9-4, right) representing more than 125 families, and 750 genera. Two varieties of the orchids growing in the garden are shown in Figure D9-2. The 40-acre valley hosts a true tropical rainforest and is a natural greenhouse with fertile volcanic soil that is protected from the strong trade winds. Nature trails meander throughout the valley and provide beautiful views of waterfalls (see Figure D9-3) and the rugged coastline (Figure D9-4, left). Figure D9-2: Two of the many varieties of orchids in the Hawaiʻi Tropical Botanical Garden. Photo credit: A.D. MacTavish (2012). Figure D9-3: Onomea Falls, Hawaiʻi Tropical Botanical Garden. Photo credit: A.D. MacTavish (2012). 134 �Figure D9-4: Twin Rocks, Onomea Bay (left photo); Giant Spider Lily (Amaryllidaceae) (right photo); Hawaiʻi Tropical Botanical Garden. Photo credits: A.D. MacTavish (2012). Stop D9-77. ‘Akaka Falls and Hakūnā Falls (alternate). Data sources: Hazlett and Hyndman (2007); Robinson (2010). • UTM 274620E, 2196745N; parking lot. The trail to the falls (see Figure D9-5) is on the southwest side of the parking lot. This is a very popular and busy spot; if there is nowhere to park along the access road within a reasonable walking distance then this stop can be skipped. The walking roundtrip to both waterfalls takes about 30 minutes (not including the oohs and aahs). ‘Akaka Falls on Kolekole Stream is the tallest single waterfall in the state of Hawai‘i at 442ft (135m). Hakūnā Falls, located about 1800ft (550)m south of ‘Akaka Falls, is 400ft (122m) high, but was formed on a tributary flowing into Kolekole Gulch. Both formed as water flowed over resistant Hāmākua lava flows at the head of Kolekole Gulch. The view of Hakūnā Falls from the trail is disappointing due to tree growth and can be easily skipped. A careful look at the walls of the gorge after a rainfall will reveal a network of small falls resembling delicate shreds of lace on the cliffs. Figure D9-5: ʻAkaka Falls. Photo credit: Will Seaborn (2016), willseaborn.com website. 135 �Stop D9-78. Laupāhoehoe Point. Hazlett and Hyndman (2007); Robinson (2010); AGI Glossary of Geology, 4th Edition (1997). • UTM 265625E, 2212210N; parking area; good, if somewhat windy, lunch stop At Laupāhoehoe Point (see Figure D9-6) is a late-stage Mauna Kea ‘a‘ā flow of hawaiite that erupted from a vent well upslope that flowed into Laupāhoehoe Gulch and then into the sea. Hawaiite is a postshield stage, alkaline olivine basalt midway in composition between alkali olivine basalt and mugearite and is gradational into both (AGI Glossary of Geology, 4th Edition, 1997). This area was devastated on April 1, 1946 by a series of 6 tsunami waves, between 30 and 37ft (9 and 11m) high, that originated in the Aleutian Islands after a large megathrust earthquake. These waves destroyed a school and killed 21 students and 3 teachers. This same tsunami also struck Hilo and killed 159 people and destroyed more than 1300 homes and businesses (Tsunami Park). Figure D9-6: Laupāhoehoe Point from the 656 to 985ft (200-300m) high cliffs located to the south (left photo). Wave-washed Laupāhoehoe hawaiite ʻaʻā flow with the cliffs of the northeastern coastline in the background (right photo). Photo credits: A.D. MacTavish (2019). Stop D9-79. Waipi‘o Valley Overlook. Data sources: Hazlett and Hyndman (2007); Robinson (2010); Easton and Easton (1995). • UTM 229878E, 2226615N; parking area for viewpoint. The almost 1mi (1.6km) wide, ~6mi (9.6km) long Waipi‘o Valley (see Figure D9-7) is the largest and southernmost of 7 similar valleys that wrap around the eastern side of the extinct Kohala volcano. This valley was once 300 to 400ft (100 to 130m) deeper and was cut during the Pleistocene to a point about 360ft (110m) below present sea level (the Lualualei Stand). As continental glaciers receded on the northern continents the valley was gradually flooded by rising sea levels and is now gradually being infilled by sediment from Lālākea and Waipi‘o streams, which have created a broad valley floor with a very fertile floodplain that was once an important old Hawai‘i population centre. This lookout provides a spectacular view of the valley, its 1500 to 2000ft (460-610m) high northern wall, and the dark grey sandy beach. The exposed valley walls are composed of the slightly alkaline basalts of the 400,000 to 150,000yr old upper Pololū Formation which comprised the last eruptive activity of Kohala volcano. Steep sea cliffs and deep amphitheater-headed valleys are typical of the windward coasts of the Hawaiʻian Islands. There are several signboards near the lookout wall that provide information on the valley. 136 �Figure D9-7: Head of the Waipiʻo Valley and beach, from the overlook to the south, with the 1500 to 2000ft (460 to 610m) high sea cliffs in the back ground. Photo credit: A.D. MacTavish (2020). Stop D9-80. Waipi‘o Valley Road and Beach, Alternate (no formal stops). Data sources: Hazlett and Hyndman (2007); Robinson (2010). • UTM 228915E, 2226815N; road junction on valley floor. No formal stops are set for this road since there is little in the various guidebooks and the authors have not walked the road. Hazlett and Hyndman (2007) state that a walk along the road to the valley floor reveals several flows containing large, white, weathered, soft and crumbly plagioclase crystals up to 1in (2.5cm) in length (phenocrysts). Such a concentration of large crystals suggests that they accumulated near the top of a stagnant magma chamber after the main phase of shield activity had ceased. Lālākea Stream enters the valley over 2, narrow, 300 ft (90m) high waterfalls which can be easily seen from where the road reaches the valley floor. The cobbles in the stream bed contain many large phenocrysts of black pyroxene and green olivine. The mouth of the valley acts as a tsunami funnel which raises the waves to towering heights. The 1946 tsunami that devastated Hilo and the Laupāhoehoe school was about 40ft (12m) high at the beach and swept inland over ½ a mile (800m). The right fork in the road at the valley floor leads east to Waipi‘o Valley Beach. Stop D9-81. Saddle between Kohala and Mauna Kea Volcanoes. Data sources: Hazlett and Hyndman (2007); Robinson (2010); MacTavish (2019). • There are no formal stops on this stretch of highway due to a lack of safe parking areas. Between Honoka‘a and Waimea the Highway follows the elevated saddle and contact between the flows from Kohala and Mauna Kea. This contact is mostly buried beneath recent alluvium. Due to the funneling effect between the 2 volcanoes the velocity of the westward-blowing trade winds increases as it flows toward Waimea (known as the Venturi Effect). During the traverse from east to west the vegetation cover will change dramatically as the highway transitions from the often thickly forested (eucalyptus, swamp mahogany, cypress, iron wood pine, etc.) windward (wet) side of the island to the less-forested undulating grasslands characteristic of the leeward (dry) side of the island. Little rock is exposed along the Highway but there are some deeply-weather road cuts. As Waimea is approached the weathered and vegetated remnants of small cinder cones can sometimes be identified along both sides of the highway and on a clear day there are good views of Mauna Kea’s summit. 137 �3.10. Day 10 – Kohala Volcano – Waimea to Hāwī Day 10 examines Kohala Volcano, which, as described previously, is an extinct volcano forming the large, ridge-shaped northern peninsula of Hawaiʻi. It is the oldest volcano on the island and last erupted ~100,000 years ago. Kohala was higher in the past; however, after >100,000 years of subsidence and erosion it is now 5480ft (1670m) high. It is mainly covered by Hāwī Formation alkalic cinder cones and lava flows which overlie the older, late shield stage, ~400,000yr old Pololū Formation flows. Much of the northeastern flank of Kohala slid into the ocean somewhere between 400,000 and 150,000 years ago producing the >1500ft (460m) high cliffs that characterize the volcano’s northeastern shoreline (Hazlett and Hyndman. 2007). There are innumerable streams that have yet to erode the cliffs down to sea level producing a large number of waterfalls, particularly after heavy rain. The 10 stops planned for Day 10 examine the geological, archeological, and cultural aspects of the Kohala volcano (see Figure D10-1): 82. 83. 84. 85. 86. 87. 88. 89. 90. Pu‘u Kawaiwai cinder cone (lava of alkalic Hawaiite composition), southern Northwest Rift Zone; Scenic Overlook, Northwest Rift Zone axis, benmoreite exposure; Cinder cones on the north-northwest oriented axis of the Northwest Rift Zone; driving stop; View of Cinder Cones as well as Haleakala Volcano located on the Island of Maui; Pololu Valley Lookout; Residual coastal boulders; Moʻokini Luakini Heiau and Kapakai Kokoiki (King Kamehameha birthplace); Lapakahi State Park, ruins; Mugearite flow, Hawi Volcanics; pseudodykes in colluvium 88 a 88b 87 Hawi 86 85 3 89 84 90 83 91 82 Waimea Figure D10-1: Day 10 field stops on Kohala Volcano. Figure modified after Hazlett and Hyndman (2007, p.111). 138 �Stop D9-82. Pu‘u Kawaiwai Cinder Cone. Data source: Hazlett and Hyndman (2007). • UTM 214165E, 2218875N; parking area is a pull-out on the left (west) side of Highway. A gated road leads into the Pu‘u Kawaiwai hawaiite cinder cone (see Figure 10-2); if the owners are not present the walk can be made into the quarry, otherwise it can easily be viewed from a distance. In the quarry wall nearest the road (Figure 10-2, right) the structure of the cone as it grew is visible with crossbedding, erosional unconformities, and large blocks scattered throughout. The cone’s crater was filled with spatter and cinder when the eruption shifted to 3 other vents further downslope. Good views of Mauna Kea (see Figure 10-3), Mauna Loa, and the Kona coast of Hualālai are possible from the scenic lookout on a clear day. Figure D10-2: The left photo shows Pu‘u Kawaiwai Cinder Cone from Highway 19. The right photo is a close-up of the quarry excavated into the northeastern flank of the cinder cone. Photo credits: A.D. MacTavish (2019) Figure D10-3: The northern flank of Mauna Kea seen from the Stop D9-77 parking area located on the southwestern flank of Kohala. Photo credit: A.D. MacTavish (2019). 139 �Stop D9-83. Scenic Lookout, Benmoreite Lava. Data sources: Hazlett and Hyndman (2007); AGI Glossary of Geology (1997). • UTM 211338E, 2221675N; overlook parking area. Medium-grey benmoreite of the Hawi Volcanics (see Figure 10-4) comprise the rock-cuts across from, and to the north of the parking area. These rocks were erupted upslope from near the Pu‘u Loa Cinder Cone about 140,000 years ago. The hillsides around this location host numerous clusters of cacti. Benmoreite is a rare, unusual alkalic rock which at this locality consists of black plagioclase- and amphibole-porphyritic flows and chaotic mass-flow deposits, possibly lahars (Figure 10-4). Amphibole phenocrysts are rarely observed within Hawaiʻian lava and the ones here exhibit brown haloes. The AGI Glossary of Geology (1997) describes benmoreite as a silica-saturated to silica-undersaturated igneous rock intermediate between mugearite and trachyte in composition with a differentiation index of between 65 and 75 and with K2O:Na2O <1:2. Figure D10-4: Closeup of the benmoreite mass flow deposit (lahar?) located north of the field stop parking area. Photo credit: A.D. MacTavish (2019). Stop D9-84. Cinder cones along the Highway. Data source: MacTavish (2019). • UTM 207115E, 2229755N; no suitable road stops available, view from vehicle Along this stretch of highway south of the town of Hawi, where there are fewer trees bordering the road, can be seen a series of vegetated cinder cones that erupted along a north-northwest-oriented Kohala Northwest Rift Zone axis. Stop D10-85. Cinder cones and Hāleakala Volcano. Data source: MacTavish (2019). • UTM 204798E, 2233056N; park on widened area on right (east) side of highway. This is one of few the places along this highway where vehicles can safely stop to view old Kohala cinder cones located to the south (see Figure D10-5, left). On a clear day this is also a good location to view Hāleakala Volcano on the Island of Maui which is located to the northwest (see Figure D10-5, right). This stop can be skipped if the weather is not clear enough to view Hāleakala on Maui. 140 �Figure D10-5: The left photo shows a vegetated Kohala cinder cone. Hāleakala Volcano on the Island of Maui can be seen in the distance in the right photo. Photo credits: A.D. MacTavish (2019). Stop D10-86. Pololū Valley Scenic Lookout and trail. Data source: Hazlett and Hyndman (2007); Love Big Island website; Big Island Hikes website. • UTM 214210E, 2236565N; parking area and trailhead. This busy spot provides a good view of the scenic Pololū Valley and the associated coastline (see Figure D10-6). This valley is the northernmost of 7 large erosional valleys (gulches) located along the eastern coastline of Kohala. It is a large, flat-floored, amphitheatre-headed valley, like Waipi‘o Valley, and during the last ice age it was much deeper. The valley head, ~4mi (6.4km) inland (see Figure D10-6, right), is filled with a mugearite flow that flowed over fault scarps ~140,000 years ago. This spot provides good views of Kohala’s northeastern coastal cliffs (Figure D10-6, left), which are the headwall of a massive slide that carried debris into the Hawaiian Deep located 75mi (120km) away. In April 1946 a tsunami devastated the valley. The initial wave was 55ft (16.7m) high at the beach when it hit. The short, but steep, slippery when wet, 0.5mi (1km) long Pololū Trail (or Āwini Trail) switchbacks from the overlook down to the valley floor and a black sand beach. There is an elevation change of 490ft (150m) and the hike takes ~20-25min. The boulder strewn black sand beach is backed by lush tropical forest and is flanked by ~500ft (150m) cliffs. This segment of the Pololū Trail is apparently the first part of an extended trail that leads southeast to the Honokane Nui valley, but there is little information available on the extension. Only the beach is public land, the rest of the valley is privately-owned Figure D10-6: On the left is the northeastern coastline of Kohala near the Pololū Valley Scenic Lookout. On the right is the head of the Pololū Valley located south of the Scenic Lookout. Photo credits: A.D. MacTavish (2019). 141 �Stop D9-87. Residual Boulders. Data source: MacTavish (2019). • UTM 200345E, 2243410N, parking a short distance upslope from exposure. At this location are a large number of alkali basalt boulders resting on the surface of a highly-weathered alkalic basalt. These boulders were not formed by beach wave action and are certainly not erratic glacial boulders (see Figure D10-7). By looking closely, it can be easily seen that these are residual boulders remaining after fracture surfaces within the original flows were deeply weathered in the moist tropical Hawaiʻian environment. This produced a saprolite that was then eroded away (see Figure D10-8). At this location the saprolite process is incomplete with the softer saprolitized rock weathering away to leave the relatively less-altered rock core intact. The weathering here consisted of a combination of rainwater, wind, and possibly the occasional high wave. What remains are many rounded surface boulders and numerous partially exposed boulders within variably-weathered saprolite. Figure D10-7: Coastline residual boulders at field trip Stop D9-87 located at the northern tip of the island. Photo credit: A.D. MacTavish (2019). Figure D10-8: Residual boulders eroded out of saprolitized alkalic basalt flows. Note that many of the boulders remain imbedded within the strongly weathered flow in the photo on the right. Photo credits: A.D. MacTavish (2019). 142 �Stop D10-88a. Moʻokini Luakini Heiau, Kohala Historic Sites Monument. Data source: NPS Website. • UTM 199455E, 2242575N; drive east from Stop D10-87; park and walk to site if road too wet. Moʻokini Luakini Heiau (see Figure D10-9, left) is one of the oldest and most sacred ‘heiau’ (places of worship) in the Hawaiʻian Islands and is considered a living spiritual temple. The ancient Hawaiʻians had many types of heiau, each with their own distinct function and use. Heiau ranged in size from single upright stones to massive, complex structures. Larger heiau were built by ali'i (chiefs), but the largest and most complex luakini heiau (sacrificial temples), could only be constructed and dedicated by an ali'i 'ai moku. Luakini heiau were reserved for human or animal sacrifice rituals and were usually dedicated to the war god Ku. Rituals performed at these sites highlighted the ali'i 'ai moku's spiritual, economic, political, and social control over his lands and his authority over the life and death of his people. Mo'okini Heiau was a luakini heiau built in the shape of a parallelogram: the west wall is 267ft (81.3m) long, the east wall 250ft (76.2m), the north wall 135ft (41.1m), and the south wall 112ft (34.1m). Tapered, dry-stacked, mortarless stone walls that are 10ft (3m) wide at their base and between 7 and 14ft (2.1-4.3m) high enclose the heiau. Oral tradition says the rocks forming the walls were passed hand to hand along a line of thousands of men from the Niuli'i area 10mi (16km) to the east. Inside the northern end of the heiau is a large stone platform with smaller platforms scattered throughout the site that once supported thatched temple buildings. Outside, on the north side, is Papa-nui-o-leka, a stone on which human flesh was separated from bones after ritual sacrifice (see Figure 10-9, right). According to tradition, Mo'okini Heiau was the primary place of worship of the northern part of the Island. The site was active through the early part of the 19th century and was the war temple of King Kamehameha I, housing the war god of his family, ‘Ku-ka-'ili-moku’, before the transfer of the god to Kamehameha's new war temple, ‘Pu'ukohola Heiau’, located 21mi (33km) south near Kawaihae. Kamehameha’s son and heir Liholiho also used Mo'okini Heiau. In 1819, after his father's death, Liholiho ended kapu and abolished that part of the Hawaiian religion that depended on heiau. In spite of royal orders that they be destroyed, Mo'okini and several other large heiau were spared. In 1978, ‘Kahuna Nui’ (High Priestess) Leimomi Moʻokini Lum lifted the kapu (taboo) forbidding anyone but ali'i and kahuna from entering Mo'okini Heiau and also rededicated the heiau to the children of the land. In 1994, she again rededicated the heiau, this time to the children of the world. Visitors to the site often bring a flower or a lei to leave at the heiau as an offering of respect. Figure D10-9: The left photo shows the ruins of Moʻokini Luakini Heiau from the east. The flat-topped stone, shown in the right photo, is thought to be a where humans were ceremonially sacrificed and then skinned. Photo credits: A.D. MacTavish (2020). 143 �Stop D10-88b. Kapakai Kokoiki, King Kamehameha I Birthplace. Data source: NPS Website. • UTM 198805E, 2242410N; alternate stop. Approximately 2,000ft (610m) south of Mo'okini Heiau, is Kapakai Kokoiki (Royal Housing Complex) and the birthplace of King Kamehameha I (see Figure D10-10). It was typical for the housing complex of an ali'i 'ai moku to be near, and associated with, a luakini heiau. This is one of the few places in the Hawaiʻian Islands where historians know the exact location of a housing complex and its associated heiau. Over the centuries Kapakai served as the residence of ali'i 'ai moku when ceremonies were conducted in Mo'okini Heiau. Religious ceremonies lasted several days and nights and during this time, ali'i 'ai moku and high priests would leave the heiau for short periods to return to Kapakai. Kamehameha I was born in the Kapakai Royal Housing Complex and later stayed there while conducting ceremonies in Mo'okini Heiau. Figure D10-10: Western wall of Kapakai Kokoiki. Photo credit: Donnie B. MacGowan, lovebigIsland.com website Stop D10-89. Lapakahi State Park (Old Hawaiian Village Ruins). Data source: bigislandhikes.com. • UTM 197345E, 2233475N; park entrance; turn right (west) from Highway 270 to access road. • UTM 187140E, 2233525N; small parking area in front of small park building and picnic area. Lapakahi State Historical Park is a large area of ruins from an ancient Hawaiian village (see Figures D1011 and -12). The area offshore from the ruins is now a Marine Life Conservation District. Lapa kahi means "single ridge" and refers to the ancient ahupua'a (land subdivision) that existed here some 600yrs ago. The village was a place of maka‘āinana where fisherman and farmers lived and worked together. The farmers grew kalo (taro) and ‘uala (sweet potato). There are many kinds of ancient structures and artifacts to be viewed along the short, easy 1mi (1.6km) hike, including individual houses, large residential complexes, canoe storage houses, salt-making pans, kukui nut lampstands, and even a few kōnane games. Walk the trail through the village in a clockwise direction after taking a guide pamphlet from the building at the edge of the parking lot. The guide will describe what you are seeing at the various numbered stops. This trail takes 45-60min to complete. 144 �Figure D10-11: The left photo shows the southwestern portion of the ancient village. The right photo looks north along the northwestern coastline of the island from the southwestern end of the village. Photo credits: A.D. MacTavish (2019). Figure D10-12: The NPS Marker 8 in the left photo denotes a hollow stone used to make salt from sea water. The white stones are bleached coral. The right photo shows a stone ‘board’ used to play the game of Kōnane, which is similar to checkers. Photo credits: A.D. MacTavish (2019). Stop D10-90. Mugearite Flow, Hawi Volcanics. Data sources: Hazlett and Hyndman (2007); AGI Glossary of Geology (1997); MacTavish (2019). • UTM 197488E, 2232149; park on gravel road on right side of highway about 260ft (80m) north of field stop. • UTM 197520E, 2232065N; field stop, on east side of highway. This low outcrop is composed of vesicular, plagioclase feldspar-porphyritic mugearite which also seems to contain reddish to greenish grains which could be olivine partially altered to iddingsite. The AGI Glossary of Geology (1997) describes mugearite as an extrusive or hypabyssal alkaline igneous rock consisting of oligoclase with subordinate alkali feldspar and mafic minerals, often with olivine more abundant that clinopyroxene. Although generally nepheline-normative the rock may contain normative hypersthene, or even quartz and will exhibit a 45-65 differentiation index with normative plagioclase more sodic than An30. 145 �Stop D10-91: Pseudodykes in colluvium (alternate). Data source: Easton and Easton (1995); AGI Glossary of Geology (1997). • UTM 202825E, 2219965N; parking on gravel road, right (west) side of highway, ~395ft (120m) north of outcrop. • UTM 202890E, 2219855N; outcrop. This large highway road cut provides a good view of an outcrop consisting mainly of colluvium overlain by an ‘a‘ā flow. Colluvium is defined by the AGI Glossary as any loose, heterogeneous, and incoherent mass of soil and/or rock fragments deposited by rain-wash, sheetwash or slow, continuous downslope creep at or near the base of slopes or hills. What makes this outcrop interesting are the at least 5 pseudodykes developed in the colluvium on the seaward (west) side of the highway (see Figure D10-13). The authors have been unable to find a satisfactory or consistent definition in the literature of the Easton and Easton (1995) ‘pseudo-dykes’. Most suggest that the ‘pseudodykes’ are not intrusive, but are similar to clastic dykes seen in purely sedimentary environments. What do you think? Safety Warning: This is a very busy highway. Stay well to the right while walking south along the paved shoulder from the vehicles and while viewing the pseudodykes at the field stop. It is preferrable to view the pseudodykes from across the highway from the road cut that hosts them, so take considerable care when crossing the road. pseudodykes Figure D10-13: Colluvium outcrop with 3 of the 5 pseudodykes highlighted. Photo credit: Google Earth. 3.11. Day 11 – Waimea to Kailua-Kona Most of Day 11 (the final day of the field trip) will be examining the rocks of Hualālai Volcano. The stops planned for Day 11 are (see Figure D11-2): 92. 93. 94. 95. Mauna Kea western rift zone cinder cone field Hāpuna beach; basaltic ankaramite lava with pyroxene and olivine phenocrysts Hualālai and 1859 Mauna Loa flows contact Hualālai trachyte flows with large Pu‘u Wa‘awa‘a trachytic cinder/pumice cone to the southeast; and 96. Scenic Lookout; Kaʻūpūlehu Flow, Hualālai Northwest Rift Zone; alkalic basalt (1801-1802, from last known eruption), mafic/ultramafic intrusive xenoliths; cinder cones along rift to summit. 146 �Waimea 93 92 94 95 96 Kailua-Kona Figure D11-2: Location of Day 11 field trip stops. Figure modified from Hazlett and Hyndman (2007, p.106). Stop D11-92. Mauna Kea Western Rift Zone cinder cone field. Data source: MacTavish (2019). • UTM 218840E, 2206290; widened gravel shoulder on right side of Highway 200, about 1300ft (400m) past the junction with Highway 190. This location provides a good view of a large, vegetated, breached cinder cone (see Figure D11-3) and other, generally smaller cones associated with the Mauna Kea Western Rift Zone Cinder Cone Field. The only turn-around spot about is ~4600ft (1400m) southeast along Highway 200 on the right where a gated road leads into the large cinder cone. After turning around drive back to Highway 190 and turn right. Drive to the junction with Highway 19 in Waimea and turn west on Highway 19 and drive to the Hapuna Beach access road near the western coastline of the island. 147 �Figure D11-3: Breached and vegetated cinder cone associated with the Mauna Kea Western Rift Zone Cinder Cone Field. Photo credit: A.D. MacTavish (2019). Stop D11-93. Ankaramite Lava, Hapuna Beach State Park. Data source: Hazlett and Hyndman (2007); AGI Glossary of Geology (1997); MacTavish (2019). • UTM 204450E, 2212930N; parking lot. • UTM 204310E, 2213195N; take access walkway from parking lot to beach then walk north along beach to the outcrop on the east side of the beach at this location. Near the north end of the beach, just before it passes into the resort area to the north, are rock ledges composed of the rare basaltic lava ankaramite (see Figure D11-4). The AGI Glossary of Geology (1997) describes ankaramite as an olivine-bearing basanite containing numerous olivine and pyroxene phenocrysts. At this location the flow forms an irregular steep-sided ledge composed of a dark grey, thick, variably vesicular massive flow overlain by a well-developed and defined blocky ‘a‘ā flowtop. A close look at the base of the flow top shows blocks that are still partially connected to the underlying mass of the flow. A closer look shows that the flow contains numerous green, to yellowish-green olivine crystals and glomerocrysts and fewer black pyroxene grains, which are often associated with the olivine grains. Figure D11-4: The left photo shows the ankaramite outcrop at Hapuna Beach. The right photo is a closeup of the flow-top breccia at the top of the massive ankaramite flow. Photo credits: A.D. MacTavish (2019). 148 �Stop D11-94. 1859 Mauna Loa ‘a‘ā flow. Data source: Hazlett and Hyndman (2007). • UTM 204505E, 2104910N; parking in gravel lot on right side of highway. This partially vegetated 1859 Mauna Loa ‘a‘ā flow (see Figure D11-5) is the youngest Mauna Loa flow on this side of the island and it overlies older Hualālai flows. Figure D11-5: Partially vegetated 1859 Mauna Loa ʻaʻā flow. Photo credit: A.D. MacTavish (2019). Stop D11-95. Pu‘u Wa‘awa‘a Cinder Cone State Park. Data source: Pu‘u Wa‘awa‘a Ahupuaʻaʻ Ōhiʻa Cone Trail System Visitor Guide; AllTrails website. • UTM 202159E, 2192490N; automatic entrance gate; drive through gate and along road to the left to an information kiosk. Stop D11-95a. Information Kiosk. • UTM 202255E, 2192375N; information kiosk with parking. Pu‘u Wa‘awa‘a Cinder Cone (see Figure D11-6) is a large trachyte cinder and pumice cone that erupted from Hualālai about 100,000 years ago producing a set of thick trachyte flows. It is considered the largest cinder cone on the island and the oldest feature on Hualālai Volcano. Cone Trail guides and maps are available at the kiosk. The sharp curve on the highway just before the turnoff to the cone curves around exposed expressions of the trachyte flows and underlies most of the town of Pu‘uanahulu and the golf course at the Big Island Country Club. Stop D11-95b, Pu‘u Wa‘awa‘a Cinder Cone Trail (Alternate). • UTM 202965E, 2188455N, Summit The popular, moderate difficulty, 6.5mi (10.5km) long Pu‘u Wa‘awa‘a Cinder Cone Trail takes ~4hrs to complete and leads to the summit of Pu‘u Wa‘awa‘a at 3967ft (1209m). The summit peak provides an excellent view of the surrounding area to the sea and along the coast. Along the hike there is also a chance to see the native Hawaiʻian owl known as pueo and the native hawk known as ‘io. The cone once hosted an obsidian mine and is still part of a working ranch. 149 �Figure D11-6: Pu‘u Wa‘awa‘a Cinder and Pumice Cone. Photo credit: A.D. MacTavish (2019). Stop D11-96. Kaʻūpūlehu Flow, Hualālai 1801 to 1802 Alkalic Lava Flow. Data source: Hazlett and Hyndman (2007). • UTM 192870E, 2188715E; scenic lookout at widened highway shoulder just before the bridge that spans part of the flow. Please be very careful of traffic on the highway at this stop. This stop overlooks the partially vegetated alkalic Kaʻūpūlehu Flow field from the last known Hualālai eruption that took place in 1801 and 1802 (see Figure D11-7, left). These flows emanated from the Northwest Rift Zone located upslope to the southeast and contain a large number of dunite, gabbro, and peridotite xenoliths which will look like angular dark to light green chunks within the dark grey lava. This ~0.9mi (1.5km) wide flow field consists of both pahoehoe and ‘a‘ā flows, several well-developed surface flow channels, and some partially collapsed lava tubes (see Figure D11-7, right) Figure D11-7: The left photo shows the partially vegetated 1800 to 1801 Hualālai Kaʻūpūlehu alkaline lava flow field that heads downslope to the sea. The right photo shows a partially collapsed lava tube (skylight?) in the 1800-1801 flow field. Photo credits: A.D. MacTavish (2019). Safety Note: This a narrow pull-out along a very busy road, so be very careful of highway traffic. THIS IS THE FINAL STOP OF THE FIELD TRIP. 150 �4. Glossary of Volcanic Terms; (© G. J. Hudak, NRRI University of Minnesota, 2020) ‘A’ā lava: A Hawaiian term for lava that has a rough, jagged, spiny, and often clinkery surface. In thick aa flows, the surface comprises rubble composed of loose, rough lapilli and blocks that generally hides a thick, more massive flow interior (Tilling et al., 1987). The thickness of the surface crust of aa lavas is controlled by cooling (Kilburn, 2000, p. 291). Active volcano: A volcano that is currently erupting, one that has erupted during recorded history, or one that has erupted during recorded history and is likely to erupt again (Foxworthy and Hill, 1982). Accessory fragment: A lithic fragment composed of country rock that has been explosively ejected during an eruption (Cas and Wright, 1987, p. 54). Accessory fragments within pyroclastic deposits may be difficult to distinguish from accidental fragments. In general terms, referred to as a xenolith. Accidental fragment: A clast picked up locally by pyroclastic flows and surges (Cas and Wright, 1987, p. 54). Accidental fragments may be difficult to distinguish from accessory fragments. In general terms, referred to as a xenolith. Accretionary lapilli: Spherical aggregates (commonly with a concentric structure) formed by the accretion of moist ash in eruption clouds (White and Houghton, 2000, p. 495). Also used for all ash aggregates, including mud lumps (Houghton et al., 2000, p. 513). Achnelith: A type of juvenile fragment characterized by smooth, glassy moulded surfaces formed from lava spray from extremely fluid mafic eruptions (Walker and Croasdale, 1972). Agglomerate: A course, pyroclastic deposit composed of a large proportion of fluidal-shaped volcanic bombs that are formed, in the strictest sense, by a fall deposit in the immediate vicinity of a volcanic vent. It is best applied to describe bomb and scoria deposits that build strombolian cones, and should never be used as a non-generic term for a “volcanic breccia” (Cas and Wright, 1987, p. 359). A key component of identifying an agglomerate is that many bombs will plastically deform and will become agglutinated. Agglutinated: Melted together to form a single solid mass upon cooling (Hazlett and Hyndman, 1996). Aerosol: Fine liquid or solid particles suspended in the atmosphere. Aerosols composed of tiny droplets of sulfuric acid are commonly formed during explosive volcanic eruptions. Airfall: Volcanic ash that has fallen through the air from an eruption cloud. Airfall deposits are characteristically well-sorted and well-layered, and typically exhibit mantle bedding (Foxworthy and Hill, 1982; Cas and Wright, 1987). Alkalis: The elements potassium and sodium (Hazlett and Hyndman, 1996). Alkalic Basalt: Basalt-like rock compositions that are enriched in the alkali element sodium. Examples include nephelinites, hawaiites, and ankaramites (Hazlet and Hyndman, 1996). Alteration (see Hydrothermal Alteration). Alteration mineral assemblages: Mineral assemblages found in rocks that result from chemical reactions between the original rock and an agent of alteration (for example, hot volcanic vapors or hydrothermal fluids). Amygdaloidal: A volcanic texture comprising vesicles (rounded holes resulting when magma cools around gas bubbles) which have been subsequently filled by secondary minerals. Amygdule: An individual vesicle which has been subsequently filled-in by secondary minerals. 151 �Andesite: A grey to grey-green colored volcanic rock containing 53% to 63% silica (compositionally between basalt and dacite). Minerals commonly found in andesite include intermediate composition plagioclase and hornblende. Andesite magma: A magma with a chemical composition ranging from 53% to 63% which, upon crystallization, forms an andesite. Ankaramite: An alkalic basalt containing many large, black pyroxene crystals and a lesser number of green olivine crystals (Hazlett and Hyndman, 1996). Armoured lapilli: A type of accretionary lapilli composed of a crystal, pumice, or lithic fragment core which is surrounded by a rim of fine to coarse ash (McPhie et al., 1993, p. 29). Ash: A textural term for volcanic fragments less than 2mm in diameter (Fisher, 1966; Schmid, 1981). Ash is a common product of explosive volcanic eruptions. Ash cloud: A cloud of ash produced during pyroclastic eruptions (Miller, 1989). These clouds can result from rapid rising of the hot, buoyant ash-rich eruptive plume, or can be derived by elutriation at the top of a pyroclastic flow (Cas and Wright, 1987). Ash Cone: A low, broad volcanic cone enclosing a wide, shallow crater (Hazlett and Hyndman, 1996). Ash flow: A type of pyroclastic flow comprising dominantly ash-sized particles. Hot ash flows may be called “glowing avalanches” or “nuee ardentes”, and if their volume is large enough, may eventually form deposits known as welded tuffs. These types of flows are extremely dangerous and historically have killed hundreds of thousands of people. Asthenosphere: A zone of soft, nearly molten rock within the earth’s upper mantle. The tectonic plates of the earth ride on top of the asthenosphere (Hazlett and Hyndman, 1996). Atmospheric shock wave: A strong compressional shock wave caused by a combination of volcanic ejecta and sonic waves. Avalanche: A large mass of material or mixtures of materials (e.g., snow, ice, rock, soil, etc.) that is falling or sliding rapidly due to the force of gravity. Debris avalanches are avalanches composed of a mixture of earth materials (Foxworthy and Hill, 1982). Ballistic fragment: An explosively ejected rock fragment that follows a ballistic (arced) trajectory. Basalt: A dark colored (usually dark grey, dark green, or black), low silica content (45% to 53% SiO2) volcanic rock. Minerals commonly found in basalt include intermediate to calcium-rich plagioclase, pyroxene, and commonly olivine. Accessory minerals commonly include ilmenite and magnetite. Basaltic magma: A low viscosity, low silica (45% to 53% silica) magma that, upon crystallization, forms the volcanic rock basalt. Basanite: A variety of basalt that contains small crystals of plagioclase and pale gray nepheline (Hazlett and Hyndman, 1996). Base surge: A turbulent, low-density cloud of rock debris, water, and/or steam that moves over the ground surface at extremely high speeds. Base surges are commonly the result of directed volcanic explosions. Base surge deposits are commonly composed of cross-bedded deposits comprising ash and lapilli. Bimodal: A term used to describe a material composed of two distinctly compositionally and/or texturally different components. Commonly used to describe volcanic terrains that have nearly equal proportions of felsic and mafic volcanic rocks. 152 �Blocks: Fragments of solid rock greater than 64 millimeters in diameter that are ejected during volcanic eruptions. Blocks are commonly composed of accessory fragments made up of crystallized magma associated with the eruption (e.g., pieces of a lava dome). Blocky lava: Lava flows that are characterized by highly fractured surfaces which contain fragments of debris (usually flow fragments) up to several meters in diameter. The size of the surface fragments in blocky lavas is controlled by the rheology of the lava in the interior of the flow (Kilburn, 2000, p. 291). Boiling lake: A lake which has a temperature of nearly 100°C. Examples include the “Boiling Lake” on Dominica and a lake of mud on Saint Lucia (Bardintzeff and McBirney, 2000, p. 159). Bombs: Juvenile fragments of semi-solid or plastic magma ejected during a volcanic eruption. Based on their shapes after they hit the ground and cool, bombs are given various textural names including breadcrust bombs, cow-dung (cow pie) bombs, spindle bombs (fusiform bombs) and ribbon bombs. Bomb Sag: A depression in an ash layer made by the impact from a fragment deposited in the ash (Hazlett and Hyndman, 1996) Caldera: Large, circular to elongate, volcanic collapse depressions that form from the rapid extrusion of magma form a shallow subterranean magma chamber. In general, the diameter of a caldera is much greater than any of its individual volcanic vents (Williams and McBirney, 1979, p. 207). Caldera cycle: A commonly observed evolutionary sequence recognized in many caldera complexes. From oldest to youngest, the seven stages of the caldera cycle are: 1) regional tumescence and generation of ring fractures; 2) ignimbrite (pyroclastic) eruption(s); 3) caldera collapse; 4) pre-resurgent volcanism and intra-caldera sedimentation; 5) resurgent doming; 6) major ring fracture volcanism; and 7) terminal fumarolic and/or hot spring activity. Cinders: A term to describe generally highly vesicular, mafic lava lapilli. Cinder cone: A small, generally conical-shaped volcano formed by accumulation of ejected cinders and other volcanic debris that falls back to the earth close (proximal) to the location of the volcanic vent (Gardner et al., 1995). Clay (minerals): A group of aluminum-bearing hydrous phyllosilicate minerals (for example, kaolinite). Clay (textural): A sedimentary grain size classification for particles less than 1/256 mm in diameter, regardless of mineralogy. Cognate lithic fragment: Non-vesiculated juvenile magmatic fragments that have silicified from the erupting magma (Cas and Wright, 1987, p. 54). Columnar jointing: A type of fracture pattern resulting from the thermal contraction of hot volcanic rocks after their crystallization which commonly is expressed in elongate, pentagonal or hexagonal columns oriented perpendicular to the cooling surface. Columnar jointing is common in all compositions of lava flows, although it is generally best developed in mafic (basalt) lava flows and in felsic welded tuffs. Composite volcano: A generally steep sided volcano composed of a mixture of lava flows, pyroclastic deposits, and volcaniclastic sedimentary deposits. Composite volcanoes commonly have increasing slopes toward their summits since they generally have mainly lava flows and sedimentary deposits near their base and pyroclastic (tephra) deposits near their summits. Conduit: The underground passage or passages through which magma makes it way to the earth’s surface. 153 �Cooling unit: A group of hot pyroclastic deposits (ignimbrites) that cools at more or less the same time. A deposit from a single eruption that shows simple variations in the degree of welding is known as a simple cooling unit. When many ignimbrites occur over an extremely short period of time, each individual ignimbrite may be deposited, and start to weld over a previous deposit or group of deposits that are cooling and undergoing welding. The resulting deposits have several zones of partial and dense welding, and since they more or less cool together, are known as compound cooling units (Cas and Wright, 1987, p. 253-255). Coulée: A type of rhyolite lava flow that forms when lava issues from one side of a volcanic vent and produces a lava flow which is elongate in plan-view (Cas and Wright, 1987, p. 81). Crater: A steep sided, usually bowl or funnel shaped depression that commonly occurs at the top of a volcanic cone, and is often a vent for eruptions (Lipman, 2000, p. 643). Volcanic craters may be formed by either explosion or collapse in the vicinity of the volcanic vent. Crossbeds: Layers within sedimentary and/or volcaniclastic rocks that are inclined relative to the major bedding structures within the unit. Curie point: The temperature at which a body loses (by heating) or preserves (by cooling) its permanent magnetization. As rocks cool, the electromagnetic field aligns magnetic minerals in the magma, and their orientation is preserved as the rocks cool below the Curie point. Dacite: A generally light-colored, relatively silica rich (65% to 68 % SiO2) volcanic rock (extrusive equivalent of a quartz diorite or a tonalite). Dacitic magmas have a relatively high viscosity, and their associated volcanic eruptions may produce thick, muffin-shaped lava flows (lava domes) or, commonly, may be explosive and produce abundant tephra resulting in ash falls, ash flows, and surges. Dacites typically contain intermediate plagioclase (andesine or oligoclase) and quartz (>10%) with pyroxene and/or hornblende with minor biotite and/or sanidine (volcanic K-feldspar). Debris flow: A type of mass flow comprising a dense, cohesive, flowing mixture of sediment (mud through boulder sized materials, generally >50% by volume), water, and commonly, organic debris. Debris flows generally move downslope in laminar fashion due to the force of gravity (Vallance, 2000, p. 601; Carey, 2000, p. 627). Debris flows generated at volcanoes are commonly referred to as lahars. Decompressive melting: Melting that occurs when rocks undergo a decrease in pressure. This commonly occurs in the vicinity of hot spots as mantle rocks rise to shallower levels in the earth due to convective rise and upwelling (Sigurdsson, 2000, p. 15). Melting occurs as a result of decreasing pressure, not increasing temperature. Deposit: Earth materials that have accumulated by some natural process (Gardner et al., 1995). Deposits may be the result of volcanic (e.g., lavas or pyroclastic), sedimentary (either clastic or chemical), or hydrothermal (precipitation) processes. Devitrification: The solid-state transformation of volcanic glass into crystalline materials (AGI, 1976, p. 117). Devitrification tends to be more prevalent in densely-welded tuffs, but may also occur in less densely-welded or unwelded pyroclastic and/or volcaniclastic deposits. The main products of devitrification are cristobalite (SiO2) and alkali feldspar (KAlSi3O8) (Cas and Wright, 1987, p. 258). Diatreme: A funnel-shaped, pipe-like volcanic conduit, usually filled with volcaniclastic debris, emplaced by the explosive energy of gas-charged magmas. Diatremes are believed to result from hydrovolcanic fragmentation and subsequent wall rock collapse (Vespermann and Schminke, 2000, p. 683), and may reach depths up to 2500 meters. Diamond-bearing diatremes are economically important and are referred to as kimberlite pipes. 154 �Dike: A discordant, sheet-like body igneous body formed from the injection of magma into a fracture within the brittle crust of the earth (Carrigan, 2000, p. 219: Marsh, 2000, p. 191). Generally, a tabular igneous body which cross-cuts the planar structures in the adjacent rocks. Directed blast: A hot, low-density mixture of gas, rock debris, and ash that is propelled by a volcanic eruption and generally moves along the ground at high speeds (Miller, 1989). Dome (aka Lava Dome): A steep-sided mass of lava that is generally formed immediately above the volcanic vent from which it was extruded. Domes are generally circular in plan and have a relatively small surface area relative to other types of lava flows. Domes may be spiny, rounded, or flat on top, and often have rough, blocky surfaces formed by the fragmentation of the dome’s crust during intrusion. Domes may grow by extrusion of lava onto the outer surface of a previously formed dome (exogenous dome) or may be formed by inflation of a pre-existing dome (endogenous dome). Domes are most commonly the result of extrusion of viscous lava (primarily of the composition of rhyolite and dacite, but andesite may occur as well). Dormant volcano: A volcano that is not currently erupting, but is thought to be likely to erupt in the future. Downsag caldera: A type of caldera characterized by inward sloping topography, inward tilted wall rocks, and an apparent absence of large displacement caldera bounding faults (Lipman, 1997). Downsag calderas are believed to result from small volume eruption from a deep-seated subvolcanic intrusion. Dunite: A plutonic rock composed primarily of olivine. More specifically “A dunite is an ultrabasic igneous rock dominated by essential olivine (>90% volume), often with accessory clinopyroxene, orthopyroxene, spinel, ilmenite, and magnetite. Dunite is usually coarse- to medium grained and is a peridotite.” (http://www.alexstrekeisen.it/english/pluto/dunite.php). Epithermal mineralization: A mineral deposit formed from relatively low temperature (generally <350° C) hydrothermal solutions at shallow levels (<2km) in the earth’s crust. Epithermal mineralization is a common feature on many volcanoes. Eruption: The expulsion of volcanic materials (magma, volcanic gases) from a vent or fissure at the earth’s surface. In a general sense, eruptions are considered to be relatively large explosions which result in the expulsion of volcanic materials at or onto the earth’s surface. Extinct volcano: A volcano that is not presently erupting and is unlikely to do so in the future (Foxworthy and Hill, 1982). Extrusion: The eruption of molten rock (Hazlett and Hyndman, 1996). Facies: A part of a rock body that can be differentiated from another part of a related rock body by textural or compositional variations. The general appearance or composition of one part of a rock body as contrasted with other parts (AGI, 1976, p. 155). Facies changes: The textural and compositional changes that occur laterally and/or vertically within related rock bodies. Fire fountain: A spray of lava emitted from a vent or a fissure composed of a highly fluid mixture of basaltic magma and gas (Vespermann and Schminke, 2000, p. 683: Spudis, 2000, p. 697). Deposits from fire fountains produce mantling deposits composed of dense, plastic juvenile fragments and ash known as “agglomerates”. Fissure: A fracture or crack in the earth with an open separation. 155 �Flow banding: A foliation commonly observed in intermediate and felsic lavas, that results from shearing of the lava during laminar flow (Cas and Wright, 1987, p. 78). In rhyolite flows, flow banding is commonly exhibited by alternating bands comprising volcanic glass and spherulites (small, radiating bodies of devitrified glass). Fuel-coolant interaction: The interaction of magma (fuel) with external water (coolant) that may result in thermal explosions (Vespermann and Schminke, 2000, p. 683). Fumarole: A vent which releases volcanic gases. These include steam (H2O), carbon dioxide (CO2), sulfur dioxide (SO2), hydrogen sulfide (H2S), as well as other volatile gases emitted from subterranean magmas. Fumarolic activity: Volcanic gas emissions, with or without an accompanying change in the temperature or compositions of the gasses/fluids emitted (USGS Glossary of Volcano and Related Terminology). Gabbro: A phaneritic mafic igneous rock that is chemically equivalent to basalt. In detail, “Gabbros can contain: 25-50% of mafic minerals (Augite, Hypersthene, Olivine, Hornblende) and 45-70% of plagioclase (Labradorite or bytownite). If the plagioclase is less calcic than labradorite, the rock belongs in the Diorite family. Some low silica, dark-colored rocks containing olivine and plagioclase of the andesine range are by some petrologist included as gabbros.” (http://www.alexstrekeisen.it/english/pluto/quartzgabbro.php). Geyser: A special type of hot spring characterized by intermittent discharged of water and volcanic gases brought about by expansion of a vapor phase (generally steam) in the subsurface. Graben: An elongate crustal block that has moved downward relative to bounding fault systems (Foxworthy and Hill, 1982). Hawaiite: A type of alkalic basalt (Hazlett and Hyndman, 1996). Heterolithic: A clastic (volcaniclastic) deposit containing of a variety of different types of rock fragments. Hot spot: An area, generally located in the middle of a lithospheric plate, characterized by anomalous heat flow. Mantle material rises toward the earth’s surface and undergoes decompressive melting at hot spots which may form volcanoes (as in Hawaii) or cause partial melting of the overlying crust which leads to the formation of volcanoes (e.g., Yellowstone region). Hot spring: A thermal spring containing water at a higher temperature than the human body (98°F/37°C) Hydrothermal: Pertains to hot water or the action of hot water which has been heated by or in association with magma (Gardner et al., 1995). Hydrothermal alteration: Changes in rocks or minerals brought about by metasomatism with hydrothermal fluids (generally hot water). Hydrothermally altered: Minerals or rocks that have undergone hydrothermal alteration. Hydrothermal system: The system comprising the rocks, fluids, vapors, and conduits associated with hydrothermal activity. In general, hydrothermal systems have the following components: 1) a shallow magma chamber or cooling intrusion (provides the heat for the system); 2) fluids which can be of magmatic, meteoric, or connate origin, that are heated by the intrusion and flow through the rocks adjacent to (or sometimes within) the heat source; 3) fractures or high permeability zones which allow transfer of fluids from one part of the system to another part of the system. In most cases, this transfer 156 �is believed to be the result of buoyancy contrasts between the colder and warmer fluids within the system. Hydrovolcanic eruptions: A general term for eruptions caused by the mixing of magma with water (Vespermann and Schminke, 2000, p. 683). Encompasses hydroclastic, hydromagmatic, and phreatomagmatic eruptions. Hyaloclastite: A deposit comprising small, angular glass fragments formed by nonexplosive shattering of lava or magma flowing into water, ice, or water-saturated sediment (Batiza and White, 2000, p. 361: Schmidt and Schmincke, 2000, p. 383). Igneous: Refers to the processes associated with magma, or the rocks formed via the solidification of magma. Igneous rock: A variety of rock formed via crystallization from a magma. The two major classes of igneous rocks are volcanic (crystallized at or near the earth’s surface, for example, basalt) and plutonic (crystallized at depth within the earth, for example, gabbro). Ignimbrite: A term used for pyroclastic flow deposits, that is synonymous with “ash tuff” (Lipman, 2000, p. 643). According to Cas and Wright (1987, p. 98), the term should only be used to describe pumiceous pyroclastic flow deposits. Island arc: A curved chain of islands, generally convex towards the open ocean, which is bounded on its convex side by a deep oceanic trench (typically a subduction zone) and generally a deep-sea basin (AGI, 1976, p. 234). Jokulhlaup: The Icelandic term for “glacial outburst floods” which are commonly caused by subglacial volcanic eruptions. Juvenile fragment: Glassy or partially crystallized fragments which represent samples of an erupting magma. These include fragments such as pumice, scoria, reticulate, achneliths (Pele’s tears, Pele’s hair), and various types of volcanic bombs (Cas and Wright, 1987, p. 47-53). Kipuka: An area of older land surrounded by younger lava flows (Hazlett and Hyndman, 1996). Lahar: The Indonesian term for a debris flow or a mudflow originating on a volcano (Harris, 2000, p. 1301). Lahars are generally composed of volcanic materials, but can contain significant amounts of nonvolcanic materials derived from erosion during flow. Most volcanologists prefer this term to be used for the process and not the sedimentary deposits that it forms, but unfortunately, this distinction has been largely ignored in the geological literature. Many lahars are composed of sand and coarser materials, and thus, can be distinguished from “mudflows” which predominantly contain silt- or clay-sized grains (Rodolfo, 2000, p. 973). Landslide: A general term for relatively dry, gravity-induced movements of rock, sediment and/or soils (commonly with associated organic debris and/or human-made construction materials (e.g., houses, buildings, roads, etc.)) that are perceptible to the human eye. Lapilli: A textural term for fragments in volcanic rocks and volcanic deposits that range from 2mm to 64mm in diameter (Fisher, 1966; Schmid, 1981). Lateral blast: A volcanic eruption which is directed horizontally instead of vertically. Lateral blasts may be caused by sudden decompression of a shallow magma chamber residing within the flanks of a volcano (for example, the 1980 eruption of Mt. St. Helens), or along the base or side of a lava dome (for example, the 1902 eruption of Mt. Pelée in Martinique) (Nakada, 2000, p. 945). 157 �Laterite: A type of soil that forms in regions with warm, moist climates. Lateritic soils are commonly composed of kaolin clay, aluminum oxide, and iron oxide. Lateritic soils are commonly red in color (Hazlett and Hyndman, 1996). Lava: The term used for magma that has been erupted on to a planet’s surface. Lava flow: An outpouring of lava from a vent or fissure that spreads along the ground surface, as well as the crystallized rock resulting from solidification of the outpouring (Peterson and Tilling, 2000, p. 957). Lava lake: A region typically within the summit of a shield volcano which contains partially crystallized or molten lava which lies immediately above a volcanic conduit which joins the lava lake to the magma chamber. Strong magma convection within volcanic conduits sustains lava lakes within their respective volcanic vents (Walker, 2000, p. 285). Lava tube: A hollow region, commonly found within crystallized pahoehoe lava flows, which was filled with hot, flowing lava during a volcanic eruption. Lava tubes are formed when the top surface of a channelized lava flow crystallizes, and the magma flowing in the interior of the lava flow drains during and/or immediately following a volcanic eruption. Levées: Walls of lava that form at the margins of a lava flow. Lherzolite: “A lherzolite is an ultrabasic igneous rock dominated by essential Olivine and clinopyroxene and orthopyroxene in equal proportions. Accessory minerals include plagioclase, spinel, garnet, ilmenite, chromite and magnetite. Lherzolites are a peridotite and the main component of the upper mantle. Their aluminous phases change with pressure, with plagioclase present at low pressures, spinel at intermediate pressure and garnet at high pressure.” (http://www.alexstrekeisen.it/english/pluto/lherzolite(tl).php). Lithophysae: Radial aggregates of fibrous crystals which have formed around an expanding vesicle in a melt which is capable of flowing (Cas and Wright, 1987, p. 84). Lithophysae are commonly the result of vapor-phase crystallization within a rhyolitic magma. They should not be confused with spherulites, which are similar-shaped structures formed from devitrification of volcanic glass. Lithic: Fragments of previously-formed rocks or dense fragments that occur within volcaniclastic deposits. Lithic fragments may be accessory fragments, accidental fragments, or juvenile fragments. Lithospheric plates: The series of rigid slabs that comprise the earth’s lithosphere (crust and upper mantle. This term is synonymous with tectonic plates. Littoral: An adjective describing physical features or processes associated with shorelines of oceans, seas, or lakes (Peterson and Tilling, 2000, p. 957). Lobate lava: A submarine lava comprising elongate, flattish lobes with smooth, outer glassy skins (Batiza and White, 2000, p. 361). Maar: A type of monogenetic volcano, generally formed by subterranean phreatic or phreatomagmatic eruptions that occur as magma explosively interacts with ground water or subsurface moisture. Maar craters are cut into the surrounding country rock, vary from 10-500 meters deep, and range from a few hundred meters to 3 km in diameter. Maar volcanoes are generally surrounded by low, shallowly outward-dipping beds of well-bedded volcanic ejecta that rapidly decrease in thickness away from the vent. The volcanic deposits are mainly emplaced by base surges and fallout, and commonly contain very little (or in the case of phreatic eruptions, no) juvenile volcanic materials (Vespermann and Schminke, 2000, p. 685: Cas and Wright, 1987, p. 376-377). 158 �Mafic: A compositional term for igneous rocks which contain 45%-55% SiO2 (by weight). Mafic rocks are generally dark colored, and are characterized by mineralogy including pyroxene and calcium-rich plagioclase, variable amounts of olivine, and accessory minerals such as ilmenite and magnetite. Examples of mafic rocks include basalt and gabbro. Mafic lava: A lava with a silica content (by weight) ranging from 45-55% (AGI, 1976, p. 447; Peterson and Tilling, 2000, p. 957). Magma: A term used to describe subsurface molten rock (Jeanloz, 2000, p. 41). Magmas are generally considered to be silicate melts (Grove, 2000, p. 133; Wallace and Anderson, 2000, p. 149), but may also be composed of carbonatitic liquids (Spera, 2000, p. 171). Magmas are composed of up to three components (liquid, crystalline solids, and gas (or supercritical fluid) bubbles; Grove, 2000, p. 133), and may be fully liquid or partially crystalline. Lavas are magmas that have erupted on to a planet’s surface. Magma chamber: A subterranean region composed of magma that may have a conduit or set of conduits leading to a volcanic vent or vents on a planet’s surface. Magnetic polarity: The direction of the magnetic poles (either normal or reversed) that is preserved in igneous rocks after they cool below their Curie temperature (USGS Glossary of Volcano and Related Terminology) Magnitude: A numerical measure of the size of an earthquake based on the amount of seismic energy released. The magnitude of an earthquake is determined by measuring the highest-amplitude waves and correcting for distance and the type of seismometer used (McNutt, 2000, p. 1015). The seismic magnitude scale is logarithmic, with each increase in one unit on the scale equivalent to a tenfold increase in the wave amplitude. Mantle: The part of the earth’s interior lying above the outer core and below the Mohoroviĉić discontinuity. The mantle is commonly divided into three parts: the upper mantle (depths down to ~400 km), the transition zone (~400-670 km depth), and the lower mantle (~670-2900 km depth). Mantle bedding: Pyroclastic deposits generated by ash fall which maintain a uniform thickness and drape over all but the steepest topography (Cas and Wright, 1987, p. 96). Mantle plume: An elliptical, drop-shaped mass of mantle that ascends toward the earth’s crust due to its relatively lower density relative to the adjacent mantle. The density contrast is commonly the result of higher heat content of the plume, but may also be the result of chemical anomalies within the mantle (Perfit and Davidson, 2000, p. 89: Sigurdsson, 2000, p. 271). Mantle plumes are associated with intraplate rifting and volcanism. Mantle plumes are the hypothetical cause of hot spots (Hooper, 2000, p. 345). Melilite: A group of minerals that commonly form in the place of feldspar in silica-deficient, sodium-rich alkalic volcanic rocks (Hazlett and Hyndman, 1996). The melilite group is “A group of tetragonal sorosilicates with a disilicate anion (Si2O7)6- or an Al or B-bearing derivative thereof and the general formula given above, where M denotes a small- to medium-sized divalent or trivalent cation (mostly Mg and Al, or rarely Fe, B, Zn, Be, Si, etc.) and X is Si, Al or rarely Be or B. In general, Al or B replace one Si atom when M is a trivalent ion, but the charge can also be balanced by coupled substitution of Ca2+ with a monovalent ion, especially Na, and M3+ with M2+, such as in Alumoåkermanite, where Al3+ is still dominant on the M-site but the mineral is far from end-member composition. In petrology "melilite" usually refers to minerals in the åkermanite-gehlenite series, by far the most abundant members of the group.” (https://www.mindat.org/min-29310.html). 159 �Megabreccia: Coarse, heterolithic breccia deposits formed during caldera collapse, which contain fragments which are generally greater than one meter in diameter (Lipman, 1976). Megabreccia fragments may be so large that individual fragments may not be readily recognizable on the scale of an outcrop. Mesa lava: Generally rhyolitic in composition, a lava flow with an approximately circular plan which forms a biscuit-shaped body (Cas and Wright, 1987, p. 81). Mesobreccia: Heterolithic breccia deposits formed during caldera collapse which contain fragments that are generally less than 1 meter in diameter (Lipman, 1976). Metamorphic rock: In the strictest sense, rocks that have formed in the solid state in response to pronounced changes in temperature and/or pressure without any change in the bulk chemical composition of the rock. Metamorphic processes are generally confined to regions within the earth below the zones of weathering, cementation, and diagenesis. Metamorphism: In the strictest sense (isochemical metamorphism), the process by which consolidated rocks undergo textural and mineralogical changes brought about by changes in temperature and/or pressure. The textural and/or mineralogical changes associated with metamorphism are thermodynamic responses to the physical conditions present in the metamorphic environment. In general, increasing metamorphism results in dehydration of the rocks, as well as an increase in the grain size of the rocks. Metasomatism: A type of metamorphism characterized by the exchange of chemical species between rocks and their associated altering fluids and/or vapors. Moat sediments: A general term for sedimentary deposits that occur between the topographic walls and the resurgent central cores of the calderas. In felsic caldera systems, moat sediments are commonly intruded by, and associated with, lava domes. Monogenetic volcano: A volcano that erupts only once (Walker, 2000, p. 283). Monolithic: A type of volcaniclastic deposit in which all the clasts present are of the same composition. Moraine: A topographic feature or landform composed of an accumulation of sediment that has been carried and subsequently deposited by a glacier. Mudflow: A flowing mixture composed of water and mud (clay- and silt-sized sediments). The term should be used exclusively for mud-dominated mass flows, and should not be used as a substitute for the term “lahar” (Rodolfo, 2000, p. 973-974). Mudflows are common in both volcanic and non-volcanic environments. Mugearite: An “orthoclase-bearing oligoclase basalt, with major olivine, accessory apatite, and opaque oxides. Pyroxene may or may not be present.” (https://www.mindat.org/glossary/mugearite). Nested caldera: A type of caldera which is found within a larger, older caldera structure. Nueés ardente: The term used for a “glowing avalanche” resulting from a small-volume block and ash flow produced by the collapse of an actively growing lava dome (LaCroix, 1904). In recent years, the term has unfortunately been more widely used as a synonym for “ignimbrite”. Its use should be restricted to the original definition of LaCroix (Cas and Wright, 1987, p. 225). Orogeny: A term which describes the process of forming mountains, particularly by folding and thrusting (AGI, 1976, p. 308). 160 �Outwash: Sediments deposited by glacial meltwater beyond the active glacial ice. Outwash sediments are commonly characterized by poorly bedded gravels interlayered with well-bedded (and commonly cross-bedded) sands. Pahoehoe lava: A Hawaiian term to describe lava flows with smooth, continuous surfaces (Kilburn, 2000, p. 291). Pahoehoe flows may have a variety of surfaces described as smooth, ropy (characterized by rope-like, commonly braided flow folds on the lava flow’s surface) , or shelly (vesicular and cavernous; Cas and Wright, 1987, p. 66-67). Pahoehoe toes and lobes form when largely degassed mafic magma issues from tubes relatively far from the erupting vent. Palagonite: A yellow clay that forms from the hydration of basalt glass (Hazlett and Hyndman, 1996). Piecemeal caldera: A type of caldera characterized by an internal structure composed of several individual fault-bounded blocks (Lipman, 1997). Piecemeal calderas may result from non-uniform subsidence of a caldera formed from a single eruption, or may be the result of subsidence following a series of large eruptions (multicyclic; Lipman, 1997; Lipman, 2000, p. 655-656). Pele’s hair: A type of achnelith composed of thin, hair-like strands of volcanic glass. These thin, cylindrical strands of volcanic glass are commonly golden in color, have diameters between 1-500m in diameter, and may be up to 1 meter in length. They are formed from stretched magma droplets emitted into the atmosphere during fire fountaining and strombolian eruptions (Vergniolle and Mangan, 2000, p. 447). Pele’s tears: A type of achnelith composed of small droplets of shiny black volcanic glass that have been ballistically moulded and quenched during flight into spherical, dumbbell, or tadpole shapes. These droplets generally range from a few millimeters to a few centimeters in size, are generally dense, but locally may be quite vesicular (Vergniolle and Mangan, 2000, p. 447). Pele: The mythological Polynesian goddess of volcanoes. In Hawaii, this temperamental goddess makes her home in Kilauea’s fiery vent, Halemaʻumaʻu (Sigurdsson and Lopes-Gautier, 2000, p. 1297). Pelean eruption: A type of volcanic eruption characterized by a ground hugging glowing avalanche (pyroclastic flow) resulting from a mixture of hot volcanic gases, ash, and incandescent lava fragments. Pelean eruptions may occur when pyroclasts are blown out of a central volcanic vent and then collapse onto the earth’s surface to form a pyroclastic flow (Tilling, 1985). Pelean eruptions may also occur as a result of the explosive disintegration of a lava dome (as was the case for the lava dome on Mt. Pelée, Martinique in 1902). Peperite: A genetic term for a rock formed by in-situ disintegration and mixing of molten magma or lava with wet, poorly consolidated sediment (Batiza and White, 2000, p. 361). A breccia-like deposit formed from the extrusive or intrusive mixture of lava or magma with wet sediment (Schmidt and Schminke, 2000, p. 383). Peridotite: “Peridotites are a group of ultrabasic igneous rocks containing more than 40 vol% olivine with or without orthopyroxene and clinopyroxene. Accessory phases include garnet, spinel, plagioclase, ilmenite, chromite and magnetite. Peridotites comprise the bulk of the Earth’s upper mantle and are present as xenoliths within a wide range of mantle-derived magmas and within the mantle sequences of ophiolites.” (http://www.alexstrekeisen.it/english/pluto/peridotites.php). Perlite: Hydrated obsidian, generally light grey in color, that is commonly characterized by rounded, onion-skin-like fractures (perlitic cracks). Apache’s tears are unhydrated clumps of fresh obsidian that are commonly found within regions containing perlite. 161 �Phreatic eruption: A steam eruption, commonly associated with water, mud, and other earth materials, that is caused when groundwater, heated by a magma, flashes (and explosively expands) into steam (Harris, 2000, p. 1301). Phreatic eruptions expel no juvenile (magmatic) material, and are commonly the precursor to magmatic eruptive activity. Phreatomagmatic eruption: A type of explosive volcanic eruption that occurs when water (groundwater or surface water) comes in contact with hot magma. The quenching of the magma by the water causes the magma to violently fragment into juvenile (cognate) particles that are bounded by fracture surfaces and by rounded walls of broken vesicles. Due to the moisture present, accretionary lapilli are also common in volcanic deposits resulting from phreatomagmatic eruptions (Williams and McBirney, 1979, p. 247-248). Pillow breccia: A mixture of coarse, typically glassy fragments and broken to whole pieces of pillow lava formed from the shattering of pillow lava crusts (Batiza and White, 2000, p. 361). Pillow breccias commonly form in areas where pillow lavas are not strong enough to maintain their competence along steep submarine slopes or scarps. Pillow hyaloclastite: Hyaloclastite deposits immediately surrounding, and intimately associated with, pillow lavas. Pillow lava: A type of submarine lava flow consisting of interconnected, elongated lava tubes. Crosssections of individual lava tubes resemble pillows with convex upper surfaces and flat or concave lower surfaces (Schmidt and Schminke, 2000, p. 383). Both radial and concentric cooling fractures may be present along the margins of individual pillows, and these fractures are brought on by thermal contraction during cooling. Growth of the pillow tubes takes place as the outer, commonly striated outer glassy surface of the pillow tube fractures, and a new tube “buds” from the fracture in a manner similar to the way that toothpaste is squeezed out of a tube. Pipe-like alteration zone: A type of narrow, cylindrical or inverted-cone shaped, discordant hydrothermal alteration zone that is typically confined to a narrow region in close proximity to a synvolcanic structure (e. g. synvolcanic fault). Pipe–like alteration zones are commonly formed by the highest temperature hydrothermal fluids within a hydrothermal cell (Morton and Franklin, 1987). Plate (piston)-type caldera: A type of caldera in which the caldera floor subsides more or less evenly as one coherent block. Plate-(piston)-type calderas are believed to result from single, large volume pyroclastic eruptions from relatively shallow depth (hypabyssal) magma chambers. Plinian eruption: Named for Pliny the Younger (who witnessed the destruction of Pompeii by eruptions from Mt. Vesuvius), a type of violently explosive volcanic eruption that ejects large volumes of tephra high into the atmosphere (Harris, 2000, p. 1301). Pluton: A body of rock which has formed beneath the earth from crystallization and consolidation from a magma (AGI, 1976, p. 334). Plutons may be considered extinct magma chambers (Marsh, 2000, p. 191). Large plutons (>40 square miles in area) are called “batholiths”. Ponded flow: A term used to describe a lava flow that has ponded within a depression or a volcanic vent. A lava lake is a specific type of ponded flow that occurs a volcanic conduit. Pumice: Solidified fragments of quenched, highly vesicular (>60%) silicic magma or lava (Cashman et al., 2000, p. 421). The highly vesicular nature of pumice results from large volumes of gas rapidly expanding within a rapidly cooling magma. The low density of pumice commonly permits it to float on water for extended periods of time. Hot pumice, however, has been shown experimentally to sink rapidly upon interacting with water (Whitham and Sparks, 1986). 162 �Pyroclastic: Refers to processes resulting from the explosive fragmentation of a magma or lava. May also be used to describe the deposits formed by explosive volcanic activity and directly deposited by transport processes resulting directly from this activity (Cas and Wright, 1987, p. 8). Pyroclastic is a Greek term which means “fire-broken” (Harris, 2000, p. 1301). Pyroclastic fall: The “rain-out” of pyroclastic material through the atmosphere from an eruption jet or eruption plume during an explosive volcanic eruption (Wilson and Houghton, 2000, p. 545; Houghton et al, 2000, p. 555). Pyroclastic fall deposit: Volcaniclastic (pyroclastic) deposits formed from the rain-out of clasts through the atmosphere from an eruption jet and/or plume during an explosive eruption (Houghton et al., 2000, p. 555). Fall deposits typically exhibit mantle bedding, are well sorted, and commonly show welldeveloped planar stratification (Cas and Wright, 1987, p. 95-96). Pyroclastic flow: A dense, hot, dry, high particle concentration mixture of gas and hot rock fragments (ash, pumice, blocks, etc.) that travels along the ground surface, typically at high velocity (generally on the order of hundreds of feet or meters per second; Harris, 2000, p. 1301) away from a volcano. The high speeds of pyroclastic flows are possible because they flow over a thin layer of hot, commonly expanding and escaping gases. Most of the material within a pyroclastic flow is contained within concentrated particle dispersion located at the flow’s base (Wilson and Houghton, 2000, p. 545). Pyroclastic flow deposit: Pyroclastic (volcaniclastic) deposits that are left by pyroclastic flows (Cas and Wright, 1987, p. 96). The deposits are usually topographically controlled (infilling valleys and topographic depressions), massive, and poorly sorted. Depending upon their thickness and heat retention, pyroclastic flow deposits may coalesce into welded tuffs. Pumice-rich pyroclastic flow deposits are often called “ignimbrites”. Pyroclastic surge: A type of turbulent, low density (low particle concentration) pyroclastic cloud or pyroclastic density current. Being more dilute than pyroclastic flows, surges can sweep over ridges, hills, and other topographic boundaries. Two kinds of surges are known: wet surges have temperatures <100°C and contain steam that condenses into water droplets that surge along the ground surface with gas and pyroclasts; and dry surges, which have temperatures >100°C, and form by either hydrovolcanic eruptions with low water/magma ratios, or by magmatic eruptions driven solely by expanding magmatic gases (Valentine and Fisher, 2000, p. 571). Pyroclastic surge deposit: Pyroclastic deposits that are left by pyroclastic surges. These deposits mantle topographic features but also generally thicken within topographic depressions. These deposits are generally well-sorted, and are enriched in crystals and lithic fragments relatively to pyroclastic flow deposits. Surge deposits commonly exhibit unidirectional sedimentary bedforms, including low angle cross-bedding, dune forms, climbing dune forms, pinch and swell structures, and chute and pool structures (Cas and Wright, 1987, p. 98). Quenching: The rapid cooling of magma to form glass (Batiza and White, 2000, p. 361). Fuel-coolant interactions commonly lead to quenching. Abrupt quenching may cause a rapid volume decrease which leads to fragmentation of the glass (cooling-contraction granulation). Reticulite: An exceptionally porous type of scoria containing porosities ranging from 95-99%98% (Vergniolle and Mangan, 2000, p. 447; McPhie et al., 1993, p. 27). Commonly referred to as “threadlace” scoria, reticulite is made up of a honeycomb-like network of thin glass fibers. Rhyolite: A volcanic rock containing greater than 68% silica (by weight). Rhyolites are composed primarily of alkali feldspars (sanidine and orthoclase) and quartz (>10% by volume), with lesser amounts of sodic plagioclase (albite, oligoclase), hornblende, or biotite. Accessory minerals include zircon, 163 �apatite, and tourmaline. Due to their high silica content (and thus high degree of polymerization), rhyolite lavas are very viscous and commonly form lava domes, mesa lavas, or coulees. Rhyolitic magmas with high gas contents typically explode violently to form pyroclastic flows, pyroclastic surges, and pyroclastic falls. Rhyolite magma: A magma which contains greater than 68% silica by weight. Rift: A linear topographic feature formed by crustal extension. Rift structures associated with volcanism are commonly composed of a graben with a central high region, which is usually the site of active volcanism (for example, along the mid-ocean ridges). Rift Zone: A zone of fissures and volcanic vents that commonly form along the flanks of volcanoes (Hazlett and Hyndman, 1996). Ring fracture/Ring fault: The arcuate bounding faults upon which caldera (cauldron) subsidence takes place. Ring fractures (faults) define the structural limits of calderas. Most observed ring faults are nearly vertical or dip steeply inward (toward the center of the caldera), and this is thought to be a result of doming of the caldera structure following its initial formation (Lipman, 2000, p. 649-650). Ropy Pahoehoe: A type of pahoehoe lava characterized by flexible crusts that are bent into tight folds as lava flows. These tight folds form lava surfaces that appear to be made up of a series of braided ropes (Kilburn, 2000, p. 295). Satellite vent: A secondary vent on a volcano, commonly located on the volcano’s flank. Scoria: Solidified fragments of quenched, highly vesicular (>60%) mafic magma or lava (Cashman et al., 2000, p. 421). The highly vesicular nature of scoria results from rapid cooling of gas-rich lava. Scoria cone: Small volcanic landforms formed from focused (single-vent) subaerial strombolian eruptions of basalt or basaltic-andesite magma. These features have an inverted cone-shaped profile and are generally circular in plan, although elongate scoria cones can be formed from multiple-vent volcanic eruptions (Cas and Wright, 1987, p. 371-372). Seismic wave: A term for elastic earth waves formed by either earthquakes or explosions. Seismic waves include both surface waves as well as body waves. Seismicity: The phenomenon of earth movement or seismic activity. Seismograph: A scientific instrument used to detect and record seismic waves. Semiconformable alteration zone: A regional zone of hydrothermal alteration typically characterized by a sheet-like or cloud-like geometry. Semiconformable alteration zones are generally quite extensive in permeable rock units (e.g., tuffs, medium- to coarse-grained clastic sediments and sedimentary rocks), and are generally patchy in less permeable rock units (e.g., lavas, intrusions). These zones commonly are found along the periphery of “pipe-like” alteration zones, which are generally confined to regions in close proximity to synvolcanic structures (e.g., synvolcanic faults zones) (Morton and Franklin, 1987). Shelly pahoehoe: A type of pahoehoe lava characterized by highly vesicular, extremely fragile crusts that form over hollow lava blisters. The surfaces of these blisters break easily when stepped upon, giving the impression of walking on eggshells (Kilburn, 2000, p. 295). Shield volcano: A broad, low-relief volcano constructed by flows of relatively fluid lava (e.g., basalt: Spudis, 2000, p. 698). Flank slopes on shield volcanoes are typically < 5° (Zimbelman, 2000, p. 771). Silica: The chemical compound silicon dioxide, SiO2. 164 �Silicic: A term used to describe silica-rich volcanic rock or magma (Miller, 1989). A chemical classification for a type of rock or magma containing >62% SiO2 (Peterson and Tilling, 2000, p. 958) or 63% SiO2 (Cas and Wright, 1987, p. 16) by weight. Silicic lava: A lava with a silica content greater than 62% (by weight). Synonymous with the term “felsic lava” (Peterson and Tilling, 2000, p. 957). Sinter: A type of fragile, commonly white or grey rock formed by precipitation of silica from cooling hydrothermal solutions at or near a hydrothermal vent. Precipitation of siliceous sinter (often with associated sulfide minerals and precious metals) commonly occurs in neutral and acid hydrothermal systems under the influence of biogenic agents such as algae and bacteria (Cas and Wright, 1987, p. 316). Slabby pahoehoe: A type of pahoehoe lava with a surface composed of slabs of broken lava crust that are up to meters across and up to several centimeters thick (Kilburn, 2000, p. 295). Solfatara: A type of steam vent or dry fumarole that is characterized by quiet discharge (<20 m/s), and that precipitates a significant amount of sulfur (Hochstein and Browne, 2000, p. 850-851). Spatter: Fragments of fluid lava that are thrown out of a vent during an eruption (Hazlett and Hyndman, 1996). Spatter bomb: A glassy pyroclast greater than 64mm in diameter that takes on a fluidal shape by the force of ejection (Vergniolle and Mangan, 2000, p. 447). Spherulite: Typically rounded, radiating arrays of crystal fibers produced by the high temperature devitrification of volcanic glass. In felsic rocks, the crystal fibers are generally composed of alkali feldspar and a silica polymorph (either quartz or cristobalite), whereas in mafic rocks the fibers commonly consist of plagioclase and/or pyroxene. Spherulites typically have diameters of 0.1-2.0 cm, but can be much larger (commonly up to 20 cm). Isolated spherulites are generally spherical, but adjacent spherulites may impinge upon one another to produce long chains that are often aligned with flow foliation (McPhie et al., 1993, p. 24-25). Spreading center/Spreading ridge: Places on the ocean floor characterized by active volcanism and where separation of lithospheric plates takes place. Stratovolcano: A generally steep sided volcano composed of alternating layers of lava flows, pyroclastic deposits, and commonly, volcaniclastic sedimentary deposits (Walker, 2000, p. 283). Stratovolcanoes commonly have increasing slopes toward their summits since they generally have mainly lava flows and sedimentary deposits near their base and pyroclastic (tephra) deposits near their summits. Also called a “composite volcano”. Stony rhyolite: Very finely crystalline rhyolite lava (Cas and Wright, 1987, p. 84). Strombolian eruption: Volcanic eruptions of basaltic magma, slightly more violent than Hawaiian eruptions, that produce large amounts of scoria and ash around a central vent to form a cone. Strombolian eruptions are typically pulsating and have periods of several seconds (Wolf and Sumner, 2000, p. 321). The deposits consist of lava spatter, vesicular bombs, scoria lapilli, and mafic ash (Vespermann and Schminke, 2000, p. 683). Named after Stromboli, an Italian volcano. Subduction zone: A sloping region at collisional plate boundaries where one tectonic plate overrides another tectonic plate. In most regions, continental crust overrides oceanic crust which is then consumed in the subduction zone (continental – oceanic plate boundary), but in many areas, oceanic crust may be overridden by another plate of oceanic crust (oceanic – oceanic plate boundaries). Deep 165 �oceanic trenches commonly occur as the surface landform associated with subduction zones. Melting of the subducting slab commonly produces magma which rises to the earth’s surface to produce volcanic arcs. Surtseyan eruption: Hydrovolcanic eruptions dominated by jets of wet tephra (scoria and ash) that result in the formation of tuff cones. The term “surtseyan” is generally used for volcanoes erupting through seawater. Named after Surtsey, a volcano which emerged from the sea off the coast of Iceland in 1963 (Vespermann and Schminke, 2000, p. 683). Synvolcanic: A term used to describe a process or feature that was active or produced during volcanic activity. Synvolcanic fault: A fault or geological structure present or produced at the time of volcanic activity. Tectonic: A general term used to describe the forces involved in the deformation of the earth’s crust. Commonly used to also describe the geological structures or features produced by such deformation. Tectonic plate: One of the large segments of the earth’s lithosphere (crust and upper mantle, up to 250km thick) that comprise the earth’s outer shell. At the present time, there are 16 major tectonic plates that “float” on top of the asthenosphere, the plastic layer in the earth’s mantle. Tephra: A general term used by volcanologists to describe all fragmental volcanic ejecta produced during explosive volcanic eruptions (Dehn and McNutt, 2000, p. 1271). This includes ash (<2mm diameter fragments), lapilli (2-64 mm diameter fragments and fragments greater than 64 mm in diameter known as bombs (semi-solid or plastic ejecta) or bombs (solid ejecta) (Tilling et al., 1987). Thread-lace scoria: See “reticulite”. Trap-door caldera: A type of caldera formed when one part of the caldera floor subsides to a greater depth than the other side of the caldera floor. In general, trap-door calderas have a partial ring fracture (associated with the side of greatest caldera collapse) and a hinge area (associated with the side of least collapse). Trap-door calderas may represent either calderas that have undergone incomplete collapse, or calderas formed from eruptions from shallow asymmetrical magma chambers (Lipman, 1997: Lipman, 2000, p. 654. Tremor: A continuous vibration of the ground around active volcanoes (Vergniolle and Mangan, 2000, p. 447). Tremors defined on seismographs may have either a regular sine-wave appearance (harmonic tremor) or an irregular, pulsating appearance (spasmodic tremor) (McNutt, 2000, p. 1015). Tuff: A lithified volcaniclastic rock composed primarily of ash, with up to minor volumes of lapilli and/or blocks and bombs (Fisher, 1966). Originally used as a non-genetic rock name, common use today typically implies (incorrectly) that the tephra comprising the rock was deposited while hot. Similar deposits that have no indication of being hot while deposited are commonly referred to as “tuffaceous” (McPhie et al., 1993, p. 8). Tuff cone: A type of hydroclastic volcano that is generally higher than (generally >50 m high), and has steeper external flanks (commonly >25°) than tuff rings or maars (Vespermann and Schminke, 2000, p. 684). Craters within tuff cones are generally higher in elevation than the adjacent land surface. Tuff cones are made up primarily of juvenile clasts deposited from lateral surges, airfall, and associated volcaniclastic remobilization processes. Tuff ring: A type of hydroclastic volcano, generally <50m high, defined by craters with low depth/width ratios that sit at or above the elevation of the adjacent land surface. The rims around tuff rings are 166 �composed of juvenile and accidental clasts and are deposited in beds with dips <25° (Vespermann and Schminke, 2000, p. 684). Tumescence: The doming or uprising of a volcano commonly due to inflation of a shallow magma chamber. Regional tumescence commonly occurs prior to a major pyroclastic eruption, but may also occur following an eruption as less volatile magma is emplaced into the shallow crust (Smith and Bailey, 1968). Tuya: A flat-topped, steep-sided volcano that erupted into a lake thawed into a glacier by volcanic heat (Smellie, 2000, p. 403). Commonly referred to as a “table mountain”. Unconformity: A surface of erosion that separates younger strata from older rocks (AGU, 1976, p. 448). Variolite: A spherulite-like radiating aggregate composed of feathery, needle-like crystals of plagioclase and pyroxene that occur in mafic volcanic rocks (typically basalt). Variolites may result from devitrification, but are commonly believed to be formed in subaqueous rocks by quench-induced crystallization (Cas and Wright, 1987, p. 420). VEI index: The Volcanic Explosivity Index, which is a measure of the size of an eruption based on its magnitude, intensity and destructive power. The VEI is measured on an eight-point scale, where “8” is the most destructive and powerful eruption (Cioni et al., 2000. p. 477). Vent: A surface opening through which volcanogenic materials are erupted (Davidson and DeSilva, 2000, p. 663). Typically thought of as a hole in a planet from which volcanic products (magma, ash, etc.) are erupted (Spudis, 2000, p. 697). Vesicle: A frozen bubble in a volcanic rock. Vesicles are formed when magma crystallizes around a gas bubble (Spudis, 2000, p. 697). Vesicular: A textural term describing volcanic rocks filled with frozen gas bubbles (vesicles). Vesicular tuff: Tuffs containing millimeter to centimeter-sized, irregular to round vesicles which are interpreted to form during trapping of air or vapor in wet ash deposits (Vespermann and Schminke, 2000, p. 683). Vesuvian eruption: Commonly used as a synonym for a “Plinian” eruption (e.g., Tilling, 1985), , but also used to describe basaltic eruptions which involve long-sustained gas streaming with little ash being released (as in the 1906 eruption of Vesuvius; Cas and Wright, 1987, p. 130). Viscosity: A measurement of the ratio of shear stress to the rate of shear strain in a fluid (Williams and McBirney, 1979, p. 20). In common language, how easily a fluid will flow. Considered the most important physical property of a magma because it largely determines eruptive style as well as volcano morphology. Magma viscosity generally increases as the silica content of the magma increases (due to silica polymerization) and as the temperature of the magma decreases. Magma viscosity may also be affected by the presence of trace elements (e.g., Ti) or volatiles (e.g., H2O, CO2, SO2, etc.). In general, common magmas increase in viscosity in the following order: komatiite, basalt, andesite, dacite, rhyodacite, rhyolite. Volcaniclastic: A non-genetic term used to describe any fragmental aggregate of volcanic parentage (Cas and Wright, 1987, p. 8). Rocks formed by the fragmentation of volcanic materials (either magma or volcanic rocks) irrespective of the method of fragmentation. Pyroclastic rocks and epiclastic rocks are both considered to be “volcaniclastic”. Volcanic bomb: Juvenile fragments of semi-solid or plastic magma ejected during a volcanic eruption. Based on their shapes after they hit the ground and cool, bombs are given various textural names 167 �including breadcrust bombs, cow-dung (cow pie) bombs, spindle bombs (fusiform bombs) and ribbon bombs Volcanic cycle: A general term used to describe a period of increased volcanic activity. Volcanic field: A region comprising a large number of volcanic edifices. Volcanic fields are usually associated with basaltic volcanism, and commonly comprise a number of small, monogenetic volcanoes (e.g., cinder cones, maars, tuff cones, tuff rings, small shield volcanoes, lava domes). Fields may form in linear trends associated with tectonic structures (such as faults), on the flanks of larger composite or shield volcanoes, or within calderas (Connor and Conway, 2000, p. 331). Volcanic landslide: A landslide that occurs along the flank of a volcano. Volcano: A mound, hill or mountain constructed by the extrusion of lava and/or pyroclastic material from beneath the ground (Fisher et al., 1997, p. 43). A vent in the earth’s crust from which molten lava, pyroclastic materials, volcanic gases, etc. issue (AGU, 1976, p. 457). Vulcanian eruption: An explosive volcanic eruption generally expelling less than 1km3 of material, but with an eruption column that may reach heights of up to 10-20km (Nakada, 2000, p. 945). These eruptions last on the order of seconds to minutes (Morrissey and Mastin, 2000, p. 463). Welding: The sintering together of hot, glassy fragments, irrespective of shape and size, by compactional lithostatic load (Cas and Wright, 1987, p. 165). Welded tuff: A hard pyroclastic rock compacted by internal heat and pressure from overlying pyroclastic deposits. 4.1. Glossary References AGI, 1976. Dictionary of Geological Terms: American Geological Institute, Anchor Press, Garden City, New York, 472 pages. Bardintzeff, J.-M., and McBirney, A, R., 2000. Volcanology, 2nd Edition: Jones and Bartlett Publishers, Sudbury, Massachusetts, 268 pages. Batiza, R., and White, J. D. L., 2000. Submarine Lavas and Hyaloclastite, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 361-381. Carey, S. D., 2000. Volcaniclastic Sedimentation Around Island Arcs, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 627-642. Carrigan, C. R., 2000. Plumbing Systems, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 219-235. Cas, R. A. F., and Wright, J. V., 1987. Volcanic Successions: Modern and Ancient: Allen and Unwin, London, 529 pages. Cashman, K. V., Sturtevant, B., Papale, P., and Navon, O., 2000. Magmatic Fragmentation, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 421-430. Cioni, R., Marianelli, P., Santacroce, R., and Sbrana, A., 2000. Plinian and Subplinian Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 477-494. Connor, C. B., and Conway, F. M., 2000. Basaltic Volcanic Fields, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 331-343. Davidson, J., and DeSilva, S., 2000. Composite Volcanoes, in Sigurdsson, H., 2000, Encyclopedia of 168 �Volcanoes: Academic Press, San Diego, California, p. 663-681. Dehn, J., and McNutt, S. R., 2000. Volcanic Materials in Commerce and Industry, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 1271-1282. Fisher, R. V., 1966. Rocks composed of volcanic fragments and their classification: Earth Science Reviews, v. 1, pp. 287-298. Fisher, R. V., Heiken, G., and Hulen, J. B., 1997. Volcanoes: Crucible of Change: Princeton University Press, Princeton, New Jersey, 317 pages. Foxworthy and Hill, 1982. Volcanic Eruption of 1980 at Mount St. Helens: The First 100 Days: USGS Professional Paper 1249. Gardner et al., 1995. Potential Volcanic Hazards with Regard to Siting Nuclear Power Plants in the Pacific Northwest: USGS Open-File Report 87-297. Grove, T., 2000. Origin of Magmas, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 133-147. Harris, S. L., 2000. Archaeology and Volcanism, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 1301-1314. Hazlet, R. W., and Hyndman, D. W., 1996, Roadside Geology of Hawaii: Mountain Press Publishing Company, Missoula, MT, 304 p. Hochstein, M. P., and Browne, P. R. L., 2000. Surface Manifestations of Geothermal Systems with Volcanic Heat Sources, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 835-855. Hooper, P. R., 2000. Flood Basalt Provinces, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 345-359. Houghton, B. F., Wilson, C. J. N., Smith, R. T., and Gilbert, J. S., 2000. Phreatoplinian Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 513-525. Jeanloz, R., 2000. Mantle of the Earth, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 41-54. Kilburn, R. J., 2000. Lava Flows and Lava Fields, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 291-305. LaCroix, A., 1904. La Matagne Pelee et ses eruptions. Paris, Masson. Lipman, P. W., 1976. Caldera collapse breccias in the western San Juan Mountains, Colorado: Geological Society of America Bulletin, v. 87, p. 1397-1410. Lipman, P. W., 1997. Subsidence of ash-flow calderas: relation to caldera size and magma chamber geometry: Bulletin of Volcanology, v. 59, p. 198-218. Lipman, P. W., 2000. Calderas, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 643-662. Marsh, B. D., 2000. Magma Chambers, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 191-206. McNutt, S. R., 2000. Volcanic Seismicity, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 1015-1033. 169 �McPhie, J., Doyle, M., and Allen, R., 1993. Volcanic Textures: A guide to the interpretation of textures in volcanic rocks: University of Tasmania Centre for Ore Deposit and Exploration Studies, Hobart, Tasmania, 198 pages. Miller, 1989. Potential Hazards from future volcanic eruptions in California: United States Geological Survey Bulletin 1847. Morrissey, M. M., and Mastin, L. G., 2000. Vulcanian Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 463- 475. Morton, R. L., and Franklin, J. M., 1987. Twofold classification of Archean volcanic-associated massive sulfide deposits: Economic Geology, v. 82, p. 1057-1063. Nakada, S., 2000. Hazards from Pyroclastic Flows and Surges, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 945-955. Perfit, M. R. and Davidson, J. P., 2000. Plate Tectonics and Volcanism, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 89-113. Peterson, D. W. and Tilling, R. I., 2000. Lava Flow Hazards, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 957-971. Rodolfo, K. S., 2000. The Hazard from Lahars and Jokulhlaups, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 973-995. Schmid, R., 1981. Descriptive nomenclature and classification of pyroclastic deposits and fragments: recommendations of the IUGS Subcommission on the Systematics of Igneous Rocks: Geology, v. 9, pp. 41-43. Schmidt, R., and Schminke, H.-U., 2000. Seamounts and Island Building, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 383-402. Sigurdsson, H., 2000a. The History of Volcanology, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 15-37. Sigurdsson, H., 2000b. Volcanic Episodes and Rates of Volcanism, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, 271-279. Sigurdsson, H., and Lopes-Gautier, R., 2000. Volcanoes and Tourism, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, pp. 1283-1299. Smellie, J. L., 2000. Subglacial Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, pp. 403-418. Smith, R. L. and Bailey, R. A., 1968. Resurgent Cauldrons: Geological Society of America Memoir 116, pp 613-662. Spudis, P. D., 2000. Volcanics on the Moon, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 697-708. Spera, F., 2000. Physical Properties of Magmas, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 171-190. Tilling, 1985. Volcanoes: United States Geological Survey General Interest Publication. Tilling, Heliker, and Wright, 1987. Eruptions of Hawaiian Volcanoes – Past, Present, and Future: United States Geological Survey General Interest Publication. 170 �USGS Glossary of Volcano and Related Terminology: United States Geological Survey / Cascades Volcano Observatory, Vancouver, Washington, http://vulcan.wr.usgs.gov/Glossary/ volcano_terminology.html Vallance, J. W., 2000. Lahars, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 601-616. Valentine, G. A., and Fisher, R. V., 2000. Pyroclastic Surges and Blasts, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 571-580. Vergniolle, S. and Mangan, M., 2000. Hawaiian and Strombolian Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 447-461. Vespermann, D., and Schminke, H.-U., 2000. Scoria Cones and Tuff Rings, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 683-694. Walker, G. P. L., 2000. Basaltic Volcanoes and Volcanic Systems, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, pp. 283-289. Walker, G. P. L., and Croasdale, R., 1972. Characteristics of some basaltic pyroclastics: Bulletin of Volcanology, v. 35, p. 303-317. Wallace, P., and Anderson, A. T., 2000. Volatiles in Magmas, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 149-170. White, J. D. L., and Houghton, B., 2000. Surtseyan and Related Phreatomagmatic Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 495-511. Whitham, A. G., and Sparks, R. S. J., 1986. Pumice: Bulletin of Volcanology, v. 48, p. 209-223. Williams, H., and McBirney, A. R., 1979. Volcanology: Freeman, Cooper and Co., San Francisco, California, 397 pages. Wilson, C. J. N., and Houghton, B. F., 2000. Pyroclastic Transport and Deposition, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 545-554. Wolf, J. A., and Sumner, J. M.,2000. Lava Fountains and Their Products, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 321-329. Zimbelman, J. R., 2000. Volcanism on Mars, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 771-783. 171 �5. Field Guide References Aciego, S.M.; Jourdan, F.; DePaolo, D.J.; Kennedy, B.M. , Renne, P.R.; and Sims, K.W.W. 2010 Combined U-Th/He and 40 Ar/39 Ar Geochronology of Post-shield Lavas from the Mauna Kea and Kohala volcanoes, Hawaii; Geochimica et Cosmochimica Acta, 74(5), p.1620-1635. AllTrails website, www.alltrails.com . Babb, J.L., Kauahikaua, J.P., and Tilling, R.I. 2011. The story of the Hawaiian Volcano Observatory—A remarkable first 100 years of tracking eruptions and earthquakes: U.S. Geological Survey General Information Product 135, 60p. Big Island Hikes website, www.bigislandhikes.com 135 Clague, D.A. and Dalrymple, G.B. 1987. The Hawaiian-Emperor volcanic chain, Part 1, Geologic Evolution; in Decker, R.W., Wright, T.L., and Stauffer, P.H., eds., Volcanism in Hawaii: U.S. Geological Survey Professional Paper 1350, v.1, Chapter 1, p.5-54. Clague, D.A. and Dixon, J.E. 2000. Extrinsic controls on the evolution of Hawaiian ocean island volcanoes; Geochemistry, Geophysics, Geosystems (G3), v.1, no.4, 1010, 9 p. Clague, D.A., and Moore, J.G. 1991. Geology and petrology of Mahukona Volcano, Hawaii: Bulletin of Volcanology, v. 53, no. 3, p. 159–172. Clague, D.A.; Paduan, J.B.; Caress, D.W.; Moyer, C.L.; Glazer, B.T.; and Yoerger, D.R. 2019. Structure of Lō‘ihi Seamount, Hawai’i and lava flow morphology from high-resolution mapping; Frontiers in Earth Science 7, Article 58, p.1-17. Clague, D. R. and Sherrod, D.R. 2014. Growth and Degradation of Hawaiian Volcanoes; in Characteristics of Hawaiian Volcanoes, Poland, M.P., Takahashi, T.J., and Claire M. Landowski, C.M., editors; U.S. Geological Survey Professional Paper 1801, p.97-146. Dana, J.D. 1890. Characteristics of volcanoes, with contributions of facts and principles from the Hawaiian Islands; New York, N.Y., Dodd, Mead; 399p. and Co. Eakins, B.W.; Robinson, J.E.; Kanamatsu, T.; Naka, J.; Smith, J.R.; Takahashi, E.; and Clague, D.A. 2003. Hawaii's volcanoes revealed: U.S. Geological Survey Geologic Investigations Series Map I-2809, 1 plate, https://pubs.usgs.gov/imap/2809/. Easton, R.M. 1978. Stratigraphy and Petrology of the Hilina Formation: The oldest exposed Lavas of Kilauea Volcano; unpublished M.Sc. Thesis, University of Hawaii at Manoa, Honolulu, Hawaii, 274p. Easton, R.M. 1987. Stratigraphy of Kilauea Volcano; in Decker, R.W., Wright, T.L., and Stauffer, P.H., eds., Volcanism in Hawaii: U.S. Geological Survey Professional Paper 1350, v.1, Chapter 11, p.243260. Easton, R.M. and Easton, M.G. 1995. Highway Geology of the Hawaiian Islands; Easton Enterprises; 168p. Easton, R.M. and Garcia, M.O. 1980. Petrology of the Hilina Formation, Kilauea Volcano, Hawaii; Bulletin Volcanologique 43; p.647-673. Foulger, G.R. and Anderson, D.L. 2006. The Emperor and Hawaiian Volcanic Chains: How well to they fit the plume hypothesis? www.MantlePlumes.org website. Frey, F.A. and Clague, D A. 1983. Geochemistry of diverse basalt types from Loihi Seamount, Hawaii: petrogenetic implications; Earth Planet. Sci. Lett. 66, p.337–355. 172 �Garcia, M.O., Hanano, D., Flinders, A., Weis, D., and Kurz, M. 2012. Age, geology, geophysics, and geochemistry of Mahukona Volcano, Hawaiʻi; Bulletin of Volcanology 74, p.1445-1463. Garcia, M.O.; Kurz, M.D.; and Muenow, D.W. 1990. Mahukona: the missing Hawaiian volcano; Geology, 18: p.1111-1114. Garcia, M.O., Muenow, D.W., Aggrey, K.E., and O’Neil, J.R. 1989. Major element, volatile, and stable isotope geochemistry of Hawaiian submarine tholeiitic glasses; J. Geophys. Res. 94, p.10525–10538. Gazdar, Nasir. 2003. Hawaii’s Volcanoes Revealed; PowerPoint Presentation, revise 2018. Hawai‘i Tropical Botanical Garden website – www.htbg.com . Hazlett, Richard. 2014. Explore the Geology of Kīlauea Volcano; Hawaiʻi Pacific Parks Association, Revised Edition 2014; 146p. Hazlett, R.W. and Hyndman, D.W. 2007. Roadside Geology of Hawaiʻi; Mountain Press Publishing Company, MT, 5th Edition (originally published in 1996), 304p. Head, J.W. and Wilson, L. 1989. Basaltic pyroclastic eruptions: Influence of gas-release patters and volume fluxes on fountain structure, and the formation of cinder cones, spatter cones, rootless flows, lava ponds, and lava flows; J. Volcanol. Geotherm. Res. 37: p.261-271. Holcomb, R.T. 1976. Preliminary map showing products of eruptions, 1962-1974 from the upper east rift zone of Kilauea volcano, Hawaii: US Geological Survey Map MF-811, 1:24,000. Holcomb, R.T. 1987. Eruptive history and long-term behavior of Kilauea volcano; in Decker, R.W., Wright, T.L., and Stauffer, P.H., eds., Volcanism in Hawaii: U.S. Geological Survey Professional Paper 1350, v.1, Chapter 12, p.261-350. Jackson, E.D., Clague, D.A., Engleman, E., Friesen, W.F., and Norton, D. 1981. Xenoliths in the alkalic basalt flows of Hualalai Volcano, Hawaii; U.S. Geological Survey Open File Reports, 81-1031; 33p. Kirby, S.H. and Green, H.W. 1980. Dunite xenoliths from Hualalai Volcano: Evidence for mantle diapiric flow beneath the island of Hawaii; American Journal of Science, 280-A: p.550-575. Lockwood, J.P. and Hazlett, R.W. 2010. Volcanoes: Global Perspectives; John Wiley & Sons, 541p. Love Big Island website, www.lovebigisland.com . Macdonald, G.A., Abbott, A.T., and Peterson, F.L. 1983. Volcanoes in the Sea – The Geology of Hawaii: Second Edition, University of Hawaii Press, Honolulu, HI, 517p. MacTavish, A.D. 2019. Unpublished geological field notes, July 30 to August 9, 2019. MacTavish, A.D. 2020. Unpublished geological field notes, February 11 to 21, 2020. Malahoff, A. 1987. Geology of the summit of Loihi submarine volcano; in R.W. Decker, T.L. Wright, and P.H. Stauffer, eds., Volcanism in Hawaii: USGS Professional Paper 1350, v.1, Chapter 6, p.133–144. Mattox, S. 1994. A teacher’s guide to the geology of Hawaii Volcanoes National Park. (Activity 4.2). Honolulu, HI: Hawaiʻi Natural History Association. Merguerian, C. and Okulewicz, S. 2007. Geology of Hawaii; Hofstra University, Geology 280F – Field Trip Guidebook, Summer Session Two – 23 July to 02 August 2007; 137p. Moore, J.G. and Clague, D.A. 1992. Volcano growth and evolution of the island of Hawaii; Geological Society of America Bulletin 1992; 104, no. 11, p.1471-1484. 173 �Moore, J.G., Clague, D.A., Holcomb, R.T., Lipman, P.W., Normark, W.R., and Torresan, M.E. 1989. Prodigious submarine landslides on the Hawaiian Ridges; Journal of Geophysical Research, 94: p.17465-17484. Moore, J.G., Normark, W.R., and Lipman, P. W. 1979. Loihi Seamount – a young submarine Hawaiian volcano; in Proceedings of the Hawaii Symposium on Intraplate Volcanism and Submarine Volcanism, Hilo, 127p. Morgan, W.J. 1971. Convection plumes in the lower mantle; Nature 230, p.42-43. Hawai‘i Volcanoes National Park Illustrated Trails Map. 2010. National Geographic Trails Illustrated Map 230. Peterson, D.W. and Moore, R.B. 1987. Geologic history and evolution of volcanic concepts, Island of Hawaii; in Decker, R.W., Wright, T.L., and Stauffer, P.H., eds., Volcanism in Hawaii: U.S. Geological Survey Professional Paper 1350, v.1, Chapter 7, p.149-189. Pu‘u Wa‘awa‘a AhupuaʻaʻŌhiʻa Cone Trail System Visitor Guide. Robinson, J.E. and Eakins, B.W. 2006. Calculated volumes of individual shield volcanoes at the young end of the Hawaiian Ridge; J. Volcanol. Geotherm. Res., 151: p.309-317. Robinson, R.C. 2010. Illustrated Geological Guide to the Island of Hawaii; Santa Monica College; 287p. Robinson, R.C. 2012. Hawaii Volcanoes National Park, A Geologic Guide; 70p. Rowland, S.K. and Munro, D.C. 1993. The 1919-1920 eruption of Maunaiki, Kilauea: chronology, geologic mapping, and magma transport mechanism; Bulletin of Volcanology, v.55, p.190-203. Schmincke, H-U. 2004. Volcanism; Springer-Verlag, 324p. Seach, J. 2022. www.volcanolive.com website, list compiled by John Seach. Stearns, H.T. 1946. Geology of the Hawaiian Islands; Hawaii Div. Hydrogr. Bull. 8: 105p. Swanson, D.A. and Christianson, R.L. 1973. Tragic base surge in 1790 at Kilauea volcano; Geology, v.1, p.83-86. Temblor.net Website. 2018. USGS Volcanic Hazard Map from Lava Flows (2010). Tilling, R.I., Heliker, C., and Swanson, D.A. 2010. Eruptions of Hawaiian Volcanoes – Past, Present, and Future: U.SW. Geological Survey General Information Product 117, 63p. USGS Hawaiian Volcano Observatory (HVO) Website. Evolution of Hawaiian Volcanoes. www.usgs.gov/observatories/hvo US National Park Service (NPS), Hawaiʻi Volcanoes National Park. • Kīlauea Iki Trail Guide. • Kipukapuaulu Trail Guide • Mauna Ulu Eruption Guide. Will Seaborn, Visualization Artist and Photographer. 2016. www.willseaborn.com website. Walker, G.P.L. 1990. Geology and volcanology of the Hawaiian Islands; Review Article, Pacific Science, vol. 44, no. 4: p.315-347. White, J. D. L. and Houghton, B. 2000. Surtseyan and Related Phreatomagmatic Eruptions, in Sigurdsson, H., 2000, Encyclopedia of Volcanoes: Academic Press, San Diego, California, p. 495-511. 174 �Wikipedia. 2022. Evolution of Hawaiian Volcanoes; 7p. Dark Tourism website, www.darktourism.com . 175 � Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource Institute on Lake Superior Geology Dublin Core The Dublin Core metadata element set is common to all Omeka records, including items, files, and collections. For more information see, http://dublincore.org/documents/dces/. Title A name given to the resource The Volcanoes of the Island of Hawaii Description An account of the resource The Volcanoes of the Island of Hawaii: Field Trip Guide Institute on Lake Superior Geology Special Publication 3 Allan MacTavish, M.Sc., P.Geo. George Hudak III, Ph.D., P.Geo. December 2023 Creator An entity primarily responsible for making the resource Institute on Lake Superior Geology Date A point or period of time associated with an event in the lifecycle of the resource 2023-12 Format The file format, physical medium, or dimensions of the resource PDF Language A language of the resource English Type The nature or genre of the resource Text