1 10 75 https://digitalcollections.lakeheadu.ca/files/original/72cc0c183e5d4fc8ab0811c66edf140d.pdf 7527b93a85865027caf582e7cb0434fb PDF Text Text 72nd Annual Meeting Thunder Bay, Ontario - May 21-22, 2026 Institute on Lake Superior Geology Part 1 – Program and Abstracts �Thank you to our sponsors! �65th Annual Meeting Institute on Lake Superior Geology May 21-22, 2026 Thunder Bay, Ontario HOSTED BY: Mark Puumala and Peter Hinz Co-Chairs Ontario Geological Survey (Retired) Proceedings - Volume 72 Part 1 – Program and Abstracts Compiled and edited by Pete Hollings & Mark Smyk Cover Photos: Top: Amethyst veins in Rossport Formation at the Blue Points Amethyst Mine, north of Highway 11-17 near Big Pearl Lake.Middle: Neoarchean mafic metavolcanic rocks, Highway 102 at the intersection with Mud Lake Road. Bottom: Corestones of the McKenzie Granite at the Archean-Paleoproterozic unconformity, Highway 11-17 near Crystal Beach. All photos courtesy Mark Puumala. �72nd Institute on Lake Superior Geology Volume 72 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trips 1 & 4: “Classic” Geological Sites in the Thunder Bay Area Trip 2: Geology of the Quetico Supprovince North of Thunder Bay Trip 3: Gold Deposits of the Shebandowan Greenstone Belt Trip 5: Structural Geology and Gold Mineralisation of the Mine Centre Area Trip 6: Amethyst Deposits of Thunder Bay Reference to material in Part 1 should follow the example below: Akin, K. and Swanson-Hysell, N., 2026. Constraining the 3-D Geometry of the Duluth Complex, MN, Using Magnetic Fabrics and Paleomagnetic Data. In; Hollings, P. and Smyk,, M., (Eds.), Institute on Lake Superior Geology Proceedings, 72nd Annual Meeting, Thunder Bay, Ontario, Part 1 - Abstracts and Proceedings. v.71, part 1, 1-2. Published by the 72nd 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 72nd ILSG Annual Meeting - Part 1 Table of Contents Institutes on Lake Superior Geology, 1955-2026............................................................... ii Sam Goldich and the Goldich Medal................................................................................. iv Goldich Medal Guidelines................................................................................................. iv Institute on Lake Superior Geology Goldich Medal............................................................v Goldich Medalists.............................................................................................................. vi Sam Goldich and the Goldich Medal................................................................................ vii Goldich Medal Guidelines............................................................................................... viii Goldich Medal Committee ................................................................................................ ix 2026 Goldich Medal Recipient.......................................................................................... ix Citation for Goldich Medal Recipient..................................................................................x Honoring the Pioneers of Lake Superior Geology............................................................. xi In Memoria........................................................................................................................ xii Report of the Chair of the 71st Annual Meeting .............................................................. xvi Eisenbrey Student Travel Awards.................................................................................... xix Joe Mancuso Student Research Awards.............................................................................xx Doug Duskin Student Paper Awards..................................................................................xx 2026 Student Paper Awards Committee........................................................................... xxi Board of Directors............................................................................................................ xxi Local Committee.............................................................................................................. xxi Field Trip Leaders and Guidebook Authors.................................................................... xxii Index..................................................................................................................................88 -i- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Institutes on Lake Superior Geology, 1955-2026 # 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 17 1971 Duluth, Minnesota D.M. Davidson 18 1972 Houghton, Michigan J. Kalliokoski - ii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 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 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 - iii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 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 & A. Blaske 60 2014 Hibbing, Minnesota J. Miller & M. Jirsa 61 2015 Dryden, Ontario R. Cundari & P. Hinz 62 2016 Duluth, Minnesota J. Miller, C. Schardt & D. Peterson 63 2017 Wawa, Ontario A. Pace, A. Wilson & T.J. Bornhorst 64 2018 Iron Mountain, Michigan L. Woodruff, W. Cannon & E.K. Stewart 65 2019 Terrace Bay, Ontario P. Hollings & M.C. Smyk 66 2020 Meeting cancelled Cancelled by the COVID-19 pandemic 67 2021 Virtual meeting M. Jirsa, M. Smyk & P. Hollings 68 2022 Sudbury, Ontario R.M. Easton & W. Bleeker 69 2023 Eau Claire, Wisconsin R. Lodge, E.K. Stewart, & C. Ames 70 2024 Houghton, Michigan T.J. Bornhorst, E. Vye, P. Cobin, & J. Degraff 71 2025 Mountain Iron, Minnesota A. Radakovich, A. Severson, E. Nowariak, S. Saari, A.C. Hirsch 72 2026 Thunder Bay, Ontario P. Hinz and M. Puumala - iv - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Institute on Lake Superior Geology Goldich Medal -v- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Goldich Medalists 1979 Samuel S. Goldich 1996 David L. Southwick 2012 James D. Miller 1980 not awarded 1997 Ronald P. Sage 2013 Tom Waggoner 1981 Carl E. Dutton, Jr 1998 Zell Peterman 2014 Laurel Woodruff 1982 Ralph W. Marsden 1999 Tsu-Ming Han 2015 Rodney J. Ikola 1983 Burton Boyum 2000 John C. Green 2016 Mark A. Jirsa 1984 Richard W. Ojakangas 2001 John S. Klasner 2017 Philip Fralick 1985 Paul K. Sims 2002 Ernest K. Lehmann 2018 Val W. Chandler 1986 G.B. Morey 2003 Klaus J. Schulz 2019 Mark Severson 1987 Henry H. Halls 2004 Paul Weiblen 2020 not awarded 1988 Walter S. White 2005 Mark Smyk 2021 Allan MacTavish 1989 Jorma Kalliokoski 2006 Michael G. Mudrey 2022 Terrence J. Boerboom 1990 Kenneth C. Card 2007 Joseph Mancuso 2023 Peter Hollings 1991 William Hinze 2008 Theodore J. Bornhorst 2024 Suzanne W. Nicholson 1992 William F. Cannon 2009 L. Gordon Medaris, Jr 2025 Robert Michael Easton 1993 Donald W. Davis 2010 William D. Addison & 2026 William (Bill) Rose 1994 Cedric Iverson 1995 Gene La Berge Gregory R. Brumpton 2011 Dean M. Rossell - vi - �Proceedings of the 72nd 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 - vii - �Proceedings of the 72nd 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. 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. - viii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 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. Goldich Medal Committee Serving through the meeting year shown in parentheses Marcia Bjornerud, Academic member - Chair (2023-2026) Robert Cundari, Government member (2024-2027) Phil Larson, Industry member (2025-2028) 2026 Goldich Medal Recipient William (Bill) I. Rose Michigan Tech University, Houghton, Michigan - ix - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Citation for Goldich Medal Recipient William I. Rose It is my heartfelt honor to present the late William I. Rose (Bill) with the 2026 Institute on Lake Superior Geology’s Goldich Medal. Bill has made tremendous contributions to the field of geoheritage and to increasing public understanding of the value and global importance of Lake Superior geology. This highly significant phase of his career, despite being retired for most of it, came from a genuine desire to encourage people to “get outside and love it”, to increase their Earth science literacy, and to deepen their love of Lake Superior. This work is strongly aligned with the criteria and spirit of this distinguished award. Bill served for 41 years as a professor of geology and volcanology at Michigan Technological University, working alongside scientists from around the world. He mentored countless graduate students, many of whom became close friends and respected colleagues. He took immense pride in their accomplishments and in his role advancing global volcano research. The volcanology program he helped build at Michigan Tech has become one of the world’s leading departments, drawing students from around the globe and producing leaders in the field. In his transition to retirement, Bill’s research focus shifted to geoscience education and outreach. This new direction was rooted in his dedication to K-12 educators through projects like the Michigan Teachers Excellence Program (MiTEP) and other teacher professional development in the Keweenaw that focused on Lake Superior geology. Bill held teachers in very high esteem, recognizing them as multipliers and the heart of essential knowledge growth. Working with educators helped launch Bill’s commitment to the field of geoheritage, inspiring the multitude of initiatives and learning resources that he developed for both formal and informal learners within the Keweenaw and Lake Superior regions. The thoughtful design of these programs yielded a vast inventory of Keweenaw geosites that could be used to explore the ways Lake Superior geology guides and influences our lives and culture. Bill shared countless “geostories” with the Keweenaw community - an expression he coined, along with “geopoetry”. His enthusiasm and energy never waned, and he never told a story the same way twice. His stories have made the global significance of Lake Superior geology accessible to people and have helped them to see how geology has shaped their own identity, history, and culture - the very essence of geoheritage. These stories resonated with people, inspiring a sense of pride rooted in the geology of our place and understanding just how fascinating Lake Superior geology is. His stories have inspired others to share their own geostories in the Keweenaw community, such as the Keweenaw National Historical Park and the Carnegie Museum. Bill shared every geoheritage outreach resource he created for zero profit in order to help the Keweenaw thrive and to promote greater understanding of Lake Superior geology. Signage, books, geotours, boulder gardens, museum exhibits, concerts in the belly of an abandoned copper mine, geologic contributions to federal grant applications to support local conservation efforts, and the labyrinth Keweenaw Geoheritage website - all of these were given freely to support our community shift from an extractive economic past and to be forward-thinking and supportive of conservation, education, recreation tourism opportunities - all rooted in our rich geology. This generosity is punctuated by the family gift of Silver Island to the Keweenaw Land Trust - an example of Bill’s strong advocacy for the protection of Lake Superior geosites and the promise of continued public access and education. -x- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 At the very heart of Bill’s education and outreach work is community. Through this work he fostered relationships at the local, national, and global level. At the local level his efforts have united teachers, artists, scientists, outdoor recreation enthusiasts, conservation organizations, and tourists, all drawn together by a common curiosity of Lake Superior geology. Nationally, Bill played a vital and formative role in shaping the vision for geoheritage in the United States, contributing to numerous workshops hosted by the U.S. Committee for Geoheritage and Geoparks and the Geological Society of America. At the global level, the Keweenaw has achieved recognition as a leader in the US geoheritage movement through Bill’s pursuit of prestigious global designations. He spearheaded the designation of the Jacobsville Sandstone as one of the first Global Heritage Stone Resources in the world and the first in the United States, recognized by the International Union of Geological Sciences (IUGS) and the UNESCO’s International Geoscience Program. Bill also promoted the Keweenaw as a strong candidate to become the first UNESCO Global Geopark in the United States. Within both global and national communities, the Keweenaw is largely viewed as a Geopark. Bill’s active membership with the ILSG served as a bridge between the professional geoscience community and the broader public. He was a first or co-author on numerous abstracts and field guides presented at ILSG meetings, including the Geological Field Trip, Eastern Isle Royale, Michigan (2013) and the Self-guided geological field trip to the Keweenaw Peninsula, Michigan (1994). Bill was visionary and big thinking; this is clearly reflected in his research and many contributions to the training and education of both geoscientists and the broader public. Bill’s leadership in geoheritage and passion for education and outreach has deepened public understanding, appreciation, and desire to protect Lake Superior geology. I am brimming with gratitude to see Bill’s service honored with the prestigious Goldich Medal award. Submitted by Erika Vye Great Lakes Research Center, MTU 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 the inception of the Institute in 1955. Biographical sketches of those pioneers will be presented at future annual meetings so that all may 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 summarize the contribution of the nominee. 2) The Organizing Committee will select one or two individuals to be highlighted at the next Annual meeting and submit those names to the Board for approval. - xi - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 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) 2024 Roland Duer Irving (1847-1888) 2025 Robert Bell (1841-1917) In Memoria William Ingersoll Rose (1944-2025) William Ingersoll Rose, aged 81, died at his home in Eagle Harbor, Michigan, on July 17, 2025. Born in Detroit, Bill moved with his family at age five to New Mexico, where his love of rocks and the Earth began. Bill spent his childhood exploring the desert, riding horses, swimming in the neighborhood pool, and working at a local television station. New Mexico planted the seeds of a lifelong fascination with geology. After high school, Bill attended Dartmouth College, where he received Bachelor’s and Ph.D. degrees. Professor Dick Stoiber, one of the pioneers of volcano research, recognized potential in the unpolished young man and offered him an opportunity to study volcanoes in Guatemala—a pivotal experience that would shape Bill’s life. Bill and his wife, Nanno, settled in Houghton in 1970 where Bill joined the faculty of Michigan Tech. Bill’s work as a volcanologist took him across the globe and occasionally, Nanno and his two sons were able to come along. Following that first trip to Guatemala, Bill developed a deep passion for understanding volcanic eruptions. His adventures throughout Central America, along with his love of its people and landscapes, led him to speak Spanish and immerse himself in local cultures. He devoted himself to forecasting volcanic eruptions to help protect people living near volcanoes. Bill served for 41 years as a Professor of geology and volcanology at Michigan Tech, working alongside scientists from around the world. He mentored countless graduate students, many of whom became close friends and respected colleagues. He took immense pride in their accomplishments and in his role advancing global volcano research. The volcanology program he helped build at MTU has become one of the world’s leading departments, drawing students from around the globe and producing leaders in the field. He was instrumental in establishing signature programs such as the International Masters in Volcanology and Geotechniques and the Peace Corps Master’s International program in Mitigation of Geologic Natural Hazards. In retirement, Bill remained active and engaged. He developed geoheritage materials, led tours of Isle Royale and the Keweenaw Peninsula, and supported teachers, artists, kayakers, hikers, bicyclists, and tourists in learning - xii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 about the region’s rich geological and cultural history. The Keweenaw has achieved recognition as a leader in the US geoheritage movement through Bill’s pursuit of prestigious global designations. He spearheaded the designation of the Jacobsville Sandstone as one of the first Global Heritage Stone Resources in the world and the first in the United States, recognized by the International Union of Geological Sciences and the UNESCO’s International Geoscience Programme. Bill also promoted the Keweenaw to become the first UNESCO Global Geopark in the United States. Due to his efforts, within both global and national communities, the Keweenaw is viewed as a Geopark by definition. Many of these geoheritage initiatives were presented at ILSG. His Field Trip Guidebook for Isle Royale: Keweenawan Rift Geology, co-authored with Justin Olson, is one of ILSG’s Special Publications. For his remarkable work in the Lake Superior region, his teaching, supervisory and outreach efforts, and support of ILSG, Bill was posthumously awarded the Samuel S. Goldich Medal in 2026. Bill treasured time with his children and grandchildren, especially during family vacations in Eagle Harbor. Always curious, he took the road less-traveled and delighted in whatever he discovered along the way. Richard Wayne (Dick) Ojakangas (1932 - 2025) Dr. Richard (Dick) Wayne Ojakangas died peacefully in his sleep on December 16, 2025 at the age of 93. Dick was born November 20, 1932, in Moose Lake, Minnesota, and grew up in Kettle River and Warba. He was very proud of his 100% Finnish heritage. After graduating from Grand Rapids High School, he enrolled as a business major at the University of Minnesota Duluth (UMD), intending to take over the family store in Warba after graduation. However, during his senior year, he took an introductory geology class from Dr. Robert Heller, and switched his major after the first lecture to geology. He joined the Reserve Officer’s Training Corps because he felt it was his patriotic duty to serve his country. After graduation, he was assigned to the USAF base in Upper Heyford, England. He married Finnish beauty Beatrice (Peaches) Luoma and they moved to England within one week after their wedding. Due to his geological expertise, Dick was assigned to be a Petroleum Supply Officer, fueling jets with highly toxic JP4 jet fuel. After two years in the Air Force, he continued his studies in geology, earning a master’s degree from the University of Missouri, and a PhD from Stanford University. Returning to Duluth, “Dr. OJ” enthusiastically taught geology at UMD for 38 years, where he was beloved by many hundreds of students. Dick was renowned as an entertaining and exceptional geology professor. He began each lecture with a Finn joke, and the punchlines were meticulously written on his calendar. His colleagues in the Geology Department were also his extended family and lifelong friends. He retired in 2002. Dick wrote or collaborated on more than 60 scientific publications and several books, including the highly acclaimed Minnesota’s Geology, and Roadside Geology of Minnesota. He was awarded the prestigious Horace T. Morse Award for Distinguished Teachers from the U of M, and received an honorary PhD from the University of Helsinki, Finland. A long-time member of the ILSG, Dick was awarded the Samuel S. Goldich Medal by the Institute in 1984. He was a fixture at annual meetings, leading field trips and giving presentations, either as himself or as one of his alter egos, like the Old Prospector or Herr Dr. Direktor Professor Wolfgang von Schlummerklutz from the World Panzerenkotklotzen Institute in Europe! Dick was a passionate traveler and photographer, doing geological research on all seven continents. His work in Antarctica as part of the United States Antarctic Research Program resulted in having Mount Ojakangas being named after him. In India, he found evidence of the first Archaean glaciation ever discovered. In Finland, he and a colleague were the first to determine the direction that glaciers moved through northern Europe. As a worldrespected geologist and an engaging, highly understandable speaker, he spread his enthusiasm for science by giving lectures on cruise ships from 1978 to 2017, feeding his obsession with traveling the world. Curious and inquisitive, Dick entertained many interests and hobbies. He lived an active life - running several - xiii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Grandma’s Marathon’s (all without training) and cross-country skiing (doing the Birkebeiner 54 km race in Wisconsin many times, also without training). Dick was an avid mushroom hunter on all continents. His wife said that he could ‘find mushrooms, whether they were there or not!’ Known for his generosity, humor, and warm personality, Dick made sure everyone felt included and always sought out strangers, who promptly became his friends. His trademark greeting “Hiya!” and farewell “Cheers!” are remembered fondly by his family and friends. Paul Willard Weiblen (1927 – 2025) Professor Paul Willard Weiblen, 98 years old, died peacefully in the presence of family on Tuesday, December 23, 2025 in St. Paul, Minnesota. PW, to friends, colleagues and students, was born in Miller, South Dakota in 1927. After graduation from high school in 1945, he entered the U.S. Army. After military service he returned to college and earned a B.A. degree at Wartburg College in Waverly, Iowa (1950), and an M.A. in History at the University of Minnesota (1952). PW came into geology in a roundabout way. Apparently, he was in Istanbul, Turkey working as a travel agent for American Express when he encountered a geologist exploring the world for uranium deposits. Consequently, he returned to the University of Minnesota in 1959 to pursue geology. He focused on the metamorphism of the Paleoproterozoic Thomson Formation of east-central Minnesota for his Master’s thesis (1962) and on the geology and petrology of the Bald Eagle intrusion of the Duluth Complex in northeastern Minnesota for his Ph.D. (1965). He stayed at the University in the Geology and Geophysics Department as an Assistant Professor (1965), Associate Professor (1969), Professor (1980), and Professor Emeritus (1997), teaching Minnesota geology and characterizing the minerals of the Duluth Complex with the Minnesota Geological Survey. He was hired specifically to organize and supervise the Department’s new Electron Microprobe Laboratory (1965-1980) in the Space Science Centre. He also served as Curator of the petrology collection (1970-1997) and supervisor of the scanning electron microscope facility (1970-1997). Over his 32 years as a faculty member, Paul’s analytical expertise and unbridled curiosity led him to pursue, and engage others, in many areas of research. The principal focus of his research was on the petrology and mineral deposits of the Duluth Complex. A highlight was a 1980 American Journal of Science paper, co-authored by Minnesota Geological Survey Chief Geologist G.B. Morey, that summarized the stratigraphy, petrology and structure of the Duluth Complex. Another significant area of interest in Paul’s career was lunar petrology. In the early 1970s, he and Edwin Roedder (USGS) confirmed the phenomenon of silicate liquid immiscibility by examining lunar glasses, and terrestrial basalts. In 1978-79, Paul served as lead curator of NASA’s Washington, D.C. lunar sample collection. The focus of Paul’s research in the latter part of his academic career and into his retirement was building and promoting the electric pulse disaggregator (EPD, or “the Zapper”). He was introduced to the EPD and its inventor, Nikolay S. Rudashevsky, during a visit to Russia in 1991.   Recognizing the potential of this instrument to create ultraclean mineral separates for a variety of applications, PW built and installed an EPD at U of M in 1992. He actively promoted it to other scientists who have used it to prepare samples for detrital zircon dating, mineral liberation analyses, and microfossil studies. As a teacher and student advisor, PW was engaging, approachable, and deeply committed to his students. He routinely taught undergraduate and graduate level Igneous Petrology and Optical Mineralogy and offered hands-on classes on electron microprobe analysis. His annual petrology field trips up the Gunflint Trail were legendary. As a graduate advisor, he was open to letting his 11 PhD and 13 MS students develop their own thesis projects, with topics that included igneous petrology; volcanology; structural, metamorphic and field geology; geochemistry, mineralogy, petrography, and economic geology. A particular source of pride was that all but one graduate thesis was based on the geology of Minnesota. PW was awarded the Goldich Medal from the Institute on Lake Superior Geology in 2004 for his lifelong commitment to promoting geologic studies of Minnesota. - xiv - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Ronald Parker Sage (1938 - 2026) Ronald Parker Sage, 87, of Kingsford, MI, passed away peacefully on January 28, 2026, after a battle with numerous health conditions. Ron was born on August 4, 1938, in Pontiac, Michigan. Ron graduated in 1960 with a BSc degree in geological engineering from Michigan Technological University in Houghton. While at MTU, Ron spent most of his free time collecting rocks and minerals in the copper and iron mining districts, earning him the nickname “Rocky”. He was a student of Kiril Spiroff, the “Mad Russian”. Ron held his Alma Mater in high esteem. In the early 1960s, Ron worked for three years as an engineer for the Shell Oil Company in west Texas. His duties included well siting, well logging, well workovers, and other production-related activities. In 1966, Ron graduated with a Master’s degree in Geology from the Colorado School of Mines. It was here that he was first exposed to alkalic rocks, his study topic and thesis being “Geology and Mineralogy of the Cripple Creek Syenite Stock, Teller County, Colorado”. This led to employment with Anaconda American Brass Ltd to investigate CuNi-PGE minerals in the Port Coldwell alkalic complex near Marathon, Ontario. During the summer of 1967, Ron searched for base metals in the Ely greenstone belt in northern Minnesota for Bear Creek Mining. It was in that year that he first met two other greats of Lake Superior geology, Ned Eisenbrey and Gene LaBerge. In 1969, Ron again worked for Anaconda American Brass Ltd., this time north of Lake Superior in the Schreiber greenstone belt, searching for gold and base metals. In the fall of 1969, Ron joined the Ontario Geological Survey, his professional home for over 30 years. Ron’s work for the OGS took him to many parts of the province, but never very far, and never for very long, from Lake Superior. He spent four years on a helicopter reconnaissance in Northern Ontario, then mapped the Slate Islands in Lake Superior, and next worked on a multi-year program to study alkalic rocks north of Port Coldwell along the northern extension of the Trans Superior Tectonic Zone and along the Kapuskasing Structural Zone. In 1978, Ron was assigned to the Michipicoten greenstone belt. Here he spent 10 years mapping Archean supracrustal rocks over approximately 540 square miles, with emphasis on the gold and base metal potential. In 1993, Ron was assigned to a province-wide program of kimberlite documentation to stimulate diamond exploration. Some of this work was again in the Michipicoten area, where diamond-bearing rocks had recently been discovered. Despite his busy professional schedule, Ron was able to complete his PhD degree in 1986 for a thesis submitted to Carleton University in Ottawa, entitled “Alkalic Rock Complexes and Carbonatites of Northern Ontario, and their Economic Potential.” Ron was a long-time ILSG supporter, giving presentations, leading field trips and Co-Chairing the annual meetings in 1987, 1997 and 2006. In 1997, Ron was recognized for his many contributions when he received the Samuel S. Goldich Medal from the Institute on Lake Superior Geology. Charles Edward (Charlie) Blackburn (1940 - 2026) Charlie passed away on Friday March 6, 2026 at the Royal Jubilee Hospital, Victoria, BC. Although he identified himself as a Welshman, having grown up in a small village near Cardiff, Wales, by an accident of fate during the early days of World War II, Charlie was actually born in Kidderminster, England. He went to Swansea University to complete his Bachelor of Science degree in geology. Charlie loved the summer field mapping excursions in Northern - xv - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Norway working towards his goal of becoming a professional geologist. Geology was always his passion and drawing maps, his gift. Charlie emigrated from Wales to Canada and undertook a Master of Science degree at the University of Western Ontario, London, Ontario where he met his wife of sixty years Christine (nee Spence) – also a recent UK immigrant. It was love at first sight. The couple were married after a two-month courtship. They left for Italy where Charlie studied the metamorphic puzzle of the Seisia-Lanzo zone in the Valle d’Aosta at the University of Padua, Italy. He liked that his office was right opposite the Cappella Della Scrovegni – famous for its Giotto frescoes. His time in Italy left him with an enduring love for the people and culture there. Charlie returned to Canada in 1969 and in 1970 accepted a position as a mapping geologist with the Ontario Geological Survey (OGS) in Toronto. As a field geologist, Charlie spent summers of his early career in the bush of northern Ontario and didn’t see too much of his family. As the children grew, Charlie saw the need to be with his family more and so accepted the position of Resident Geologist in Kenora, Ontario. Many of his over 75 OGS publications resulted from his mapping efforts in the Archean greenstone belts of the western Wabigoon Subprovince and other areas in the Kenora District. He was the lead author of the seminal review of the Wabigoon in the 1991 compendium, Geology of Ontario. In the early 90’s, he took a sabbatical from his Resident Geologist duties to return to mapping the Separation Lake area. Charlie retired from the OGS at age 60 after 30 years of service and, with his wife Christine, became co-founder of their consulting company, Blackburn Geological Services. Charlie Chaired the 1985 ILSG annual meeting in Kenora and was on the organizing committee for the 2002 annual meeting, also held in Kenora. He delivered papers and chaired sessions at many ILSG meetings and led field trips to the Separation Rapids rare-element pegmatite field and other locations in the Kenora District, of which he had an encyclopedic knowledge due to his years of mapping and documenting mineral occurrences in the Superior Province. Report of the Chair of the 71st Annual Meeting Amy Radakovich, Allison Severson, Eric Nowariak, Aaron Hirsch, Stacy Saari Mountain Iron, Minnesota The 71st Institute on Lake Superior Geology (ILSG) was held May 14 to 17, 2025 in Mountain Iron, Minnesota at the Mountain Iron Community Center. The meeting was sponsored by the State of Minnesota’s Iron Range Resources and Rehabilitation agency, Bayside Geoscience, the Geological Society of Minnesota, the Mesabi Range Geological Society, George Hudak Geosciences, PLLC, and the University of Minnesota Duluth’s (UMD) Swenson College of Science and Engineering Earth and Environmental Sciences department, as well as individual contributors Roger Anderson, Allan MacTavish, Dave Dahl, Tom Erickson, and Barry Frey. The meeting was cochaired by Amy Radakovich, Allison Severson, Eric Nowariak, and Aaron Hirsch of the Minnesota Geological Survey (MGS), and Stacy Saari of the Minnesota Department of Natural Resources (MNDNR). Patrice Cobin and Julie Stark of Michigan Technological University served as registrars for the meeting. The institute was attended by a total of 137 participants of which 25 were students. Generous donations from the following individuals helped provide a reduced registration and field trip price for students: Kate Clover, Jim and Isabel DeGraff, Tom Erickson, Tom Fitz, Aaron Hirsch, Paula Leier-Engelhardt, Bob Mahin, Vince and Susan Matthews, Jim Miller, Allison Severson, Mark and Lauri Severson, John Verhoeven, and Gerry White. The 71st meeting consisted of two full days of technical sessions, which ran from Thursday morning, May 15 through Friday afternoon, May 16th. The meeting also held pre-and post-meetingfield trips on May 14th and May 17th. A total of 51 presentations were subdivided into 8 technical sessions; 6 technical sessions for 26 oral presentations (of which 1 was presented by a student), and 2 poster technical sessions with a total of 23 poster presentations (of which 14 were presented by students). The chairs continued the previous meeting’s precedent of including two poster sessions to allow both attendees and judges more time to review the posters. The first th - xvi - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 presentation of the technical sessions was given by Mark Smyk (OGS - retired; Goldich Medalist in 2005) who gave the citation for Robert Bell, the 2025 Pioneer of Lake Superior Geology. Bell is the 6th person to be recognized for their contributions to Lake Superior Geology prior to the initiation of the ILSG. The technical sessions of the 71st annual meeting of ILSG were published in 2025 as Part 1 of Proceedings Volume 71 (95 pages). Five Doug Duskin Best Student Paper Awards were given for student oral and poster presentations as judged by the 2024 Student Paper Awards Committee chaired by Aaron Hirsch (MGS). PhD student poster awards were given to Zsusanna Allerton and Madelyn Banks. Undergraduate student poster awards were given to Celia Cortopassi and Lyndsie Vickers. Omar Khali Droubi received the best oral presentation award. The 71st ILSG also awarded 12 Eisenbrey Student Travel and Participation Awards to help defray the cost of travel to and participation in the ILSG professional meeting for undergraduate and graduate students. The awardees were Drew Casper, Haley Johannesen, Mary Elizabeth Shalifoe, Linsey Hula, Omar Khalil Droubi, Samara Gries, Renee Jeutter, Aidan Kwiatkowski, Celia Cortopassi, Lyndsie Vickers, Zsuzsanna Allerton, and Bekah Thomson. As usual, field trips were a highlight of the 71st ILSG. Mountain Iron’s close proximity to exposures of Archean, Paleoproterozoic, and Mesoproterozoic rocks made it a prime location to run numerous excellent field trips. The meeting offered 8 field trips which included 4 pre-meeting trips on Wednesday May 14, and 4 post-meeting trips on Saturday May 17. Seven field trips focused on the varied Precambrian geology of northeastern Minnesota, and one trip highlighted the unique Quaternary features of the region. Seven of the eight field trips were able to run, with one cancelled due to active wildfires in the field trip area. The remaining 7 field trips were well attended. There were 130 registrants for the field trips, excluding leaders, representing over 100 different individuals (some registrants took multiple trips). Pre-meeting trip 1 was a “Transect through the Quetico subprovince of northern Minnesota,” led by Eric Nowariak (MGS) and Mark Jirsa (MGS-retired). Pre-meeting trip 2 was led by Mark Severson (Natural Resources Research Institute, Teck - retired), Cullen Phillips (New Range Copper Nickel), and Kevin Boerst (Twin Metals Minnesota) and highlighted “Drill Core from three Cu-Ni deposits of the Duluth Complex.” Pre-meeting trip 3 asked the question “How do you make iron and/or manganese in Proterozoic iron formation?” and was led by Alex Steiner and Dean Peterson (Big Rock Exploration) and Latisha Brengman (University of Minnesota Duluth [UMD]). Pre-meeting trip 4 was led by George J. Hudak (University of Minnesota; George Hudak Geosciences, P.L.L.C) and Zsuzsanna Allerton and Annia Fayon (University of Minnesota) and highlighted “New geological insights into the genesis of iron ores at Lake Vermillion-Soudan Underground Mine State Park.” Post-meeting trip 5 traveled to numerous “Neoarchean alkalic intrusions in the Wawa and Quetico subprovinces” and was led by Terry Boerboom (MGS-retired) and Amy Radakovich (MGS). Mark (NRRI, Teck - retired), Allison (MGS), and Lauri (earth science teacher - retired) Severson planned to lead post-meeting trip 6 focused on a “Unique Keweenawan inclusion (Colvin Creek) in the Duluth Complex.” However, the trip was cancelled due to wildfire conditions, and participants were invited to join other trips or receive a refund. Post-meeting trip 7 led by Dean Peterson (Big Rock Exploration) and George Hudak (University of Minnesota; George Hudak Geosciences, P.L.L.C) visited numerous “Classic outcrops of northeastern Minnesota” Field trip 8 was led by Phil Larson (Vesterheim Geoscience, PLC), Andrew Breckinridge (University of Wisconsin - Superior), and Howard Mooers (UMD) and focused on Glacial Lake Norwood and the Koochiching Lobe.” Field trip guides were published in 2025 as Part 2 of the Proceedings Volume 71 (200 pages). A catered welcome reception was held at the Mountain Iron Community Center on Wednesday evening, May 14, after all of the pre-trips returned. The event was well attended, and offered a chance for meeting attendees to reconnect with colleagues and friends prior to the start of technical sessions. Steve Solkela provided entertainment for a portion of the evening. The annual ILSG social and banquet were hosted at the Mountain Iron Community Center on Thursday evening, May 15, 2025. Ninety-three people were in attendance at the sold-out banquet. After introductions - xvii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 and announcements, Mark Puumala announced the location of the 2026 meeting as Thunder Bay, Ontario. The program continued with ILSG awarding the prestigious Goldich Medal to the very deserving Robert Michael (Mike) Easton (Ontario Geological Survey), who unfortunately could not be present at the meeting. Wouter Bleeker (Geological Survey of Canada) provided the citation for Mike, highlighting Mike’s long tenure with the OGS, his impressive publication record, and his contributions to ILSG. Another highlight of the banquet was the keynote presentation by Pete Kero, P.E., Senior Environmental Engineer with Barr Engineering Co and visionary behind the award-winning Redhead Mountain Bike Park in Chisholm, Minnesota. His fascinating talk entitled “Mine to Mountain Bike Mecca: The story of the Redhead Mountain Bike Park” detailed the transformation of ten idled open pit iron mines in northeast Minnesota into a world-class recreation destination for mountain biking, hiking, and paddling. Kero fielded many questions from the engaged audience and sold and autographed his book Minescapes: Reclaiming Minnesota’s Mined Lands after the keynote presentation, which ended the banquet program. The Institute’s Board of Directors met on Thursday May 15, 2025 to discuss ILSG business and approve the 2026 meeting location. The meeting was attended by Amy Radakovich (Board Chair and Assistant Treasurer), Ted Bornhorst, Carsyn Ames, Peter Hollings (Secretary), and Mark Jirsa (Treasurer). Guests at the meeting were the meeting co-chairs Allison Severson, Eric Nowariak, Aaron Hirsch, and Stacy Saari and also Mark Puumala, the Chair of the proposed 2026 Thunder Bay meeting (approved by the board - see below). Michael Easton was unable to attend. Institute’s Board of Directors meeting notes were taken by ILSG Secretary Hollings, which are as follows: 1. Accepted report of the Chairs for the 70th ILSG, as published in the Proceedings volume, and minutes of last Board meeting, May, 2024 (Hollings). 2. Received and discussed 2024-2025 ILSG Financial Summary (Jirsa/Radakovich). Final approval tabled for Email vote after necessary revisions are made to balances as listed 3. Received, discussed, and accepted 2024-2025 report of the Secretary (Hollings). 4. Approved Alli Severson as on-going ILSG Board member and Pete Hinz and Mark Puumala as coChairs. 5. Discussed and approved appointing Amy Radakovich as the Institute Treasurer. This was subsequently approved by the Membership. Mark Jirsa was thanked for his 31 year service to the Institute. 6. Discussed and approved replacing Dean Peterson as the “member from industry” on Goldich Committee (end of term 2025) with Phil Larson. 7. Approved Thunder Bay as the site for the 72nd annual ILSG meeting. The meeting will be Chaired by Pete Hinz and Mark Puumala with tentative dates of May 19 to 23. 8. A number of future meeting locations were discussed including Grand Marais (Jim Miller), Baraboo (Esther Stewart & Carsyn Ames) and Marquette. 9. The revised Eisenbrey guidelines were discussed and approved with edits. Changes expand the list of expenses which are eligible for reimbursement from the Eisenbrey award (ex: registration fees, meals, lodging, and transportation are all now included) 10. There was discussion over the format and page limits for the abstracts. It was agreed that the two page limit would be maintained. 11. The cost of hosting the meeting registration through MTU was discussed. MTU currently charges 12% of the total registration sales as their fee. It was agreed that the hosts of each meeting would evaluate possible hosting options and pick the one that worked best for them. Puumala indicated that next year the hosts would likely go with a Canadian registrar so that registration fees could be charged in Canadian dollars 12. The cost of printing the Proceedings and Field Guide volumes was discussed. It was agreed that future meeting Chairs would explore the possibility of making the full printed volumes a paid option for participants - xviii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 and providing only the guides for individual trips. 13. The ongoing storage of ILSG poster boards and easels was discussed. MTU has stored them for the last ~10 years but can no longer offer to do that. Boards and easels were stored for the past year at the Minnesota Geological Survey, but there is no room for permanent storage there. It was suggested that the storage and transport of the posters and easels become the responsibility of the meeting hosts, such that after each ILSG meeting, the boards and easels would leave with the host of the following year’s meeting. This way storage and transport costs can be built into the next year’s meeting costs. Thunder Bay hosts do not need boards next year and did not want to take them across the border given recent border crossing issues. It was suggested that ILSG perhaps have two sets of boards and two sets of easels - one that resides in Canada and one that resides in the USA. Carsyn Ames volunteered to store the boards and easels at the Wisconsin Geological Survey for the next year, delaying the need to make a final decision. Our large, five-person committee allowed us to divide-and-conquer the innumerable tasks to make The 71st annual ILSG meeting a great success. We were proud to continue the long-standing tradition of bringing people together from many states and provinces to share and learn about the fascinating geology of the Lake Superior region, both in the meeting and ‘on the rocks.’ The co-chairs would like to thank the many people and organizations who made the meeting possible, including the Mesabi Range Geological Society and UMD students who ran the registration table and helped with merchandise sales, and the numerous individuals who offered to drive rental or personal vehicles on our fieldtrips. The Sawmill supplied all meeting and field trip food, Caribou provided coffee and tea for the field trips, and Peplinjack’s Bakery supplied the delicious field trip pastries. Lastly, we would like to thank the numerous generous donors who donated hundreds of rock and mineral specimens, books, and maps that made up the biggest and most profitable book sale and silent auction in ILSG memory. The sale and auction netted a total of approximately $4,500 which will be used to fund student participation at subsequent meetings. We look forward to seeing everyone next year in Thunder Bay! Respectfully submitted, Amy Radakovich, Allison Severson, Eric Nowariak, Aaron Hirsch, and Stacy Saari Co-chairs, 71st Institute on Lake Superior Geology 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 end of the annual meeting. 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. - xix - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 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. 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 2025, the ILSG Board of Directors selected two students to be granted research funding of $500 each from the Joe Mancuso Student Research Fund. The awardees were: Kathryn Akin, University of Minnesota- Twin Cities Alyssa Hellrung, University of Wisconsin 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. - xx - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 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. 2026 Student Paper Awards Committee Emily Smyk - Bayside Geoscience Justin Jonsson - Ontario Geological Survey Nick Swanson-Hysell - University of Minnesota Board of Directors Board appointment continues through the close of the meeting year shown in parentheses, or until a successor is selected Peter Hinz and Mark Puumala, Co-Chairs (2026-2029) - Ontario Geological Survey, Retired Alli Severson (2025-2028) - Minnesota Geological Survey Ted Bornhorst (2024-2027) - Michigan Tech, Houghton Carsyn Ames (2023-2026) - Wisconsin Geological & Natural History Survey, Madison Amy Radakovich, Treasurer (2025-2028) - Minnesota Geological Survey Peter Hollings, Secretary (2024-2027) - Lakehead University Local Committee Chairs Peter Hinz and Mark Puumala - Ontario Geological Survey, Retired Organising Committee Robert Cundari - Ontario Geological Survey, Thunder Bay, Ontario Al MacTavish - Thunder Bay, Ontario Mark Smyk - Lakehead University, Thunder Bay, Ontario Pete Hollings - Lakehead University, Thunder Bay, Ontario Jim Miller - Thunder Bay, Ontario - xxi - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Field Trip Leaders and Guidebook Authors Field trips have been the mainstay of the ILSG since its inception 72 years ago. We give special thanks to the field trip leaders and guidebook authors who volunteered their time and talent in carrying that tradition forward. Trips 1 & 4: Classic” Geological Sites in the Thunder Bay Area - Mark Smyk (Lakehead University) and Mark Puumala (Geological Consultant) Trip 2: Geology of the Quetico Subprovince and Shebandowan greenstone belt north of Thunder Bay - Riku Metsaranta and Gaetan Launay (Ontario Geological Survey) Trip 3: Geological assemblages, regional structural framework and tectonic evolution of the Neoarchean Shebandowan greenstone belt - Dorothy Campbell, Justin Jonsson and Vittoria D’Angelo (OGS Resident Geologist Program) Trip 5: Archean Geology and Metallogeny of the Rainy Lake Wrench Zone - K. Howard Poulsen (Geological Consultant) Trip 6: Amethyst Deposits of Thunder Bay - Steve Kissin (Lakehead University) and Greg Paju (OGS Resident Geologist Program) - xxii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Technical Program Wednesday May 20 (Parking Lot G14, Lakehead University) 8:00 a.m. Field Trip 1: “Classic” Geological Sites in the Thunder Bay Area Leaders: Mark Smyk and Mark Puumala 8:00 a.m. Field Trip 2: Geology of the Quetico Subprovince and Shebandowan greenstone belt north of Thunder Bay Leaders: Riku Metsaranta and Gaetan Launay 8:00 a.m. Field Trip 3: Geological assemblages, regional structural framework and tectonic evolution of the Neoarchean Shebandowan greenstone belt Leaders: Justin Jonsson and Vittoria D’Angelo 5:00 p.m. Return of Trips 1-3 4:00 p.m. - 8.00 p.m. Registration (Faculty Lounge, Lakehead University) 6:00 p.m. - 9.00 p.m. Ice Breaker Social, Poster Setup and Core Shack (Faculty Lounge, Lakehead University) Thursday May 21 7:30 a.m. - 4:00 p.m. Registration (Faculty Lounge, Lakehead University) 8:30a.m. - 9:00 a.m. Introductory Remarks (Room UC0050, Lakehead University) Technical Session I NOTE: Asterisk * denotes a student eligible for a Best Student Paper Award Session Chairs: Mark Puumala and Jim Miller 9:00 a.m. Stephan, T., Phillips, N., and Hollings, P. Timing and conditions of magmatism, metamorphism, and strain partitioning in the western Shebandowan Greenstone Belt (Superior Province) 9:20 a.m. MacDonald, P., Hastie, E., Malegus, P., Kamo, S., Hamilton, M. and Marsh, J. Implications of recent geochronology on the regional geology and timing of gold mineralization in the Red Lake greenstone belt, Ontario 9:40 a.m. Hollings, P., Vrzovski, J., Cooke, D. and Gorner, E. Using epidote and chlorite mineral chemistry to extend the alteration footprint around the Hemlo Au deposit, N. Ontario 10:00 a.m. - 10:30 a.m. Coffee Break, Poster Session and Core Shack 10:30 a.m. Tiitto*, H., Phillips, N., and Stephan, T. Deformation processes in a mid-crustal strike-slip shear zone: Insights from the Archean Quetico Shear Zone, Superior Province, Canada 10:50 a.m. Sheshnev*, V., Hollings, P., Tolley, J., Angombe, M., Deller, M. and Stern, R. Whole Rock and Mineral Chemistry of the Eagle’s Nest Intrusion, McFaulds Lake Greenstone Belt, Ontario, Canada: Insights into the Origin and Paragenesis - xxiii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 11:10 a.m. Carlton*, K., Tikoff, B. and Nachlas, W. An introduction to the northwestern Huron Mountains of the Upper Peninsula, Michigan: field relations and preliminary structural interpretations 11:30 p.m. - 1:00 p.m. Lunch Break, Poster Session and Core Shack (ILSG Board Meeting by invitation) Technical Session II Session Chairs: Esther Stewart and Phil Larson 1:00 p.m. Salerno, R., Cannon, W. F., Thompson, J., Souders, A., Vervoort J. and Hillenbrand, I. Reassessing variations in metamorphism across the Penokean orogen in Northern Michigan: Part 1, new Pressure-Temperature-Time-Deformation constraints 1:20 p.m. Cannon, W. F., Salerno, R., Drenth, B. and Bedrosian, P. Reassessing variations in metamorphism across the Penokean orogen in Northern Michigan: Part 2, Reinterpreting metamorphic nodes 1:40p.m. Hirsch, A. Can we improve the bouguer gravity resolution in the Cuyuna Range? Increasing gravity measurements in a region of high gravity station density. 2:00 p.m. - 2:30 p.m. Coffee Break, Poster Session and Core Shack 2:30 p.m. Allerton, P. and Hudak, G. Characterization of hematite ore from former Ely mines, NE Minnesota 2:50 p.m. Steiner, R.A., Watson, N., Riley, J., Hammer, M., Thole, J., Feinberg, J., Sandri, H. and Savage, B. Oxidation to Ores: Petrological Insights into Supergene Manganese Enrichment at the Emily Deposit, Minnesota 3:10 p.m. Hagedorn, G. Ice flow history, surficial geology, and till composition of Georgia Lake area, northwestern Ontario Poster Session 3:30 - 5:00 p.m. 6:00 p.m Annual Banquet and Award Presentation (Faculty Lounge, Lakehead University) Announcement of 73rd Annual Meeting Location 2026 Goldich Award Presentation to Bill Rose 2026 Quiz night Meeting participants not registered for the banquet are welcome to attend the quiz night - xxiv - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Friday May 22 9:00 a.m. - 12:00 p.m. Registration Technical Session III Session Chairs: Shannon Zurevinski and Therese Pettigrew 8:30 a.m. Beyer, S., Cutts, J., Hnatyshin, D., Powell, J., Camacho, A., Cawood, T. and Drever, G. Preliminary geochronology of lithium pegmatites and host rocks, Archean Superior Province, northwestern Ontario 8:50 a.m. Quigley, A., Mahin, R., and Gamet, N. Critical Mineral Potential of the Watersmeet Gneiss Dome, MI USA 9:10 a.m. Bleeker, W. and Wodicka, N. Improved Precision and Better Accuracy: SHRIMP-II Detrital Zircon Analysis of Samples Across the Stratigraphy of the Midcontinent Rift 9:30 a.m. Easton, R.M. and Kamo, S. The Badgerow complex, a Midcontinent Rift-related REE-Zr-rich peralkaline intrusion in the Grenville Province near Verner, Ontario 9:50 a.m. - 10:20 a.m. Coffee Break, Poster Session and Core Shack 10:20 a.m. Nitescu, B., Torres, D., and Gaona, J.. Models of the regional gravity and magnetic anomalies associated with the Nipigon Embayment 10:40 a.m. Bain, W. and Hollings, P. Coeval silicate melt and PGE-bearing salt melt inclusions in the Thunder and Seagull intrusions, Ontario: An overview of evidence and data processing challenges 11:00 a.m. Drost, A. and Heggie, G. A new look at the Seagull mafic-ultramafic Intrusion and potential hydrogen and helium accumulations 11:20 a.m. Swanson-Hysell, N., Zhang, Y., Mohr, M. and Schmitz, M. Linking the Southwestern Laurentia large igneous province and rapid Duluth Complex emplacement through mantle plume dynamics 11:40 p.m. - 1:00 p.m. Lunch Break, Poster Session and Core Shack Technical Session IV Session Chairs: Wouter Bleeker and Peter Hinz 1:00 p.m. Smith, J., Kaski, K., Tschirhart, V. and Enkin, R. Integrating petrophysical data with full tensor magnetic gradiometry for improved interpretation and modelling of remanently magnetized intrusions in the Midcontinent Rift 1:20 p.m. Peterson, D., Steiner, A., Sweet, G. and Boucher, C. Physical Magmatic System Interpretation of the Marathon Cu-Pd Deposit, Coldwell Complex, Ontario 1:40 p.m. Smyk, E., Dolega, S., Churchley, J. and Flank, S. Optimizing data collection for better geological interpretations and adding value to your project - xxv - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 2:00 p.m. Lizzadro-McPherson, D., Vye, E., Degraff, J., and Rose, W. Interactive Geospatial Geoheritage: Efforts to Support Place-based Exploration and Digitally Preserve Keweenaw’s Geoheritage 2:20 p.m. - 2:50 p.m. Coffee Break, Poster Session and Core Shack 2:50 p.m. Degraff, J., Hiltunen, L., Lafreniere, D., Lizzadro-McPherson, D., Vye, E., Cowling, B., Bornhorst, T. and Rose, W. Digital Preservation and Enhanced Utility of Exploration Core Descriptions from the Keweenaw Copper District, Michigan: Progress toward a Map-based Web Portal 3:10 p.m. Stone, A., Lizzadro-McPherson, D. amd Vye E. Rocks and Roots: The Role of Geoheritage in Biodiversity Stewardship 3:30 p.m. Smyk, M., Hodge, J. and Robillard, C. Pukaskwa Redux: Revisiting and Reconnecting with Superior’s Wild North Shore 3:50 p.m Presentation of Best Student Paper Award and Eisenbrey Awards 5:00 p.m. Field Trip 5: Archean Geology and Metallogeny of the Rainy Lake Wrench Zone Leader: Howard Poulsen Parking Lot G14, Lakehead University Poster Presentations Akin*, K. and Swanson-Hysell, N. Constraining the 3-D Geometry of the Duluth Complex, MN, Using Magnetic Fabrics and Paleomagnetic Data Angombe, M., Phillips, N., Hollings, P., Stephan T., Sheshnev, V., Deller, M. and Smith, A. Decoding Shear Zone Evolution in the McFaulds Lake Greenstone Belt, Ontario: Constraints on CrystalPlastic Deformation and Timing from in-situ Titanite U–Pb Thermochronology Bilboe*, M., Zurevinski, S. and Conly, A. Quartz Trace Element and TEM Analysis of Selected Economic LCT Pegmatites Buchholz, T., Falster, A. and Simmons, W. Update to: a complex F-rich alkalic pegmatite in the pyroxene syenites of the Stettin Complex, Wausau Complex, Marathon County, Wisconsin Chaisson*, A., Smyk, M. and Zurevinski, S. Petrography and Geochemistry of the Mound Lake Pluton, Northwestern Ontario Duffy*, P., Brengman, L. and Eyster, A. Integrated X-Ray Diffraction and Petrography Document Carbonate Mineral Heterogeneity and Hematite Mineralization in the Upper Biwabik Iron Formation, MN Ellison*, K ., Cisneros, J., Eyster, A. and Brengman1, L. Comparing mineralogy along a surface to depth transect of the ~2.7 Ga North Limb Soudan Iron Formation, NE Minnesota - xxvi - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Erickson, S., Fayon, A., Allerton, Z. and Hudak, G. Middle school virtual field trip lessons materials for Archean formations of Lake Vermilion-Soudan Underground Mine State Park Gilberg*, N., Fralick, P. and Li, Z. Geochemical Constraints on Mn Cycling in the Paleoproterozoic Gunflint Formation Gosai*, M., Fralick, P. and Li, Z. Modified Sequential Iron Extraction Method for Analyzing Rare Earth Elements in Banded Iron Formations Grauch, V. and Heller, S. Time-to-depth conversion of seismic-reflection data from eastern Lake Superior and implications for the eastern arm of the Midcontinent Rift Harding*, M. and Hollings, P. Geochemistry, Petrogenesis, and Mineralization of the Makwa Deposit, Bird River Sill Hellrung*, A., Droubi, O., Ruggles, C. and Bonamici, C. Using Anisotropy of Magnetic Susceptibility and U-Pb Geochronology from the Bush Lake Granite, Florence County, WI to Understand Post-Penokean Continental Growth Jonsson, J. and Li, Z. Petrographic Study of Granular Iron Formation in the Gunflint Formation: Evidence for Well-Oxygenated Surface Waters Marin López*, V., Brengman, L., Eyster, A., Mitchell, J., Pu, X., Mangum, J. and Walker, P. Quantitative analysis of iron mineral composition and crystal sizes in the contact metamorphosed Biwabik iron formation and the Bald Eagle intrusion, NE, MN, USA Nowak*, R., Deering, C. and Essig, E. Origin of the World-Class Eagle, Eagle East, and Tamarack Ni-Cu-PGE Deposits and comparative analysis with other Midcontinent Rift- and Siberian Trap-related intrusions Nowariak, E. and Severson, A. Bedrock Geology of the Ericsburg NW, Ericsburg NE, Ray SW, and Ray SE Quadrangles, St. Louis and Koochiching Counties, Minnesota Paliewicz, C., Post, S. and Thakurta, J. Petrographic, geochemical, and mineralogical analyses of manganiferous iron formations and associated lithologies at the Cuyuna Range, central Minnesota Saini-Eidukat, B., Chittick, S. and Nesheim, T. Current geologic and geophysical research on the Precambrian basement of eastern North Dakota, USA Stewart, E., McNall, N., Hart, D., Ames, C., Chase, P., Stewart, E. and Graham, G. Subsurface mapping of the late Ordovician Maquoketa Group in eastern Wisconsin using airborne electromagnetic and well data Tolley, J. and Hollings, P. Variations in Olivine Major Element Composition Across the Midcontinent Rift System NOTE: Asterisk * denotes a student eligible for a Best Student Paper Award - xxvii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Saturday May 23 (Parking Lot G14, Lakehead University) 8:00 a.m. Field Trip 4: “Classic” Geological Sites in the Thunder Bay Area Leaders: Mark Smyk and Mark Puumala 8:00 a.m. Field Trip 6: Amethyst Deposits of Thunder Bay Leaders: Steve Kissin and Greg Paju 5.00 p.m. Return of Trips 4 & 6 Sunday May 24 (Parking Lot G14, Lakehead University) 5.00 p.m. Return of Trip 5 - xxviii - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Constraining the 3-D Geometry of the Duluth Complex, MN, Using Magnetic Fabrics and Paleomagnetic Data AKIN, Kathryn1 and SWANSON-HYSELL, Nicholas1 1 Department of Earth and Environmental Sciences, University of Minnesota, Minneapolis, MN, USA The Midcontinent Rift developed within the interior of Laurentia during a period of extension and magmatism from 1109 Ma to 1084 Ma (Swanson-Hysell et al., 2019). Emplaced during the development of the Midcontinent Rift, the Duluth Complex is interpreted as the second-largest exposed mafic intrusive complex on Earth. The Duluth Complex is composed of an anorthositic series and a layered series of gabbro and troctolite cumulates (Figure 1; Miller et al., 2002). Many studies have been conducted on the geology, mineralization, structure, timing, and mechanisms of emplacement of the Duluth Complex and nearby Beaver Bay Complex and North Shore Volcanic Group, but there is still some uncertainty surrounding the thickness, and therefore overall volume, of the Duluth Complex. Figure 1: Map of the Duluth Complex field location in northeastern Minnesota. Red diamonds represent sampling locations from the August 2025 field season. Geological map data from Bauer (2022). The tilt of the Duluth Complex is not well-constrained in the anorthositic series, given the absence of macroscopic igneous foliation, so this research is focused on developing data on the magnetic fabrics of the Duluth Complex along a transect to constrain the igneous foliation and to use these data to develop new estimates of the tilt and thickness of the intrusion. Anisotropy of magnetic susceptibility (AMS) is sensitive to changes in mineral alignment and, therefore, is used to constrain igneous foliation, especially in samples that do not display an obvious macroscopic fabric in the field (Schmidt et al., 2007). Remanent magnetization data collected and compared with the expected directions of contemporaneous volcanics can also provide further insight into tilt. Together, the new susceptibility and remanence data will provide important petrophysical information for interpreting upcoming USGS aeromagnetic surveys currently being flown in -1- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 northeastern Minnesota. New constraints on the intensity of remanent magnetization and its ratio with susceptibility (the Koenigsberger ratio) will be added to the Rock Properties database maintained by the Minnesota Geological Survey (Chandler et al., 2011). REFERENCES Bauer, E.J., Jirsa, M.A., Block, A.R., Boerboom, T.J., Chandler, V.W., Peterson, D.M., Wagner, K.G., McDonald, J.M., Dengler, E.L., Meyer, G.N., and Hamilton, J.D., 2022, C-54, Geologic Atlas of Lake County, Minnesota: Minnesota Geological Survey: University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/254822. Chandler, V.W., and Lively, R.S., 2011, Density, Magnetic Susceptibility, and Natural Remanent Magnetization of Rocks in Minnesota: An MGS Rock Properties Database: Minnesota Geological Survey, https://hdl.handle.net/11299/175580 Miller, J.D., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., and Wahl, T.E., 2002, RI-58 Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota:, https://hdl.handle. net/11299/58804. Schmidt, P.W., McEnroe, S.A., Clark, D.A., and Robinson, P., 2007, Magnetic properties and potential field modeling of the Peculiar Knob metamorphosed iron formation, South Australia: An analog for the source of the intense Martian magnetic anomalies? Journal of Geophysical Research: solid Earth, v. 112, doi:10.1029/2006JB004495. 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, vol. 131, pp. 913–940, doi:10.1130/b31944.1. Swanson-Hysell, N.L., Hoaglund, S.A., Crowley, J.L., Schmitz, M.D., Zhang, Y., and Miller Jr., J.D., 2021, Rapid emplacement of massive Duluth Complex intrusions within the North American Midcontinental Rift: Geology, vol. 49, pp. 185-189, https://doi.org/10.1130/G47873.1. -2- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Characterization of hematite ore from former Ely mines, NE Minnesota ALLERTON, P. Zsuzsanna1 and HUDAK, J. George1,2,3 1 2 3 Earth and Environmental Sciences, University of Minnesota, Minneapolis, MN 55455, USA Earth and Environmental Sciences, University of Minnesota, Duluth, MN 55812, USA George Hudak Geosciences P.L.L.C., Duluth, MN 55804, USA The hematite ore deposits located in the Vermilion Range in Ely, northeastern Minnesota, represent some of the highest-grade iron ores ever mined in the United States. These deposits occur within Neoarchean (~2.7 Ga) Algoma-type banded iron formations (BIFs) in the Ely Greenstone belt which is dominantly composed of greenschist facies metamorphosed volcanic, sedimentary and intrusive rocks. The ore bodies, exploited in underground mines such as the Zenith, Pioneer, Sibley and others, consist of steeply dipping, tabular to lens-shaped masses of massive hematite that replace jaspilitic BIF. These bodies are enclosed within greenstone wall rocks and are often localized along brecciated zones within a complex regional fold structure. Machamer’s 1968 study of the Zenith mine details the textural varieties of high-grade hematite ore formed by hypogene hydrothermal replacement of jaspilitic BIF. His petrographic and field descriptions identify five prominent ore textures that reflect stages of replacement, brecciation, cementation, and zoning. The characterization and documentation of these five textures at the Pioneer and Sibley mines are the focus of this research. Hematite ore samples utilized for this study were obtained from the Minnesota DNR Hibbing Core Library. Zenith mine ore samples were not available for re-analysis. Ore texture types described are consistent with the nomenclature developed by Machamer (1968). Type 1, the most abundant texture, is a dense, uniform material composed almost entirely of crystalline hematite, representing the primary massive replacement ore (Figure 1A). Type 2 texture consists of brecciated fragments of type 1 ore cemented by a later generation of secondary crystalline hematite, which commonly contains minute vugs lined with small hematite crystals and appear in a reticulated pattern resembling a boxwork (Figure 1B). Type 3 texture is similar to type 2 but features a cement composed dominantly of carbonate minerals (primarily ankerite or siderite) rather than hematite (Figure 1C). Type 4 texture is composed largely of carbonate minerals; it may contain fragments of earlier type 1 hematite material as well as earlier-formed carbonates, reflecting deeper or more advanced carbonate replacement (Figure 1D). Type 5 texture consists principally of magnetite with variable amounts of carbonate minerals, hausmannite (manganese oxide, Mn3O4) and pyrite; this type is generally non-merchantable due to its lower iron content or higher sulfur. The great bulk of the ore mined at Zenith (and similarly at Sibley mine) consisted of types 1 and 2, with lesser amounts of type 3. Many of the types preserve faint layering parallel to the ore-body walls, produced by alternating textural variations in hematite or by interlayering of massive hematite with more porous hematite or carbonates. Texture types 4 and 5 become more abundant with depth in the Zenith mine (Machamer, 1968). These five textures record a progressive hypogene upgrade of BIF to hematite ore that is generally similar to what has been observed in recent research at the Soudan mine (Allerton, 2025; Allerton et al., 2025). Upgrade processes include initial silica replacement by massive hematite, followed by repeated brecciation and multi-stage cementation, and downward transition to carbonate assemblages, producing the dense, low-impurity ore that made the Ely deposits economically significant (Machamer, 1968). -3- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Hematite ore textures from the Sibley and Zenith mines, Ely, MN. A) Type 1 texture showing dense crystalline “matrix” with primary hematite aggregates (white to off white, Hem 1) and vug spaces (black, V). B) Type 2 texture exhibiting brecciated Type 1 material (Hem 1 fragment outlined with white dashed line) cemented by secondary hematite crystals (Hem 2) with reticulated pattern and occasional minute silicates (light gray). C) Type 3 texture displaying brecciated Type 1 material (Hem 1) cemented by mostly carbonates (patchy dark gray, Crb) and some silicates (light gray, Sil). D) Type 4 texture presenting mainly carbonates (patchy light and dark gray, Crb), sporadic silicates (light gray, Sil), and hematite aggregates (Hem 1 fragment outlined with white dashed line) and stingers. REFERENCES Allerton, Z.P., 2025. Thermal and hydrothermal effects of Proterozoic events on Archean rocks in northeastern Minnesota, USA: University of Minnesota ProQuest Dissertations & Theses [Ph.D. thesis]. Allerton, Z.P., Courtney-Davies, L., Danišík, M., Hudak, G.J., Teyssier, C., Mitchell, J.T., and Larson, P., 2025. Hematite double-dating defines Proterozoic mineralization and thermal history of Archean banded iron formations in northeastern Minnesota, USA: Geology, https://doi .org /10.1130 /G53517.1. Machamer, J. F., 1968. Geology and origin of the iron ore deposits of the Zenith Mine, Vermilion District, Minnesota (Special Publication SP-2). Minnesota Geological Survey. -4- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Decoding Shear Zone Evolution in the McFaulds Lake Greenstone Belt, Ontario: Constraints on Crystal-Plastic Deformation and Timing from in-situ Titanite U–Pb Thermochronology ANGOMBE, Moses1, PHILLIPS, Noah2, HOLLINGS, Pete1, STEPHAN, Tobias1, SHESHNEV, Vlad1, DELLER, Mathew3 and SMITH, Andrew3 1 Department of Geology, Lakehead University, 955 Oliver Rd, Thunder Bay, P7B5E1, ON, Canada Department of Earth Sciences, University of Southern California, 3651 Trousdale Pkwy, Los Angeles, 90089, California, United States of America 2 3 Wyloo, 1127 Premier Way unit 1, Thunder Bay, 90089, P7B 0A3, ON, Canada Constraining deformation conditions, kinematics and timing of shear zone activity is essential for determining whether mechanical processes concentrate and localize metal deposits. The McFaulds Lake Greenstone Belt in northern Ontario hosts some of Canada’s most prospective mineralization, including magmatic sulphide, chromite and volcanogenic massive sulphide (VMS)–type deposits. A robust reconstruction of the belt’s deformation history is hindered by an understudied, poorly exposed, arcuate, regionally extensive, dextral shear system including, the Webequie, Triple‑J, and McFaulds shear zones. This study integrates field-based structural observations, microstructural analysis, and in-situ titanite U–Pb geochronology to (1) resolve the kinematic architecture of the major shear zones, (2) constrain the crystal-plastic deformation mechanisms, and (3) determine the temperature and timing of deformation. Newly acquired kinematic results derived from field outcrop‑scale S–C fabrics and asymmetrically rotated porphyroclast microstructures indicate that the NW‑striking Webequie Shear Zone accommodated dextral‑reverse displacement, while the NE‑striking McFaulds and Triple-J Shear Zone are characterized by a dextral‑normal sense of shear. Deformed quartz in phyllonites and mylonites from all shear zones exhibits fine‑grained polygonal aggregates with a few subgrains and a weak crystallographic preferred orientation. These textures indicate that shearing was accommodated predominantly through diffusion‑creep–assisted grain‑boundary sliding processes. Five deformed titanite grains from mylonitic tonalite associated with the Triple‑J shear zone yielded U–Pb dates of ~2775 Ma and Zr‑in‑titanite temperatures of 530–640 °C. In contrast, eighteen euhedral to subhedral titanite grains yield dates between ~2768 and ~2812 Ma and Zr‑in‑titanite temperatures of 650–900 °C. All analyzed titanite grains show no significant difference in temperature or U–Pb dates between rims and cores. We infer that the younger U–Pb dates (~2775 Ma) recorded in deformed titanite constrains the timing of crystal‑plastic deformation, whereas the older, higher‑temperature dates (~2768–2812 Ma) from intact titanites reflect either metamorphic or crystallization. The overlap in deformation and crystallization ages for both deformed and undeformed titanites suggests that shearing in the McFaulds Lake Greenstone Belt was broadly synchronous with emplacement of the regional tonalite suite. These preliminary results show that both shear deformation and magmatism play a critical role in forming the McFaulds Lake Greenstone Belt and its critical mineral deposits. -5- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Coeval silicate melt and PGE-bearing salt melt inclusions in the Thunder and Seagull intrusions, Ontario: An overview of evidence and data processing challenges. BAIN, Wyatt1 and HOLLINGS, Pete 2 1 2 Department of Earth Sciences, Western University, 1151 Richmond St, London, ON N6A 5B7 Canada Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada The Seagull (~90km north-northeast of Thunder Bay) and Thunder (~12 km north-northwest of Thunder Bay) intrusions are two magmatic sulphide-bearing mafic-to-ultramafic intrusions formed during the early stages of Midcontinent rift (MCR) formation. Investigation of olivine crystals from both intrusions reveals abundant assemblages of polycrystalline silicate inclusions and coeval assemblages of hypersaline inclusions. Both inclusion types occur along primary growth zones in their host crystals and undergo partial homogenization at >700 °C. This indicates that these inclusions contain primary, orthomagmatic fluids trapped at magmatic conditions (i.e., immiscible silicate and salt melt). Scanning electron microscope (SEM) analysis shows that the silicate melt inclusions from both intrusions have similar bulk chemistry and host assemblages of feldspar-apatite-phlogopitebiotite-ilmenite-pyrrhotite with a coexisting volatile phase. Similarly, salt melt inclusions from both intrusions also had similar bulk compositions and comprise mixtures of NaCl-KCl with variable amounts of C- and B-bearing salts. The time-resolved laser ablation-inductively coupled plasma mass spectrometry (LA-ICPMS) Figure 1: a., b., Photomicrograph of olivine-hosted coeval assemblages of silicate (SMI) and salt melt (HIS) inclusions from the Thunder (a) and Seagull (b) intrusions. c. Annotated backscatter electron (BSE) image of a silicate melt inclusion exposed at the surface of an olivine crystal. Alb=Albite; Bio=Biotite; Phl=Phlogopite; Apt=Apatite; Hbl=Hornblende; Po=Pyrhotite; Ill=Illmenite; Ol=Olivene d. BSE image of a salt melt inclusion exposed at the surface of an olivine crystal and accompanying energy dispersive spectroscopy maps showing the distribution of selected elements for the same area (right). -6- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 signals from unhomogenized salt melt inclusions from both intrusions consistently showed unambiguous, overlapping peaks for the following element groups: K-P-La-Ce-Ta-U-Th-Nb-RbSr-Ba-Nd-Li, Co-Ni-Cu-Zn-Ag-Pb-S, and Pd-Pt-Au-Sb-Bi. The overlapping peaks for base metals and S likely reflect the presence of crystalline sulphides. Likewise, the overlap of the PGE+Au and Sb-Bi suggests the presence of PGE-bearing antimonide and bismuthide minerals (i.e. PGM). This indicates that salt melts coexisted with silicate melts during the emplacement of both intrusions and were significantly enriched in base metals and PGEs. This data, along with observations of salt melt inclusions in other mafic-ultramafic intrusions (Mcfall et al., 2021; 2023), suggests that these fluids may be important transport media for Ni-Cu-PGE in orthomagmatic environments. Salt melt compositions derived from LA-ICPMS data had unusually high PGE concentrations in the 10s to 100s of ppm. These results should be treated critically, as reducing data from salt melt inclusions presents several technical challenges. These include uncertainty in determining a major element internal standard for salt inclusions and matrix mismatch between the inclusions and the external standard. ICPMS systems are also typically limited in their ability to analyze halogens, C, and S, which are typically major element components of salt melt inclusions (e.g. Xu et al, 2024; Bain et al., 2022). This talk will provide an overview of the geology of the Seagull and Thunder intrusions, present textural and geochemical data from coeval polycrystalline silicate melt and salt melt inclusion in both and discuss the various data reduction schemes being used on this data set. This talk will also discuss a general workflow for salt melt analysis using SEM and LA-ICPMS techniques. REFERENCES Bain, W.M., Lecumberri-Sanchez, P., Marsh, E.E., and Steele-MacInnis, M., 2022. Fluids and melts at the magmatichydrothermal transition, recorded by unidirectional solidification textures at Saginaw Hill, Arizona, USA. Economic Geology, doi:10.5382/econgeo.4952 McFall, K.A., McDonald, I., Yudovskaya, M.A., Kinnaird, J., Hanley, J.J., Kerr, M., and Tattitch, B., 2023. Carbonatedominated hypersaline brines and their importance for metal transport in magmatic and magmatic-hydrothermal critical mineral systems. AGU Fall Meeting, San Francisco, Volume of Abstracts, V44A-08 McFall, K.A., McDonald, I., Yudovskaya, M.A., Kinnaird, J., Hanley, J.J., Kerr, M., and Tattitch, B., 2022. High temperature (> 800° C) brine and sulphide melt interaction during the formation of Northern Bushveld magmatic sulphide Cu-NiPGE deposits. Goldschmidt Conference, Hawaii, Volume of Abstracts, #9496 Xu, X., Bain, W.M., Tornos, T., Hanchar, J.M., Lamadrid, H.M., Lehman, B., Xu, X., Steadman, J.A., Bottrill, R.S., Soleymani, M., Rajabi, A., Li, P., Tan, T., Shihong Xu, S., Locock, A.J., Steele-MacInnis, M., 2024. Magnetiteapatite ores record widespread involvement of molten salts. Geology. 52, 417-422. doi:10.1130/G51887.1 . -7- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Preliminary geochronology of lithium pegmatites and host rocks, Archean Superior Province, northwestern Ontario BEYER, Steve1, CUTTS, Jamie1, HNATYSHIN, Danny1, POWELL, Jeremy1, CAMACHO, Alfredo2, CAWOOD, Tarryn3, and DREVER, Garth4 1 2 3 4 Natural Resources Canada, Geological Survey of Canada, 601 Booth Street Ottawa, ON K1A 0E8 Canada University of Manitoba, 125 Dysart Rd Winnipeg, MB R3T 2N2 Canada University of British Columbia-Okanagan, 3247 University Way Kelowna, BC V1V 1V7 Canada Frontier Lithium Inc., 2614 Belisle Drive Val Caron, ON P3N 1B3 Canada With a combined resource estimated at 50 million tonnes Li grading 1.6% Li2O [1], the PAK and Spark lithium-cesium-tantalum (LCT) pegmatites in northwestern Ontario represent a major potential source of Li, as well as other rare metals such as Nb, Sn, Ta, Rb, and Cs. Together with other Li pegmatite showings in the region (Fig. 1), this suggests high Li prospectivity for the northwestern Superior Province. Better understanding of these significant but understudied pegmatites, together with their peripheral peraluminous granites and other host rocks, will help refine models of rare-metal-enriched pegmatite formation in Archean terranes, and lead to improved discovery success. Here we present multi-mineral geochronological data for the pegmatites and host rocks to clarify connections between pegmatite emplacement and regional tectonics. The crystallization ages of pegmatites and host rocks were investigated using U and Pb isotopes in zircon and monazite measured by SHRIMP. The oldest rock in the area is gabbro that hosts the Spark pegmatite, in which zircon gives an age of 2861 ±3 Ma. Although this unit is mapped as the 2925 Ma Setting Net assemblage of the Favourable Lake greenstone belt, the age is instead within error of the younger 2858 ±5 Ma Eastern Trout assemblage [2]. Zircon in the Pakeagama Lake granite, a biotite-muscovite-garnet peraluminous granite that hosts the PAK pegmatite, gives an age of 2727 ±4 Ma, the first reported age for this pluton. Zircon from coarse K-feldspar-muscovite-apatite-quartz pegmatite at PAK, and zircon from tonalite that hosts the Pennock Lake pegmatite 20 km northwest of PAK, yield ages of 2727 ±1 and 2728 ±4 Ma, respectively, which are the same age as the Pakeagama Lake granite within error. Similar Th/U ratios, indistinguishable ages, and some textural evidence suggests that PAK pegmatite zircon may be inherited from the Pakeagama Lake granite. An overgrowth on one zircon in Spark gabbro gives an age of 2683 ±6 Ma. Isotopes of Hf are used to trace the source of the melt from which the zircon crystallized, and were measured in situ using LA-MC-ICPMS in the same location as the SHRIMP spots. Zircon from gabbro hosting the Spark pegmatite have the most radiogenic εHf values of 5.30 ±0.33, intersecting the value of depleted mantle at 2.86 Ga. Zircon from the PAK pegmatite and tonalite hosting the Pennock Lake pegmatite are less radiogenic, having εHf values of 2.21 ±0.35 and 1.25 ±0.25, respectively, possibly suggesting mixing with older continental crust. Lastly, we examine the thermochronology of muscovite in pegmatite zones, and biotite and hornblende in host rocks and contact zones using Ar-Ar isotope systematics. Step heating age spectra for muscovite (n=10) in the PAK, Spark, and Pennock pegmatites, and the Pakeagama Lake granite, are all disturbed and yield integrated ages between 2532 and 2174 Ma. Hornblende (n=1) in gabbro at Spark gives a slightly disturbed age spectrum with a pseudo-plateau age of 2805 ±4 Ma. Biotite (n=3) in metavolcanics at Spark, and at the contact between the Spark pegmatite and metavolcanics, yield pseudo-plateau ages of 2447 and 2446 ±1 Ma, respectively, whereas biotite in the Pakeagama Lake granite yields a pseudo-plateau age of 1955 ±10 Ma, possibly suggesting partial disturbance of Ar systematics during the Trans-Hudson orogeny. In situ Ar-Ar ages in transects from grain edge to center in muscovite from the Spark pegmatite range from 2687 ± 16 Ma to 1933 ±42 Ma, the oldest age indistinguishable from the U-Pb zircon overgrowth age of 2683 ±6 Ma in Spark gabbro. It is possible this age (~2685 Ma) represents the emplacement of the Spark pegmatite. Taken collectively, these data indicate that the host rocks comprise both ~2861 Ma gabbro and ~2727 Ma granite and tonalite. Although pegmatite emplacement has not yet been directly constrained, it may have occurred together with a thermal pulse at ~2685 Ma, as recorded by Ar-Ar dates from muscovite in the Spark pegmatite, -8- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 and a zircon overgrowth in the host rock. Figure 1. Map showing the location of LCT pegmatites in northwestern Ontario and their host rocks. The area shown in the main map is indicated by the red box in the location map. LCT = lithium-cesium-tantalum; NRCan MRDEM = Natural Resources Canada medium resolution digital elevation model REFERENCES Accad, E., Bisaillon, C., Gagnon, D., Ibrango, S., Liskovych, V., Prévost, G., Sellars, E., and Vasquez, L., 2025. NI 43-101 Technical Report Feasibility Study – PAK Lithium Project, Mine and Mill in Northwestern Ontario, Canada. DRA Americas Inc. Corfu, F., Davis, D.W., Stone, D., and Moore, M.L., 1998. Chronostratigraphic constraints on the genesis of Archean greenstone belts, northwestern Superior Province, Ontario, Canada. Precambrian Research, 92, 277–295. -9- �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Quartz Trace Element and TEM Analysis of Selected Economic LCT Pegmatites BILBOE, Michael1, ZUREVINSKI, Shannon1, and CONLY, Andrew1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1, Canada. This study assesses geochemical and textural trends of economic LCT-type pegmatitic quartz using different analytical applications, namely laser ablation- inductively coupled plasma mass spectrometry (LA-ICP MS) trace element geochemistry, microscope-based laser induced breakdown spectroscopy (LIBS) and high-resolution transmission electron microscopy with- energy dispersive X-ray (TEMEDX) analyses. The study utilized samples from well-documented economic LCT pegmatites (Northwestern Ontario and Manitoba) to assess a variety of modern questions relating to trace element geochemistry. Specifically, this study observes geochemical trends in trace element composition of quartz that display spodumene-quartz intergrowth (SQUI) textures, extrapolates classification of SQUI textures (after Breasley, 2025) to the economic Pakeagama pegmatite (Ontario) and utilizes HR-TEM techniques to image potential Li-bearing nano inclusions hosted within quartz. Breasley (2021) outlined varieties of SQUI originating from unique crystallization sequences. In this study, SQUI from the Pakeagama pegmatite was compared to the recent proposed classifications to ensure consistency in texture classification can be met in different pegmatite systems. Few studies have targeted quartz trace element trends in Group 1 SQUI-bearing pegmatites. Trace element trends in SQUI should be properly understood to avoid improper conclusions when inferring mineralization trends outlined by Müller et al. (2021). Trends in SQUI-associated quartz trace elements were analyzed and compared with non-SQUI pegmatite quartz trace element trends using LA-ICP-MS and LIBS. It was found that few groups of trace elements, particularly Na and Ge, show weak to moderately depleted values with respect to the ratio of Li/Al specifically in quartz grains associated with SQUI (Figure 1). This is interpreted to be the result of trace elements present in the parent mineral (petalite) preferentially incorporating into spodumene rather than quartz during SQUI formation. Additionally, LIBS analysis suggests that elevated concentrations of Li are incorporated into micas and feldspars in the North Aubry sample, likely related to elevated trace element incorporation seen in quartz. Nanoinclusions (fluid and mineral) are thought to be a major contributor to trace element incorporation in quartz (Shah et al., 2022). TEM-EDX analysis was conducted to document and image potential nanoinclusions hosted in quartz. The analyzed portion of the North Aubry sample did not host nanoinclusions displaying any detected Li signatures, however, a decrepitated nanofluid inclusion, with detected sodium and chlorine, was identified (Figure 2). The results suggest that nanoinclusions, while present, may not necessarily contribute significantly to trace element concentrations of Li, Ti, Ge or Be in pegmatitic quartz (possibly due to their presence below detection limits), however, the observed nanoinclusions could suggest the potential Li-brine fluid fluid inclusions and this may be contributing to well-documented quartz trace element concentrations in quartz. REFERENCES Breasley, C. (2021). Lithium aluminosilicate formation and textural origins in evolved pegmatites: Insights from the Tanco Pegmatite, Manitoba and Prof Pegmatite, British Columbia. Doctoral Thesis, University of British Columbia. Müller, A., Keyser, W., Simmons, W. B., Webber, K., Wise, M., Beurlen, H., Garate-Olave, I., Roda-Robles, E., & Galliski, M. Á. (2021). Quartz chemistry of granitic pegmatites: Implications for classification, genesis and exploration. Chemical Geology, 584, 120507. Shah, S. A., Shao, Y., Zhang, Y., Zhao, H., & Zhao, L. (2022). Texture and Trace Element Geochemistry of Quartz: A Review. Minerals, 12(8), 1042. Young, T. (2023). Trace Element Geochemistry of Pegmatitic Quartz from the Superior Province, ON HBSc thesis, Lakehead University. - 10 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Ge (PPM) Versus Li/Al ratios in analyzed samples. Samples Pakeagama, Tanco and Frontier are SQUI-hosted quartz analyses. Data from three additional non-SQUI samples, Seymour, Georgia and Mavis Lake, were included to better highlight the role SQUI has on quartz trace element incorporation (Seymour, Georgia and Mavis Lake data from Young, 2024). Figure 2: TEM image of a nanoinclusion in quartz, identified in the North Aubry sample. EDX mapping of the inclusion detected Na and Cl. - 11 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Improved Precision and Better Accuracy: SHRIMP-II Detrital Zircon Analysis of Samples Across the Stratigraphy of the Midcontinent Rift BLEEKER, Wouter1 and WODICKA, Natasha1 1 Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8, Canada As part of on-going research on the evolution of North America’s Midcontinent Rift (MCR), its stratigraphy, and the detailed setting of its mineral systems, we continue our efforts to improve the age constraints on key geological features of the rift. In addition to many new and improved U-Pb ages on igneous units [e.g., 1,2], we are also undertaking detrital zircon dating of key stratigraphic units across the MCR stratigraphy, from bottom to top (Fig. 1), to resolve remaining questions of depositional ages and sediment provenance. We do so by using the SHRIMP-II ionprobe at the GSC in Ottawa (Fig. 2). With typical spot sizes of ~13x16 μm, fewer corrections during data processing, no down-hole parent-daughter fractionation, and the ability to do multiple, carefully placed spots (away from cracks Figure 1: Generalized stratigraphy of the MCR. Many key ages and mineral systems are indicated. Small red squares identify our detrital samples analyzed by SHRIMP. Figure 2: The SHRIMP-II lab at the Geological Survey of Canada, Ottawa. (SHRIMP: sensitive highresolution ion microprobe.) or other complexities) on grains of particular interest, the SHRIMP-II ionprobe yields significantly more precise and accurate data than more rapid laser ablation analysis, and a more rigorous check on concordancy [3,4]. A typical sample run will analyze 80–100 grains, with >90% of the results falling within the 95–105% concordancy interval (accuracy) used in final interpretation. With multiple spots (n=3–5) on key grains, the 2s uncertainty of weighted mean ages can be improved to ±5–15 Ma (precision). All of this does take a fair amount of machine time, with a typical spot analysis taking ~15 mins, and an entire sample run, including calibration on well-characterized zircon reference materials, more than 24 hrs. Analysis is done on polished grain mounts that are imaged in both BSE (backscatter) and CL (cathodoluminescence) mode prior to analysis to guide grain selection and spot location. Here we briefly discuss some initial results. One such result, on the high-energy “event layer” near the top of the Gunflint Formation, was presented at an earlier ILSG meeting [5]. It confirmed that this layer contains ejecta material from the Sudbury target area in the form of ca. 2460–2450 Ma zircons from the Creighton Granite and Copper Cliff Rhyolite. In the next sample up (Rove Formation greywackes), our results fail to identify any age peaks younger than ca. 1845 Ma, which we consider the maximum depositional age for the Rove Formation [cf. 6], i.e. entirely a Penokean foreland basin. - 12 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 There are hints of some younger grains, perhaps to as young as ca. 1805 Ma, but this requires further work. Interestingly, in addition to some Archean input, there is also one grain at ca. 2311 Ma of the reworked felsic ash material known from the upper Huronian Supergroup [see also ref. 5]. Thin sandstone layers intercalated with the Pillar Lake Volcanics basalt flows, near Armstrong, show youngest grains at ca. 1500 Ma, similar to our Sibley Group sandstone samples, and do not contain any of the abundant younger grains (and peaks) prominent in the basal MCR sandstones discussed below (see Fig. 3). This confirms our interpretation that these thin sandstone beds and the Pillar Lake Volcanics are part of basal Sibley Group rift volcanism and sedimentation at ca. 15001480 Ma, not a northern outlier of ca. 1.11 Ga MCR stratigraphy sensu stricto [cf. 7]. Three samples of the sandstone/quartzites (Bessemer, Nopeming, and Puckwunge formations) immediately below the onset of “Early Stage” basaltic volcanism yield generally similar results with youngest grains in the 1135–1100 Ma age range (weighted means), and strong peaks (modes) at ca. 1125 Ma, 1160–1140 Ma and various older ages (e.g., 1470 Ma, Wolf River Batholith), all the way to 3.3 Ga (Fig. 3). These are just some initial results and a full and complete analysis of all 12 samples will be presented elsewhere. SOME REFERENCES [1] [2] [3] [4] [5] [6] [7] Bleeker, W., Smith, J., Hamilton, M., Kamo, S., Liikane, D., Hollings, P., Cundari, R., Easton, M., and Davis, D., 2020. Geological Survey of Canada, Open File 8722, p. 7–35. DOI: 10.4095/326880. Smith, J., Bleeker, W., and Hamilton, M., 2026. GSA Bulletin, v. 138(3–4), p. 1419–1438. DOI: 10.1130/B37649.1. Stern, R.A., 1997. Geological Survey of Canada, Current Research 1997-F, p. 1–31. DOI: 10.4095/209089. Stern, R.A., and Amelin, Y., 2003. Chemical Geology, v. 197, p. 111–146. DOI: 10.1016/S0009-2541(02)00320-0. Bleeker, W., Wodicka, N., Kamo, S., Hamilton, M., Emon, Q., and Smith, J., 2024. 70th ILSG Meeting, Proceedings & Abstracts, Part I, p. 11–12. Heaman, L., and Easton R.M., 2006. Ontario Geological Survey, Miscellaneous Release, MRD-191, 78 p. Hollings, P., Smyk, M., Bleeker, W., Hamilton, M., Cundari, R., and Easton, M., 2021. Canadian Journal of Earth Sciences, v. 58(10), p. 1116–1131. DOI: 10.1139/cjes-2021-0012. Figure 3: Example of our SHRIMP-II detrital zircon results: probability density plot for the Bessemer Quartzite (BNB-18022), sampled just below the onset of basalt flows. Inset: images of youngest grains with 3 spots (repeated analyses at the same locality), yielding a weighted mean age of 1101±14 Ma. - 13 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Update to: a complex F-rich alkalic pegmatite in the pyroxene syenites of the Stettin Complex, Wausau Complex, Marathon County, Wisconsin BUCHHOLZ, Thomas1, FALSTER, Alexander2, and SIMMONS, William2 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494, 1 MP2 Research Group, Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217, USA 2 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 composed of various syenite phases. This abstract is an update to studies of this dike reported in ILSG 2024 and 2025; interested readers are referred to those abstracts. As noted by Buchholz et al. (2025) relatively common soft, pale yellow to creamy to brown grains typically contain high Ti-Ce-Fe contents with traces of other elements, and were suspected to consist of an unidentified Ti-Ce4+-Fe phase. Hand-picked grains from several visually identical samples were analyzed using powder XRD to determine crystalline phases present. Results indicate the presence of only three crystalline phases: arfvedsonite, lucasite-(Ce), (CeTi2(O,OH)6) and cerianite(Ce), (Ce4+,Th)O2. To balance charges in lucasite-(Ce), Ce is likely present as Ce4+ and OH probably absent or negligible. The altered grains may have originally been a LREE-Ti-rich mineral such as chevkinite-(Ce) or aeschynite-(Ce) that were subsequently altered under oxidizing conditions, removing LREE3+ and Si (and altering Ce3+ to Ce4+), thus allowing the crystallization of lucasite-(Ce) and cerianite-(Ce). Oxidation states appear to have fluctuated during pegmatite crystallization, as Ce3+ rich minerals such as synchysite/parisite, britholite-group minerals, monazite-(Ce) and indeed sparse remnants of chevkinite-(Ce) are present in later crystallizing portions of the dike. The potential for britholite-group minerals was discussed by Buchholz et al. (2025), and since then two group minerals have been identified: fluorbritholite-(Ce) and britholite-(Ce). Both form small pale pink to whitish rounded masses in pockets and vugs. At a minimum EDS analysis is required to distinguish these two species, as well as distinguish them from visually similar synchysite/parisite series minerals. Although bismuthinite is known from thin veinlets crosscutting the pegmatite (Buchholz et al., 2025), native Bi has subsequently been found as masses in small interior zone vugs in the pegmatite. Standards-based EDS indicates the Bi contains small amounts of Te; approximately 2-3.5 wt. %. The Te (as Te2-) is probably present as small admixed grains of a Bi-Te mineral such as tellurobismutite, hedleyite or another Bi-Te species. Possible nacareniobsite-(Y) was found as an inclusion in a small aggregate of fergusonite-(Y). Standards-based EDS data show good agreement with the published composition of the species, but the small size of the grain (approx. 25 µm) and the scarcity of the mineral suggest more examples should be sought to confirm this data. Nacareniobsite-(Y) was first described in 2023 and is so far a one-locality mineral, suggesting this may be the second locality for this species. Recent thorough cleaning of fresher exposures has revealed that parallel joints or fractures are closely spaced across much of the pit exposure. All are parallel, near-vertical and roughly oriented WNW-ESE. Possible displacement is unknown at this time, but they suggest a degree of oriented stress may have affected portions of the pluton late in its cooling history or at sometime thereafter. REFERENCES: Buchholz, Thomas, Falster, Alexander, and Simmons, Wm, 2024. Preliminary mineralogy of a pegmatite in the pyroxene syenites of the Stettin Complex, Wausau Complex, Marathon County, Wisconsin (Abstract): Institute on Lake Superior Geology, 70th Annual Meeting, Part I, Program and Abstracts, 19-20. - 14 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Buchholz, Thomas, Falster, Alexander, and Simmons, William, 2025. A complex F-rich alkalic pegmatite in the pyroxene syenites of the Stettin Complex, Wausau Complex, Marathon County, Wisconsin (abstract): Institute on Lake Superior Geology, 71st Annual Meeting, Part I, Program and Abstracts, 15-16. Van Wyck, N. 1994. The Wolf River A-type magmatic event in Wisconsin: U/Pb and Sm/Nd constraints on timing and petrogenesis (abstract): Institute on Lake Superior Geology, 40th Annual Meeting, Part 1, Program and Abstracts, 81-82. - 15 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Reassessing variations in metamorphism across the Penokean orogen in Northern Michigan: Part 2, Reinterpreting metamorphic nodes CANNON, W. F.1, SALERNO, R.1, DRENTH, Benjiman J.2 and BEDROSIAN, Paul A.2 1 2 U.S. Geological Survey, Reston, VA U.S. Geological Survey, Denver, CO A classic study of regional metamorphism (James, 1955) documented variations in metamorphic grade in Paleoproterozoic sedimentary rocks across the Upper Peninsula of Michigan. James interpreted the spatial variations of index minerals as four discrete nodes of metamorphism with concentric zones, defined in pelitic rocks, ranging from chlorite to sillimanite grade (Figure 1). Those isograds are widely used up to the present day to characterize the Penokean metamorphism of the region. These concentric nodes imply localized sources of heat across the region rather than a more widespread source related to regional orogenic processes. Figure 1. Map showing isograds interpreted by James (1955) and distribution of metamorphic index minerals from James and later studies. Compilation of metamorphic index minerals in northern Wisconsin indicates that the high-grade metamorphism extends well west of the Watersmeet node as mapped by James. Widespread garnet occurrences observed in core drilled through Paleozoic cover rocks also show that metamorphism to at least garnet grade extends far east of the exposed Peavy node. Patterned area is proposed allochthon(s) which include the Iron River-Crystal Falls and Menominee iron ranges. Gray shaded region in SE is area of Paleozoic cover. We propose an alternative interpretation for the Watersmeet and Peavy metamorphic nodes and their implied discrete heat sources. The index mineral occurrences in Figure 1 show a belt, at least 250 km long, of metamorphism to garnet or higher grade including scattered occurrences of kyanite to about 50 km north of the Niagara fault. That belt is broken by a gap of about 50 km between the Watersmeet and Peavy nodes where rocks are mostly chlorite-grade sedimentary rocks. We suggest that the gap is a result of post-metamorphic northward emplacement of allochthons of low-grade rocks over the more highly metamorphosed rocks, and that the belt of high-grade rocks is continuous beneath the allochthons. The belt of high metamorphic grade rocks, thus, is a result of regional tectonic burial to mid- to lower crustal depths, and related heating during the climactic closing phase of the Penokean orogeny, rather than to largely speculative individual heat sources. More localized - 16 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 heating by contemporaneous intrusions likely caused some magnification of regional heating such as in the Peavy node (Roy, et al., 2025) The allochthonous nature of the Paleoproterozoic rocks was first proposed by Sims (1992) and supported by more recent work (i.e. Cannon and Ottke, 1999), and recently acquired aeromagnetic and electromagnetic data. The very close spacing of isograds inferred by James (1955), such as the southeastern edge of the Watersmeet node and southwestern edge of the Peavy node, would require extreme lateral temperature gradients that are difficult to reconcile with progressive heating from a central source. Those abrupt lateral changes in metamorphic temperatures are more consistent with a tectonic contact between the high-grade rocks and overthrust low-grade rocks. If that is correct, it has significant implications for the age of allochthon emplacement and the nature of post-Penokean tectonism in the region. The peak metamorphism of the Watersmeet and Peavy nodes is well constrained to 1837-1825 Ma at depths of 30-35 km (Roy, et al., 2025: Salerno, et al., in press). Emplacement of allochthons with low metamorphic grade directly atop these mid- to lower-crustal rocks implies that the high-grade rocks were largely exhumed before emplacement, and that overthrusting must have been a post-Penokean event. Rapid exhumation of active orogens has been documented in many places globally with rates measured in kilometers/million years, so exhumation observed in Michigan could have been accomplished in 10 million years or less. Thus, the suggested overthrusting could be only slightly younger than the generally accepted ~1830 Ma date for termination of Penokean deformation, nevertheless recording continued post-Penokean regional compressive tectonism in the region. REFERENCES Cannon, W.F., and Ottke, D., 1999. Preliminary digital geologic map of the Penokean (early Proterozoic) continental margin of Northern Michigan: U.S. Geological Survey Open-File report 99-547. James, H.L., 1955. Zones of regional metamorphism in northern Michigan: Geological Society of America Bulletin, v. 66, p. 1465-1488. Roy, Supratik, Holder, R.M., Jahandar, R., Brenner, D.C., Nelson, L.L. and Viete, D.R., 2025. Mantle heating drove shortduration Barrovian-type regional metamorphism during the Penokean orogeny, Michigan (USA) Geological Society of America Bulletin, https://doi.org/10.1130/B38653.1 Salerno, R., Cannon, W.F., Thompson, J., Souders, A., Vervoort, J., and Hillenbrand, I., in press. Unraveling protracted modification of Archean and Paleoproterozoic crust in central Laurentia, Penokean orogen, with garnet and accessory mineral geochronology and microstructural analysis: Geological Society of America Bulletin. Sims, P.K., 1992. Geologic map of Precambrian rocks, southern Lake Superior region, Wisconsin and northern Michigan: U. S. Geological Survey Miscellaneous Investigations Map I-2185, scale 1:500,000. - 17 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 An introduction to the northwestern Huron Mountains of the Upper Peninsula, Michigan: field relations and preliminary structural interpretations CARLTON, Kenz M.1, TIKOFF, Basil1, and NACHLAS, William O.1 University of Wisconsin–Madison, Department of Geoscience, 1215 West Dayton Street, Madison, Wisconsin 53706, USA 1 The Huron Mountains of the Upper Peninsula, Michigan, are part of a granite-greenstone terrane and likely represent part of the southern extent of the Superior Craton. Recent field mapping and microstructural analysis indicate the existence of an amphibolite basement intruded by compositionally variable granitoids. The amphibolite basement is a banded schist with a high amphibole content that may represent a strongly metamorphosed mafic protolith. The two plutons of this site each have rapidly varying appearances and expressions of fabrics, banding that varies from non-existent to thick gneissic, and variable compositions from monzogranite to quartz-rich tonalite lithologies. The contacts between the schist and granitoid plutons of this site vary in expression over relatively short distances and, in some cases, can be traced from a planar feature into a 50+ m wide transition zone. The relation between the granitoid and the amphibolites is intrusive, as a range of sizes of amphibolite inclusions can be found within the plutons, usually near the contacts. Mafic and felsic dikes are both abundant. Ongoing work to analyze bulk and trace element geochemistry and U-Pb geochronology will constrain the timeframe of geologic events, the tectonic origin of the groundmass (i.e., which terrane, protolith), and the source of plutonism. The pervasive regional fabric displays a general northwest strike/northeast dip; however, the foliation expression in outcrop is frequently inconsistent, with tens of degrees of difference in both strike and dip possible within 30 meters or less. In general, traceable exposures of the schist-pluton contacts are parallel or subparallel to foliation. Additional structures found in outcrops include mesoand micro-scale faults and meso-scale or larger shear zone features. In thin section, microstructures indicate solid-state deformation, including myrmekite, cuspate-lobate grain boundaries, and internal grain deformation. These analyses support the model of emplacement of quartz-rich plutons into a meta-mafic basement during regional shearing, in the northwestern Huron Mountains. - 18 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Petrography and Geochemistry of the Mound Lake Pluton, Northwestern Ontario CHAISSON, Amy1, SMYK, Mark1, and ZUREVINSKI, Shannon1 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Mound Lake lies approximately 90 km north-northeast of Thunder Bay and 25 km northwest of Nipigon. The Mound Lake Pluton is a 7 km-wide, ovoid, muscovite-bearing granite that has intruded Quetico metasedimentary rocks. It was first described by Hart (2005) and Hart et al. (2005) and thought to be a prospective fertile granite, capable of hosting or spawning rare metal mineralization (cf. Breaks et al., 2005). This study documents the petrography, mineral chemistry, and whole-rock geochemistry of the Mound Lake granitic rocks. The study utilized samples from an initial geochemical and geological reconnaissance program (Smyk, 2022). Thirty-one samples were collected from the pluton and another was collected from from a granitic pegmatite dyke in andalusite schist from the shore of Frazer Lake. Analytical methods included transmitted light microscopy, major and trace element whole-rock geochemistry, and quantitative mineral compositional analyses using Scanning Electron Microscopy- Energy Dispersive X-ray Spectroscopy (SEM-EDX) with Back Scattered Electron (BSE) imaging to characterize mineral textures and compositions. The Mound Lake granitic rocks host irregular pegmatitic patches and miarolitic cavities containing quartz and large, drusy K-feldspar crystals. Massive, medium-grained granitic rocks are crosscut by a variety of aplitic and pegmatitic dykes. Petrographic and mineral compositional analysis identifies the pluton as a two-mica granite, composed of K-feldspar, quartz, muscovite, biotite, and plagioclase, with accessory zircon, apatite, monazite, tourmaline and thorite. Plagioclase, whose compositions range from albite to oligioclase, locally exhibited Na-rich, albite rims. Biotite compositions were found to represent annite/siderophyllite endmembers. Perthitic exsolution and granophyric intergrowths exemplify late-stage crystallization, while sericitization and chlorite alteration are related to post-magmatic hydrothermal activity. The presence of granophyric intergrowths suggests that at least some portions of the magma experienced pronounced undercooling during the final stages of crystallization. Geochemical data confirm a peraluminous, S-type affinity (Alumina Silica Index of 1.05–1.31) with trace element signatures plotting in the Volcanic Arc Granite (VAG) and syn-collisional fields. Granitic rocks display moderate LREE enrichment and HREE depletion, with variable Eu anomalies reflecting the relative role of plagioclase fractionation and accumulation. The Mound Lake Pluton shows increased Li and Cs (+ Ce, Ta and Be) concentrations along its northern contact (Figure 1). Elevated Ce concentrations correlate with samples with higher monazite content. The consistently peraluminous nature and mineralogy (muscovite + biotite, + garnet) support the contention that the pluton is a product of metasedimentary melting, likely triggered by thermal relaxation following oblique accretion in the Superior Province (Chappell, 1999). A spodumene-bearing, granitic pegmatite dyke, discovered in 2023 (https://www.geologyontario.mines.gov.on.ca/mineral-inventory/ MDI000000003501), approximately 3 km north of the northern contact of the pluton, attests to the fertility of local granitic rocks. - 19 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1. Li (left) and Cs concentrations (right) across the Mound Lake pluton (data from Smyk, 2022). REFERENCES Breaks, F. W., Selway, J. B., and Tindle, A. G. (2005). Fertile Peraluminous Granites and Related Rare-Element Pegmatites, Superior Province of Ontario. Short Course Notes, 17, pp.87–125. Chappell, B. W. (1999). Aluminium saturation in I- and S-type granites and the characterization of fractionated haplogranites. Lithos, 46(3), pp.535–551. https://doi.org/10.1016/S0024-4937(98)00086-3. Hart, T.R. 2005. Precambrian geology of the southern Black Sturgeon River and Seagull Lake area, Nipigon Embayment, northwestern Ontario; Ontario Geological Survey, Open File Report 6165, 63p. Hart, T.R., Whaley, A.G. and Pace, A. J. 2005. Precambrian Geology of the Southern Black Sturgeon River–Seagull Lake– Disraeli Lake Area, Nipigon Embayment, Northwestern Ontario; Ontario Geological Survey, Preliminary Map P.3562, scale 1:50 000. Smyk, M. C. (2022). NI 43-101 Early-Stage Exploration Property Report, Mound Lake Property, Thunder Bay District, Ontario, Canada; Technical Report, 107p. https://www.geologyontario.mines.gov.on.ca/persistentlinking?assessment=20000022160. - 20 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Digital Preservation and Enhanced Utility of Exploration Core Descriptions from the Keweenaw Copper District, Michigan: Progress toward a Map-based Web Portal DeGRAFF, James, HILTUNEN, Lindsay, LAFRENIERE, Don, LIZZADRO-McPHERSON, Dan, VYE, Erika, COWLING, Bob, BORNHORST, Theodore, J., and ROSE, William (deceased) Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 U.S.A. The Michigan copper rush starting in 1843 at Copper Harbor (Fig. 1) led to 150 years of mining that produced ~7.5 x 106 MT of copper (1), 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. Companies invested significantly in trenching, coring, and mining operations that generated an enormous body of geologic information. The U.S. Geological Survey (USGS) compiled much of this information in the 1950s as bedrock geology maps with supporting cross sections and reports. Available online in digital form, these map products are derived in large part from a substantial quantity of detailed paper records that are not easily accessed, including core descriptions from exploratory holes drilled from 1899 through the 1970s. Drilling records produced after the 1950s generally have not been used in later investigations also because of difficulty of access. Paper records and microfiche that degrade with time are stored at various locations (2-4), further complicating their use. A few years ago, we began a volunteer project to identify and gather such information into a digital image repository, to extract it into tabular databases, and to explore how to make it available (5) for use by scientists, industry, land-use planners, and the general public (Fig. 2). These early efforts led to a two-year project funded by a Save America’s Treasures grant (ST-256897-OMS-24) through the National Park Service, focused on drilling records in the Michigan Figure 1: Michigan’s native copper mining district with exploratory diamond-drill holes (DDHs) coded by information that is available. TBD – to be determined; WUP – Western Upper Peninsula. Technological University Archives. The current project has three phases: 1) scan all paper records of core descriptions, drafted vertical sections, and drilling metadata; 2) convert scanned records to character data and store in files with - 21 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 tabular formats; 3) create an online, GIS-based, search tool to provide access to the materials. Phase 1 of the project, now complete, has produced scanned core logs for 801 diamond-drill holes from 64 series. After the project was terminated in April 2025 and then reinstated in June, we prioritized Phase 3 to develop the online GIS-based search and delivery tool for scanned files in case funding was lost again. Functional design work is complete and implementation is being tested. Drill hole locations for the GIS-based map were digitized from USGS maps of the Keweenaw Peninsula and supplemented with data from Michigan’s EGLE website. A drillhole attribute table contains positional data, hole direction, total depth, and drilling metadata. Phase 2 of the project is ongoing and involves extracting character data from PDF files and creating tabular data for each core description. We are investigating optical character recognition to extract character data combined with AI tools to organize the data into prescribed tabular formats. This has proven successful for high-fidelity records but requires human checking and editing to ensure the accuracy of extracted data. Less well preserved records may require humans to transcribe them and manually enter characters into the tables. Upon making these MTU records available to others in an online format, we hope to extend this work to similar records in Acknowledgements: We thank the U.S. National Park Service for the grant that makes this work possible. Casey Koch and Gwen Martin performed nearly all of the document scanning. This work is possible because of the foresight of many late geologists who gathered and preserved the original paper records. Figure 2: Potential uses of the database upon completion. the other archives. REFERENCES 1. 2. 3. 4. 5. 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 Midcontinent of North America: Geological Society of America Field Guide 24, p. 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. Michigan Technological University Archives, 2025, Major Mining Company Collections MS-001, MS-002, MS080, MS-635: J. Robert Van Pelt and John and Ruanne Opie Library, Houghton, Michigan (accessed December 2025). DeGraff, J.M. and Rose, W.I., 2020, Digital capture and preservation of historic mining data from the Keweenaw copper district, Michigan: GSA Abstracts with Programs, v. 52, no. 5, doi: 10.1130/abs/2020NC-348035. - 22 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 A new look at the Seagull mafic-ultramafic Intrusion and potential hydrogen and helium accumulations DROST, Abraham 1 and HEGGIE, Geoff 2 1 2 Rift Minerals Inc. 1113 Jade Court, #102, Thunder Bay, Ontario P7B-6V3 Canada Pursuit Geosciences, 245 Nicholetts Road, Murillo, Ontario P0T-2G0 Canada The mafic-ultramafic Seagull intrusion located approximately 80km northeast of Thunder Bay, Ontario and forms part of the Paleoproterozoic 1.1 Ga Midcontinental Rift which extends in an arcuate shape from Iowa through Lake Superior into Michigan (Fig. 1). The intrusion was intruded into the Archean Quetico Metasedimentary Terrain and transects a portion of the Sibley Group Metasedimentary rocks. The Quetico Terrain is dominated by deep water turbidites accumulated in a forearc basin between adjacent volcanic terranes, that underwent inversion during crustal accretion. Partial melting of the Quetico Terrane at depth resulted in the generation and emplacement of S-type melts at shallower levels with both uranium occurrences and LCT pegmatites present (Fig. 2). Figure 1. Geological and geophysical interpreted extent of the 1.1 Ga Midcontinent Rift centered on Lake Superior. Distribution of major rock types shown along with location of Seagull Project (Rift Minerals) and Topaz Project (helium: Pulsar Helium) Figure 2. Geology map of the Lake Nipigon area. Archean basement terrains shown in the legend. 1.1Ga Midcontinent Rift rocks shown in purple with early olivine bearing intrusions outlined in red. Uranium occurrences identified are demarked by yellow and orange circles from Ontario OMI database. Historic exploration between 1998 and 2012 on the Seagull Intrusion included airborne and ground geophysical surveys and approximately 20,000m of diamond drilling. The geology of the Seagull intrusion is characterized by mafic-ultramafic rocks, with in-excess of 700 m of variously serpentinized olivine cumulate rocks, predominantly lherzolites and pyroxenites (Fig. 3). This exploration work identified disseminated to semi-massive sulphide mineralization containing nickel, copper and platinum group elements along parts of the intrusion’s basal contact and as reef-type mineralization. Additionally, the exploration operator at the time reported the presence of naturally occurring gases at pressure. Histoically, the intrusion was targeted for orthomagmatic mineralization, without attention being paid to the presence of gas. With the discovery of an unconventional helium reservoir within the MCR, the prospectivity of the area has pivoted, resulting in new ideas in explored areas. Serpentinization is well known as an alteration process that generates hydrogen. The presence of ubiquitous uranium and LCT pegmatite occurrences in the Archean basement metasedimentary rocks of the Quetico Terrain are a potential source of helium (Fig. 2). Lithostatic pressures, structural plumbing and concentration gradients can potentially result in downward migration of generated gases (Strauch et al, 2023). - 23 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 3. East-West cross-section through the Seagull intrusion as interpreted from diamond drilling. Modified from East West Resources (2002). Figure 4. Cross section through inversion model of the Ambient noise tomography (ANT) survey completed by Sisprobe (2024). Historical drill traces shown in white. Cross section at AZ of 027° facing NNW. In 2024, Rift Minerals completed an ambient noise tomography (ANT) survey with Sisprobe to refine the internal geometry of the Seagull intrusion and to identify subsurface velocity contrasts interpreted to reflect lithological and alteration variations. Integrated interpretation of drilling and geophysical data sets, including ANT velocity modelling, has been used by Rift to refine the interpreted geometry of the Seagull intrusion and underlying basement. The ANT velocity section (Fig. 4) is of high statistical quality and agrees well with stratigraphic variations identified in drilling. An unexplained low velocity interval within or beneath high velocity Quetico basement rocks below the Seagull Intrusion, topping at ~1250m, is being targeted for high pressure gas reservoir potential (Fig. 4). Rift Minerals and its funding partner Anteros Metals Inc. initiated a drill program in 2026 to test the deep lower velocity feature with drill hole RM26-01. The drill hole intersected disseminated to locally weakly net-textured, orthomagmatic sulphide mineralization in the basal cumulate sequence of the Seagull intrusion grading:* • 7.25 metres from 587.00 to 594.25 m grading 1.58 g/t Pt+Pd (0.72 part per million Pt and 0.86 ppm Pd), with 294 ppm copper and 2,168 ppm nickel; • 1.00 m from 606.25 to 607.25 m grading 2.27 g/t Pt+Pd (1.02 ppm Pt and 1.25 ppm Pd), with 1,660 ppm Cu and 2,080 ppm Ni. * Weighted-average results using a 0.5-gram-per-tonne-platinum-plus-palladium cut-off During the drilling of hole RM26-01 pressurized gas was encountered at a depth of approximately 877m within a narrow fault zone in the Quetico basement rocks. The 877-metre occurrence is located approximately 100m southwest from drill hole WM01-08, which reportedly encountered pressurized and flammable gas at a similar stratigraphic level when drilled in 2001. The significance, continuity and composition of the gas remain under evaluation. REFERENCES Strauch, B., Pilz, P., Zimmer, M and Hierold, J., 2023. Hydrogen Migration through natural rocks – an experimental approach. Harvard University – EGU23, the 25th EGU General Assembly, held 23-28 April, 2023 in Vienna, Austria (https://egu23.eu) - 24 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Integrated X-Ray Diffraction and Petrography Document Carbonate Mineral Heterogeneity and Hematite Mineralization in the Upper Biwabik Iron Formation, MN DUFFY, Paige1, BRENGMAN, Latisha1, and EYSTER, Athena2 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Heller Hall 1114 Kirby Drive, Duluth, MN 55812, USA 1 Department of Earth and Climate Sciences, Tufts University, Lane Hall, 2 North Hill Road, Medford, MA 02155, USA 2 Core LWD-99-1 preserves the ~1.9 Ga Biwabik Iron Formation located near the Virginia horn, outside the contact metamorphic aureole associated with the intrusion of the Duluth Complex ca. 1.1 Ga. In this study, X-ray diffraction (XRD) and petrographic data are used to: (1) characterize carbonate mineral heterogeneity; and (2) evaluate depositional and post-depositional mineral assemblages. Emphasis was placed on identification of Fe2+ bearing carbonates (e.g., ankerite, siderite) and hematite-magnetite relationships. To minimize contamination and weathering effects, outer surfaces were removed during processing, and veins were avoided. Samples were cut, dried, then crushed to a uniform powder using a SPEX ShatterBox. XRD analyses were performed using a PANalytical X’Pert diffractometer, with data collected in θ-2θ geometry over a range of 5-65 degrees, sufficient to capture all minerals of interest. Analysis of diffraction data was done using X’Pert HighScore (Malvern PANanalytical) software. XRD analysis of 24 samples documents the presence of multiple different carbonate and oxide minerals throughout core LWD-99-1. Carbon was detected in 92% of analyzed samples, with 77% associated with carbonate minerals. All carbonate phases identified petrographically and with scanning electron microscopy (Duncanson et al., 2024) were also detected by XRD, indicating strong agreement between methods. Siderite is the most common carbonate phase, occurring in 41.7% of samples, followed by ankerite at 33.3%, dolomite in 20.8%, kutnohorite in 16.7%, and calcite in 6.3%. Carbonate mineral distribution greatly varies by informal stratigraphic unit. Siderite is prevalent in the Lower Slaty, Lower Cherty, and Upper Slaty, whereas kutnohorite (a calcium manganese carbonate) only occurs in the Upper Cherty. Calcite is restricted to the uppermost part of the Upper Slaty while dolomite and ankerite are most abundant in the Upper Cherty but also appear once in the Lower Cherty and twice in the Upper Slaty. Overall, carbonates are more abundant in the Upper Slaty and Upper Cherty compared to the Lower Slaty and Lower Cherty. Within the Upper Slaty, siderite occurs in 42.9% of samples, ankerite in 28.9%, dolomite in 28.6%, and calcite in 14.3%. In the Upper Cherty, siderite and ankerite each occur in 50% of the samples, dolomite in 20%, and kutnohorite in 40%. These distributions highlight a clear variation of carbonate minerals in the upper portions of the stratigraphy. Additionally, preliminary XRD and petrographic observations of iron oxide minerals suggest an overall increase in hematite occurrence in the Upper Cherty and the Lower Cherty, with petrographic data indicating magnetite is more prevalent in slaty units. Such transitions towards increasing carbonate mineral diversity and increasing hematite up section could link to depositional changes in the system or post-depositional oxidation reactions. Post-depositional mineral reactions and accounting of ferrous: ferric iron ratios are of critical interest as they preserve a record of fluid: rock interaction driven by multiple geologic events. Some redox reactions that involve siderite and magnetite are of broader interest for tracking hydrogen production or stimulation potential (Geymond et al., 2023; 2025). Ongoing work includes detailed accounting and mapping of mineral distributions and mineral reactions across the lateral extent of the iron formation. - 25 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: XRD analysis of LWD-99-01 Sample MIR 17-11 taken from the Upper Cherty showing variation in carbonate mineralogy (ankerite (00-033-0282), dolomite (00-036-0426), kutnohorite (00-043-0695), and siderite (00-029-0696)). REFERENCES Duncanson, S., Brengman, L., Johnson, J., Eyster, A., Fournelle, J., Moy, A., 2024. Reconstructing diagenetic mineral reactions from silicified horizons of the Paleoproterozoic Biwabik Iron Formation, Minnesota. American Mineralogist, 109, 339-358. Geymond, U., Briolet, T,. Combaudon, V., Sissmann, O., Martinez, I., Duttine, M., Moretti, I., 2023. Reassessing the role of magnetite during natural hydrogen generation. Front. Earth Sci. 11, 1169356. Geymond, U., Truche, L., Sissmann, O., Kubaniova, D., Recham, N., Martinez, I., 2025. Mineralogical changes and H2 generation yield during hydrothermal alteration of a magnetite-siderite assemblage. Journal of Geophysical Research: Solid Earth, 130, 8. - 26 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 The Badgerow complex, a Midcontinent Rift-related REE-Zr-rich peralkaline intrusion in the Grenville Province near Verner, Ontario EASTON, Robert Michael1 and KAMO, Sandra L.2 Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, retired, 933 Ramsey Lake Road, Sudbury, Ontario P3E 4W1 1 Jack Satterly Geochronology Laboratory, Department of Earth Sciences, University of Toronto, Toronto, Ontario M5S 3B1 2 The Badgerow complex (Lumbers 1975; Easton 2025) is located approximately 8.5 km north of the community of Verner within the northern Nepewassi domain of the Grenville Province (Easton 1992). The main part of the complex is roughly circular, approximately 4.5 by 3.5 km in size (Figure 1), and is only weakly deformed, with a narrow gneissic margin and a massive to slightly foliated interior. It consists predominantly of pink weathering, medium-grained monzogranite with less than 5% mafic minerals (sample 24RME-3047). Near the eastern margin of the complex, fine-grained monzogranite veins crosscut medium-grained gabbro of the complex containing relict pyroxene cores rimmed by amphibole. The monzogranite was sampled for geochemistry and U-Pb geochronology because of the relatively undeformed nature of these rocks, and the fact that the complex is the only near-circular pluton within Nepewassi domain. Approximately 600 m northeast of the near-circular body, Lumbers (1975) included an approximately 6 km long, up to 1 km wide, lens of gneissic syenite as part of the Badgerow complex. Well-exposed along Highway 575 (Figure 1), the lens is a homogeneous, medium-grained, gneissic amphibole syenite (sample 24RME-3052) hosted by migmatites. Given its mineralogy, and its greater degree of deformation, the lens was assumed to be older than the granitic rocks. It is unclear why Lumbers (1975) included it in the Badgerow complex. Preliminary geochemical results from the complex were reported in Easton (2025, 2026). Sample 24RME-3052 (Figure 2) is peralkaline and has niobium, yttrium, zirconium and total rare earth Figure 1. Simplified geological map of the Badgerow complex in the Grenville Province north of Verner (from Easton 2025). Sites sampled for geochemistry and for U/ Pb geochronology are indicated. Figure 2. Chondrite-normalized rare earth element plot for granitoid samples mentioned in the text (from Easton 2025). Remember the y-axis scale is logarithmic, so the difference between samples 24RME-3047 and 24RME-3052 is larger than it might appear (e.g., La normalized is 97 ppm for sample 24RME-3047 but 1866 ppm for sample 24RME-3052). Sample 24RME-1114 is an undeformed monzogranite exposed near Noelville. Normalizing values of Sun and McDonough (1989) were used. - 27 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 contents (164, 155, 4800 and 1936 ppm, respectively) that are some of the highest recorded for any igneous rock sample from the Grenville Province of Ontario. There are some similarities between sample 24RME-3052 and the West Bay migmatitic monzonite body to the south of Verner (Easton 2014). Key differences are that the West Bay body samples do not show a europium anomaly (Figure 2) nor are they peralkaline. It is unclear if the gneissic syenite was originally an intrusive or a volcanic rock. If volcanic, it has a comendite composition. In contrast, granite sample 24RME-3047 has a much lower total rare earth content (Figure 2). Preliminary U–Pb chemically abraded-isotope dilution thermal ionization mass spectrometric results on zircons have been obtained from samples 24RME-3047 and 24RME-3052. Zircons from sample 24RME-3047 are discordant and lie along a reference line anchored between 1097 and 2700 Ma. The most concordant zircon from sample 24RME-3052 gives a 207Pb/206Pb age of 1106 Ma (igneous based on Th/U). This age is older than Grenvillian metamorphism in Nepewassi domain (1030-980 Ma, Easton 2026), but similar to the Early Stage of Midcontinent Rift magmatism (11101104 Ma, Smith et al. 2026) and the Rb-Sr age of a mantle-xenolith bearing lamprophyric breccia at Elliot Lake (1112.8±4.95 Ma, Legros et al. 2024). Both these potential Midcontinent Rift-related intrusions lie along the northwest-southeast rifting trend of the Early Stage of magmatism, despite their location east of any previously described Midcontinent Rift magmatism. These new results suggest that other Midcontinent Rift-related intrusions may be present in the Sault Ste-Marie to North Bay area. REFERENCES Easton, R.M. 1992. The Grenville Province; Chapter 19 in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 2, p.713-904. ——— 2014. Geology and mineral potential of the Nepewassi domain, Central Gneiss Belt, Grenville Province; in Summary of Field Work and Other Activities, 2014; Ontario Geological Survey, Open File Report 6300, p.16-1 to 16-12. ——— 2025. Zirconium and rare-earth element potential of a Grenville Province gneiss north of Verner, northeastern Ontario; in Summary of Field Work and Other Activities, 2025; Ontario Geological Survey, Open File Report 6421, p.10-1 to 10-7. ——— 2026. Geological, geochemical, geophysical and petrographic data from the Wanup area, Grenville Province, northeastern Ontario; Ontario Geological Survey, Miscellaneous Release—Data 397. Legros, H., Czas, J., Luo, Y., Woodland, S., Sarkar, C., Shirey, S.B., Schulze, D, and Pearson, D.G. 2024. Post‑Archean Nb‑REE‑U enrichment in the Superior craton recorded in metasomatised mantle rocks erupted in the 1.1 Ga Midcontinental Rift event; Mineralium Deposita, v.59, p.373-396. Lumbers, S.B. 1975. Burwash area, districts of Nipissing, Parry Sound and Sudbury; Ontario Department of Mines, Geological Report 116, 158p. Accompanied by Map 2271, scale 1:126 720. Smith, J.W., Bleeker, W. and Hamilton, M. 2026. The 1093 Ma Crystal Lake Intrusion: A nickel-copper mineralized intrusion emplaced during the younger southwest–northeast rift phase of the Midcontinent Rift (North America); Geological Society of America, Bulletin, published online Oct 15, 2025, 20p. Sun, S-S. and McDonough, W-F. 1989. Chemical and isotopic systematics of oceanic basalts: Implications for mantle compositions and processes; in Geological Society of London, Special Publication No.42, p.313-345. - 28 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Comparing mineralogy along a surface to depth transect of the ~2.7 Ga North Limb Soudan Iron Formation, NE Minnesota. ELLISON, Kimberly1, CISNEROS, John Alex1, EYSTER, Athena2, and BRENGMAN, Latisha1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Heller Hall, 1114 Kirby Drive, Duluth, MN 55812, USA 1 Department of Earth and Climate Sciences, Tufts University, Lane Hall, 2 North Hill Road, Medford, MA 02155, USA 2 The Lake Vermilion-Soudan Underground Mine State Park in Northeast Minnesota is home to the classic 2.7 Ga Algoma-type Banded Iron Formation - a type of authigenic, chemical sedimentary rock known to record past ocean chemistry. Here, we compare multiple generations of mineralization in the Soudan Iron formation to evaluate the relative timing of oxidation reactions. Recent U-Pb and (U-Th)/He hematite geochronology places new age constraints on iron mineralization of microplaty hematite, documenting that this generation of hematite post-dates initial deposition by over 1 billion years (Allerton et al., 2025). Distinguishing between initial mineral formation and later overprinting is critical for reconstructing paleowater-rock interactions within the Soudan system. The goal of this work is to compare drill core and outcrop records from the north limb of the Soudan fold to samples from the mineralized portion of the Soudan mine, with a focus on building a spatial map documenting these oxidation reactions. To evaluate mineral reactions in the Soudan Iron formation, we combine transmitted and reflected light petrography with X-Ray Diffraction (XRD), focusing on a vertical transect of samples from drill core 26501 from the north limb of the Soudan Iron Formation, comparing these samples to nearby surface outcrop samples, and mine samples from the fold hinge to the west. Preliminary results indicate shallow core samples (28.5 to 95 feet) contain dominant mineral assemblages of quartz, calcite, magnetite, hematite, and minimal iron silicates, while deeper samples (129 to 394 feet) mainly contain quartz, magnetite, carbonate, chalcopyrite, and iron silicate assemblages, lacking visible hematite. To compare optical data to XRD data, we prepared powdered whole rock samples using standard cutting and crushing techniques for XRD scanning at positions ranging from 2θ = 5° to 2θ = 65°. Resulting peaks were matched to mineral reference patterns provided by the X’Pert HighScore analysis software and compared to observed mineralogy of thin section samples from drill core 26501. XRD results confirm the presence of quartz, magnetite, and hematite in shallow drill core samples, and an assemblage of quartz, magnetite, and iron silicates in deeper drill core samples. Combined, petrographic data and XRD data indicate hematite is confined to shallow drill core samples. This observed trend continues in petrographic data from surface outcrop samples near the same location, which also contain abundant hematite. The absence of hematite at depth in drill core samples, combined with the top-down nature of the hematite distribution, could indicate minor amounts of hematite locally formed from surface oxidation distal to the mine site. Next steps include more detailed paragenesis work in combination with larger-scale mapping of the spatial distribution of hematite in drill core along the north limb of the Soudan iron formation towards the historic mine site which sits at the fold hinge. Mapping the extent of multiple generations of oxidation reactions can help document past fluid-rock interactions and allows for identification of preserved ferrous iron-containing assemblages at depth in the iron formation. Such ferrous-iron-containing phases may record depositional information and are of interest for potential natural or stimulated hydrogen generation. - 29 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1.) Petrographic photomicrographs and X-ray Diffraction patterns of the Soudan iron formation from core 26501. (A) Reflected light image of sample 26501-28.5, a shallow (28 feet depth) banded iron formation sample with quartz and hematite. (B) Cross-polarized light image of sample 26501-191, a deeper iron formation sample (191 feet), that contains quartz, calcite, magnetite, and iron-silicate mineral phases. (C) XRD analysis of sample 26501-28.5 (28 feet depth) with peaks that match mineral reference patterns of quartz, magnetite, and hematite. (D) XRD analysis of sample 26501-242 (242 feet depth) with peaks that match mineral reference patterns of quartz, magnetite, and iron silicates. REFERENCES Duncanson, S., Brengman, L., Johnson, J., Eyster, A., Fournelle, J., Moy, A., 2024. “Reconstructing diagenetic mineral reactions from silicified horizons of the Paleoproterozoic Biwabik Iron Formation, Minnesota”. Mineralogical Society of America, Volume 109, Number 2, American Mineralogist, https://doi.org/10.2138/am-2022-8776. Geymond, Ugo, Briolet, T., Combaudon, V., Sissmann, O., Martinez, I., Duttine, M., Moretti, I., 2023. “Reassessing the Role of Magnetite during Natural Hydrogen Generation”. Frontiers in Earth Science, Volume 11, Frontiers, 10.3389/ feart.2023.1169356. Zsuzanna, P. Allerton, Courtney-Davies, L., Danisik, M., Hudak, G., Teyssier, C., Mitchell, J., Larson, P., 2025. “Hematite double-dating defines Proterozoic mineralization and thermal history of Archean banded iron formations in Northeastern Minnesota, USA”. Geology, Volume 53, page 11, Geological Society of America, https://doi. org/10.1130/G53517.1. - 30 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Middle school virtual field trip lessons materials for Archean formations of Lake VermilionSoudan Underground Mine State Park ERICKSON, Stephanie S.1, FAYON, Annia2 , ALLERTON, Zsuzsanna1,2, and HUDAK, George2,3,4 1 2 3 4 Curriculum and Instruction, University of Minnesota School of Earth and Environmental Science, University of Minnesota, Minneapolis, MN 55455, USA School of Earth and Environmental Science, University of Minnesota, Duluth, MN 55812, USA George Hudak Geosciences P.L.L.C., Duluth, MN 55804, USA The Lake Vermilion-Soudan Underground Mine State Park located in St. Louis County, Minnesota provides unique opportunities to learn about Archean geology and mineral resources of northern Minnesota. Archean rocks exposed in the park consist of a series of mafic lava flows and intrusive rocks interlayered with classic banded iron formation, iron ore, felsic tuffs, and chloritesericite schists (Hudak et al., 2014, Hudak and Peterson, 2014; Peterson et al., 2016) and record deformation associated with the accretionary growth of the Superior craton. A cross-section through the stratigraphy can be observed along a trail through part of the east side of the park. The trail is in the planning stages and is in collaboration with the state park. The purpose of this project is to enhance formal and informal Earth science education in Minnesota. After consultation with local secondary teachers the project expanded to include a virtual field trip with an accompanying lesson as part of the formal education portion of the project. In 2019 Minnesota revised their science standards (Minnesota Department of Education, 2019). These changes marked a significant change in the pedagogical practices aligned with national trends such as Next Generation Science Standards (NGGS) (NGSS Lead States, 2013). There are a number of shifts in instruction teachers are challenged to make when implementing these standards including using phenomenon based instruction (BSCS Science Learning, 2017; Reiser et al., 2021). Phenomenon based instruction engages students in a series of lessons arranged in a cohort storyline around a real world, observable events. An additional challenge facing Minnesota educators was moving Earth science in from 8th grade to 6th grade. According to survey data collected from the Minnesota Earth Science Teachers Association many 6th grade teachers did not feel prepared to teach Earth science content. A combination of lack of high quality instructional materials for Minnesota phenomenon and gaps in the required background knowledge are some factors contributing to these findings. This project provided teachers with high quality instructional materials that are aligned with the 2019 Minnesota State Science Standards for 6th grade teachers. Three, 45-minute lessons were designed to address the stratigraphy standard. The goal for the students is to tell the geological story of the park. The first lessons take students on a virtual walk through the park stopping at six significant outcrops along the way (Figure 1). At each stop students are making observations of the rock outcrops and hand samples while also asking questions. The second lesson, using information about rock formation, processes the map and picture of core samples taken from locations in the park (Figure 2) while applying principles of deformation and stratigraphy. formations and thus the early geologic history of the Earth. The lessons conclude with students writing a story of the Archean formations and thus the early geologic history of the Earth. - 31 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: The first two stops orient students to the park and where they learn about the park’s iron mining history including an open pit and a trip down to a deep mine. After emerging from the underground mine they make three stops at outcrops: the classic BIF outcrop, the schist in BIF outcrop, and Ely Greenstone pillow basalts. The final stop is to make observations of tuff and the lower section of the Ely greenstone from rocks found on the “ground.” (after Peterson et al., 2016 ) Figure 2: Virtual core samples that students use to correlate and deduce the order the rocks are formed in. Each core sample comes from a point of the map in figure 1. These are not actual core samples rather simplified samples that allow students to correlate the stratigraphy of the area. REFERENCES BSCS Science Learning. (2017). Guidelines for Assessing Instructional Materials that Exemplify the NGSS. https://bscs.org/ reports/guidelines-for-assessing-instructional-materials-that-exemplify-the-ngss/ Hudak, G. J., and Peterson, D. M., 2014, Non-Ferrous Mineralization Associated with the Wawa-Abitibi Terrane and Duluth Complex Cu-Ni-PGM Deposits, Northeastern Minnesota: Society of Economic Geologists, Guidebook Series, v. 47, 150 p. Hudak, G. J., Radakovich, A., Pignotta, G., and Schwierske, K., 2014, Field Trip 2 – A Walk in the Park – Neoarchean Geology of Lake Vermilion State Park: Institute on Lake Superior Geology, Proceedings Volume 60, Part 2 – Field Trip Guidebook, p. 37-75. Minnesota Department of Education. (2019). 2019 Minnesota Academic Standards in Science. https:// education.mn.gov/mdeprod/idcplg?IdcService=GET_FILE&dDocName=MDE086711 &RevisionSelectionMethod=latestReleased&Rendition=primary NGSS Lead States. (2013). Next generation science standards: For states, by state. The National Academies Press. Washington D.C. Peterson, D.M., Hudak, G.J., Radakovich, A., Pignotta, G., and Schwierske, K., 2016, Geologic Map of Lake Vermilion/ Soudan Underground Mine State Park: Precambrian Research Center Map PRC/Map-2016-01, 1:10,000 scale. Reiser, B. J., Novak, M., McGill, T. A. W., & Penuel, W. R. (2021). Storyline units: An instructional model to support coherence from the students’ perspective. Journal of Science Teacher Education, 32(7), 805–829. https://doi.org/10 .1080/1046560X.2021.1884784 - 32 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Geochemical Constraints on Mn Cycling in the Paleoproterozoic Gunflint Formation GILBERG, Nolan1, FRALICK, Philip1, and LI, Zhiquan1 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Iron formations (IFs) are iron-rich (>15% Fe) and siliceous (>20 wt.% SiO2) chemical sedimentary rocks that precipitated from seawater. Most IFs were deposited between 2.80 and 1.85 Ga during the Neoarchean and Paleoproterozoic, followed by a near one-billion-year hiatus before reappearing in the Neoproterozoic. The Gunflint Formation in the Animikie Basin, overlain by the siliciclastic Rove Formation, is composed mainly of IFs, chert, carbonates, and minor siliciclastic sediments deposited during the Paleoproterozoic (~1.88 Ga), and represents the final major episode of IF deposition. Therefore, investigating the source materials and redox conditions of the Gunflint Formation is key to understanding this transitional period in the marine environment. This study conducts a high-resolution stratigraphy and chemostratigraphy study of a 142.9-meterdeep drill hole (MC-1-89), located south of Thunder Bay in the Gunflint Iron Range. Samples were taken in short intervals of ~1-5 meters along the drill core, where 55 samples were analyzed for major, trace and rare earth (REE+Y) element concentrations through ICP-OES and MS. All samples from drill core MC-1-89 consists of IFs (often magnetite, hematite rich, or jaspilite), chert, carbonates, and siliciclastic rocks (often fine sandstone and argillaceous mudstone). IFs contain a total Fe content ranging from 15-36%. MnO values are enriched in the upper and lower portion of the hole (0.30, 0.57 wt.% respectively), while depths 30-90m show an average of 0.10 wt.%. Samples with >1 wt.% Al2O3 and >0.1 wt.% TiO2 are excluded for REE+Y analysis due to potential detrital contamination. The rest of the samples do not show correlation of REE+Y with Al2O3 + TiO2 (R2 < 0.1), suggesting the REE+Y system is authigenic. All samples display positive Eu/Eu* (1.18 – 3.14, average ~1.83), suggesting a strong hydrothermal input. Moreover, most of the samples display a depletion of LREE, enrichment of HREE, along with high Y/Ho ratios (average of 30.4), suggesting marine signatures. All these features are typical of global Paleoproterozoic IFs. A key distinction between the Gunflint Formation and other Paleoproterozoic IFs is the presence of positive Ce anomalies in many samples, which contrasts with most Archean and Paleoproterozoic IFs. Ce/Ce* values decline with depth (0.35 – 1.92, average =1.32). The elevated Ce might be related to the cycling of Mn oxides in the water column, but further detailed work is still needed to better constrain the mechanism. - 33 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Modified Sequential Iron Extraction Method for Analyzing Rare Earth Elements in Banded Iron Formations GOSAI, Meghna, FRALICK, Philip, and LI, Zhiquan Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Banded iron formations (BIFs) are chemically precipitated sedimentary rocks characterized by alternating iron-rich and silica-rich layers, formed predominantly in Precambrian marine environments. Rare earth elements (REEs) are among the most used geochemical tools for understanding the origin and deposition of iron formations and other iron oxide–rich sedimentary rocks, because the precipitation of ferric iron oxyhydroxides can adsorb signatures from the water column and thus preserve a seawater REE signature. However, BIFs that formed in shallow-marine settings often contain detrital material, thereby affecting the bulk rock geochemistry. For instance, detrital input may elevate light REEs and suppress yttrium (Y) anomalies, complicating interpretation. Sequential extraction of different iron phases (e.g., magnetite, iron carbonates, and iron sulphides), developed by Poulton and Canfield (2004), was used to accurately determine the composition of ironbearing minerals without interference from detrital materials. However, the chemical solutions used in this process introduce additional dissolved ions, thereby increasing total dissolved solids (TDS) and making it difficult to analyze low REE concentrations using ICP-MS. Therefore, this study aims to develop a method to reduce the introduced TDS while still extracting enough REEs for detection by ICP-MS. Six concentrations (10%, 20%, 40%, 60%, 80%, and 100%) of an ammonium oxalate monohydrate and oxalic acid solution were used for sequential extraction. This solution was used to selectively extract magnetite from two types of samples: (1) magnetically selected magnetite grains, and (2) bulk rock powder from the same sample. The extracted iron solutions were then analyzed for REE anomalies for interpretation. Sample patterns were compared to determine the minimum concentration required to introduce additional elements into the solution without resulting in a high dilution factor. The patterns were also compared with those from Dolega’s (2018) bulk-rock acid digestion to assess any improvements in REE patterns. The results and comparison indicate that the REE patterns show the greatest improvement at a solution concentration of 40%. However, one concern is the absence of a positive Y anomaly, which differs from the original bulk rock data (Dolega, 2018). It is likely that reprecipitation causes the interference with Y, but further work is still needed for this investigation. REFERENCES Dolega, S., 2018. Geochemistry of Shallow and Deep Water Archean Meta-Iron Formations and Their Post-depositional Alteration in Western Superior Province, Canada. Unpbl. MSc thesis, Lakehead University, Department of Geology. Poulton, S.W., and Canfield, D.E., 2005. Development of a Sequential Extraction Procedure for Iron: Implications for Iron Partitioning in Continentally Derived Particulates. Chemical Geology, 214, 209–221. - 34 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Time-to-depth conversion of seismic-reflection data from eastern Lake Superior and implications for the eastern arm of the Midcontinent Rift GRAUCH, V.J.S.1, and HELLER, Samuel J.2 1 2 U.S. Geological Survey, MS 973, Federal Center, Denver, CO 80225 U.S. Geological Survey, MS 939, Federal Center, Denver, CO 80225 Seismic-reflection data were acquired in the mid 1980s along several lines across eastern Lake Superior by industry and the Great Lakes International Multidisciplinary Program on Crustal Evolution (GLIMPCE) (Fig. 1). The lines form part of a larger network of crossing lines over the entire lake, which can be used to develop three-dimensional geologic models of the Mesoproterozoic Midcontinent Rift that lies below. To better interpret these lines, we developed velocity models to convert seismic reflections versus two-way travel time (TWTT) to reflections versus depth. In addition, the velocity models themselves provide insights into the structure of the Midcontinent Rift by recognizing common velocity ranges for certain rock types (Grauch, 2023). Figure 1. Seismic-reflection lines overlain on Bouguer gravity for eastern Lake Superior. Gravity map from Anderson and Grauch (2018) is displayed in color shaded-relief, with illumination from the northeast. Lake Superior shores are outlined in black. Digital data are publicly available for lines A, F, and G, collected as part of GLIMPCE. Digital data were derived for the industry lines (LS-15, LS-25, LS-26, and LS-36) by scanning published images from McGinnis and Mudrey (2003) and estimating the location parameters. Velocity model development was guided by (1) bathymetric data, providing thickness of the lowvelocity water column; (2) previous shallow seismic-reflection studies targeting the top of bedrock below glacial till and lake sediments; (3) previous refraction studies, which provide information on depth and compressional velocity at interfaces of large velocity contrast; and (4) correlations across multiple lines, allowing independent constraints on individual lines to influence modeling on crossing lines. In addition, digital data for the GLIMPCE lines were analyzed using common midpoint gathers to check the accuracy of the modeled velocities. Gravity anomalies provided qualitative guidance on broad velocity variations. The velocity models consist of intervals of constant velocity bounded by prominent horizons recognized in the seismic-reflection TWTT sections before time-to-depth conversion. After - 35 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 conversion, the resulting reflection sections versus depth show similar overall geometry compared to the TWTT sections, although structural relief is more subdued. Thus, several qualitative aspects of the results are similar to those observed by previous workers (e.g., Cannon et al., 1989; Mariano and Hinze, 1994; Samson and West, 1994). For example, lines that cross the lake from SW to NE are interpreted to show a symmetric basin of fairly uniform basalt thickness except at the edges of the basin, where the basalts rise and thin and are expressed by pronounced gravity highs (Fig. 1). The thickness of the overlying sedimentary section increases toward the middle of the basin to 7–9 km and the underlying volcanic section is locally folded. In contrast, the velocities derived from the modeling indicate different rock types than anticipated from the previous interpretations at the edges of the basin. The upturned basalt edges have been previously interpreted as basalt layers thrust over the younger Jacobsville Sandstone, with sharply rounded reflection patterns considered as thrust rollovers on lines LS-26 and LS-36 between the crossings with LS-15 and LS-25 (Mariano and Hinze, 1994). Where these authors interpret Jacobsville Sandstone under thrust faults, the models indicate velocities on the order of 6.0 km/s instead of the expected velocity range of 3.0–4.5 km/s for this unit (Grauch, 2023). The higher velocities are consistent with those of igneous or basement rocks instead. An alternate interpretation is that the upturned edges represent the vestiges of magmatic feeder zones and the sharply rounded reflection patterns represent igneous intrusions. The zones may be faulted and folded due to the later compressional regime that affected the region. 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. Cannon, W.F., Green, A.C., 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.H., and Spencer, C., 1989, The North American Midcontinent rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305–332. doi: 10.1029/TC008i002p00305. Grauch, V.J.S., 2023, Compressional-wave seismic velocity, bulk density, and their empirical relations for geophysical modeling of the Midcontinent Rift system in the Lake Superior region: U.S. Geological Survey Scientific Investigations Report 2023-5061, 60 p. https://doi.org/10.3133/sir20235061. 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. 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. https://wgnhs.wisc.edu/pubs/000480/. Samson, C., and West, G. F., 1994, Detailed basin structure and tectonic evolution of the Midcontinent Rift System in eastern Lake Superior from reprocessing of GLIMPCE deep reflection seismic data: Canadian Journal of Earth Sciences, v. 31, p. 629–639. - 36 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Ice flow history, surficial geology, and till composition of Georgia Lake area, northwestern Ontario HAGEDORN, Grant1 Ontario Geological Survey, Ministry of Energy and Mines, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5 Canada 1 During the last glaciation, the Lake Superior basin was covered by the Laurentide Ice Sheet. The ice sheet advanced over the landscape, eroding the substrate and depositing a variety of sediments including till (a common sample medium for mineral exploration) and glaciofluvial sand and gravel (a common source of aggregates). During deglaciation, glacial lakes inundated the landscape depositing successions of silt and clay, which can act as a barrier for mineral exploration and infrastructure development. As such, the Ontario Geological Survey completed a three-year field mapping program which measured striations and landforms to decipher different ice flow directions, mapped the surficial geology around the Georgia lake pegmatite, and collected regional scale till samples to identify mineral prospectivity (Figure 1). These data hold broad applications for regional mineral exploration and land use planning / resource management decisions for local communities. Striation and landform mapping were used to determine the relative age and direction of ice flow over the region. A southwest flow is pervasive across mafic uplands, suggesting this was the paleoflow direction during thickest ice cover (Arrows labeled 1 in Figure 1). As the ice sheet thinned, it became more topographically-controlled resulting in southward ice flow in lowlands, and westward ice flow on mafic uplands (Arrows labeled 2 in Figure 1). Finally, a late-stage re-advance out of the Lake Superior basin created northwestward striations and landforms in the areas around Thunder Bay (Arrows labeled 2 in Figure 1). Surficial mapping completed in the Georgia Lake area indicate more sediment than previously identified although the sediments are mostly thin (>2 m). Till is common at surface and many new small eskers have been mapped. Glacial lake sediments are present, and at a higher elevation than previously indicated. Postglacial organic accumulations are also abundant over the landscape, specifically over poorly-drained substrates like till and glaciolacustrine silt and clay. Till samples were also collected as part of the project and analyzed for till matrix geochemistry and indicator minerals. Till compositions indicate two units differentiated based on bedrock provenance. One till contains southwest transported carbonate material while the other contains locally sourced bedrock material. Lithium material transported southwest from the Georgia Lake pegmatite is also clearly identified in both the geochemistry and indicator mineral data. Further work is being completed on the till samples to indicate prospectivity of the region for other deposits. - 37 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Study area for the project. Highways and towns are labeled. Ice flow directions indicated by white arrows with the corresponding flow event as the number beside (1: older, 2: younger). Surficial geology mapping area is indicated by the dash box. Till sample locations are circles. - 38 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Geochemistry, Petrogenesis, and Mineralization of the Makwa Deposit, Bird River Sill HARDING, Myles1 and HOLLINGS, Pete1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, On P7B 1J4, Canada. The Maskwa West-Dumbarton layered mafic-ultramafic intrusion is located approximately 145 km northeast of Winnipeg, Manitoba and is host to the Ni-Cu-PGE Makwa Deposit. The intrusion is related to the 2743 ± 0.5 Ma Bird River Sill (BRS; Scoates and Scoates, 2013) which is approximately 15-25km long and is made up of several separated ~800m thick differentiated mafic-ultramafic intrusive bodies. The Maskwa West-Dumbarton intrusion is emplaced into the mafic metavolcanic MORB-type massive to pillowed basalt Northern Lamprey Falls Formation (Mealin, 2008, Duguet et al., 2009). After the discovery of the Maskwa deposit in 1975, a year later 332,000 tonnes of nickel copper ore was mined in a shallow open pit (Grid Metals Corporation, 2024). In 2004 Mustang Minerals (now Grid Metals Corporation) acquired the property and have since completed extensive drilling and geophysical surveys targeting PGE mineralization. The approximately 5 km long Maskwa West-Dumbarton intrusion is composed of a ~500m thick upper gabbro-anorthositic section and a ~500m thick lower section of metaperidotite-metapyroxenite (Mustang Minerals Corp., 2014). The intrusion has been metamorphosed to the lower amphibolite facies (Coats and Buchan, 1979) with primary igneous textures Maskwa West-Dumbarton obscured or completely overprinted by alteration. The Makwa deposit is a conventional basal accumulation type magmatic sulphide deposit with the highest grade mineralization hosted within the lowest portion of the ultramafic series (Grid Metals Corporation, 2024). The deposit is comprised of a magmatic assemblage of disseminated to net textured and semi-massive pyrrhotite-pentlandite-chalcopyrite as well as low sulphide platinum group minerals (PGM) mineralization (Grid Metals Corporation, 2024). The open pit resources at Makwa are indicated to be 14.2 million tonnes with 0.48% nickel, 0.11% copper, 0.02% cobalt, 0.37 g/t palladium, and 0.10 g/t platinum (Grid Metals Corporation, 2024). The most recent up to date resource estimate for the high-grade zone indicates 4.8 million tonnes with a grade of 0.89% nickel and a 1.26% nickel equivalent (Grid Metals Corporation, 2024). The purpose of this project is to characterize the stratigraphy of the Maskwa-Dumbarton body and Ni-Cu-PGE mineralization. Assess the effects of alteration on the mineralogy, trace element geochemistry, and ore remobilization. A fence of five drill holes covering the stratigraphy of the intrusion were selected for this project where 151 core samples were collected. Forty polished thin sections were cut in representative areas for petrographic and scanning electron microscope (SEM) analysis. 141 of those samples were selected for whole rock geochemical analysis. A combination of petrographic and geochemical analysis was used to characterize the Makwa mafic and ultramafic rocks. Primary mineralogy is almost entirely replaced (Fig. 1) therefore preserved relict cumulus textures along with whole rock geochemistry are utilized to determine primary mineralogical composition. The Makwa ultramafic samples dominantly plot as Mg-rich cumulates within the komatiite field (Fig. 2) displaying a trend of Fe-enrichment highlighting strong fractionation. The results of this study will be used to determine the evolution, geotectonic setting, and sulfur source of the sulphides. REFERENCES Coats, C. J. A., & Buchan, R. (1979). Petrology of serpentinized metamorphic olivine, Bird River Sill, Manitoba. Canadian Mineralogist, 17, 847–855. Duguet, M., Gilbert, H.P., Corkery, M.T. and Lin, S. (2009): Geology and structure of the Bird River Belt, southeastern Manitoba (NTS 52L5 and 6): reprinted with revisions; in Report of Activities 2006, Manitoba Science, Technology, Energy and Mines, Manitoba Geological Survey, p. 170–183. Grid Metals Corp. – Combined Makwa and Mayville Project, Technical Report NI 43-101 – June 14, 2024 - 39 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Mealin, C. A., & University of Waterloo. Department of Earth Sciences. (2008). Geology, geochemistry and Cr-Ni-Cu-PGE mineralization of the Bird River sill evidence for a multiple intrusion model. University of Waterloo. Mustang Minerals Corp. – Combined Makwa and Mayville Project, #2098 Technical Report NI 43-101 – April 30, 2014 Scoates, J. S., & Scoates, R. F. J. (2013). Age of the Bird River Sill, southeastern Manitoba, Canada, with implications for the secular variation of layered intrusion-hosted stratiform chromite mineralization. Economic Geology and the Bulletin of the Society of Economic Geologists, 108(4), 895–907. Figure 1. Photomicrograph (XPL) of Makwa peridotite displaying mesh-textured serpentine replacing metamorphic blade shaped olivine in net-textured sulphides. Figure 2. Jensen Cation Plot highlighting Makwa Mg-rich cumulates dominantly within the komatiite field displaying Fe and Al-enrichment trends highlighting strong fractionation. - 40 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Using Anisotropy of Magnetic Susceptibility and U-Pb Geochronology from the Bush Lake Granite, Florence County, WI to Understand Post-Penokean Continental Growth HELLRUNG, Alyssa1, DROUBI, Omar Khalil1, RUGGLES, Claire1, and BONAMICI, Chloë1 Department of Geosciences, University of Wisconsin-Madison, 1215 W. Dayton Street, Madison, Wisconsin, 53706, USA 1 The Bush Lake granite in Florence County, Wisconsin, is well suited to constrain the timing of granitic magmatism relative to Proterozoic deformation events as the youngest intrusion in the Dunbar Gneiss Dome. The Dunbar Gneiss Dome is south of the Niagara fault zone, which marks the suture of the Pembine-Wausau terrane to the Superior craton during the 1.85 Ga Penokean orogeny (Schulz and Cannon, 2007). This suture may have been reactivated during later orogenic events, such as the ca. 1.75 Ga Yavapai orogeny, the ca. 1.65 Ga Mazatzal orogeny, and/or the ca. 1.45 Ga Baraboo orogeny. Emplacement and deformation of the Bush Lake granite determined through U-Pb geochronology, microstructural analysis, and anisotropy of magnetic susceptibility (AMS) fabric data provides insight into the tectonic history of the region. The Bush Lake granite is a weakly peraluminous biotite granite that contains quartz, plagioclase, megacrystic alkali feldspar, and accessory allanite, zircon, titanite, and apatite. Microstructures in the Bush Lake granite indicate variable solid-state deformation, including interlobate grain boundaries and undulose extinction in quartz, as well as grain size reduction of quartz and feldspar. Magnetic mineralogy, which informs the AMS fabric, is dominated by paramagnetic biotite with trace magnetic oxides. AMS fabrics generally record NW-SE striking foliations and moderately plunging to subvertical lineations (Figure 1), which are consistent with predominantly NE-SW shortening at a high angle to the Niagara fault zone and associated vertical thickening of the crust. Based on solid-state deformation microstructures, this magnetic fabric formed after emplacement and crystallization of the unit and records a younger period of deformation than previously thought. Cathodoluminescence (CL) imaging shows that most Bush Lake zircons preserve oscillatory zoning of likely magmatic origin, though many zircon crystals also have irregular, disturbed zoning and low-CL regions consistent with alteration. The Bush Lake granite was previously interpreted to have intruded at ~1835 Ma as a late-stage intrusion of the Paleoproterozoic Penokean orogeny, coeval with other nearby granites (Sims et al., 1985). Based on U-Pb SIMS analyses of zircon, the Bush Lake granite is interpreted to have emplaced at 1749 ± 1 Ma, making it coeval with the 1754 ± 11 Ma Amberg granite, ~28 km southwest (Holm et al., 2005), rather than the more proximal ~1835 Ma granites in the Dunbar Gneiss Dome. Solid-state deformation recorded by the Bush Lake granite may signify a broader regional deformation event in northern Wisconsin after 1750 Ma, possibly related to re-activation of the Niagara Fault Zone during the Yavapai orogeny or later events. REFERENCES Holm, D. K., Van Schmus, W. R., MacNeil, L. C., Boerboom, T. J., Schweitzer, D., and Schneider, D., 2005. U-Pb zircon geochronology of Paleoproterozoic plutons from the northern midcontinent, USA: Evidence for subduction flip and continued convergence after geon 18 Penokean orogenesis. Geological Society of America, 117(3/4), 259-275. Schulz, K. J., and Cannon, W. F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research, 157(14), 4-25. Sims, P. K., Peterman, Z. E., and Schulz, K. J., 1985. The Dunbar Gneiss-granitoid dome: Implications for early Proterozoic tectonic evolution of northern Wisconsin. Geological Society of America Bulletin, 96, 1101-1112. - 41 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Simplified geologic map of the Bush Lake pluton (pink) in Florence, Wisconsin, with sample locations plotted and colored by average magnetic susceptibility [SI]. Lower hemisphere equal area net projections bordering the map show the AMS foliation plane and lineation at each site for each specimen. At each site, ≥ 2 rock samples are collected from different parts of the outcrop to test for slumping. Sample BL07 is an example of a failed test with two distinct sample groupings and does not provide reproducible data. - 42 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Can we improve the bouguer gravity resolution in the Cuyuna Range? Increasing gravity measurements in a region of high gravity station density. HIRSCH, Aaron1 1 Minnesota Geological Survey, University of Minnesota, 2609 Territorial Road, St. Paul MN 55114 In East-central Minnesota, the Cuyuna-Penokean orogen is made up of deformed Precambrian rocks of the Penokean-Fold-Thrust belt and adjacent terranes. This complexly folded and thermally overprinted region hosts the 2nd largest known manganese occurrences in the US (Cannon et al., 2017) and has been mined intermittently since the early 1900s. Despite decades of mining, mapping the geology is difficult with most of the bedrock overlain by thick glacial sediments from multiple glacial advances. Mapping of this critical resource and the surrounding region has relied on very limited outcrops, historical mining records, drill core, and geophysical datasets. The Minnesota Geological Survey (MGS) houses state-wide aeromagnetic, gravity, and rock property geophysical datasets that are a key tool in mapping the bedrock geology. The MGS gravity database consists of over 60,000 variably spaced measurements (Chandler et al., 2010). In the Cuyuna-Penokean area, specifically the areas around the Emily District, North Range, and parts of the South Range, the average gravity measurement spacing is ~1.6km with select areas at 0.8-1km. Station spacing of this density is generally considered very good coverage for regional geologic modeling. Due to the complex geology of the area, the MGS set out to determine if increased gravity data will further improve the geophysical resolution and subsequent geologic mapping. Over three field seasons, as part of an Earth Mapping Resources Initiative (Earth MRI) funded project, 210 new gravity points were measured, processed, and added to the gravity database. Gravity stations were tied to an existing base station, and three new field base stations were created in the area to reduce gravity loops. Measurements were prioritized along five transects perpendicular to structure: 2 North-South and 3 Northwest-Southeast profiles. Due to the varying age and accuracy of the gravity database and base stations, tie-point measurements were made at existing gravity station locations for comparison and if any corrections were needed. Multiple comparisons were made between the original and updated datasets with raw 2D Bouguer gravity profiles and gridded Bouguer gravity and second vertical derivatives analyzed (Blakely, 1996). An increase in gravity measurement density resulted in variable differences along profiles resulted in less smoothing and small shifts in slope in some regions but little to no difference in others. Both Bouguer and 2nd vertical derivative gravity grids showed significantly less variability due to inherent smoothing from the minimum curvature gridding process. Two-dimensional modeling was also performed to assess the impact of increased gravity measurement density to geologic mapping. REFERENCES Blakely, R. J.,1996, Potential Theory in Gravity and Magnetic Applications (441 p.). Cambridge: Cambridge University Press. Cannon, W.F., Kimball, B.E., and Corathers, L.A., 2017, Manganese, in chapter L of Schulz, K.J., DeYounge, J.H., Jr., Seal, R.R., II, and Bradley, D.C., eds., Critical mineral resources of the United States – Economic and environmental geology and prospects for future supply: USGS Professional Paper 1802, p. L1–L28. Chandler, V. W., Lively, R. S., and Wahl, T. E., 2010, Gravity and Aeromagnetic Data Grids of Minnesota, Minnesota Geological Survey, http://purl.umn.edu/92939 - 43 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Bouguer gravity map of the Emily District, North Range, and South Range. Black dots are the existing gravity stations. Triangles are the new gravity stations. Circle in bottom left corner is the base station used for this study. Gravity values range from -14.9 - -67.8 mGals. - 44 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Using epidote and chlorite mineral chemistry to extend the alteration footprint around the Hemlo Au deposit, N. Ontario HOLLINGS, Pete1, VRZOVSKI, Joseph1, COOKE, David2, and GORNER, Emily1 1 2 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, P7B 5E1, Canada CODES, University of Tasmania, Private Bag 79, 7001, Hobart, Australia The Hemlo deposit is a world class Archean Au deposit situated in Northern Ontario, Canada with historic production of >21 Moz of Au over 35 years of continuous operation. The deposit has a strike length of ~3 km with a well-documented alteration footprint surrounding mineralization. LA-ICPMS analyses of epidote, chlorite and pyrite from within and surrounding the deposit (Fig. 1) have identified major and trace element variations in mineral chemistry that allow for the discrimination of deposit-proximal and deposit-distal signatures. Epidote compositions vary with distance from Hemlo, with the highest concentrations of As and Sb in epidote proximal to the mineralized zones. Anomalous trace element compositions in epidote can be detected up to 1.5 km further than the mapped alteration footprint. Chlorite also displayed variation in trace elements with deposit-proximal chlorite displaying exponentially higher Ti/Sr and V/Co values than deposit-distal and intrusion-related chlorite. The Ti/Sr ratio for chlorite expanded the geochemical footprint of the Hemlo deposit by up to 1 km. Pyrite displayed anomalous enrichments in a number of elements, with Au, Te and As proving to be the most effective pathfinder elements in pyrite as they were detected at anomalous concentrations up to 2.5 km from the deposit. Several post-mineralization intrusions that surround the deposit were evaluated using epidote and chlorite chemistry to assess whether they generated any false positive geochemical anomalies. The distal post-mineralization intrusions have epidote with consistently low As and Sb concentrations and elevated Fe/Al values relative to deposit-related epidote and can be easily distinguished. Intrusion- Figure 1. Location of samples collected for this study - 45 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 related chlorite displayed low Ti/Sr and V/Co values relative to the deposit chlorite and was also found to be more enriched in Fe relative to deposit-proximal chlorite. These results indicate that the postmineralization intrusions did not produce false positive mineral chemistry anomalies. Variations in chlorite Fe-Mg content can be tracked spectrally using the position of the diagnostic 2250 nm absorption feature. Chlorite displays a range of wavelengths from 2240 – 2256 nm throughout the Hemlo district. Chlorite with lower wavelengths (< 2248 nm) display lower average Fe/Mg (<1) values whereas chlorite with longer wavelengths (> 2252 nm) display higher Fe/ Mg (>1) values. Spectral variations 1550 nm absorption feature of epidote can be used to track compositional variations between the Fe-(epidote) and Al-(clinozoisite) epidote group endmembers. Epidote throughout the Hemlo area displayed a range of wavelengths from 1540 – 1564 nm. These variations in spectral features of epidote could be correlated to epidote major element variations with wavelengths > 1550 nm having on average lower Fe/Al values (< 0.8), whereas wavelengths < 1448 nm displayed average Fe/Al values of ~1. The systematic variations in syn-mineralisation epidote and chlorite compositions around Hemlo suggests that methods developed for investigating geochemical footprints defined by green rock alteration around porphyry systems may also be applicable to Archean orogenic gold deposits. - 46 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Petrographic Study of Granular Iron Formation in the Gunflint Formation: Evidence for WellOxygenated Surface Waters JONSSON, Justin1 and LI, Zhiquan2 Ontario Geological Survey, Ministry of Energy and Mines, Suite B002, 435 James St. South Thunder Bay, ON P7E 6S7 Canada 1 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Granular iron formation (GIF) exhibits distinct features compared to banded iron formation (BIF), being characterized by granule-rich textures and commonly interpreted as detrital, with some grains derived from sedimentary reworking of iron-rich clays, mudstones, arenites, and even stromatolites. Other granules consist of concentric hematite cortices that likely precipitated from Fe(II)-rich waters upon interaction with oxygenated shallow seawater. Previous studies have demonstrated that GIF provides valuable insights into shallow marine environments, as physical energy from waves, tides, and storms is largely restricted to depths above the storm wave base. The 1.88 Ga Gunflint Formation comprises both BIF and GIF, along with chert, carbonates, and minor siliciclastic materials, deposited on a storm-dominated continental shelf. In this study, we examine the petrography of GIF from the lower Gunflint Formation to identify evidence for redox variations in a shallow marine environment. Thin sections of the GIF commonly exhibit oolitic textures, with subordinate peloids and oncoids. Ooids and oncoids are typically composed of hematite, whereas peloids commonly consist of a chert core with hematite rims. Most granules display well-developed concentric hematite cortices, suggesting that iron oxides were directly precipitated from an Fe(II)-rich water column. The grains are not uniformly in contact with one another; instead, many appear to be suspended within the matrix, indicating co-deposition of granules with silica gel. Approximately 30% of the matrix consists of carbonate material, which is randomly distributed within the chert matrix. Hematite grains in the GIF exhibit platy to needle-like morphologies, with grain sizes generally less than 15 µm. Most grains fall within the 1–5 µm range, suggesting an authigenic origin. Some ooids contain manganese carbonates within their inner rims, similar to those observed in the matrix, indicating Mn enrichment in bottom sediments. Our findings suggest that during the early depositional stage of the Gunflint Formation, bottom sediments of the surface water were enriched in Mn, indicating that surface waters were sufficiently oxidizing to promote the precipitation of Mn oxides. However, subsequent early burial of organic matter may have facilitated Mn reduction. Importantly, redox conditions in the shallow marine environment appear to have been oxidizing enough to preserve Fe oxides, but not sufficiently oxidizing to retain Mn oxides. The occurrence of bacterial reduction suggests an increase in organic carbon burial during this time, potentially associated with enhanced primary productivity; however, further investigation is required. - 47 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Interactive Geospatial Geoheritage: Efforts to Support Place-based Exploration and Digitally Preserve Keweenaw’s Geoheritage LIZZADRO-MCPHERSON, Daniel J.1, VYE, Erika C.1, 2, DeGRAFF, James M.2, and ROSE, William I.2 The Great Lakes Research Center, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 USA 1 Department of Geological and Mining Engineering Sciences, Michigan Technological University, 630 Dow Environmental Sciences, 1400 Townsend Drive, Houghton, MI 49931 USA 2 The Keweenaw Peninsula, renowned for many superlatives – world’s largest native copper deposit, first major industrial mining complex in the United States – continues to inspire scientists, historians, and the general public. Ongoing geoheritage efforts enable these groups to explore the deep connections between the underlying geology, landforms, mining industry, and the people working and living on this land for over a millennia. Geoheritage uses a structured approach to identify, manage, and protect geosites and areas with geologic features of significant scientific, educational, cultural, or aesthetic value. Grassroots efforts, spearheaded by Bill Rose, have raised awareness and elevated the prestige of Keweenaw Geoheritage on the global stage despite lacking any formal designation. Bill’s efforts with others to create the first U.S. Geoheritage Park is still in the development stage, while other efforts led by Michigan Technological University (MTU) personnel are helping to bring Bill’s dream to fruition through two geospatial projects: 1) the Keweenaw Geoheritage Geodatabase and companion webGIS-viewer; and 2) Preservation, Indexing, and Enhanced Utility of Historic Copper Mining Drill Hole Records. The Keweenaw Geoheritage geodatabase and webGIS-viewer serve as a living atlas designed to facilitate ways of understanding relationships people hold with the Keweenaw’s geology. The publicly accessible interactive map explores how geology influences education, conservation, and sustainable economic development initiatives in the region (Fig. 1). Each geosite provides a) a brief description of how the site contributes to Keweenaw’s Geoheritage, b) a 360-view, and c) a description of the scientific, educational, cultural, economic, and aesthetic significance of the site (Lizzadro-McPherson & Vye, 2024). This effort supports the co-stewardship of cultural heritage, restoration of legacy mining sites, conservation issues, and the development of economic opportunities based on the region’s globally significant geology. The diamond drill hole (DDH) project aims to digitally preserve at-risk paper core logs, map DDH locations and details, and produce a robust database with a webGIS-based finding aid. The DDH core records document the more recent history of exploratory drilling by the copper mining industry (1899-1970) and contain information still relevant to geological research and exploration for critical minerals. The inventory of records is a tabular database of transcriptions of down-hole data from each scanned core log. An interactive webmap-based finding aid with PDF records and tables of interval descriptions on an open access data portal is in development. These innovative, interactive, geospatial resources aim to enhance scientific inquiry and broaden public engagement and exploration of Keweenaw’s iconic geologic landscape. REFERENCES Lizzadro-McPherson, D.J., and Vye, E.C. (2024). Keweenaw Geoheritage Geodatabase. Michigan State Geological Survey; U.S. Geological Survey, National Cooperative Geologic Mapping Program (Award #G23AC00285 FY23). - 48 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Fig. 2: Diamond drill hole record (left) and mapped surface location with metadata for Suffolk Exploration drilling campaign (right). Fig. 1: Keweenaw Geoheritage Viewer with pop-up displaying the core geoheritage values of the geosite at Great Sand Bay, Keweenaw County, MI. Fig. 2: Diamond drill hole record (left) and mapped surface location with metadata for Suffolk Exploration drilling campaign (right). - 49 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Implications of recent geochronology on the regional geology and timing of gold mineralization in the Red Lake greenstone belt, Ontario MACDONALD, Peter1, HASTIE, Evan1, MALEGUS, Paul2, KAMO, Sandra3, HAMILTON, Mike3 and MARSH, Jeff4 Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, 933 Ramsey Lake Rd, Sudbury, ON P3E 6B5, Canada 1 2 Resident Geologist Program, Ontario Geological Survey, 227 Howey St, Red Lake, ON P0V 2M0, Canada Jack Satterly Geochronology Laboratory, Department of Earth Sciences, University of Toronto, 22 Ursula Franklin St, Toronto, ON, M5S 3B1, Canada 3 Mineral Exploration Research Centre, Harquail School of Earth Sciences, Laurentian University, 935 Ramsey Lake Rd, Sudbury, ON P3E 2C6, Canada 4 As part of the Ontario Geological Survey’s Red Lake bedrock mapping compilation project, geochronology samples were collected from the Red Lake gold camp to improve the ages of volcanic assemblages, sedimentary units and intrusive suites. Eighteen samples were analyzed using ID-TIMS and LA‑ICP‑MS uranium/lead methods on zircon grains. The new ages suggest significant revisions to the geographic presence and/or stratigraphy of the Balmer, Ball, Trout Bay and Confederation assemblages; as well as expanding the regional presence of the Huston conglomerates and identifying the presence of English River terrane sedimentation in the Uchi Subprovince. Newly dated intrusions from throughout the belt refine the timing of synvolcanic, syntectonic, and post‑tectonic magmatism, along with improving the known timing of early gold mineralization and later remobilization. Geochronology from the LP Fault highlights a sequence of felsic and porphyritic intrusive magmatism that is coeval with known gold mineralizing events in the main camp. - 50 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Quantitative analysis of iron mineral composition and crystal sizes in the contact metamorphosed Biwabik iron formation and the Bald Eagle intrusion, NE, MN, USA. MARIN LÓPEZ, Valentina1, BRENGMAN, Latisha1, EYSTER, Athena,2 MITCHELL, Jennifer3, PU, Xiaofei4, MANGUM, John4, and WALKER, Patrick4 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Heller Hall, 1114 Kirby Drive, Duluth, MN 55812, USA 1 2 Department of Earth and Climate Sciences, Tufts University, Lane Hall, 2 North Hill Road, Medford, MA 02155, USA Characterization Facility and the Department of Earth and Environmental Science, University of Minnesota, Twin Cities, S-104 John T. Tate Hall, 116 Church Street Se, Minneapolis, MN 55455, USA 3 4 The National Laboratory of the Rockies, 15013 Denver West Parkway, Golden, CO 80401 Integrated experimental, theoretical, and field data demonstrate the potential viability of hydrogen production via subsurface fluid-rock interaction in systems with significant ferrous iron content (Mayhew et al., 2018; Ellison et al., 2021; Geymond et al., 2022; 2023; 2025; Templeton et al., 2024). As olivine is a key mineral of interest for hydrogen generation either through natural water-rock interaction, or engineered production, we focus on quantifying mineral compositions, crystal size distributions, and modal mineralogy in lithologic units from northeast Minnesota to enable future quantification of hydrogen production feasibility. Units of focus are the troctolitic portion of the Bald Eagle Intrusion (BEI; drill core LOD-6, n=16 samples), and the olivine-rich contact metamorphosed Biwabik iron formation (drill cores 8041 and 8016, n=12 and 13 samples respectively). Olivine and serpentine crystal size distributions (CSD) were quantified using image-based analysis. Combining 2D CSD measurements from BEI depths 970, 1091, and 1212.5 feet (n = 407 crystals from 3 samples; Figure 1A) yielded 8.0% partially serpentinized olivine, and 35.2% fully serpentinized olivine, with the remaining 56.8% of the sample composed of plagioclase, oxides, and minor phases external to olivine crystals. Olivine compositions (Fo76) are similar across BEI samples from multiple depths. In addition to olivine, BEI samples contain pyroxenes, labradorite, and titanium-bearing magnetite and ilmenite, with serpentine-group minerals present along key fracture sets. Reaction boundaries between olivine and serpentine were observed using transmission electron microscopy (TEM; Figure 1B, C). Serpentines are either amorphous or nano-crystalline with variations in crystallinity dependent on orientation in the fracture. Banding was observed in both Focused Ion Beam sections within serpentines proximal to olivine edges (Figure 1C). In metamorphosed Biwabik iron formation samples, olivine compositions are iron-rich (Fo12). In addition to olivine, meta-iron-formation samples contain quartz, oxides, sulfides, pyroxenes, amphiboles, chlorite, mica, calcite, with minor amounts of garnet, plagioclase, serpentine, and accessory phases. CSD analysis of metamorphosed iron formation sample 8016-271 (n = 190 crystals from 1 sample) yielded 55.3% olivine. Next steps include comparison of 2D CSD analyses with 3D X-ray computed tomography data. Overall, the presence of abundant olivine indicates the area could be of interest for future hydrogen generation. To quantify hydrogen generation potential, heterogeneity between serpentinized and un-serpentinized zones should be quantified to extend data from the mineral to the intrusion and formation scale, in addition to connecting hydrologic, geomechanical, and geochemical parameters to mineral data. Next steps include workflow modifications to improve scalability, and application of the workflow to the contact metamorphism Biwabik iron formation. - 51 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1. Image-based CSD analysis and Transmission electron microscopy images for sample LOD-6-1212.5. A) Traces of olivine and serpentine crystals in thin section with mineral proportions calculated using CSD analysis after Higgins, 2000. B) STEM image of reaction boundary between olivine and serpentine. C) TEM image of serpentine and olivine boundary. Top right diffraction pattern of serpentine with a green oval around planes (001) and (002) where the brightest area shows direction of growth of serpentine. Green arrows show crystal orientation. Bottom right diffraction pattern of olivine. REFERENCES Ellison, E. T., Templeton, A. S., Zeigler, S. D., Mayhew, L. E., Kelemen, P. B., Matter, J. M., et al. (2021). Low-temperature hydrogen formation during aqueous alteration of serpentinized peridotite in the Samail ophiolite. J. Geophys. Res. Solid Earth 126, e2021JB021981. doi:10.1029/2021JB021981. Geymond, U., Briolet, T., Combaudon, V., Sissmann, O., Martinez, I., Duttine, M., & Moretti, I. (2023). Reassessing the role of magnetite during natural hydrogen generation. Frontiers in Earth Science (Lausanne), 11. https://doi.org/10.3389/ feart.2023.1169356 Geymond, U., Truche, L., Sissmann, O., Kubániová, D., Recham, N., & Martinez, I. (2025). Mineralogical changes and H2 generation yield during hydrothermal alteration of a magnetite-siderite assemblage. Journal of Geophysical Research: Solid Earth, 130, e2024JB030724. https://doi.org/10.1029/2024JB030724 Higgins, M. (2000). Measurement of crystal size distributions. American Mineralogist , 85 (9): 1105–1116. https://doi. org/10.2138/am-2000-8-901 Mayhew, L. E., Ellison, E. T., Miller, H. M., Kelemen, P. B., and Templeton, A. S. (2018). Iron transformations during low temperature alteration of variably serpentinized rocks from the Samail ophiolite, Oman. Geochimica Cosmochimica Acta 222, 704–728. doi:10.1016/j.gca.2017.11.023 Templeton, A. S., Ellison, E. T., Kelemen, P. B., Leong, J., Boyd, E. S., Colman, D. R., & Matter, J. M. (2024). Low-temperature hydrogen production and consumption in partially-hydrated peridotites in Oman: implications for stimulated geological hydrogen production. Frontiers in Geochemistry, 2. https://doi.org/10.3389/fgeoc.2024.1366268. - 52 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Models of the regional gravity and magnetic anomalies associated with the Nipigon Embayment NITESCU, Bogdan1, TORRES, David Santiago1, and GAONA, Jorge Mario1 1 Department of Geosciences, Universidad de los Andes, Cra. 1 Nº 18A - 12 Bogotá, Colombia The Nipigon Embayment, a region dominated by Proterozoic rocks around Lake Nipigon, extends northward for approx. 150 km into the Superior craton from the Nipigon/Thunder Bay region on the northern shore of Lake Superior. The Embayment is characterized by the presence of intruded maficultramafic rocks and diabase sills dating from the early magmatic stage of Keweenawan rifting in Lake Superior (Heaman et al., 2007). The relationship between the Nipigon Embayment and the MCR has long been a topic of scientific investigation. Various researchers proposed that the Nipigon Embayment represents a viable candidate for a possible third branch of the MCR system (e.g., Hinze and Chandler, 2020), based on various lines of evidence, such as the existence of mafic-ultramafic igneous rocks in the upper crust with geochemical and geochronological similarities to the MCR rocks (e.g., Heaman et al. 2007; Hollings et al., 2007), and the anomalous upper mantle beneath the region reflected in weak seismic anisotropy (Ola et al., 2016), low velocity (Frederiksen et al., 2007; 2013; Foster et al., 2020), and electrical resistivity (Ferguson et al., 2005). However, some investigators suggest that the Nipigon Embayment is related to pre-existing structures, arguing against this region representing a third branch of the MCR due to its lack of Keweenawan extensional features (e.g., Hart and MacDonald, 2007). In this contribution, the gravity and magnetic regional anomalies associated with parts of the Nipigon Embayment are evaluated, both qualitatively, using various filters, and quantitatively, using 2.5D forward modelling. The positive mass anomalies that account for the regional gravity highs in the area covered by the Nipigon sills are equivocal and could be related either to Nipigon magmatic rocks or to covered older rocks bodies, such as Archean mafic-ultramafic intrusions and greenstone belts. If it is assumed that some of these anomalies are related to the Nipigon magmatic rocks, then the gravity models suggest the existence of structures that may have acted as feeders for the emplacement of the Nipigon Embayment mafic-ultramafic intrusive bodies and diabase sills. The Figure 1: 2.5D forward model of the Bouguer gravity anomaly along an W-E profile in the northern part of the Nipigon Embayment, assuming that the cause of the anomaly is related to Nipigon magmatic rocks. Density values: Nipigon magmatic rocks 2.89 g/cc; greenstone belt mafic rocks 2.9 g/cc; granitoid and tonalite 2.64 g/cc; background 2.67 g/cc. - 53 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 magnetic models of the regional magnetic anomalies indicate the presence of a significant subsurface volume of highly magnetic rocks within the Nipigon Embayment crust. These results are compatible with the interpretation of this region as a segment of the crust affected by magmatism in the initiation stage of the MCR, possibly as an incipient, undeveloped part of the rift controlled by pre-existing structures. REFERENCES Ferguson, I.J., Craven, J.A., Kurtz, R.D., Boerner, D.E., Bailey, R.C., Wu, X., Orellana, M.R., Spratt, J., Wennberg, G., Norton, M., 2005. Geoelectric response of Archean lithosphere in the western Superior Province, central Canada. Phy. Earth Planet Int. 150, 123–142. https://doi.org/10.1016/j.pepi.2004.08.025 Foster, A., Darbyshire, F., Schaeffer, A., 2020. Anisotropic structure of the central North American Craton surrounding the Mid-Continent Rift: Evidence from Rayleigh waves. Prec. Res. 342, 105662. https://doi.org/10.1016/j. precamres.2020.105662. Frederiksen, A.W., Miong, S.K., Darbyshire, F.A., Eaton, D.W., Rondenay, S., Sol, S., 2007. Lithospheric variations across the Superior Province Ontario, Canada: Evidence from tomography and shear wave splitting. J. Geophys. Res-Earth 112, 1–20. https://doi.org/10.1029/2006JB004861. Frederiksen, A.W., Bollmann, T., Darbyshire, F., van der Lee, S., 2013. Modification of continental lithosphere by tectonic processes: A tomographic image of central North America. J. Geophys. Res-Earth 118, 1051–1066. https://doi. org/10.1002/jgrb.50060. Hart, T.R., 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. Can. J. Earth Sci. 44, 1021–1040. https://doi.org/10.1139/e07-026. Heaman, L.M., Easton, R.M., Hart, T., MacDonald, C.A., Hollings, P., Smyk, M., 2007. Further refinement to the timing of Mesoproterozoic magmatism Lake Nipigon region, Ontario. Can. J. Earth Sci. 44, 1055–1086. https://doi. org/10.1139/e06-117. Hinze, W.J., Chandler, V.W., 2020. Reviewing the configuration and extent of the Midcontinent rift system. Prec.Res. 342, 105688. https://doi.org/10.1016/j.precamres.2020.105688. Hollings, P., Hart, T., Richardson, A., MacDonald, C.A., 2007a. Geochemistry of the Mesoproterozoic intrusive rocks of the Nipigon Embayment, northwestern Ontario: Evaluating the earliest phases of rift development. Can. J. Earth Sci. 44, 1087–1110. https://doi.org/10.1139/e06-127. Ola, O., Frederiksen, A.W., Bollmann, T., van der Lee, S., Darbyshire, F., Wolin, E., Revenaugh, J., Stein, C., Stein, S., Wysession, M., 2016. Anisotropic zonation in the lithosphere of Central North America: Influence of a strong cratonic lithosphere on the Mid-Continent Rift. Tectonophysics 683, 367–381. https://doi.org/10.1016/j.tecto.2016.06.031. - 54 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Origin of the World-Class Eagle, Eagle East, and Tamarack Ni-Cu-PGE Deposits and comparative analysis with other Midcontinent Rift- and Siberian Trap-related intrusions NOWAK, Robert1, DEERING, Chad1 , and ESSIG, Espree1 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931, USA 1 The 1.1 Ga Mesoproterozoic Midcontinent rift hosts the Eagle, Eagle East, and Tamarack Ni-CuPGE deposits and Embayment Prospect. These deposits are hosted by ultramafic igneous rocks and have some of the highest Ni-Cu grades on Earth. We use new bulk-rock data and published datasets (bulk-rock, mineral chemistry, and isotopic analyses) to examine major, minor, and trace element trends of both Midcontinent rift-related alkaline and tholeiitic intrusions (Nowak et al., 2025). In addition, we compare the geochemical data to local kimberlite-hosted lower-crustal xenoliths and local igneous (Archean) and sedimentary (Paleoproterozoic) country rocks. We found the peridotite magma compositions dominantly consist of primitive mantle compositions with varying abundances of subduction-related components, alkaline-transitional melts, and local country rock contaminates (e.g., Baraga and Animikie Basin sediments). The subduction-related components are interpreted to be derived from previous Archean and Paleoproterozoic subduction events and likely hosted within the sub-continental lithospheric mantle. Importantly, these subduction-related components are also interpreted to have acted as oxidizing agents within the melt, stabilizing sulfate (+2 FMQ (fayalite–magnetite–quartz) to FMQ) while inhibiting sulfide crystallization as the magma ascended through ~50 km of the Superior craton. This study largely corroborates the previous findings with respect to the contribution of local country rock contamination to the Eagle–Tamarack peridotite host rocks, which is estimated to be minimal (<5%). However, the incorporation of <5% reductive pelitic siltstone contamination results in strong shifts in the oxygen fugacity of the peridotite melt, from +2 FMQ to slightly below FMQ, as determined from spinel Fe3+/∑Fe ratios (Figure 1). This shift in oxygen fugacity resulted in the transition from total sulfate (+2 FMQ) to sulfate + sulfide (<+2 FMQ to FMQ) to total sulfide (<FMQ). This shift in oxygen fugacity is a key contributor to the formation of Ni-Cu-PGE-rich massive sulfides within the Eagle peridotite. This study presents an expanded geochemical interpretation for the exploration of Midcontinent rift-related Ni-Cu-PGE deposits to include peridotites with subduction-like signatures and contaminated via <5% reductive sedimentary country rocks. Based on these findings, we also comparatively analyze geochemical samples from Figure 1: Downhole profiles of drillhole 03EA034 of Fe3+/∑Fe ratios of spinel (Ding et al., 2010); with oxygen fugacity estimates (relative to FMQ) this study), Ni, Cu, and S (all in wt%; this study), and the relative proportion (%) of compositions (eclogite, amphibolite, subducted sediment, alkaline-transitional, and Baraga Basin sediments) used to reconstruct the multielement compositions of Eagle peridotite. Analytical error (accuracy; 1σ) is estimated to be smaller than the symbol size. - 55 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Midcontinent rift-related prospective intrusions and Siberian-Trap-related intrusions in order to better determine economic vs. subeconomic host rock signatures. REFERENCES Ding, X., Li, C., Ripley, E.M., Rossell, D., Kamo, S., 2010, The Eagle and East Eagle sulfide ore-bearing mafic-ultramafic intrusions in the Midcontinent Rift System, upper Michigan. Geochronology and petrologic evolution. G3 Geochem. Geophys. Geosyst., 11, p.1-22. Nowak, R., Deering, C., and Essig, E., 2025, Origin of the World-Class Eagle, Eagle East, and Tamarack Ni-Cu-PGE Deposits. Minerals, 15, 871. https://doi.org/10.3390/min15080871 - 56 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 BEDROCK GEOLOGY OF THE ERICSBURG NW, ERICSBURG NE, RAY SW, AND RAY SE QUADRANGLES, ST. LOUIS AND KOOCHICHING COUNTIES, MINNESOTA NOWARIAK, Eric and SEVERSON, Allison Minnesota Geological Survey, University of Minnesota – Twin Cities, 2609 Territorial Road St. Paul, MN, USA New geologic mapping presented here portrays the Precambrian bedrock geology and tectonic history of the axial zone of the Quetico subprovince across four 7.5’ quadrangles in portions of eastcentral Koochiching County and far western St. Louis County, Minnesota. The map records the Neoarchean deposition, deformation, metamorphism, and migmatization of turbiditic sediments, along with the intrusion of the granitic rocks of the Vermilion Granitic Complex during the accretion of the Wawa subprovince to the southern margin of the Superior Province, and continuing through the intrusion of the Paleoproterozoic Fort Frances dike swarm. The metasedimentary rocks of the Quetico subprovince, now predominantly biotite schist, granofels, and migmatite, have been subject to at least four successive contractional and transpressional deformation styles documented in the map area. The map pattern and dominant structural grain of bedrock is controlled by structures associated with D2 and D3 deformation. D2 deformation produced map-scale, tight to isoclinal F2 folds with well-developed ENE-WSW-striking subvertical S2 axial-planar foliation. D3 deformation coincides with the development of ENE- and NW-trending ductile shear zones with dextral motion. F3 folds are coaxial to F2 folds and manifest as isoclinal refolds and reorientations of D2 structures. D4 deformation post-dates the dominant D2 and D3 deformational events and is represented by steeply plunging broad, open folds and NNW-trending fault and fracture zones. New geochemical analyses illustrate rocks of the Vermilion Granitic Complex are generally calkalkaline, weakly peraluminous to metaluminous, magnesian granitoids with minor amphibole-rich dioritic to gabbroic rocks. Based on geological and geochemical features, the Vermilion Granitic Complex can be subdivided into groups with distinct lithologies, geochemistry, and magma sources. Tonalites, trondhjemites, and granodiorites (TTG) of the Early Magmatic Suite are distinctly more sodic than the younger Lac La Croix Suite granitoids. Compared to the Early Magmatic Suite, Lac La Croix Suite granitoids are relatively more alkalic, more aluminous, and have steeper REE profiles. Quetico metasedimentary rocks and the Vermilion Granitic Complex have been subject to at least two metamorphic events recording the burial, uplift, and intrusive history of the subprovince. M1 metamorphism is likely contemporaneous with D2 and early D3 deformation, based on the presence of syn- and post-kinematic, inclusion-rich porphyroblasts. Peak metamorphic conditions reached amphibolite facies during M1 and have been constrained to 525-575°C with pressures exceeding 6 kbars based on phase equilibrium modeling outside the thermal influence of the Lac La Croix Suite. Proximal to the Lac La Croix granite, M1 metamorphic features have been overprinted by a hightemperature, low-pressure event, M2, presumably due to the intrusion of voluminous granitoids of the Lac La Croix Suite during the waning stages of D3 deformation. M2 metamorphism manifests as inclusion-poor garnet, sillimanite-, cordierite-, and andalusite-bearing assemblages in metasediments and granitic orthogneisses. - 57 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1. A. Plutonic rock classification of igneous rocks in this study, after Enrique and Esteve, 2019. B. 2ACNK (2* molar Al2O5/(CaO+Na2O+K2O), Na2O/K2O, 2 FMSB (2*(FeOtot+MgO)wt.%*(Sr+Ba)wt.%) source identification diagram, of Laurent and others (2014). Fields for TTG (T), continental or C-type (C), and metasomatized mantle or M-type (M) granitoids from Moyen (2019) have been added. C. Alumina Saturation plot after Barton and Young (2002) for all intrusive units within the map area. - 58 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Petrographic, geochemical, and mineralogical analyses of manganiferous iron formations and associated lithologies at the Cuyuna Range, central Minnesota PALIEWICZ, Cory1, POST, Sara1, and THAKURTA, Joyashish1 Natural Resources Research Institute (NRRI), University of Minnesota Duluth, 5013 Miller Trunk Hwy, Duluth, MN 55811 USA 1 The Paleoproterozoic Cuyuna Range of central Minnesota hosts one of two significant manganese deposits in the United States and contains anomalously high manganese concentrations (up to ~50 Wt. % Mn) when compared to other Banded Iron Formations in the Lake Superior region (Cannon et al., 2017). The area was highly deformed and metamorphosed during the Penokean Orogeny and encompasses the Emily District at edge of the Animikie Basin to the north, and the North and South ranges which occur within the older fold and thrust belt to the south (Boerboom and Chandler, 2004; Southwick et al., 1988; Morey, 1990). Although the area has a rich history of iron mining and ongoing manganese exploration, many questions remain regarding the occurrence, nature, and mechanisms of manganese mineralization. This work includes new petrographic, lithogeochemical, and mineralogical data collected and analyzed from 201 drill core samples from 37 drill holes across the Emily District, North Range, and South Range (Figure 1). The regional pilot study is part of a larger USGS Earth Mapping Resources Initiative to map and better-constrain the mineral potential of the region. We emphasize the lithologic variability of mineralized iron formations throughout the range, but especially within the Emily District, which from past studies is known to be most-enriched in Mn-content. Cuyuna iron formations generally range from cherty to slaty (thick bedded to thin bedded / granular to non-granular) with manganiferous units extending from enriched (5-10% Mn), manganiferous (>10 % Mn) and highly manganiferous (>35 % Mn). Although textural and mineralogical differences of mineralized units vary widely with increasing grade, the variability and significance of non-enriched lithologies throughout the Cuyuna Range also offer insights regarding possible sources or mechanisms of mineralization, especially when taken within the context of recently integrated historic drill logs and prior works (e.g., McSwiggen et al., 1995). Textural and mineralogical variation among mineralized units exhibit many signs of hydrothermal modification during manganese enrichment. Many grains have been replaced with manganese oxides and hydroxides in both cherty and slaty iron formations and the occurrence of vugs associated with other hydrothermal accessory minerals such as carbonates, epidote, micas, and clays, along with abundant sieved and altered grains indicate that many pulses of variable hydrothermal activity likely resulted in disequilibrium of most preserved mineral assemblages. Non-mineralized units such as graywackes are typically highly altered to sericite and kaolinite and pyritic graphitic argillites have been observed to exhibit non isochemical characteristics illustrating high mobility of Fe and Mn. The occurrence of silicified and oxidized zones as they relate to variations of grade are also characterized along with variability of downhole changes of Mn, Fe, SiO2, Al2O3, and LOI plotted as split logs which also show changes of Co, Cu, and Zn to further assess the possibility of other critical minerals associated with manganese. - 59 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Geologic map after Boerboom and Chandler (2004; 2022) showing drill hole locations sampled in Crow Wing and Aitkin Counties, central Minnesota. REFERENCES Boerboom, T.J., and Chandler, V. W., 2004, Plate 2 - Bedrock Geology, in Setterholm, D. R. Geologic atlas of Crow Wing County, Minnesota, MGS County Geologic Atlas, C-16 Part A, 1:100,000. Boerboom, T.J., and Chandler, V. W., 2022, Plate 2 - Bedrock Geology, in Bauer, et al., 2022. Geologic atlas of Aitkin County, Minnesota, MGS County Geologic Atlas, C-52 Part A, 1:200,000. Cannon, W.F., Kimball, B.E., and Corathers, L.A., 2017, Manganese, in chap. L of Schulz, K.J., DeYoung, J.H., Jr., Seal, R.R., II, and Bradley, D.C., eds., Critical mineral resources of the United States—Economic and environmental geology and prospects for future supply: USGS Professional Paper 1802, p. L1–L28. McSwiggen, P.L., Morey, G.B., and Cleland, J.M., 1995, Iron-formation protolith and genesis, Cuyuna range, Minnesota: Minnesota Geological Survey Report of Investigations 45, 54 p. Morey, G.B., 1990, Geology and manganese resources of the Cuyuna iron range, east-central Minnesota: Minnesota Geological Survey Information Circular 32, 28 p. Southwick, D.L., Morey, G.B., and McSwiggen, P.L., 1988, Geologic map (scale 1:250,000) of the Penokean orogen, central and eastern Minnesota, and accompanying text: Minnesota Geological Survey Report of Investigations 37, 25 p., 1 pl. - 60 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Physical Magmatic System Interpretation of the Marathon Cu-Pd Deposit, Coldwell Complex, Ontario PETERSON, Dean1, STEINER, R. Alex1, SWEET, Gabriel1, and BOUCHER, Chanelle2 1 2 Big Rock Exploration, 2505 West Superior Street, Duluth, MN 55806. Generation PGM Inc., 100 King Street West, Toronto, ON M5X 1B1. The goal of geologic mapping and/or drill core logging in mafic magmatic ore deposits is to not just know what the lithology is at a specific outcrop and/or drill hole interval, but to know with some confidence where you are in the mineralized intrusion, i.e., within the overall magmatic system. Generation Mining (GenM) contracted Big Rock Exploration (BRE) to reevaluate the Coldwell Complex hosted Marathon Cu-Pd deposit using a magmatic system approach. Mineralized mafic intrusions are typically composed of three principal minerals, plagioclase-olivinepyroxene along with various proportions of apatite, Fe-Ti oxides, and Fe-Cu-Ni sulfides. Variations in mineralogic estimates of the three principal minerals by many geologists over decades of time can be the difference between calling a rock an anorthosite, a gabbro, a troctolite, or a peridotite. In deposit areas with decades upon decades of exploration history, these basic lithologic calls by many different geologists can directly influence how a mafic magmatic ore deposit is interpreted and/or modeled. Problems in interpretation can come to the forefront when drill hole intervals are logged strictly by lithology and subsequently digitally assigned a LithCode. Another method of logging and interpreting mineralized mafic intrusions is to approach it from the physical process side, i.e., as a magmatic system. Utilizing a magmatic system approach begins with an understanding of the initial conditions of the system. Initial conditions include the intrusive geometry and Figure 1. Schematic model of the magmatic architecture of the Marathon Cu-Pd deposit. - 61 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 flow paths, the lithology of the footwall, hangingwall and sidewall rocks, and the magmas composition, crystallinity, plagioclase-olivine phenocryst content, trace element signature and sulfide content. In general terms, mafic intrusions have slower moving marginal boundaries, which are commonly xenolithrich, surrounding a faster flowing and xenolith-poor ‘clean’ central core. Magmatic shearing is induced by the differential velocity, from margin to core, in which magmas intrude can lead to pronounced local mineralogical variability in the outcome. For example, phenocryst sorting leads to modal layering, and kinetic sieving processes raises large particles, which can be phenocrysts, autoliths and/or xenoliths, upwards in the intrusion. The rocks formed in these mafic magmatic systems, though largely governed by the initial conditions, locally can vary by associated chemical, thermal, and momentum boundaries. BRE coupled these magmatic first principals with GenM assisted field work and drillhole relogging to reevaluate the Marathon Cu-Pd deposit magmatic system. A schematic model of the interpreted magmatic system at the Marathon Cu-Pd deposit is presented in Figure 1, and stratigraphic profiles depicting the historic lithology-based coding (Lith Codes) and recently proposed magmatic systems approach coding (Unit Codes) is given in Figure 2. This talk will highlight many of BRE’s research findings on the Marathon Cu-Pd deposit magmatic system. Figure 2. The proposed magmatic system approach Unit Codes (left) versus the historic GenM Lith Codes (right) assigned to the rocks of the Marathon Cu-Pd deposit. Rectangular arrows point to how logged lithologies can be assembled into discrete magmatic system units of the Marathon Series. - 62 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Critical Mineral Potential of the Watersmeet Gneiss Dome, MI USA QUIGLEY, Ashley1, MAHIN, Robert1, and GAMET, Nolan1 1 Michigan Geological Survey, 416 Avenue C, Gwinn, MI 49841U.S.A. Precambrian gneisses and schists on the northern margin of the Watersmeet Dome in Michigan’s Upper Peninsula are unusually enriched in rare earth elements, fluorite and incompatible elements including U, Th, Hf, and Zr (Barovich et al., 1991; Sims, 1990). Rocks are mainly Archean gneisses and amphibolites although elevated REEs, fluorite and incompatible elements are associated with a gneiss and schist unit of possible Paleoproterozoic age (Barovich et al., 1991). The area is within Earth Mapping Resources Initiative (EMRI) critical mineral focus areas for both IOCG/ IOA and Magmatic REE deposits (Dicken and others, 2022). The Michigan Geological Survey (MGS) conducted detailed geologic mapping and sampling, as well as collected geophysical and geochronological data. An RS-230 BGO gamma-ray spectrometer was used to take over 600 total gamma (K/U/Th) measurements from outcrops. Additionally, an unmanned aerial vehicle (UAV), high resolution magnetic survey was flown. A previously undescribed magnetic, fine-grained schist comprised 85% of the highest total REE samples (high of 1659 ppm TREE). The schists are associated with magnetite and fluorite and coincide with a kilometer-wide central magnetic anomaly, as well as a three kilometer, roughly east-west trending, sinuous anomaly. In plots, granitoids, gneisses, and schists show three distinct populations. Group 1 clusters in the VAG-syn/COLG field, has no europium anomaly and average 72 ppm TREE. Group 2 is transitional between VAG-syn/ COLG and WPG, has a marked europium depletion, and contains an average of 136 ppm TREE. Group 3 is enriched in REE with an average 711 ppm TREE, plots in the WPG/A-Type granite field and has moderate europium depletion. All three groups are peraluminous. Group 3 rocks include enriched REE magnetic schists, magnetic granitoids, and gneisses all of which are located in proximity to each other as well as to magnetic highs. Highly fractionated, non-peralkaline felsic granites can have geochemical characteristics which overlap those for typical A-type granites Figure 1: Map showing the location of the Watersmeet project area. - 63 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 (Whalen and others, 1987). Some fractionation is indicated in Group 1 and Group 2 rocks by a semicontinuous trend of decreasing Zr, Nb, Ce, and Y. Group 3, however, displays no such evidence of fractionation, which is typical of A-Type granites. Numerous REE, F, Th, and/or U-bearing silicate, oxide and carbonate minerals including fluorite, thorite, pyrochlore, allanite, columbite, parasite, and yttrialite were identified using SEM within alteration halos along fractures and occasionally within veins. Zircons with strong pleochroic halos are common, particularly within biotite grains but also observed with amphiboles. The presence of fluorite and REE bearing-fluorocarbonates indicate that REE enrichment was facilitated, at least in part, by fluorine-rich hydrothermal fluids. Preliminary, unpublished U-Pb zircon geochronology indicate that all units are Archean and the previous proposed Paleoproterozoic ages may represent a thermal resetting event. REFERENCES Barovich, K.M., Patchett, P.J., Peterman, Z.E., and Sims, P.K., 1991. Neodymium Isotopic​Evidence for Early Proterozoic Units in the Watersmeet Gneiss Dome, Northern​Michigan. U.S. Geological Survey Bulletin 1904-G: G1-G7. ​ Dicken, C.L., Woodruff, L.G., Hammarstrom, J.M., and Crocker, K.E., 2022, GIS,​supplemental data table, and references for focus areas of potential domestic resources​of critical minerals and related commodities in the United States and Puerto Rico (ver.2.0, April 2024): U.S. Geological Survey data release, https://doi.org/10.5066/P9DIZ9N8. Pearce, Julian & Harris, Nigel & Tindle, Andrew. (1984). Trace Element Discrimination Diagrams for the Tectonic Interpretation of Granitic Rocks. Journal of Petrology. 25. 956-983. 10.1093/petrology/25.4.956.​ Sims, P.K., 1990, Geologic map of Precambrian rocks, Marenisco, Thayer, and​ Watersmeet 15-minute quadrangles, Gogebic and Ontonagon counties, Michigan, and​ Vilas County, Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map​I-2093, scale 1:62,500.​ Whalen, J.B., Currie, K.L. & Chappell, B.W. A-type granites: geochemical characteristics, discrimination and petrogenesis. Contrib Mineral Petrol 95, 407–419 (1987). https://doi.org/10.1007/BF0040220. - 64 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Current geologic and geophysical research on the Precambrian basement of eastern North Dakota, USA SAINI-EIDUKAT, Bernhardt1, CHITTICK, Steve2, and NESHEIM, Timothy2 1 2 Dept. of Chemistry and Biochemistry, North Dakota State University, Fargo, ND 58102 USA North Dakota Geological Survey, Grand Forks, ND 58202 USA In the entirety of the state of North Dakota, no crystalline basement is exposed due to Phanerozoic sedimentary cover. Regional geophysical mapping, combined with lithological data and radiometric dates, have correlated the Wabigoon and Wawa subprovinces of the Superior Craton into eastern North Dakota (Figure 1). However, understanding of the geologic and the geophysical characteristics of the basement in this region is, with some exceptions, relatively poor compared to many other areas (Figure 2). Figure 1: Map of North Dakota Precambrian geology, RRVD drill core locations, and proposed survey area (black outline). Open symbols: geochronology samples. Base map from Sims et al. (1991). Figure 2: Regional aeromagnetic map, and proposed survey area (blue outline), showing the difference in resolution between ND and MN. (The NE corner of ND does already have higher quality aeromagnetic data.) - 65 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 The North Dakota Geological Survey (NDGS), working with the Earth Mapping Resources Initiative (Earth MRI) of the U.S Geological Survey (USGS) (www.usgs.gov/special-topics/earthmri), and North Dakota State University are undertaking a renewed initiative to obtain high quality geochronologic, geochemical, geophysical, and radiometric data over eastern North Dakota. Depth to basement is on the order of a few hundred meters in eastern ND, but increases to thousands of meters westward underneath the Williston Basin. For that reason, the focus of the initiative is on the eastern region where depth to basement is less than 1000 m. As part of Earth MRI, the USGS is planning to carry out a high-resolution airborne magnetic and radiometric survey in eastern North Dakota, to be flown in 2026-27. The survey will be designed to meet complementary needs related to geologic mapping and mineral resource research. The survey design is being coordinated with the NDGS to provide complete coverage of a region that crosses the boundaries of multiple subprovinces and greenstone belts within the Archean Superior Province. The mineral systems of interest in the survey area include Mafic magmatic, Porphyry Sn, and Metamorphic. Potential critical mineral commodities include Cr, PGE, Au, Co, graphite, REE, Li, Ta, and Sn. There is additional potential for Mn, Ni, Cu, Fe, Mg, and Cs. Samples of drill core from the 1977 Red River scientific drilling project (Moore, 1978; Kelley, 1980; Beaudry et al., 2024, Pereira et al., 2024), and from other cores, will undergo geochemical, geochronological, petrological, and geophysical investigation. Portable XRF analysis for trace elements is underway, as is a gravimetric survey of eastern ND by the NDGS. REFERENCES Beaudry, C., Hess, M., Pereira, C., Saini-Eidukat, B., 2024, Petrology and geochemistry of Precambrian basement rocks in Walsh County, North Dakota. ILSG Abstr. and Proc., v.70, part 1, p. 6-7. Kelley, L.I., 1980, Kaolinitic weathering zone on Precambrian basement rocks, Red River Valley, eastern North Dakota and northwestern Minnesota. M.S. Thesis, University of North Dakota. 85 pp. Moore, W. L., 1978. A preliminary report on the geology of the Red River Valley Drilling Project, eastern North Dakota and northwestern Minnesota: Bendix Field Engineering Company Subcontract H77-059-E, 292p. https://www.osti.gov/ biblio/6538603 doi:10.2172/6538603 Pereira, C., Nesheim, T., Vervoort, J.D., and Saini-Eidukat, B., 2024, Major element geochemistry and first zircon U-Pb age dates of Precambrian basement rocks in eastern North Dakota. ILSG Abstracts and Proceedings, v.70, part 1, p.74-75. Sims, P.K., Peterman, Z.E., Hildenbrand, T.G., and Mahan, S., 1991, Precambrian Basement Map of the Trans-Hudson Orogen and adjacent terranes, northern Great Plains, U.S.A.: USGS Miscellaneous Investigations Series Map, I-2214. DOI: 10.3133/i2214 - 66 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Reassessing variations in metamorphism across the Penokean orogen in Northern Michigan: Part 1, new Pressure-Temperature-Time-Deformation constraints SALERNO, R.1, CANNON, W. F.1, THOMPSON, J. M.2, SOUDERS, A. K.2, VERVOORT J.3, and HILLENBRAND, I.2 1 2 3 U.S. Geological Survey, Reston, VA 20192, USA U.S. Geological Survey, Denver, CO 80225, USA Washington State University, Pullman, WA 99163, USA The Penokean orogeny (1890-1830 Ma) represents the earliest collisional event in a long subduction sequence active throughout the Paleoproterozoic to Mesoproterozoic along Laurentia’s southern margin. Traditionally, spatial variations in metamorphic grade in the Penokean orogenic belt were described as three “nodes” (Fig. 1) and ascribed to the main accretionary phase which ended at 1830 Ma. However, the swath of younger 40Ar/39Ar cooling ages at ~1750 Ma across the terrane suggests later collisional episodes also played an important role in modifying the Penokean orogenic belt (Schneider et al., 1996). This observation, coupled with newly mapped younger structures by recent geophysical surveys, raises questions about which features are truly Penokean in origin, and which reflect later overprinting by younger tectonic events (Drenth et al., 2021). Elucidating the causes and timing of post-Penokean modification of crust in central Laurentia is key for accurately reconstructing the outward growth of proto-North America throughout the Proterozoic. We have used multiple geochronometers and isotope systems to unravel the metamorphic evolution of the Penokean orogenic belt. New geochronology and thermodynamic modeling of metasedimentary rocks reveal variations in the timing of metamorphism and subsequent cooling histories between metamorphic nodes (Fig. 1). Directly adjacent to the Niagara fault zone, rocks in the Peavy node have garnet Lu-Hf ages of 1837±7 Ma, reflecting the age of granulite facies metamorphism in the lower crust. Overlapping garnet Sm-Nd (1830±65 Ma) and apatite U-Pb (1822±28 Ma) ages indicate rapid exhumation of these lower crustal rocks near the end of the Penokean orogeny. In contrast, rocks in the Watersmeet and Republic nodes, located farther inboard from the paleomargin, reflect later lowergrade amphibolite facies regional metamorphism after the end of the Penokean orogeny, from 1825±5 to 1782±15 Ma. Unlike the Peavy node, these samples have offset Lu-Hf and Sm-Nd ages reflecting the different closure temperatures of the two isotope systems in garnet. Dispersed Lu-Hf and Sm-Nd ages indicate prolonged residence of these rocks at mid-crustal depths and correspond with protracted cooling paths of 1-3°C/Mya, until final exhumation began at ~1750 Ma. Our results illustrate that the metamorphic nodes in the Penokean orogenic belt do not reflect the same conditions or cooling histories, and do not all represent the same tectonic event. Instead, our data reveal a sequence where early granulite facies metamorphism and rapid exhumation are linked with the end stages of the Penokean orogeny and are restricted to the belt of high-grade rocks north of the Niagara fault. Regional amphibolite facies metamorphism persisted after, requiring continued crustal thickening following both the accretionary phase of the Penokean orogeny and exhumation of deep crustal rocks. The implications of this are two-fold. 1) The metamorphic nodes in the Penokean orogenic belt are not cogenetic but rather reflect different tectonic events and different times. 2) Post-Penokean regional metamorphism followed by widespread uplift and cooling after ~1750 Ma represent significant modification of the Penokean orogenic belt throughout Geon-17. More broadly, younger overprinting on this terrane reveals that outboard tectonic activity following the Penokean orogeny played a major role in the modification of Paleoproterozoic and Archean crust in central Laurentia. - 67 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Top, generalized geologic map showing metamorphic nodes in the Penokean orogen in northern Michigan and sample locations in our study (map after Tinkham and Marshak, 2004). Bottom, temperature-time diagrams showing cooling histories of garnet-bearing rocks in three metamorphic nodes. Microstructures indicate deformation during uplift at ~1750 Ma proceeded after peak metamorphism. 40Ar/39Ar data are from previous studies and references are compiled in Salerno et al. (2026). REFERENCES Drenth, B.J., Cannon, W.F., Schulz, K.J., and Ayuso, R.A., 2021, Geophysical insights into Paleoproterozoic tectonics along the southern margin of the Superior Province, central Upper Peninsula, Michigan, USA: Precambrian Research, v. 359, doi:10.1016/j.precamres.2021.106205. Salerno, R., Cannon, W.F., Thompson, J.M., Souders, A.K., Vervoort, J., Hillenbrand, I., 2026, Unraveling protracted modification of Archean and Paleoproterozoic crust in central Laurentia, Penokean orogen, with garnet and accessory mineral geochronology and microstructural analysis: Geological Society of America Bulletin, in press. Schneider, D., Holm, D., and Lux, D., 1996, On the origin of Early Proterozoic gneiss domes and metamorphic nodes, northern Michigan: Canadian Journal of Earth Sciences, v. 33, p. 1053–1053, doi:10.1139/e96-080. Tinkham, D.K., and Marshak, S., 2004, Precambrian dome-and-keel structure in the Penokean orogenic belt of northern Michigan, USA, in Whitney, D.L., Teyssier, C., and Siddoway, C.S., eds., Gneiss Domes in Orogeny: Geological Society of America Special Paper, v. 380, p. 321-338, doi:10.1130/0-8137-2380-9.321. - 68 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Whole Rock and Mineral Chemistry of the Eagle’s Nest Intrusion, McFaulds Lake Greenstone Belt, Ontario, Canada: Insights into the Origin and Paragenesis SHESHNEV, Vlad1, HOLLINGS, Pete1, TOLLEY, James1, ANGOMBE, Moses1, DELLER, Matt2, and STERN, Richard3 1 2 3 Department of Geology, Lakehead University, Thunder Bay, Ontario, Canada Wyloo, Thunder Bay, Ontario, Canada Canadian Centre for Isotopic Microanalysis, University of Alberta, Edmonton, Alberta, Canada Orthomagmatic Ni-Cu-(PGE) deposits originate in the mantle, where source composition and degree of partial melting are the first-order controls on composition and metal fertility of the derived magmas (Naldrett, 2011). During ascent, these magmas undergo differentiation, producing more evolved compositions that reflect both the characteristics of the mantle source and subsequent magmatic processes (Barnes, 2023; Smith et al., 2024). The Eagle’s Nest intrusion is a maficultramafic, blade-shaped dike, which is host to the only known economically significant Ni-Cu-(PGE) mineralization within Meso- to Neoarchean McFaulds Lake Greenstone Belt. The Eagle’s Nest is part of the mafic to ultramafic magmatism of the Koper Lake subsuite, of the larger Ring of Fire Intrusive Suite (ca. 2736–2732 Ma; Houlé et al., 2020; Metsaranta and Houlé, 2020). Two different parental magma compositions have been proposed for the Eagle’s Nest intrusion, including a low- and a highMg komatiitic magma, both of which are inconsistent with the observed mineralogy of the intrusion (Mungall et al., 2010; Zuccarelli, 2020). To better understand the origin and nature of the Eagle’s Nest intrusion, this study integrated petrography, whole-rock geochemistry, mineral chemistry, as well as radiogenic and stable isotope systematics. The Eagle’s Nest intrusion can be subdivided into the marginal and inner zones. The marginal zone comprises mafic intrusive rock in contact with the wall rock tonalite, exhibiting the most evolved mineralogical and geochemical characteristics. The marginal zone gradationally transitions into the inner zone, which consists of ortho- to mesocumulate ultramafic rocks with more primitive compositions, reflecting the accumulation of olivine and chromite in cotectic proportions, along with variable amounts of intercumulus silicate phases and interstitial sulfides. Using the whole rock geochemistry of olivine-chromite cotectic cumulate rocks, combined with olivine and chromite mineral chemistry, a new parental magma composition was determined for the Eagle’s Nest intrusion. The new estimate suggests a komatiitic basalt magma that contained ~11 wt% FeOt and ~15 wt% MgO. The new parental magma estimate is more evolved than previously proposed compositions, however, it is consistent with the composition of identified chilled margins, associated mafic dikes, and olivine from the Eagle’s Nest intrusion. Using the newly obtained estimate, the petrographically determined crystallization sequence was recreated at low pressures, suggesting the Eagle’s Nest formed in shallow crustal levels. Whole-rock geochemistry and Sm-Nd isotopes indicate that the Eagle’s Nest magma was derived from a depleted mantle source above the garnet stability field. During transport, this magma underwent crustal contamination by the host tonalite and older supracrustal rocks. Assimilation of sulfur-bearing supracrustal material likely triggered sulfide saturation, supported by the mass-independent fractionation values of the measured Δ³³S. The intrusion’s distinct petrological and metallogenic features likely reflect both the emplacement dynamics and the parental magma composition, resulting in its unique metal endowments within the greenstone belt. REFERENCES Barnes, S.J., 2023. Lithogeochemistry in exploration for intrusion-hosted magmatic Ni-Cu-Co deposits. Geochemistry: Exploration, Environment, Analysis, vol. 23(1), pp. geochem2022–025. Houlé, M.G., Lesher, C.M., Metsaranta, R.T., Sappin, A.-A., Carson, H.J.E., Schetselaar, E.M., McNicoll, V.J., and Laudadio, A., 2020. Magmatic architecture of the Esker intrusive complex in the Ring of Fire intrusive suite, McFaulds Lake - 69 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 greenstone belt, Superior Province, Ontario: Implications for the genesis of Cr and Ni-Cu-(PGE) mineralization in an inflationary dyke-chonolith-sill complex, in Bleeker, W., and Houlé M.G. (eds). Targeted Geoscience Initiative 5, Geological Survey of Canada, Open File 8722, pp. 141–163. Metsaranta, R.T., and Houlé, M.G., 2020. Precambrian geology of the McFaulds Lake “Ring of Fire” region, northern Ontario. Ontario Geological Survey, Open File Report 6359, 260 p. Mungall, J.E., Harvey, J.D., Balch, S.J., Azar, B., Atkinson, J., and Hamilton, M.A., 2010. Eagle’s Nest a Magmatic NiSulfide Deposit in the James Bay Lowlands, Ontario, Canada, in The Challenge of Finding New Mineral Resources: Global Metallogeny, Innovative Exploration, and New Discoveries, Volume II: Zinc-Lead, Nickel-Copper-PGE, and Uranium. Society of Economic Geologists, Special Publication 15, pp. 539–557. Naldrett, A.J., 2011, Fundamentals of Magmatic Sulfide Deposits. Reviews in Economic Geology, vol. 17, pp. 1–50. Smith, W.D., Jenkins, C.M., Augustin, C.T., Virtanen, V.J., Vukmanovic, Z., and O’Driscoll, B., 2024. Layered intrusions in the Precambrian: Observations and perspectives. Precambrian Research, 50th Anniversary Invited Review, vol. 415, 107615. Zuccarelli, N., 2020. Sulfide textures, geochemistry, and genesis of the Komatiite-Associated Eagle’s Nest Ni-Cu-(PGE) Deposit, McFaulds Lake Greenstone Belt, Superior Province, Ontario. MSc Thesis, Laurentian University, Sudbury, Ontario, Canada, 108 p. - 70 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Integrating petrophysical data with full tensor magnetic gradiometry for improved interpretation and modelling of remanently magnetized intrusions in the Midcontinent Rift SMITH, Jennifer1, KASKI, Krista1, TSCHIRHART, Victoria1, and ENKIN, Randy1. 1 Natural Resources Canada, Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8 Magnetic surveys are widely used in mineral exploration to detect and delineate subsurface structures and ore-bearing systems. As near-surface, high-grade deposits become increasingly rare, exploration is shifting toward deeper targets and more complex geological settings. Full tensor magnetic gradiometry (FTMG), particularly when deployed with highly sensitive SQUID-based quantum sensors, provides high-resolution measurements of all components of the magnetic field gradient tensor, offering enhanced imaging of subtle geological structures and ore bodies that conventional total magnetic intensity (TMI) surveys may not resolve (Rudd et al., 2022). FTMG reduces the influence of regional magnetic fields, diurnal variations, and cultural noise, supporting more robust 3D inversion and geological interpretation. Despite these advantages, adoption of FTMG has been limited by logistical complexity, depth constraints, and a lack of publicly available datasets particularly in geologically complex or remanently magnetized areas. To address this, the Geological Survey of Canada is acquiring and openly disseminating precompetitive SQUID-based FTMG datasets (e.g. Fig. 1), providing real-world data for benchmarking inversion workflows and testing emerging quantum sensors. Figure 1: Maps of the total magnetic intensity (TMI) (a), and three components of the magnetic gradient tensor: Bxx (b), Byy (c) and Bzz (d) over the Escape Intrusion within the Thunder Bay North Intrusive Complex of the Midcontinent Rift. The Midcontinent Rift (MCR) provides a geologically complex environment to evaluate FTMG in remanently magnetized settings. Mafic-ultramafic conduit-type intrusions in this region, including the Escape and Current intrusions of the Thunder Bay North Intrusive Complex (TBNIC), exhibit strong remanent magnetization, generating distinct and heterogeneous magnetic anomalies (Kaski et al., 2024; Fig. 1). These characteristics make the MCR an ideal setting to assess how FTMG resolves both induced and remanent magnetic components. In this study, we integrate SQUID-based FTMG inversions with petrophysical, petrographic, and geochemical data, including magnetic susceptibility, natural remanent magnetization, and mineralogical composition, to examine how lithologic variability, serpentinization, and magnetic mineral development influence the intensity and orientation of remanent magnetization, providing a more geologically realistic framework for interpretation and modeling. Preliminary results show that integrating FTMG with rock property data improves resolution of key geological contacts and remanent magnetic sources, enabling more robust 3D modeling of conduithosted Ni-Cu-PGE systems. This study highlights the value of combining high-resolution geophysical and petrophysical datasets for interpreting complex magnetic anomalies. - 71 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 REFERENCES Kaski, K., Smith, J., Tschirhart, V.L., and Heggie, G., 2024, 3D magnetic-susceptibility and magnetization vector inversions of remanently magnetized conduit-type Ni deposits: a case study from the Thunder Bay North intrusive complex, Ontario: Geological Survey of Canada, Open File 9209, 25 p, https://doi.org/10.4095/pkwpmf1tju Rudd, J., Chubak, G., LaNier, H., Stolz, R., Schiffler, M., Zakosarenko, V., Schneider, M., Schulz, M., Meyer, M., 2022, Commercial operation of a SQUID-based airborne magnetic gradiometer: Leading Edge. https://doi.org/10.1190/ tle41070486.1 - 72 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Optimizing data collection for better geological interpretations and adding value to your project SMYK, Emily1, DOLEGA, Simon1, CHURCHLEY, Jeffrey1, and FLANK, Steven1 1 Bayside Geoscience Inc., 1179 Carrick St. Thunder Bay, ON P7B 6M3 “Data are disembodied information. Data are not the same as knowledge.” ~ W. Olsen (2012) A well-designed field or drill program is developed from the beginning to produce substantiated, appropriate and robust datasets. However, data are commonly considered interchangeable with interpretations and are often misreported to fit a geologist’s bias within the context of a project. Common instances of data distortion include: (1) identifying and classifying rocks as pre-named units with assumed occurrences; (2) designating altered rocks as separate lithologies; (3) recording qualitative descriptions rather than quantitative variables; (4) not standardizing all aspects of data collection; and (5) generating incomplete geochemical datasets in the pursuit of select geochemical data. It is a human instinct to apply human interpretations to systematic rocks and processes, but collecting purely observational, quantified geologic data can provide significantly more flexible information during later interpretation. Some findings may emerge from a dataset without being expected or predicted in advance (Olsen, 2012). More ‘expected’ findings might follow the usual predictable patterns, but unpredictable trends may be obfuscated by unintentionally engineered data biases. The most impactful approach to optimize data collection procedures is standardizing all data input for recording rock identification and descriptions, photos, and QA/QC practices. Collecting alteration, mineralization, and structural data as separate data to the lithology, rather than integrated into lithology names (e.g., carbonatized basalt), allows for separation of different datasets for multiple applications and discourages segregating single rock units due to varying characteristics. Many issues are resolved by generating mandatory fields that can only be populated by standardized terms using drop-down menus. Another approach is quantifying and binning as many descriptors as possible. A simple change is including mineral abundance ranges in mineral description fields. For example, describing weak epidote alteration as ‘Weak (2-5%)’ provides a quantitative visual cue to the core logger/mapper, ensuring consistent descriptions and binning similar mineral percentages together. Another consideration is developing sampling programs that submit all samples for consistent analytical packages. Cost-saving measures are often implemented by selectively submitting samples for different packages or only submitting samples that are anticipated to return good assay data. These practices can identify high-grade samples, but can also miss secondary, unpredictable mineralization. Without a range of geochemical data, it is impossible to assess truly elevated values from background values. Purely objective geologic data can provide new interpretations depending on the approach/aims of the geologist. Consistent and comprehensive data collection may produce unexpected results and provide a valuable final product – a strong asset that increases the value of a project, property, or deposit. REFERENCES Olsen, W. (2012). Data Collection: Key Debates and Methods in Social Research. Sage Publications Ltd. - 73 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Pukaskwa Redux: Revisiting and Reconnecting with Superior’s Wild North Shore SMYK, Mark1, HODGE, Joanna2 and ROBILLARD, Carly3 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Canadian Federation of Earth Sciences, University of Ottawa, 150 Louis Pasteur Private, Ottawa ON K1N 6N5 Canada 2 3 Parks Canada, Pukaskwa National Park, PO Box 212, Heron Bay, ON, P0T 1R0 Canada In August, 2025, the Senior Author served as Geologist-in-Residence (GIR) at Pukaskwa National Park, on Lake Superior near Marathon. The GIR program at Pukaskwa is a partnership between the Canadian Federation of Earth Sciences and Parks Canada, with volunteer expenses funded by the APGO Education Foundation. It is a two-week, volunteer position that started at Pukaskwa in 2022. The role of the GIR is to highlight Pukaskwa’s remarkable geological features and to educate park visitors and Parks Canada interpretive staff about the local geology. Guided hikes, “walk and talk” sessions, drop-in opportunities and presentations were employed to convey knowledge and messaging. As a result of the 2025 GIR program, ideas are being considered to develop a self-guided geology field trip for the readily accessible “front country” trails at Pukaskwa that expose a variety of Neoarchean supracrustal rocks of the Schreiber-Hemlo greenstone belt. Its “back country”, featuring the Coastal Hiking Trail, is underlain mainly by Neoarchean granitoids of the Pukaskwa Batholith. Archean rocks are intruded by Paleoproterozoic and Mesoproterozoic diabase dykes, the latter of which are associated with Midcontinent Rift magmatism. There are numerous features attributed to Quaternary glaciation, including prominent roches moutonnées (Figure 1), potholes and glacial polish/ striae. Modern shoreline and aeolian processes continue to redistribute sediment and create unique and critical habitats for rare and endangered plant species. The GIR program serves to remind us of the importance and value of participating in outreach activities, sharing information and underscoring the critical role that geology plays in ecological processes. The program is expanding to Fundy National Park in 2026 with the hope that further National Parks will be added in the future to provide more opportunities for geoscience outreach and education to a broader audience. Figure 1: Geologist-in-Residence, Mark Smyk, pointing out a prominent roche moutonnée at Horseshoe Beach during a guided hike of the Southern Headland Trail, Pukaskwa National Park, August, 2025 - 74 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Oxidation to Ores: Petrological Insights into Supergene Manganese Enrichment at the Emily Deposit, Minnesota STEINER, R. Alex¹, WATSON, Noa2, RILEY, Jack2, HAMMER, Mikala3, THOLE, Jeff2, FEINBERG, Josh3, SANDRI, Henry4, and SAVAGE, Brian4 ¹Big Rock Exploration LLC, 2505 W Superior Street, Duluth, MN, 55803 USA 2 3 4 Macalester College, 1600 Grand Ave, St. Paul, MN 55105 USA University of Minnesota, 116 Church Street SE, Suite 150, Minneapolis, MN 55455 USA Electric Metals (USA) Limited, 109 West 13th Street Wilmington, DE 19801 USA Electric Metals (USA) Limited’s Emily Deposit in Minnesota’s historic Cuyuna Iron Range contains zones reaching +50 wt. % manganese, making it the highest-grade manganese resource in North America and one of the highest-grade manganese deposits in the world. Manganese-oxide ores of the Emily Deposit are proposed to have formed through supergene enrichment due to deep, potentially protracted weathering of folded iron formation strata during the deposit’s 1.9-billion-year history. Weathering of manganese-bearing carbonate facies oxidizes the original rhodochrosite, drawing the manganese into solution. The manganese enriched groundwaters then migrate down-dip, along stratigraphic boundaries before redepositing manganese as oxides in the porous grainstones of the iron formation. The recent exploration drilling campaign by Electric Metals USA Limited and Big Rock Exploration provided a wealth of geologic, geochemical, and microscopic data that may be used to evaluate the hypothesized ore genesis mechanism on a deposit scale and constrain the metallurgical behavior of the ores. Here we present an analysis of a large exploration geochemical dataset using deposit-wide mass-balance calculations to determine the element mobility within the iron formation. The geochemical results are then contextualized within geology by combining optical and X-ray microscopy to identify mineral phases and phase transitions, as well as intergrowths of secondary minerals. Mass balance calculations show depletions in manganese from the weathered carbonate facies of the Emily Iron Formation and parallel enrichment of manganese into the grainstones. Integration of preliminary optical and X-ray microscopy shows a breakdown of early-formed minerals in the source carbonates and replacement by Fe-oxides and oxyhydroxides along bedding and fractures. Secondary manganese minerals appear to surround primary grains in the grainstones and may be replacing early formed ferruginous cements. These observations support the hypothesized ore-genesis model and provide the necessary information for subsequent metallurgical evaluation of the Emily Deposit including the manganese-iron-silicate mineral associations that may impact ore upgrading, grinding, and hydrometallurgical outcomes. REFERENCES Steiner, R. A., Peterson, D., Berg, T., Solie, J., Larson, M., Schaefbauer, E., Sweet, G., 2024, North Star Emily Manganese Deposit, Crow Wing County, Minnesota: Observations Interpretations, and Recommendations Following the Initial 2023 Drilling Campaign, January 17, 2024. Big Rock Exploration. - 75 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1 – Full section reflected light (above) and X-ray map showing texture of iron and manganese minerals. Areas with mixed iron and manganese minerals and pure, coarse grained manganese species are highlighted. - 76 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Timing and conditions of magmatism, metamorphism, and strain partitioning in the western Shebandowan Greenstone Belt (Superior Province) STEPHAN, Tobias1, PHILLIPS, Noah1,2, and HOLLINGS, Pete1 1 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada Department of Earth Sciences, University of Southern California, 3651 Trousdale Pkwy., Los Angeles, CA, 90089-0740, United States 2 The Shebandowan Greenstone Belt is an Archean granite–greenstone terrane within the Wawa subprovince of the Superior Province, comprising calc-alkaline to tholeiitic, felsic to ultramafic supracrustal metavolcanic rocks, synvolcanic to late intrusive suites, and felsic hypabyssal dikes and sills. Despite its economic and tectonic significance, the timing and conditions of magmatism, metamorphism, and deformation remain incompletely constrained. Here, we integrate structural geology, high-precision geochronology, metamorphic petrology, and microstructural analyses to establish a coherent tectonometamorphic framework for the western belt. Strain varies from weakly deformed domains (e.g., felsic intrusions and pillow basalts) to highstrain mylonitic zones, mainly affecting diorites and metavolcanic rocks. The orientation of the main ductile foliation orientation is relatively consistent across the study area, while stretching lineations range from shallow to steep. These variations correlate with spatial changes in vorticity, reflecting strain partitioning between high-strain shear zones and coarse-grained, feldspar-rich, and thus, mechanically strong intrusive bodies (Stephan et al. 2025). Peak metamorphic conditions of ~600–700 °C are constrained by pseudosection modeling and conventional thermometry, consistent with Zr-in-titanite temperatures (570–700 °C). Retrograde conditions of ~400–500 °C are preserved in post-kinematic assemblages. Quartz microstructures, crystallographic preferred orientations, and grain-size piezometry indicate deformation at ~400–600 °C and differential stresses of ~20–60 MPa, suggesting deformation near the brittle–ductile transition. CA-ID-TIMS U-Pb zircon geochronology identifies two magmatic phases based on concordant ages: an intrusive phase at 2718 Ma (e.g. felsic intrusion of Moss Lake Stock and Obadinaw Stock) and a younger phase at 2707 Ma (e.g. Greenwater Stock). An upper intercept age constrains volcanism at 2712 Ma in the metavolcanic sequences. In situ U–Pb titanite dates of 2711±76 Ma (2σ) and 2672±100 Ma record metamorphic events spanning greenschist- to amphibolite-facies conditions. A Re-Os molybdenite age of 2708±12 Ma overlaps with both magmatism and metamorphism, linking mineralization to tectonometamorphic processes. These results indicate synkinematic magmatism and amphibolite-facies deformation under predominantly horizontal tectonics. Strain was strongly partitioned due to competency contrasts between coarse-grained intrusive rocks and fine-grained metavolcanic units. This integrated dataset provides new constraints on the coupling between magmatism, deformation, metamorphism, and mineralization in Archean granite–greenstone belts. REFERENCES Stephan, T., Phillips, N., Tiitto, H., Perez, A., Nwakanma, M., Creaser, R., and Hollings, P. 2025. Going with the flow — Changes of vorticity control gold enrichment in Archean shear zones (Shebandowan Greenstone Belt, Superior Province, Canada). Journal of Structural Geology, 201, 105542. https://doi.org/10.1016/j.jsg.2025.105542 - 77 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Subsurface mapping of the late Ordovician Maquoketa Group in eastern Wisconsin using airborne electromagnetic and well data STEWART, Esther K.1, McNALL, Natalie1, 2, HART, Dave1, AMES, Carsyn 1, CHASE, Pete1, STEWART, Eric1, and GRAHAM, G.1 Wisconsin Geological and Natural History Survey, University of Wisconsin-Madison Division of Extension, Madison, Wisconsin 53705 1 2 Department of Geosciences, University of Wisconsin-Milwaukee, Milwaukee, Wisconsin 53211 The late Ordovician Maquoketa Group is a fine-grained unit and regional aquitard separating the upper, fractured Silurian dolostone aquifer from the deep, Cambrian-Ordovician sandstone-dolomite aquifer in eastern Wisconsin. Here, the Maquoketa Group lithostratigraphy includes, from top to bottom, the Brainard Formation (marls and shale), Ft Atkinson Formation (carbonate wackstonethrough grainstone and marls), and Scales Formation (black shales and marls). The shale-rich composition of the Maquoketa Group is readily distinguished from the overlying Silurian dolostone by airborne electromagnetic (AEM) data (Minsley et al., 2022). We undertook subsurface mapping and characterization of the Maquoketa Group to address regional issues of groundwater quantity and quality. For Wisconsin users, the resulting 3D surfaces can be used as inputs to groundwater models and aid land-use decisions by providing information on the depths, thickness, and rock properties of this aquitard. We used AEM data tied to borehole logs and core to generate raster surfaces and understand facies changes and structures across study area (Figure 1). Despite cultural interference mainly from roads, the AEM data nicely imaged the top of the Maquoketa Group aquitard. The Maquoketa Group basal surface and its internal formations were imaged by the AEM data but with greater uncertainty, and the base of the unit dipped below the penetration depth of the AEM data to the east. The Maquoketa Group extends from about 850 feet (259 m) above sea level near its western subcrop extent to 100 feet (31 m) above sea level at the eastern edge of the map area, with thicknesses between about 220 – 450 ft (67 – 137 m). The north-south strike of depth-structure elevation contours is abruptly offset in three locations, labeled on Figure 1. One of these (location 2) corresponds to the Precambrian Spirit Lake Tectonic zone (Holm et al., 2007) and fault offset of Silurian bedrock (Luczaj, 2011). Several new and existing drill core tie to the AEM data in the western study area, and lithologic variation in the cores corresponds to vertical changes in the resistivity profile of the Maquoketa Group. Internal variability in the resistivity of the Maquoketa, as imaged by the AEM data, apparently decreases to the east. Future air rotary drilling will test whether this signal is due to decreased data resolution as these units dip eastward, or whether it reflects an increase in shaley facies to the east. REFERENCES 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. Reinterpretation of Paleoproterozoic accretionary boundaries of the north-central United States based on a new aeromagnetic-geologic compilation. Precambrian Research, 157, 71-79. Luczaj, J.A., 2011. Preliminary Geologic Map of the Buried Bedrock Surface, Brown County, Wisconsin. Wisconsin Geological and Natural History Survey Open File Report 2011-02. Minsley, B.J, Bloss, B.R., Hart, D.J., Fitzpatrick, W., Muldoon, M.A., Stewart, E.K., Hunt, R.J., James, S.R., Foks, N.L., and Komiskey, M.J., 2022. Airborne electromagnetic and magnetic survey data, northeast Wisconsin. U.S. Geological Survey data release, https://doi.org/10.5066/P93SY9LI. - 78 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Figure 1: Maps showing the elevation and thickness of the Maquoketa Group (top) and an example AEM line (below). The inset map of Wisconsin (left) shows counties outlined in black and the eastern Wisconsin study area outlined in orange. Circled numbers to the left of the top Maquoketa elevation map locate offsets in depthstructure contours. The star locates the Krepline core on the map and AEM line, and formation contacts from core are tied to AEM line. Roads and railroads, symbolized above the line, cause cultural interference with the AEM signal. - 79 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Rocks and Roots: The Role of Geoheritage in Biodiversity Stewardship STONE, Abraham1, LIZZADRO-McPHERSON, Dan2, and VYE, Erika3 Michigan Natural Features Inventory, Deborah A. Stabenow Building, 1st Floor, 525 W. Allegan St., Lansing, MI 48933, United States 1 Geospatial Research Facility, Michigan Technological University, 1400 Townsend Dr., Houghton, MI 49931, United States 2 Great Lakes Research Center, Michigan Technological University, 1400 Townsend Dr., Houghton, MI 49931, United States 3 Conservation of natural surficial landforms with regional, scientific, or cultural significance has long been an intrinsic component used by scientists and educators who follow the principles of geoheritage. On the Keweenaw Peninsula, intact outcrops of Copper Harbor Conglomerate, Portage Lake Volcanics, and Jacobsville Sandstone each provide accessible learning opportunities to both students and citizens and create spaces for deeper emotional connections to the landscape. Culturally and geologically important sites are currently used in both educational tools and to increase community-wide engagement in geologic studies (Cowling et al. 2023; Lizzadro-McPherson and Vye 2023). The principles of natural heritage, hereto referred also as ‘bioheritage’, strongly overlap with that of geoheritage. As geoheritage promotes connection to landscape via geological features, bioheritage facilitates connection through valuable natural features – ecosystems, flora and fauna – and encourages the conservation of landscapes that promote biodiversity. Sites that are identified by bioheritage ecologists, botanists, zoologists, and geographers as being of regional, scientific, or cultural significance often coincide with areas of high geodiversity. Categorization of these natural features show geographies dependent on both surface geology and glacial landforms; for example, the statewide distribution of volcanic bedrock lakeshore (Fig. 1), an imperiled natural community in Michigan, is wholly limited to surface-level exposures of Keweenawan rocks (Cohen et al. 2013) and supports a series of rare plants and animals found nowhere else in the state (Albert et al. 1997; MNFI 2026). Conservation of one outcrop for geological reasoning can therefore work beneficially for bioheritage, and vice versa. In the summer of 2025, we conducted interdisciplinary research highlighting the natural connections between underlying geological formations, community ecology, and rare plant Figure 1: Portage Lake Volcanics featured prominently along a high-quality volcanic bedrock lakeshore natural community recognized under both geoheritage and natural heritage. Figure 2: Pilot data examining ecological structure of bedrock lakeshore systems. Different zones of bedrock exposure promote plant communities of unique species composition. - 80 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 distributions on volcanic bedrock lakeshores of the Keweenaw Peninsula. The project involved collaboration between botanists, geologists, geographers, and conservationists. Pilot data yielded significant plant community differences between bedrock types and microhabitats, and the summer was documented in an educational StoryMap (Stone et al. 2025) (Fig. 2). Meandering transects outlining subtle distinctions in ecosystem processes based on geological formation identified multiple new rare plant populations, including the discovery of red anemone (Anemone multifida) on the Keweenaw Peninsula (Fig. 3). The project has since led to multiple additional collaborations in the Western Upper Peninsula focused on geoand bio-education. Partnerships between bioheritage and geoheritage scientists can be valuable sources of interdisciplinary research and collaboration. Despite originating in disparate academic fields, the two disciplines can work in tandem to increase scientific understanding of our geological features while producing valuable teaching tools. Future research and educational opportunities are plentiful as the two worlds of geoheritage and bioheritage establish common ground. REFERENCES Figure 3: Red anemone (Anemone multifida), a rare plant identified during field research on the Keweenaw Peninsula. Albert, D.A., Comer, P., Cuthrell, D., Hyde, D., MacKinnon, W., Penskar, M., & Rabe, M., 1997. The Great Lakes Bedrock Lakeshores of Michigan. Michigan Natural Features Inventory, Lansing, MI. 218 pp. Cohen, J.G., Kost, A., Slaughter, B.A., & Albert, D.A., 2015. A Field Guide to the Natural Communities of Michigan. Michigan State University Press. 362 pp. Cowling, R., Lizzadro-McPherson, D.J., Verissimo, L. & Vye, E.C., 2023. Keweenaw Geoheritage Geoatlas. DOI: 10.13140/RG.2.2.30945.28005 Lizzadro-McPherson, D. J. & Vye, E.C., 2023. Keweenaw Coastal Geoheritage Story Map. DOI: 10.13140/ RG.2.2.12680.74242 Michigan Natural Heritage Database (MNFI), 2026. Michigan Natural Heritage Database. Lansing, MI. Stone, A.F., Lizzadro-McPherson, D.J., and Vye, E.C., 2025. Rocks and Roots: A Keweenawan Love Story. StoryMap. https://storymaps.arcgis.com/stories/7d9a428effe04dc4923736310182d52f - 81 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Linking the Southwestern Laurentia large igneous province and rapid Duluth Complex emplacement through mantle plume dynamics SWANSON-HYSELL, Nicholas L.1, ZHANG, Yiming1, MOHR, Michael T.2, and SCHMITZ, Mark D.2 1 2 Department of Earth and Environmental Sciences, University of Minnesota, Minneapolis, MN, USA Department of Geosciences, Boise State University, Boise, ID, USA Midcontinent Rift volcanism was protracted, spanning from ca. 1109 to 1084 Ma with major magmatic pulses separated by ~10 Myr and >30° of latitudinal plate motion (Figure 1). The long duration of magmatism and large spatial displacement of the continent are difficult to reconcile with a single stationary mantle plume beneath the rift. A corresponding question is what caused the renewal of voluminous magmatism ca. 1096 Ma that produced the massive Duluth Complex layered mafic intrusions and comagmatic lavas of the North Shore Volcanic Group after a period of relative magmatic dormancy (Miller and Vervoort, 1996), and after Laurentia had drifted >3000 km since the rift’s initiation (Swanson-Hysell et al., 2019, 2021). Figure 1: The plate motion of Laurentia reconstructed from Midcontinent Rift paleomagnetic data revealing large-scale latitudinal change between the start of early phase volcanism and the major pulse of magmatism that emplaced the Duluth Complex ca. 1096 Ma. The red dot indicates the location of the Lake Superior region in each reconstruction. High-precision ²06Pb/²38U zircon dates developed through CA-ID-TIMS geochronology have resolved temporally distinct pulses of magmatism across Laurentia’s interior. In southwestern Laurentia, the Southwestern Laurentia large igneous province (SWLLIP) encompasses >750,000 km² of ca. 1.1 Ga mafic sills, dikes, and lava flows. New dates from SWLLIP mafic rocks reveal a rapid, voluminous magmatic pulse at ca. 1098 Ma, with thick sills emplaced across Death Valley, the Grand Canyon, and central Arizona within ≤0.25 Myr (Mohr et al., 2024). Approximately 2 Myr later, the bulk of the Duluth Complex anorthositic and layered series was emplaced ca. 1096 Ma in <1 Myr (500 ± 260 kyr; Swanson-Hysell et al., 2021). Both pulses were rapid and voluminous, characteristic of plume-related large igneous provinces. The close temporal and spatial relationship between the ca. 1098 Ma SWLLIP pulse and the ca. 1096 Ma Duluth Complex pulse supports a geodynamic link through lateral plume spreading. Rates of lateral plume spread predicted by mantle plume lubrication theory (Sleep, 1997) are consistent with a model in which a plume derived from the deep mantle impinged beneath southwestern Laurentia, then spread to the thinned Midcontinent Rift lithosphere over ~2 Myr, elevating mantle temperatures and generating melt. Buoyant plume material would have been directed to the rift through “upsidedown drainage” at the base of the Laurentian lithosphere (Sleep, 1997; Swanson-Hysell et al., 2021), - 82 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 wherein material flows along the topography of the lithosphere–asthenosphere boundary from thick to thin lithosphere. This hypothesis reconciles the close temporal relationships between voluminous magmatism across Laurentia and provides an explanation for the anomalous renewal of high magmatic flux within the protracted magmatic history of the Midcontinent Rift. REFERENCES Miller Jr., J.D., and Vervoort, J.D., 1996. The latent magmatic stage of the Midcontinent rift: a period of magmatic underplating and melting of the lower crust. In: Inst. Lake Superior Geol., 42nd Ann. Mtg., Proceedings, vol. 42, pp. 33–35. Mohr, M.T., Schmitz, M.D., Swanson-Hysell, N.L., Karlstrom, K.E., Macdonald, F.A., Holland, M.E., Zhang, Y., and Anderson, N.S., 2024. High-precision U-Pb geochronology links magmatism in the Southwestern Laurentia large igneous province and Midcontinent Rift. Geology, doi:10.1130/G51786.1. Sleep, N.H., 1997. Lateral flow and ponding of starting plume material. Journal of Geophysical Research, 102, 10,001– 10,012, doi:10.1029/97JB00551 Swanson-Hysell, N.L., Hoaglund, S.A., Crowley, J.L., Schmitz, M.D., Zhang, Y., and Miller, J.D., 2021. Rapid emplacement of massive Duluth Complex intrusions within the North American Midcontinent Rift. Geology, 49, doi:10.1130/ G47873.1. 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, 131(5–6), 913–940, doi:10.1130/B31944.1. - 83 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Deformation processes in a mid-crustal strike-slip shear zone: Insights from the Archean Quetico Shear Zone, Superior Province, Canada TIITTO, Hanna1, PHILLIPS, Noah1, 2, and STEPHAN, Tobias1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7C 5E1, Canada Department of Earth Sciences, University of Southern California, 3651 Trousdale Pkwy., Los Angeles, CA, 90089-0740, United States 2 The brittle-ductile transition, where most earthquakes nucleate, occurs at ~10-15 km depth in the crust (Sibson, 1983). The structures produced at the brittle-ductile transition in active shear zones are challenging to study as they occur at depth. To further understand shear zone structures at depth, this study focuses on an analogue structure for active strike-slip systems, the Quetico Shear Zone, due to its estimated erosional depths of 10-15 kms, which exposes the Archean brittle-ductile transition zone (Percival et al., 2012). The Quetico Shear Zone is a right-lateral, strike-slip shear zone located within the Wabigoon and Quetico subprovinces and has a strike length of at least 400 km (Kennedy, 1984). This project focuses on the eastern extent of the shear zone, north of Thunder Bay, and aims to constrain the kinematics, structures, conditions, and timing of deformation processes within the shear zone and adjacent to paleo-earthquake surfaces. The extent of deformation from the shear zone was analyzed through macro- and microstructures using field mapping and microscopy. The conditions of deformation were constrained using paleopiezometry through electron backscattered diffraction of recrystallized quartz (Cross et al., 2017) and Ti-in-quartz geothermometry through secondary ion mass spectrometry measurements of recrystallized quartz (Wark and Watson, 2006). To constrain the timing of deformation, laser-ablation split-stream inductively coupled plasma mass spectrometry of apatite, monazite, titanite, and zircon was performed to produce U-Pb dates (Kylander-Clark, 2017). We found that Quetico Shear Zone deformation is characterized by increased mylonitization and brittle deformation with increasing proximity to the shear zone trace (within 500 m) where paleoearthquake surfaces (i.e., pseudotachylite veins) were found (Fig. 1). Mylonitization produces recrystallized quartz ribbons and a strong foliation unique to the Quetico Shear Zone (stronger than regional Quetico Subprovince transpressional structures), particularly in granitic units (Fig. 1C-E). Non-granitic rock units within the core of the shear zone display pervasive brittle deformation with numerous faults (Fig. 1A). Granitic rock types display more variable orientations due to the isolated quartz ribbons deforming around larger feldspar grains. The recrystallized quartz grain sizes do not correlate with increased mylonitization and proximity to the shear zone. Recrystallized quartz grain sizes remained constant within error, with calculated stress values ranging from 69 to 116 MPa, with a median of 80 MPa. The temperatures of quartz recrystallization range from 457 to 589°C, with a median of 487°C, with no clear evolution with increasing proximity to the Quetico Shear Zone trace. Apatite and titanite provided the best ages for deformation, mainly producing interpreted ages younger than the Quetico subprovince metamorphism. The interpreted Quetico Shear Zone deformation ages are approximately from 2620 to 2600 Ma. The exhumed Quetico Shear Zone appears to be deformed at a constant stress shortly after the Kenoran orogeny. REFERENCES Cross, A.J., Prior, D.J., Stipp, M., & Kidder, S., 2017. The recrystallized grain size piezometer for quartz: An EBSD-based calibration. Geophysical Research Letters, 44, 6667-6674. Kennedy, M.C., 1984. The Quetico Fault in the Superior Province of the Southern Canadian Sheild [MSc]: Lakehead University, 323. Kylander-Clark, A.R.C., 2017. Petrochronology Laser-Ablation Inductively Coupled Plasma Mass Spectrometry. Reviews in Mineralogy and Geochemistry, 83, 183-198. Percival, J.A., Skulski, T., Sanborn-Barrie, M., Stott, G.M., Leclair, A.D., Corkery, M.T., Boily, M., 2012. Geology and tectonic evolution of the Superior Province, Canada. Chapter 6 In Tectonic Styles in Canada: The Lithoprobe - 84 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Perspective. Geological Association of Canada, Special Paper 49, 321-378. Sibson, R.H., 1983. Continental fault structure and the shallow earthquake source. Journal of Geological Society, 140, 741767. Wark, D.A., and Watson, E.B., 2006. TitaniQ: a titanium-in-quartz geothermometer. Contributions to Mineralogy and Petrology, 152, 743-754. Figure 1: Quetico Shear Zone structures proximal to pseudotachylite veins: A: Plane-polarized light photomicrograph displaying pseudotachylite veins (medium brown layers cutting the white to light brown mylonitic fabric) from the core of the shear zone. Co-seismic injection veins are highlighted with white arrows. Right-lateral, late brittle faults are indicated by kinematic arrows. B: Magnified view of a pseudotachylite that has been viscously deformed. Cross-polarized light photomicrographs showing quartz microstructures of quartz-rich metamorphic rocks from increasing distance from the pseudotachylites: C: Extremely fine-grained quartz ribbons with minor feldspar porphyroclasts within a mylonite. D: Very fine-grained quartz within a quartz ribbon adjacent to fine-grained quartz in a protomylonitic granite. E: Fine- to mediumgrained recrystallized quartz within a weakly deformed granite. Note that the recrystallized grain size is consistent in C-E. - 85 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Variations in Olivine Major Element Composition Across the Midcontinent Rift System TOLLEY, James1 and HOLLINGS, Pete1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1, Canada. Olivine [(Mg,Fe)2SiO4] is an early crystallising phase in mafic–ultramafic Ni–Cu–(PGE) deposits and a sensitive recorder of mantle melting history. Its forsterite content reflects the parental melt composition, while the Ni concentration and trace element ratios can be used to constrain petrogenetic processes and the physicochemical conditions of melting. However, the plutonic nature of these deposits means primary compositions can be overprinted by sub-solidus re-equilibration and latestage fluid interaction, complicating the recovery of primary magmatic signals. Deconvoluting these signatures is critical to understanding melt generation, fractionation, and ultimately the mineralisation processes that govern the formation of these deposits. The Midcontinent Rift System (MRS) one of the most extensively mineralised large igneous provinces and renowned for its magmatic Ni–Cu–(PGE) deposits. Despite this, olivine compositional data is sparse. We present new and collated major element olivine data from multiple Ni–Cu–(PGE) deposits across the MRS to evaluate regional-scale trends in forsterite and Ni contents. We examine deposit-scale variability and explore broader implications for the underlying magmatic architecture of the rift system. This study builds on previously collected electron probe microanalyses (EPMA) of olivine from mineralised magmatic Ni–Cu–(PGE) deposits of the MRS within Canada e.g., Sunday Lake (Durán, 2025), Steepledge (Harding, 2024), Escape Lake, Current and Hele, and contributes new olivine compositional data from several unmineralized intrusions, namely Inspiration Sill, St. Ignace Island and Nipigon Sills. These data are further supplemented by olivine compositions from USA-based mineralised Ni–Cu deposits e.g., Tamarack (Goldner, 2011; Taranovic, 2015) and the Duluth Complex (Peterson, 2025). Together, this data constitutes the first regional-scale compilation of olivine chemistry across the MRS. Figure 1: Simplified geological map of the Midcontinent Rift System highlighting the distribution of major rock types. Locations of the mafic–ultramafic intrusions sampled in this study are denoted by stars (red = data collected in this study; blue = literature data). Modified after: Good et al. (2015). - 86 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 REFERENCES Ding, X., Li, C., Ripley, E. M., Rossell, D., & Kamo, S. (2010). The Eagle and East Eagle sulfide ore‐bearing maficultramafic intrusions in the Midcontinent Rift System, upper Michigan: Geochronology and petrologic evolution. Geochemistry, Geophysics, Geosystems, 11(3). Durán, K. M. (2025), Petrogenesis of the Sunday Lake Intrusion, Jacques Township, Ontario, Canada. M.Sc. thesis Lakehead University, Thunder Bay, Ontario, 222p. Goldner, B.D. (2011). Igneous petrology of the Ni–Cu–PGE mineralized Tamarack intrusion, Aitkin and Carlton Counties, Minnesota; M.Sc. thesis, University of Minnesota, Minneapolis, 156p. Good, D.J. (1992). Genesis of copper-precious metal sulphide deposits in the Port Coldwell Alkalic Complex, Ontario; unpublished Ph.D. thesis, McMaster University, Hamilton, Ontario, 203p. Good, D. J., Epstein, R., McLean, K., Linnen, R. L. and 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(4), 983–1008p. Harding, M. F., (2024). Olivine Geochemistry of the Current and Escape Lake (Steepledge) intrusions, Thunder Bay North Intrusive Complex. HBSc. Thesis, Lakehead University, Thunder Bay Ontario. Heggie, G.J. (2005). Whole rock geochemistry, mineral chemistry, petrology and Pt, Pd mineralization of the Seagull Intrusion, northwestern Ontario. M.Sc. thesis, Lakehead University, Thunder Bay, Ontario, 156p. Peterson, D.M., (2025). Compilation of electron probe microanalyses of Olivine from the Duluth Complex, Minnesota, USA [Unpublished Dataset – personal communication]. Shaw, C. S. (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. 40(2-4), 243–259. Taranovic, V., Ripley, E.M., Li, C. and Rossell, D., (2015). Petrogenesis of the Ni–Cu–PGE sulfide-bearing Tamarack Intrusive Complex, Midcontinent Rift System, Minnesota. Lithos, 212, 16–31p. - 87 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 Index AKIN, Kathryn�����������������������������������������������������1 ALLERTON, Zsuzsanna���������������������������������3, 31 AMES, Carsyn����������������������������������������������������78 ANGOMBE, Moses����������������������������������������5, 69 BAIN, Wyatt���������������������������������������������������������6 BEDROSIAN, Paul A.����������������������������������������16 BEYER, Steve�������������������������������������������������������8 BILBOE, Michael�����������������������������������������������10 BLEEKER, Wouter���������������������������������������������12 BONAMICI, Chloë���������������������������������������������41 BORNHORST, Theodore�����������������������������������21 BOUCHER, Chanelle�����������������������������������������61 BRENGMAN, Latisha����������������������������25, 29, 51 BUCHHOLZ, Thomas����������������������������������������14 CAMACHO, Alfredo��������������������������������������������8 CANNON, W. F.�������������������������������������������16, 67 CARLTON, Kenz M.������������������������������������������18 CAWOOD, Tarryn������������������������������������������������8 CHAISSON, Amy�����������������������������������������������19 CHASE, Pete������������������������������������������������������78 CHITTICK, Steve�����������������������������������������������65 CHURCHLEY, Jeffrey����������������������������������������73 CISNEROS, John Alex���������������������������������������29 CONLY, Andrew�������������������������������������������������10 COOKE, David���������������������������������������������������45 COWLING, Bob�������������������������������������������������21 CUTTS, Jamie�������������������������������������������������������8 DEERING, Chad�������������������������������������������������55 DeGRAFF, James�����������������������������������������21, 48 DELLER, Matt������������������������������������������������5, 69 DOLEGA, Simon������������������������������������������������73 DRENTH, Benjiman J.���������������������������������������16 DREVER, Garth���������������������������������������������������8 DROST, Abraham�����������������������������������������������23 DROUBI, Omar��������������������������������������������������41 DUFFY, Paige�����������������������������������������������������25 EASTON, Robert Michael����������������������������������27 ELLISON, Kimberly�������������������������������������������29 ENKIN, Randy����������������������������������������������������71 ERICKSON, Stephanie���������������������������������������31 ESSIG, Espree�����������������������������������������������������55 EYSTER, Athena������������������������������������25, 29, 51 FALSTER, Alexander�����������������������������������������14 FAYON, Annia����������������������������������������������������31 FEINBERG, Josh������������������������������������������������75 FLANK, Steven��������������������������������������������������73 FRALICK, Philip������������������������������������������33, 34 GAMET, Nolan���������������������������������������������������63 GAONA, Jorge Mario�����������������������������������������53 GILBERG, Nolan�����������������������������������������������33 GORNER, Emily������������������������������������������������45 GOSAI, Meghna�������������������������������������������������34 GRAHAM, G.�����������������������������������������������������78 GRAUCH, V.J.S��������������������������������������������������35 HAGEDORN, Grant�������������������������������������������37 HAKURTA, Joyashish����������������������������������������59 HAMILTON, Mike���������������������������������������������50 HAMMER, Mikala���������������������������������������������75 HARDING, Myles����������������������������������������������39 HART, Dave��������������������������������������������������������78 HASTIE, Evan����������������������������������������������������50 HEGGIE, Geoff��������������������������������������������������23 HELLER, Samuel J.��������������������������������������������35 HELLRUNG, Alyssa������������������������������������������41 HILLENBRAND, I.��������������������������������������������67 HILLIPS, Noah���������������������������������������������������84 HILTUNEN, Lindsay������������������������������������������21 HIRSCH, Aaron��������������������������������������������������43 HODGE, Joanna�������������������������������������������������74 HOLLINGS, Pete����������������5, 6, 39, 45, 69, 77, 86 HOMPSON, J. M������������������������������������������������67 HUDAK, George��������������������������������������������3, 31 JONSSON, Justin������������������������������������������������47 KAMO, Sandra���������������������������������������������27, 50 KASKI, Krista�����������������������������������������������������71 LAFRENIERE, Don�������������������������������������������21 LI, Zhiquan���������������������������������������������33, 34, 47 LIZZADRO-McPHERSON, Dan�����������21, 48, 80 MACDONALD, Peter����������������������������������������50 MAHIN, Robert��������������������������������������������������63 MALEGUS, Paul������������������������������������������������50 MANGUM, John������������������������������������������������51 MARIN LÓPEZ, Valentina���������������������������������51 MARSH, Jeff������������������������������������������������������50 McNALL, Natalie�����������������������������������������������78 MITCHELL, Jennifer�����������������������������������������51 MOHR, Michael�������������������������������������������������82 NACHLAS, William O���������������������������������������18 NESHEIM, Timothy�������������������������������������������65 NITESCU, Bogdan���������������������������������������������53 NOWAK, Robert�������������������������������������������������55 NOWARIAK, Eric����������������������������������������������57 OST, Sara������������������������������������������������������������59 PALIEWICZ, Cory���������������������������������������������59 - 88 - �Proceedings of the 72nd ILSG Annual Meeting - Part 1 PETERSON, Dean����������������������������������������������61 PHILLIPS, Noah���������������������������������������������5, 77 POWELL, Jeremy�������������������������������������������������8 PU, Xiaofei���������������������������������������������������������51 QUIGLEY, Ashley����������������������������������������������63 RILEY, Jack��������������������������������������������������������75 ROBILLARD, Carly�������������������������������������������74 ROSE, William���������������������������������������������21, 48 RUGGLES, Claire����������������������������������������������41 SAINI-EIDUKAT, Bernhardt�����������������������������65 SALERNO, R�����������������������������������������������16, 67 SANDRI, Henry��������������������������������������������������75 SAVAGE, Brian��������������������������������������������������75 SCHMITZ, Mark������������������������������������������������82 SEVERSON, Allison������������������������������������������57 SHESHNEV, Vlad������������������������������������������5, 69 SIMMONS, William�������������������������������������������14 SMITH, Andrew���������������������������������������������������5 SMITH, Jennifer�������������������������������������������������71 SMYK, Emily�����������������������������������������������������73 SMYK, Mark������������������������������������������������19, 74 SOUDERS, A. K�������������������������������������������������67 STEINER, R. Alex����������������������������������������61, 75 STEPHAN, Tobias������������������������������������5, 77, 84 STERN, Richard�������������������������������������������������69 STEWART, Eric��������������������������������������������������78 STEWART, Esther����������������������������������������������78 STONE, Abraham�����������������������������������������������80 SWANSON-HYSELL, Nicholas��������������������1, 82 SWEET, Gabriel�������������������������������������������������61 THOLE, Jeff�������������������������������������������������������75 TIITTO, Hanna���������������������������������������������������84 TIKOFF, Basil�����������������������������������������������������18 TOLLEY, James��������������������������������������������69, 86 TORRES, David Santiago����������������������������������53 TSCHIRHART, Victoria�������������������������������������71 VERVOORT J.����������������������������������������������������67 VRZOVSKI, Joseph�������������������������������������������45 VYE, Erika����������������������������������������������21, 48, 80 WALKER, Patrick�����������������������������������������������51 WATSON, Noa����������������������������������������������������75 WODICKA, Natasha������������������������������������������12 ZHANG, Yiming�������������������������������������������������82 ZUREVINSKI, Shannon�������������������������������10, 19 - 89 - � https://digitalcollections.lakeheadu.ca/files/original/9c27d09e8827f7b36e6389fd03c441fd.pdf ed3efa960579e362e6c602cd6bb9be41 PDF Text Text 72nd Annual Meeting Thunder Bay, Ontario - May 21-22, 2026 Institute on Lake Superior Geology Part 2 – Field Trip Guidebook �Thank you to our sponsors! �72tnd Annual Meeting Institute on Lake Superior Geology May 21-22, 2026 Thunder Bay, Ontario HOSTED BY: Mark Puumala and Peter Hinz Co-Chairs Ontario Geological Survey (Retired) Proceedings - Volume 72 Part 2 – Field Trip Guidebook Compiled and edited by Pete Hollings Cover Photos: Top - Keweenawan diabase dyke on Lake Superior shoreline near Thunder Bay, Middle Archean-Paleoproterozoic unconformity, Highway 11-17, near Pass Lake turnoff, Bottom - Colloform stromatolite, Gunflint Formation, Kakabeka Falls �72nd Institute on Lake Superior Geology Volume 72 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trips 1 & 4: “Classic” Geological Sites in the Thunder Bay Area Trip 2: Geology of the Quetico Supprovince North of Thunder Bay Trip 3: Gold Deposits of the Shebandowan Greenstone Belt Trip 5: Structural Geology and Gold Mineralisation of the Mine Centre Area Trip 6: Amethyst Deposits of Thunder Bay Reference to material in Part 2 should follow the example below: Poulsen, K.H., 2026. Archean Geology and Metallogeny of the Rainy Lake Wrench Zone. In; Hollings, P. (Ed.), Institute on Lake Superior Geology Proceedings, 72nd Annual Meeting, Thunder Bay, Ontario, Part 2 Field trip guidebook, v.72, part 2, 3-31. Published by the 72nd 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 72nd ILSG Annual Meeting - Part 2 Table of Contents Introduction - considerations and acknowledgements.........................................................1 Trips 1 & 4 - “Classic” Geological Sites in the Thunder Bay Area.....................................2 Trip 2 - Geology of the Quetico Subprovince and Shebandowan greenstone belt north of Thunder Bay...............................................................................................................44 Trip 3 - Geological assemblages, regional structural framework and tectonic evolution of the Neoarchean Shebandowan greenstone belt..........................................................67 Trip 5 - Archean Geology and Metallogeny of the Rainy Lake Wrench Zone..................82 Trip 6 - Amethyst Deposits of Thunder Bay....................................................................126 -i- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Introduction - considerations and acknowledgements Peter Hinz and Mark Puumala Resident Geologist Program, Ontario Geological Survey, Ministry of Energy & Mines, Thunder Bay, Ontario (Retired) This volume is intended to serve not only as a guide for the 72nd ILSG field trip participants but also as a reference for those interested in reprising the trips at a future date. In order to facilitate this, trip leaders have provided UTM coordinates in the NAD 83 datum for stops, as well as plain word descriptions for locating each trip stop. It should be noted that some stops are located on private land or registered mining claims. As such, individuals visiting these stops are advised to obtain the land holders’ permission prior to entering their property. If in doubt, we recommend contacting the Resident Geologist Program office in Thunder Bay for further information about current property ownership. This year’s slate of field trips include stops located either on provincial highways or busy logging roads which can create safety issues. For those participating in facilitated trips at this year’s meeting, make sure to pay attention to the trip leaders’ safety orientation at the start of the trip, and follow any stop-specific safety instructions. For individuals using the field trip guide for future private tours it is advisable to be wary of road traffic and exercise extreme caution. Please take care when crossing or parking at the sides of these roads. The organizing committee would like to thank all the field trip leaders who authored and contributed to this field guide along with those who provided comments and/or assisted with the running of the trips themselves. Field trip leaders and authors include Howard Poulsen, Riku Metsaranta, Gaetan Launay, Dorothy Campbell, Justin Jonsson, Vittoria D’Angelo, Mark Smyk, Mark Puumala, Steve Kissin and Greg Paju. The Committee thanks participating exploration companies and mine operators for their cooperation and assistance in providing access and information in regards to their properties, as well as their staff time for leading the tour participants on their respective properties. Participating companies include Delta Resources, Gold X2 Mining, Amethyst Mine Panorama and Diamond Willow Amethyst Mine. Figure 1. Map illustrating general locations of ILSG 2026 field trips. Symbols are labelled with numbers that correspond to the trip numbers (1 to 6) used in the meeting program and field trip guidebook. -1- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Trips 1 & 4 - “Classic” Geological Sites in the Thunder Bay Area Mark Smyk Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada and Mark Puumala Geological Consultant, 370 Crossbow Court, Thunder Bay, Ontario, P7G 1H5 Canada Introduction The geology of the Thunder Bay area features a variety of Archean and Proterozoic rocks of the Superior and Southern Provinces of the Canadian Shield, respectively, as well as unconsolidated deposits and landforms associated with Quaternary glacial and post-glacial processes. This field trip features examples of many of these rocks and features, providing an overview of the varied geology the area has to offer. A number of field guides (e.g., Pye, 1969; Kustra et al., 1977; Franklin et al., 1982) have covered the Thunder Bay area, including those written for the 46th (e.g. Pufahl et al., 2000; Phillips et al., 2000) and 58th Institute on Lake Superior Geology annual meetings (e.g. Fralick et al., 2012; Smyk, 2012; Phillips et al., 2012; Cundari et al., 2012). These guides contain descriptions of some of the field trip stops covered in this guide and they will be referenced appropriately. This guide also benefits from ongoing local research and mapping conducted by the Ontario Geological Survey, Geological Survey of Canada and Lakehead University. Day One of this trip features exposures north and east of Thunder Bay, while those of Day Two are located south and west of the City. Bear in mind that this trip marks the first time that many of these stops have been visited and described as part of a formal field trip. This is especially true of stops along Highway 11-17, whose expansion ca. 2010-2012 produced many remarkable new exposures. Please exercise caution when stopping and viewing roadside outcrops. Permission or admittance may need to be obtained to visit some stops; this will be outlined in the guide when necessary. Regional Geology Overview Precambrian Geology The Thunder Bay area straddles the boundary between Archean rocks of the Superior Province and Proterozoic rocks of the Southern Province (Figure 1). In the vicinity of Thunder Bay, Superior Province rocks comprise volcano-plutonic rocks of the Neoarchean Wawa Subprovince and metasedimentary and granitoid rocks of the Neoarchean Quetico Subprovince, bounding the Wawa to the north. Locally, the supracrustal rocks of the Wawa Subprovince have been subdivided into the Greenwater and Shebandowan assemblages (Williams et al., 1991). The ca. 2.72 Ga Greenwater assemblage consists of a north-younging sequence of mafic to felsic metavolcanic rocks with subordinate interbedded clastic and chemical metasedimentary rocks. Mafic metavolcanic rocks within this assemblage consist predominantly of tholeiitic to calc-alkalic pillowed flows. The intermediate and felsic metavolcanic sequences are calc-alkalic and consist predominantly of coarse-grained pyroclastic deposits and massive to feldspar-phyric flows. The ca. 2.69 Ga Shebandowan assemblage is a younger, possibly fault-bounded (Williams et al., 1991) sequence of sub-alkalic to alkalic, predominantly coarse-grained pyroclastic metavolcanic rocks with interbedded coarse- to fine-grained, commonly well-preserved, proximal metasedimentary rocks. Intrusions within the Wawa Subprovince supracrustal assemblages consist of narrow felsic dikes, syenitic to tonalitic pre- to syntectonic plutons, minor gabbro bodies and scattered narrow mafic dikes. Rocks of the Neoarchean Quetico Subprovince abut the Wawa Subprovince to the north. They consist mainly of clastic metsedimentary rocks (turbiditic wacke, arkose, quartz arenite, slate and argillite) as well as post- to syn-deformational, syenitoid to granitoid plutons (cf. Metsaranta, 2022; Metsaranta and Walker, 2019). Migmatization becomes common in the rocks towards the northern portion of the area as metamorphic grade increases (Williams, 1991). -2- The Southern Province consists of Proterozoic �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 1. Generalized geology of the Thunder Bay area; geology from Ontario Geological Survey Map 2542, Bedrock Geology of Ontario, West-Central Sheet, scale 1:1 000 000 (1991). rocks which unconformably overlie or intrude Archean basement rocks of the southern Superior Province (cf. Tanton, 1931; Pye, 1969). North and west of Lake Superior, the Southern Province comprises: 1) Paleoproterozoic (ca. 1.8 Ga) Animikie Group sedimentary and minor volcanic rocks; 2) Mesoproterozoic (ca. 1.4 Ga) Sibley Group sedimentary rocks; and -3- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 3) Mesoproterozoic (ca. 1.1 Ga) Midcontinent Rift volcanic and intrusive rocks. The Animikie Group, exposed in Ontario, Minnesota, Wisconsin and Michigan, is represented locally by the Gunflint Formation and overlying Rove Formation. These dominantly sedimentary formations constitute a largely unmetamorphosed, undeformed, homoclinal succession which dips shallowly to the southeast. The Gunflint Formation is a chemicalclastic assemblage which yielded a U-Pb age from reworked volcanic ash of 1878.3 ± 1.3 Ma (Fralick et al., 2002). Rocks containing intraformational breccias, accretionary lapilli, spherules and shocked quartz that occur near the top of the Gunflint Formation are interpreted to represent ejecta from the Sudbury impact event that occurred at circa 1850 Ma (Addison et al., 2005; Krogh, Davis and Corfu 1984). The Gunflint Formation grades upward into turbiditic sandstone and shales of the Rove Formation south of Thunder Bay. U-Pb zircon ages from ash beds in the basal Rove Formation yielded 1836+5 and 1832+3 Ma (Addison et al., 2005). A sandstone sample from the submarine fan portion of this succession yielded a youngest detrital zircon U-Pb age of approximately 1780 Ma (Heaman and Easton, 2006), but this relatively young age is widely considered to be problematic and not reflective of the true age of these rocks. Sedimentation in this part of the Animikie basin, widely thought to represent the distal foreland of the Penokean Orogen, likely ended ca. 1800 Ma or earlier. However, Maric (2006) suggested that the Rove (and correlative Virginia) Formation represents the transition from a sedimentstarved basin, with exceedingly slow deposition rates, to active deltaic progradation with sediment probably derived from the Trans-Hudson orogenic zone to the north. The Sibley Group, exposed on the Sibley Peninsula and farther north, has been subdivided into five formations; detailed descriptions of each formation have been reported previously (Franklin et al., 1980; Cheadle, 1986; Rogala, 2003; Rogala et al., 2005, 2007). The overall sedimentary environment indicates a fluctuating climatic scenario, in which the Sibley Group was deposited in a lacustrine system (Pass Lake Formation) that gradually evolved into a saline playa lake environment (Rossport Formation). As the climate progressively became drier, a sabkha-type environment developed (Fire Hill Member of the Rossport Formation). After a break in time, the Kama Hill and Outan Island formations represent outbuilding of a large deltaic complex to the north (Jones et al., 2022), and the Nipigon Bay Formation represents an aeolian environment (Rogala, 2003; Rogala et al., 2007). The depositional age for much of the Sibley Group had been constrained between ~1340 and 1450 Ma. The northern margin of the Midcontinent Rift is dominated by mafic hypabyssal rocks of the Midcontinent Rift Intrusive Supersuite (Miller et al. 2002), which intrude all Proterozoic rocks and Archean basement. South of Thunder Bay, Logan (1106.3+2.0 Ma; Smith et al., 2025) diabase sills predominate. Nipigon diabase sills (1108.2+0.9 Ma; Bleeker et al., 2020) occur in and north of the City, and form the bulk of the Nipigon Embayment. Volcanic and minor sedimentary rocks of the ca. 1108 to 1105 Ma Osler Group (Davis and Sutcliffe, 1985; Davis and Green, 1997) are exposed to the east on Black Bay Peninsula and on offshore islands in Lake Superior. While all aforementioned rocks are related to the Early Magmatic Stage (ca. 1110–1103 Ma; Miller and Nicholson, 2013) of Midcontinent Rift development, younger intrusions (ca. 1097-1092 Ma; Smith et al., 2025) are associated with a magmatic episode that followed the emplacement of the 1099 Ma Duluth Complex. Three main domains were suggested by Smith et al. (2025) in describing the northern flank of the Midcontinent Rift west of Thunder Bay and Lake Nipigon, namely, from south to north: (1) a gently south-(southeast-)tilted Midcontinent Rift margin; (2) a pronounced basement arch just north of Thunder Bay, likely representing a flexural bulge; and (3) the erosional remnant of the Lake Nipigon rift-and-sag basin, preserving the Sibley Group intruded by extensive Nipigon diabase sills (Figure 2). Quaternary Geology The first Pleistocene ice sheet in the Thunder Bay region, ca. 1 Ma, moved over and stripped a deeply weathered, relatively flat bedrock landscape (Zaniewski et al., 2020). During the Pleistocene, possibly ten or more major advances and retreats of ice took place, each with its own history of advancing and retreating lobes of ice. The region’s present landscape is the product of interplay between three major ice lobes (i.e. Patricia or Rainy River Lobe, from the north; the Hudson Bay Lobe, from the northeast; and the Superior Lobe, from the east) originating from three accumulation centers -4- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 2. West-looking cross section through the northern flank of the Midcontinent Rift, just west of Thunder Bay and Lake Nipigon, from Smith et al. (2025) during Wisconsin glaciation. Only the final event, comprising the Marquette Readvance (ca. 11 500 Ka) and its subsequent retreat, is understood in local detail (ibid). remains of the toolkit of these people other than a variety of knapped lithic tools made from taconitic chert that occurs in the local Gunflint Formation (cf. Hamilton, 1996). Starting about 11,000 years ago (Ka), Wisconsin ice melted back from its position in central Minnesota and Wisconsin, and quickly exhumed the Thunder Bay region, forming recessional moraines during brief stillstand periods (Phillips, 2004; Phillips et al., 1994). The Lake Superior basin was occupied by Early Lake Minong, the shoreline of which is found close to the 1400-foot (427m) contour in the borderland area. About 10 Ka, ice re-advanced from north of Lake Nipigon, sweeping across the Superior Basin (Marquette Readvance). As that ice began to melt, glacial lakes were formed between the moraines and the retreating ice margins. As Superior ice melted, water levels progressively lowered, forming a series of shoreline features down-slope and depositing thick lacustrine clays. Superior ice withdrew to the north of Lake Nipigon around 9.5 Ka, and for the first time since the Marquette Re-advance, the Superior basin was occupied by a single lake, Lake Minong. This lake level extended up the Kaministiquia embayment to Rosslyn, where a large delta structure was built. The Minong shoreline runs through the upper part of the city, being particularly evident in Boulevard Park where river mouth bars and terraces of the Current River are seen. The Minong shoreline in the city is strongly associated with Palaeo-Indian sites, the Cummins Site being the best-known. It is likely that as water levels fell, these early people moved down from the Arrow-Whitefish Lakes area into the Kaministiquia embayment. Little Field Trip Stop Descriptions - Day One Day One begins with visits to a number of locations northeast of the City, clustered around the northern end of Thunder Bay of Lake Superior (Figure 3) and ends near and within the City (Figure 4). This small area is underlain by a variety of rocks that record almost three billion years of local geologic history, spanning from the Neoarchean (ca. 2.7 Ga) to the Paleoproterozoic (ca. 1.8 Ga) and Mesoproterozoic (ca. 1.4 and 1.1 Ga) and perhaps to the Mesozoic (ca. 100 Ma). Unconsolidated glacial and post-glacial deposits and features attest to a long-lived, Pleistocene glaciation record. All GPS coordinates are NAD83, UTM Zone 16. STOP 1-1: Blende Lake Unconformity (0367703 E / 5383357N) This exceptional highway rock cut, like many others on this stretch of Highway 11-17, was exposed by new highway excavations ca. 2012. This ~700 m-long exposure features the unconformity between Neoarchean Wawa metavolcanic and gabbroic rocks and Paleoproterozoic sedimentary rocks of the lower Gunflint Formation (cf. Scott, 1990; Figure 5). Basement rocks here have also been described by Landman (2021) as coarse-grained amphibolite, interpreted as a mafic intrusion which has undergone amphibolite-facies metamorphism. -5- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 3. Generalized geology of the northern end of Thunder Bay of Lake Superior, showing the first 9 field trip stop locations of Day 1. Geology from Map M2232 (Carter et al., 1973). BLF – Blende Lake Fault This exposure was described by Metsaranta and Kurcinka (2022), as part of an ongoing Ontario Geological Survey (OGS) bedrock mapping project of the Animikie Basin near Thunder Bay: “…chloritized Archean felsic to intermediate intrusive rocks are locally overlain by at least 4 stromatolite mounds comprising thinly laminated black to red chert [Figure 6]. The stromatolite mounds have a height of up to 30 to 50 cm and similar widths. The top of one mound is marked by a thin stylolitic band [Figure 7]. The areas between stromatolite mounds comprise silicified grainstones that locally contain sulphide nodules up to 5 cm in diameter. The grainstones enclosing the stromatolite form medium to thick beds characterized by medium- to large-scale trough cross-stratification. At this locality, an east-dipping and roughly north-striking small displacement thrust fault puts Archean basement rocks above Gunflint Formation rocks. The fault appears to displace Midcontinent Rift–related quartz-carbonate-sulphide veins indicating Figure 4. Field trip stop location map, showing Day 1 stops 1-10 and 1-11 and Day Two stops. (See Figure 3 for map legend). -6- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 5. Paleoproterozoic Gunflint Formation sedimentary rocks unconformably overlying Archean basement, east side of Highway 11-17, STOP 1-1. that the thrusts may be related to rift inversion; however, this field relationship is equivocal.” These ferroan dolomite and siderite grainstones (medium-grained, sand-sized iron carbonates), referred to as granular iron formation, are common in the Thunder Bay region, dominating the near-shore of the Animikie Basin (see Fralick et al., 2012; STOP 1-4). The thin basal conglomerate (aka Kakabeka Member, Figure 8) of the Gunflint Formation is discontinuously distributed along the paleosurface. As described by Metsaranta and Kurcinka (2022), the conglomerate has a clast-supported texture, consisting of coarse-grained sand, granules, pebbles and rare cobbles in a sandy matrix. Quartz, pink granitoid, clastic metasedimentary and mafic metavolcanic clasts were noted. Figure 6. Mound-shaped stromatolites, with onlapping grainstones, resting on chloritized Archean basement, STOP 1-1. The stromatolite has a black chert core, and red, jaspilitic outer layers. Photo from Metsaranta and Kurcinka (2022). varies from 0 to 30 cm in thickness here, and is usually absent from the local topographic “highs” (i.e. knobs or ridges of the Archean basement), but may thicken in depressions in the paleosurface. The basal conglomerate lag may contain large, well-rounded boulders, up to 0.5 m in diameter; these have been observed only on the northwest side of the highway, opposite STOP 1-1. Black, cherty bands (0.1–1.5 cm) occur locally in parts of the Kakabeka conglomerate (ibid). Recent geochronologic study of the conglomerate shows that the main population of detrital zircons is consistent with derivation from local Neoarchean intrusions (R. Kup et al. (2025) also noted that the conglomerate Figure 7. Stylolites in Gunflint Formation grainstone, STOP 1-1, visible as a black serrated band to the right of the scale card. Figure 8: Quartz pebble-rich Kakabeka conglomerate, STOP 1-1. Photo from Metsaranta and Kurcinka (2022). -7- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Metsaranta, personal communication, 2026) like the Mackenzie granite (ca. 2672 Ma; Puumala et al., 2015). The stromatolites which occur in the basal unit of this Gunflint section (Unit 1 of Kup et al., 2025) were described by Kup et al. (2025): “Unit 1 of the Gunflint Formation is characterized by the common occurrence of cherty, domical to columnar stromatolites within the basal 2 m of the formation. Relatively large stromatolite heads tend to occur preferentially on knobs or ridges of the Archean basement, usually metres apart from one another. Smaller stromatolite heads (5–15 cm in diameter or height) are present locally, even on top of the Kakabeka conglomerate. Tabular, finely undulatory, microbialite-like structures (commonly <10 cm thick) may occur in places, or laterally connected to stromatolite heads. The microbialite-like structures may weather to a reddish colour, similar to some of the weathered stromatolite heads. Overall, the colour of the domical to columnar cherty stromatolites ranges from red, yellow, white, grey and black, depending on the iron content and the nature of weathering. Relatively large (usually fresh) stromatolite heads tend to be jet black near their centre and change to lighter grey and white toward their edges; however, edges themselves are commonly marked by red and/or yellow banding. The stromatolites are most easily seen and accessible toward the southern ends of the outcrop, directly above the road ditch level.” galena, have been noted. The fault is also exposed on the other side of the highway; Gunflint rocks are folded next to the fault there as well. The east-northeast orientation of the Blende Lake Fault is similar to other structures on and north of the Sibley Peninsula, some of which host gabbroic dykes Figure 9. Rock cut exposure of the Blende Lake Fault, east side of Highway 11-17 (STOP 1-2), separating folded Gunflint Formation rocks (left) from Neoarchean basement (right). The fault zone is cored by calcite vein / vein breccia; fault gouge flanks the vein. Field notebook for scale. Optical and SEM imaging data collected by Kup et al. (2025) suggest that well-preserved Gunflint-type microfossils (both filamentous and coccoid types) tend to occur in sporadic pockets in samples collected from this locality. attributed to the waning stages of Mesoproterozoic Midcontinent Rift magmatism. Scott (1990) noted that Gunflint rocks, normally flat-lying or gently southeastdipping, are folded and brecciated to a large extent in the area between Blende Lake and O’Connor Point on Lake Superior. Folding described by Moorhouse (1960) east of Blende Lake, was attributed to farfield Penokean fold-and-thrust deformation by Hill and Smyk (2005), prior to the recognition of the Sudbury Impact Layer and associated deformation in the Animikie Basin. Koroscil (2013) noted that thrust faults, once ascribed to Penokean deformation, cut the SIL at the Terry Fox Monument (STOP 1-10) and thus may post-date the Penokean. Landman (2021) noted that: STOP 1-2: Blende Lake Fault (0368005 E / 5383802N) The northern end of the same rock cut, approximately 500 m north-northeast of STOP 1-1, exposes the eastnortheast-striking Blende Lake Fault (cf. Scott, 1990). Rusty fault gouge occurs between Archean gabbroic rocks to the south and folded, flaser-bedded Gunflint wacke and siltstone, cored by a 3 m-wide calcite + quartz vein / vein breccia which contains Gunflint fragments (Figure 9). Base metal sulphides, including -8- “Later Proterozoic features, including the Blende Lake fault, have a common strike of eastnortheast, which aligns with the orientation of the 1.1 Ga Mid-Continent Rift in Thunder Bay. This similarity is further reflected by the Blende Lake fault being oriented subparallel to silver veins related to the Mid-Continent Rift. Similarities between orientations of brittle structures in the [Neoarchean] amphibolite and Gunflint Formation suggest that the Mid-Continent Rift �Proceedings of the 72nd ILSG Annual Meeting - Part 2 in Thunder Bay may have reactivated some Archean-aged, orogenic-related faults and shear fractures. Minor folding in the Gunflint Formation truncated by the Blende Lake fault, as well as reverse reactivation along the plane, may be evidence of compression during the later stage of the Mid-Continent Rift.” STOP 1-3: Mirror Lake Turn-off (0370744 E / 5387262N) This stop displays a number of quite enigmatic features that are still being evaluated in the context of evolving local geologic ideas. The hillside exposes iron-rich, folded and brecciated Gunflint Formation sedimentary rocks that have been intruded by a Nipigon diabase sill. Tight to isoclinal, plunging to recumbent folds have developed in certain parts of the otherwise ~flat-lying, thinly to thickly bedded, taconitic, martite(?)-bearing grainstones (Figure 10). As is the case at STOP 1-2, the cause of the deformation in the Gunflint rocks is a matter of debate. There is growing support that local folding and brecciation may be related to far-field effects generated by the Sudbury meteor impact ca. 1850 Ma., postdating the end of Gunflint deposition by perhaps ca. 20 My and preceding the onset of Rove sedimentation. Rocks interpreted as being part of the Sudbury Impact Layer were noted in geotechnical drilling ~ 1 km north of this location (P. Fralick, personal communication, 2025). Despite the fact that the Sudbury impact structure is ~660 km away and that the most dramatic deformation is usually crater-proximal (i.e. within ~5 crater radii), Addison and Brumpton (2012) noted that the Thunder Bay area would have still experienced dramatic impact-induced effects, including magnitude 10.7 earthquakes and likely tsunamis. Alternatively, it has also been suggested that some of this deformation may be related to the dominantly extensional stress regime associated with Midcontinent rifting. Local compressional (contractional) structures may form within relay zones between overlapping normal fault tips, particularly as the fault segments grow, interact, and prepare to connect. While normal fault systems are dominated by horizontal extension (pulling apart), the 3D interaction and rotation of blocks in the relay zone (or “relay ramp”) can create local stress perturbations that lead to shortening, folding, and antithetic faulting (cf. Camanni et al., 2023). Further work is required to better map the extent and character of folding and brecciation in order to suggest deformation mechanisms and causative factors. Rove shales and wackes (e.g. STOP 1-5), mapped ~2.5 km south of here by McIlwaine (1975), are not deformed. They overlie Gunflint rocks and are disconformably overlain by sandstones of the Pass Lake Formation of the Sibley Group (STOP 1-6). Figure 10. Recumbent fold in Gunflint Formation chertcarbonate rocks, STOP 1-3. Folded bedding planes are traced by dashed lines. A thin veneer of Phanerozoic(?) conglomerate (cgl) occurs on outcrop surfaces and in crevices. Two other enigmatic rocks are exposed at this location; both are conglomerates. One conglomerate occurs as thin coatings plastered on exposed outcrop surfaces and in fractures in the folded Gunflint rocks (Figures 10 and 11). It is a brown, poorly sorted, sandy, matrix-supported unit. Sibley Group (ca. 1.4 Ga) sedimentary rock clasts, ranging from sub-angular to rounded pebbles and cobbles, predominate. Most recognizable are rust-red calcareous siltstones of the Rossport Formation (with their characteristic pale reduction spots) the base of which occurs approximately 75 m stratigraphically above the Gunflint exposed here. It must also be noted that medium-grained mafic igneous clasts appear to be ca. 1.1 Ga Nipigon diabase -9- Given the presence of Sibley and Nipigon diabase �Proceedings of the 72nd ILSG Annual Meeting - Part 2 and/or enrichment (cf. Fralick and Riding, 2015). Interestingly, the exposure of stromatolitic Gunflint rocks on the west side of Highway 11-17, just opposite the Mirror Lake turnoff, only 200 m away from STOP 1-3, is notably rusty and apparently more oxidized than the vast majority of Gunflint rocks. This may represent Mesozoic paleoweathering, raising the possibility that Cretaceous deposits and paleoweathering effects may have once extended as far east as the Thunder Bay area. Any rocks that survived Pleistocene glaciation may survive in isolated patches or have been as yet unrecognized. The second conglomerate (Figures 12A and 12B) is similarly exposed as a plastered veneer draped on the exposed outcrop face. Unlike the other unit, it appears to be clast-supported and monomictic; angular, dark grey, fine-grained, shaly Gunflint fragments are cemented by calcite. This monomictic clast population suggests local derivation, perhaps a talus deposit created and cemented during the Pleistocene. Figure 11. Close-up of thin veneer of conglomerate on Gunflint substrate, showing reddish-orange Rossport Formation siltstone and Nipigon diabase fragments. clasts, the conglomerate must postdate at least Midcontinent rifting, the hitherto youngest geologic event in the local lithologic record. Although no Phanerozoic rocks have been documented in the Thunder Bay region, Cretaceous rocks have long been known to overlie the Biwabik Formation (time-correlative with the Gunflint) on the Mesabi iron range of northern Minnesota (e.g. Bergquist, 1944); “soft” iron ores there formed there during the Cretaceous. Paleomagnetic studies by Purucker (1983) in the Eldorado Beach – Nelson roads area, ~6.5 km southwest of this location, suggested that secondary enrichment of Gunflint and Mesabi iron ores took place at approximately the same time between Aptian and Cenomanian time (ca. 12594 Ma). In their study of anthraxolite in the Gunflint Formation in the Kakabeka Falls area, Hayatsu et al. Remnants of a Nipigon diabase sill form prominent, cuesta-like hills in the vicinity of this stop and around Deception and Mirror lakes (McIlwaine, 1975). Smooth, glacially polished surfaces with striae are visible at the road level in this cliffside exposure. STOP 1-4: Gunflint Formation, Blende Creek area (0369581E / 5383837N) This stop description, featuring deformed chertcarbonate units in the Gunflint Formation (Figure 13), is taken from Fralick et al. (2012): (1983) identified two very distinct macromolecular materials. These two hydrocarbon fractions were thought to represent derivation from sediments of two vastly different ages: an older one, characterized by heavier aromatic ring compounds, derived from Gunflint-aged organic remains; and another, aliphatic fraction derived from Cretaceous (or possibly Jurassic) sediments. Cretaceous microfossils were described in lateritic “buckshot” ore in the Archean Steep Rock Lake iron deposit near Atikokan (Machado, 1987) that underwent Mesozoic karstification, weathering - 10 - “Along Highway 587, rock cuts display thinly bedded, generally flat-lying sedimentary rocks of the Gunflint Formation. The outcrops we have driven past are composed of ankerite and siderite grainstones (medium-grained, sandsized iron carbonates) referred to as granular iron formation (GIF). These are common in the Thunder Bay region, dominating the near-shore of the Animikie Basin. The iron carbonate grains were produced by wave erosion of carbonate precipitates and represent storm deposits in the near-shore. The iron may have precipitated as a carbonate in this shore-proximal zone due to photosynthesizing bacteria removing CO2 from the water and thus increasing the pH and driving the carbonate phase into supersaturation. The outcrop we are looking at has these carbonate �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 12. A (left). Clast-supported, calcite- cemented conglomerate veneer over Gunflint, STOP 1-3. B (right). Close-up view of conglomerate / breccia in Figure 12A. grainstones weathering orangey-brown alternating with white chert layers. In places the chert can be seen replacing the carbonate but other layers appear to be primary chert. In the older literature an outcrop such as this would be ascribed to deeper water due to less evidence of current activity. However, because of its shore proximal location it probably formed in a Figure 13. Folded Gunflint Formation grainstones, north side of Highway 587, STOP 1-4, with locally axial-planar quartzcarbonate veins. - 11 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 quieter water location near the strand-line, i.e., a sheltered lagoonal area behind on offshore bar. These exposures are somewhat unique in that the rocks are folded; elsewhere, they are undeformed. The hinge zones, where the majority of stress is focused, are commonly fractured. These fractures may be occupied by quartzcalcite veins following the vertical axial plane. The outcrop to the west hosts numerous veins and vein breccias that strike between 40° and 45° and dip almost vertically to the southeast. These breccias contain sparry calcite, drusy quartz and also altered shale fragments, suggesting that these Rove Formation rocks likely occurred above this section during vein emplacement. A thin, northwest-dipping diabase dyke intrudes the Gunflint rocks at this location and is, in turn, cut by these veins.” Deformational features in Gunflint Formation rocks near Pass Lake have been previously ascribed to Penokean fold-and-thrust activity in the foreland (i.e. passive margin Archean basement + Gunflint Formation; Hill and Smyk, 2005). These include discrete bedding-plane faults with locally developed gouge and breccia that can be traced laterally into horizontal, hanging wall ramps with associated fault-bend folding. Previous workers had also ascribed folding to synsedimentary slumping and Keweenawan diabase sill emplacement and thought that they were attributable to local, rather than regional-scale, deformation. As introduced at STOP 1-3, there is growing support for the contention that such deformation may be related to the Sudbury impact event ca. 1850 Ma. Figure 14: Flowerpot-shaped Rove Formation concretion on wall of inactive shale quarry at STOP 1-5. a piece of organic material or other foreign object, which creates a perturbation in fluid flow with a distinct chemistry. Because the cementing agent in this case is more resistant to weathering, these concretions stand out of the soft shale and may commonly completely detach form their host rock. Groundwater and surficial water flow through the shale has led to the dissolution and subsequent precipitation of a variety of low-temperature minerals (e.g. carbonates, sulphates, hydroxides) that occur as white and yellow encrustations on the bedrock surface. One of the more unusual of these secondary minerals is yellow magnesium aluminocopiaptite ((Mg,Al) (Fe,Al)4(SO4)6(OH)2.20H2O; Resident Geologist’s Files, Thunder Bay).” STOP 1-5: “Devil’s Flower Pots” (Rove Formation concretions) 0370841E / 5382426N This stop description is taken from Fralick et al. (2012): “Just north of Highway 587, a quarry face exposure of black, fissile Rove Formation shale displays lenticular and elliptical concretions, flattened along bedding planes [Figure 14]. These structures form during diagenesis, following initial compaction and dewatering of the sediments. They represent a concentration of a cementing agent (e.g., silica, calcite) focused during the migration of fluid through the sediments. They often are nucleated around STOP 1-6: Edwards Road section (Pass Lake and Rove formations) 0371737E / 5382460N (n.b. private property; permission is required to access) This stop provides us with an excellent stratigraphic section that extends upward from the top of the Animikie Rove Formation into the Pass Lake Formation, the lowermost formation of the Sibley - 12 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 and sandstones of the Loon Lake Member, Pass Lake Formation, overlie the Rove. The conglomerate and sandstone layers are laterally discontinuous, with some conglomerates in clast-support (fluvial deposits) and some in matrix-support (sub-aerial debris-flow deposits). Successions such as this in the Sibley are typical of arid to semi-arid alluvial fans (Cheadle 1986), though this would have been a very small one. The abundant hematite probably denotes a deep water table. Clasts are locally derived from Group. The youngest detrital zircons in the Sibley Group are ca. 1.4 Ga (Rogala et al., 2007). An Rb– Sr isochron age of 1339 ± 33 Ma was determined on dolomitic mudstones from the Rossport and Kama Hill formations (Franklin 1978; Franklin et al. 1980). Recent studies of some of the concretions (quartzcarbonate + various very fine-grained impurities and inclusions) in the Pass Lake Formation, associated with late advanced diagenesis, had enough uranium to generate an age of 1483+4 Ma (W. Bleeker and H. Rochin-Banaga et al., unpublished data / personal communication, 2025; Figure 15). Together with the ca. 1500 Ma youngest detrital zircons (SHRIMP data on 3 samples; ibid), this provides a greatly improved age constraint, just marginally younger than 1500 Ma, on the deposition of the lower part of the Sibley Group. This stop description is taken from Fralick et al. (2012): “A private access road extending up the mesa provides an excellent 150 m long section exposing the disconformity between the Rove Formation and the overlying basal conglomerate and sandstones of the Pass Lake Formation [Figures 16 and 17]. The Rove shales immediately below the contact were subject to Mesoproterozoic weathering. Geochemical investigations have outlined an oxidized zone below the contact grading to a more reduced zone with abundant chlorite a few tens of centimeters lower in the section. In one area what may be a dewatering or degassing structure strongly deforms the shale. Very immature, iron oxide-rich conglomerates Figure 16. Cobble-sized clast of Gunflint Formation stromatolitic jasper/chert visible in the Loon Lake Member conglomerate exposed along the Edwards Road section, STOP 1-6. Figure 15. U-Pb concordia diagram presenting geochronological data for the Sibley Group (W. Bleeker and H. Rochin-Banaga et al., unpublished data, 2025) Figure 17. Disconformable contact (just above hammer) between weathered green Rove Formation shales and hematite-rich, basal conglomerate and sandstone of the 6-7 m thick, Loon Lake Member (Pass Lake Formation), STOP 1-6. Overlying, well-sorted, buff sandstones of the Fork Bay Member form the top of the exposure. - 13 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 the erosion of underlying units. This is sharply overlain by mature, well-sorted, medium-grained sandstones of the Fork Bay Member, Pass Lake Formation. Detrital zircon geochronology and paleocurrents (Cheadle 1986; Rogala et al. 2007) indicate that the major source of this sediment was the Trans-Hudson highlands. The travel distance accounts for its maturity compared to the locally derived underlying conglomerates. The sandstone was deposited as sheet flows into the shallow nearshore of a lacustrine system that had flooded the area (Cheadle 1986; Rogala 2003; Metsaranta 2006; Rogala et al. 2007). These sandstone layers are laterally continuous, massive to parallel-laminated, in places with trough cross-stratified or rippled tops [Figure 18]. Rare, odd features are present both in crosssectional and bedding plane views in this outcrop. These may be dewatering pipes.” exposed on this outcrop surface. These include two sets of glacial striae at 040˚ and ~060˚ and subparallel arrays of crescentic gouges and chatter marks/lunate fractures (Figure 19). En route to STOP 1-7, the highway traverses a series of baymouth bars that formed as lake levels fell from the Lake Beaver Bay stage (ca. 11 to 10.5 Ka) to the Lake Minong stage (ca. 10.5 to 8.5 Ka), connecting what had been the “island of Sibley” to the mainland near Pass Lake (Zaniewski et al., 2020; Geddes et al., 1987; see Fralick et al., 2012). This new connection formed an ideal natural trap for Palaeo-Indian hunters to use. The materials excavated at the Brohm archaeological site, on the top of the main baymouth bar, were all hunting-related projectiles and scrapers, many made onsite from chunks of jasper taconite that they carried with them from quarry sources (Zaniewski et al., 2020; MacNeish 1952). Fralick et al. (2012) noted that the matrix-supported conglomerate was probably deposited as a high-density mass-flow while the boulder-cobble, matrix-supported conglomerate probably represents a very high-viscosity mass flow as the larger clasts were suspended near the top of the flow. Upper flow regime, parallel-laminated sandstones were probably deposited by sheet-floods on an alluvial fan’s surface and are interbedded with clastsupported fluvial conglomerate. The top of the hill affords a tremendous view of Thunder Bay, Sibley Peninsula and offshore islands. A south-dipping diabase sill forms the prominent cuesta of Caribou Island. Glacial erosional features are Figure 18. Medium- to coarse-grained, well-sorted sandstone bed of the Fork Bay Member, top of hill, STOP 1-6. The majority of the bed is upper flow regime parallel laminated, with a reworked, cross-stratified top. Figure 19. Glacial erosional features exposed in the sandstone outcrop surface at the top of the hill, STOP 1-6. These include glacial striae (dashed arrows), concave up-ice crescentic gouges (CG) and concave down-ice chatter marks (CM). - 14 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 STOP 1-7: Pass Lake section (Loon Member conglomerate) 372282E / 5380560N The cliffs adjacent to the abandoned railway at Pass Lake is the type section for the Pass Lake Formation. Exposure is almost continuous for 3.2 km along the tracks and provides a ~50 m-thick stratigraphic section. Rove Formation shales, exposed at the northwestern end of the cliff exposure, disconformably underlie the Pass Lake Formation but are not exposed here. This cliff face, a popular destination for local rock climbers, exposes the basal Loon Member conglomerate and overlying, buff sandstones of the Fork Bay Member (Pass Lake Formation; Figure 20). A description was provided by Fralick et al. (2012): “The basal conglomerate thins and thickens laterally, pinching down to pebbly sandstone in places. Clasts are generally surrounded and dominated by local Gunflint Formation lithologies. The matrix is poorly sorted. The conglomerates are overlain by a thinningupward sequence of sandstone beds capped by siltstones on the top of the cliff. Individual beds are reasonably laterally continuous though sometimes lens out. They are dominated by upper flow regime parallel lamination with occasional ripples and small-scale dunes on their tops. Figure 20. Pass Lake section exposure of Loon Lake Member conglomerate overlain by Fork Bay Member sandstone at STOP 1-7. STOP 1-8: Neoarchean Pyroclastic and Clastic Sedimentary Rocks 360568E / 5379943 Both alluvial fan-braided fluvial and shallow lacustrine (Cheadle, 1986; Franklin et al., 1980, respectively) depositional environments have been proposed. The bedding organization of the conglomerates exposed here is somewhat different than those observed earlier. This opens the possibility that the conglomerates at this location were reworked by wave activity during initial lacustrine flooding. This 100 m-long rock cut on the northwest side of Highway 11-17 was exposed by highway construction ca. 2012. It features a remarkable exposure of Neoarchean pyroclastic and clastic sedimentary rocks of the Shebandowan greenstone belt that strike ~140˚ and dip steeply northeast (Figure 21). These supracrustal rocks are intruded by granitoid rocks of the McKenzie granite and may represent a large pendant within the intrusion. The exposure was the subject of an undergraduate thesis by Bjorkman (2014), from which most of the descriptions will be gleaned. The sandstone beds again represent sheetfloods, forming sand-flats in the shallow lake. The thinning- and fining-upward sequence of sandstone beds is a classic example of a transgressive succession showing decreased sand supply through time as the shoreline moves further away from the area.” Bjorkman (2014) identified 13 lithofacies/lithologic units in this complex section: Overlying red-orange Rossport Formation siltstones begin to outcrop approximately 1.2 km east of STOP 1-7 (see Fralick et al., 2012 for stop descriptions). This exposure exemplifies the close connection between Neoarchean pyroclastic activity and sedimentation (Figure 22). Bjorkman (2014) suggested that these rocks were deposited in a vent-proximal environment, a contention supported by the presence of graded ash beds, high-velocity base surge deposits and impact structures from pyroclastic bombs (Figure 23). - 15 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Lithofacies / Interpretation 1 Lahar – debris flow 2 Ash fall 3 Fluvial reworking and base surge deposits 4 Ash fall 5 Lag deposit 6 Slump 7 Lahar 8 Lahar / channelized debris flow 9 Ash tuff 10 Hornblendite 11 Hornblendite 12 Hornblendite 13 Syenite Description Unsorted, clast-supported, conglomerate; clasts are pebble-sized to small boulder-sized within a medium-grained, sand-sized matrix Fine-grained to medium-grained, sandy, continuous beds; parallel laminated with no crossbedding, with an average thickness of 1 cm or less Cross-bedded and graded, medium-grained sandstone beds, which occur in alternating sequences; this unit has a sharp basal contact with the lower Lithofacies 2 and a poorly defined contact with upper units. The transitions from the graded beds to the cross-stratified beds are distinct. Parallel-laminated, continuous layers of medium-grained, graded beds, more rarely observed to be cross-bedded at very low angles. Clast-supported conglomerate in which clasts are very uniform and commonly cobble-sized. The clasts are flattened and oval-ellipsoidal, with rounded edges. The long axes of the clasts occasionally have tail-like tips. Disturbed beds of material very similar to that found within Lithofacies 4. The parallel-laminated strata are disrupted by failure of the slope and are truncated by an angular disconformity of the overlying unit. There are rubble blocks adjacent to the truncated strata. These blocks of failed beds lie along the base of this facies, with no evidence of sorting after the failure. There is no grading in the matrix, which is massive, medium-grained sandstone. Repetitive sequences of normally graded, medium-grained sandstone beds, gravelly matrix supported beds, and non-graded massive beds composed of medium-grained sand. Discontinuous, lens-shaped beds are very common. The normal graded beds are composed of coarse-grained sandstone, which grades into fine-grained tops of beds. These often have eroded, scoured tops, with very distinctly defined bases. Cross-bedding is common. Massive graded conglomerate with angular to very rounded and moderately flattened clasts. The clasts appear monomictic and range in size between 5-15 cm, the majority being 12 cm by 7 cm. The clasts make up 70% of the total composition, while the matrix is mostly a uniform mediumgrained sandy composition, with 10% coarser sand-sized fragments. The unit is on average 3-5 m wide. Fine-grained, sand-sized matrix with medium-grained and subhedral porphyritic feldspar crystals. The weathering surface is very irregular as the feldspars stand-out from the matrix. The unit occurs sporadically, locally intruded by dark green material. The average thickness varies between 50 cm to 100 cm. There is no grading throughout this unit, and the unit conforms to the same stratigraphy as the surrounding units, which is most often Lithofacies Association 1. Medium-grained dyke which crosscuts stratigraphy. The largest of the intrusive dikes, it can be traced through the entire outcrop. It is distinguished by the very irregular shape of its contacts with the host rock. The matrix consists of green equigranular, subhedral crystals in the middle of the intrusion and lighter altered plagioclase crystals along finer-grained contacts. The body intrudes (brecciates) itself where the dike dilates. Other smaller dikes crosscut this one. Cobbleto boulder-sized xenoliths were noted. Medium-grained, green-grey mafic dyke, 5-7 cm wide, with equal amounts of mafic and felsic minerals. The dike cuts through the green intrusive veinlets. A set of dark green-grey dykes, up to 0.5 m in width, striking approximately the same direction as Lithofacies 10; may be sill-like intrusions, wispy and infiltrating intrusions which engulf clastic material. This unit commonly contains wall rock xenoliths, which are very sharp and angular. Medium- to coarse-grained, subvertical and east-southeast-striking dykes. The dykes are the youngest rock type in the outcrop and are noted regionally. They are composed of red feldspar, amphibole, biotite, and quartz. The red feldspar gives the rock a brick-red colour. A combination of subaerial and shallow subaqueous conditions likely existed at the time of deposition, with fluvial reworking and deposition occurring during periods of volcanic dormancy. Phreatomagmatic processes, similar to those that produce maar craters, likely predominated. The calc-alkalic geochemistry (Figure 24), pyroclastic volcanism and subaerial/shallow water to fluvial clastic sedimentary rocks suggest that these rocks are part of the younger Shebandowan assemblage - 16 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 21. Map of the main outcrop, STOP 1-8 (Bjorkman, 2014). Lithofacies Association 1 through 9 are (resedimented) pyroclastic and clastic sedimentary units; lithofacies association 10 through 14 are crosscutting dykes. of the Wawa subprovince. Shebandowan rocks (ca. 2690-2680 Ma), unconformably overlying the older (ca. 2720 Ma) Greenwater assemblage rocks, were deposited in fault-bounded, pull-apart basins during regional transpressive (D2) deformation. North and south of the main, supracrustal-dominated outcrop, pink granitoid rocks associated with the McKenzie granite occur (Figure 25). The McKenzie granite is approximately 22 km long (east-west) by 3.2 km wide (north-south) and has been divided into two segments that are separated by a fault (Scott, 1990). Based on the mapping of Scott (1990), and the interpretation of aeromagnetic data, Metsaranta (2015) suggested that the McKenzie granite comprises multiple distinct intrusive bodies, and referred to the western segment as the Mount Baldy intrusion. Figure 22. Resedimented pyroclastic material as conglomeratic beds and lenses, STOP 1-8. - 17 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 24: AFM plot of samples collected and analyzed by Bjorkman (2014), showing the calc-alkaline nature of the resedimented pyroclastic rocks at STOP 1-8. Figure 23: Large pyroclasts, commonly with attendant bomb sags, STOP 1-8. The Neoarchean, S-type McKenzie granite (Hughes, 2016) is primarily a peraluminous quartz monzonite, with mineral assemblages characterized by microcline-plagioclase-quartzmuscovite-biotite with minor amounts of inequigranular hornblende, chlorite, titanite and rarely calcite. The McKenzie granite exhibits a peraluminous geochemistry, with SiO2 contents ranging from 63.8 to 68.2 weight % along with enrichment in light rare earth elements and fractionated heavy rare earth elements, decreasing trends of major oxides, transition metals and high field strength elements. Scattering of the large ion lithophile elements on discrimination diagrams is likely due to remobilization during chlorite, sericite and carbonate alteration (Hughes et al., 2017). It is proposed to have formed in a similar way to the model proposed for the later stages of the Figure 25. Photo illustrating cross-cutting relationships at STOP 1-8. The granitoid dyke in the bottom half of the photo that cross-cuts all lithologies, including a hornblendite dyke (top center), is associated with the nearby McKenzie granite. genesis of the nearby Dog Lake Granite Chain, which involved partial melting of a mantle wedge beneath the Wawa-Abitibi island arc. The proposed late-stage emplacement model is consistent with recent U-Pb geochronology (Puumala et al., 2015) that indicated that the McKenzie granite was emplaced at 2672.6 ± 1.5 Ma (zircon, U/Pb thermal ionization mass spectrometry). These S-type melts, formed from the partial melting of metasedimentary rocks, may have interacted with I-type melts, allowing for the variations in geochemical and petrological data that are observed in the McKenzie granite, such as the presence of hornblende, that are not common for standard S-type granites. - 18 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 STOP 1-9: Gunflint / Archean Unconformity, Crystal Beach 358661E / 5378895N This road cut on the northwest side of Highway 1117 near Crystal Beach provides another outstanding exposure of the contact between Archean basement and unconformably overlying Paleoproterozoic Gunflint Formation, similar to that exposed at STOP 1-1. However, there are a number of unique features here that warrant description and examination. The basement at this location is the Neoarchean McKenzie granite (2672 Ma) which has been conspicuously altered 1 to 3 m below the unconformity. The original pink granite has been altered to dark green chlorite (+ clays?) up to 2 m; alteration intensity increases upward towards the unconformity. Unaltered pink, K-spar-phyric granite gives way to altered versions in which the matrix is incipiently to completely chloritized, leaving relict, unaltered K-spar phenocrysts. The phenocrysts have also been replaced (saussurite + clays + chlorite) in the most intensely altered granite, leaving only relict quartz (Figure 26). The correlation between alteration intensity and proximity to the unconformity suggests that the altered rocks may represent a regolith/saprolite. Such alteration is often interpreted as a combination of ancient subaerial weathering (true paleosols) and later fluid migration from the overlying iron formation. Geochemical studies by Yip (2016) and Fralick (personal communication, 2024) at this location suggest that iron-rich Gunflint fluids replaced and masked the geochemical signature of the original paleosol. Similar alteration characteristics were described by Kronberg and Fralick (1992), who noted that alteration of ferromagnesian minerals in felsic Archean rocks Figure 26. Selected hand samples of McKenzie granite from STOP 1-9, showing progressive alteration (chloritization) from unaltered (left) through incipient and pervasive matrix replacement (second and third from left, respectively) to complete replacement of matrix and K-spar phenocrysts (far right). southwest of Thunder Bay was apparently due to diffusion of iron-rich, Gunflint-derived fluids across the Proterozoic -Archean unconformity, consistent with slow mineral-fluid exchanges under diagenetic or low-grade metamorphic conditions. Chemical changes in mafic minerals include additions of iron, manganese, and water and losses of silica, calcium, and magnesium. They concluded that these chemical changes occurred as Gunflint fluids diffused into underlying rock over a time frame of 105-107 years. Spalling of overlying Gunflint rocks has exposed a section of smooth, bare basement paleosurface (Figure 27). The contention of Pre-Gunflint weathering is supported by the occurrence of boulder-sized, spheroidally weathered, altered granitic corestones on the paleosurface, where they are enveloped by Kakabeka Member (basal) conglomerate and saprolite/ regolith, and are draped by Gunflint grainstones (Figure 28). Kakabeka conglomerate infills depressions in the paleosurface and fractures that extend down into weathered basement. The conglomerate here consists largely of resistate quartz pebbles in a chloritic, Figure 27: Smooth, curved, bare Archean basement paleosurface (accentuated in half-shadow above yellow field notebook), exposed below overlying, draped Gunflint grainstones, STOP 1-9. - 19 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 saprolitic/regolith matrix. Although this is perhaps the first documented example of corestones in the Gunflint or Biwabik formations, Paleoproterozoic (ca. 1.85 Ga) weathering-produced corestones in the Flin FlonCreighton area of Manitoba and Saskatchewan were documented by Sindol et al. (2020). Many of the joint surfaces and fractures in this exposure are covered and infilled by vein minerals, mainly quartz/amethyst, barite, fluorite, calcite with rare base metal sulphides (pyrite, chalcopyrite, galena, acanthite). These veins constitute the 7Z amethyst occurrence (Figure 29), first explored ca. 1890 (Ontario Mineral Inventory, https://www.geologyontario.mines. gov.on.ca/mineral-inventory/MDI52A10SW00007). The following description of the occurrence is excerpted from Puumala et al. (2015). Figure 29. Amethyst-bearing vein hosted in the Gunflint Formation at the 7Z occurrence. unconformity and are hosted by both Gunflint and granitic rocks. Gunflint rocks are strongly silicified adjacent to the veins. The exposed width of the vein system is approximately 10 m. The 7Z amethyst occurrence is hosted in a vein system and/or breccia zone that strikes 050º and is located approximately at the unconformity between sedimentary rocks of the Paleoproterozoic Gunflint Formation and Neoarchean intrusive rocks of the McKenzie granite stock. The amethyst-bearing vein system has been exposed in a series of 3 historic trenches over a strike length of 180 m. The majority of the amethyst-bearing veins strike 050º (i.e., parallel to the breccia zone) with near-vertical dips. The vein widths are variable, ranging from centimetre- to metrescale. In the Gunflint Formation rocks, a nearhorizontal set of narrow veins also occurs along bedding plane fractures. A third set of narrow, approximately north-striking veins, was also observed immediately to the south of the main breccia zone in road cuts along the north side of Highway 11-17. The portion of the vein system exposed in the southwestern and central trenches is hosted by rocks of the Gunflint Formation, while the veins exposed in the northeastern trench occur at the Figure 28. Spheroidally weathered, chloritized Neoarchean granitic corestone boulders resting on the paleosurface at the Paleoproterozoic-Archean unconformity, STOP 1-9. The corestones are enveloped by saprolitic sediments and conglomerate/regolith, and overlain by draping Gunflint grainstones. The amethystine quartz in this vein system shows a wide variation in colour, ranging from light pink (i.e., rose quartz) through to deep purple. Colourless to white quartz and smoky quartz are also abundant. Veins hosted by granite tend to contain lighter coloured amethyst, while deep purple amethyst and smoky quartz are most likely to be found in the southwestern trench, which is hosted by Gunflint Formation rocks. Most amethyst crystal points are on the order of 1 cm wide. However, much larger crystals were observed in some vugs. Crystals hosted in the Gunflint Formation rocks commonly have a surface coating of hematite. Although recent sampling has reported no significant silver values, a local newspaper reported in 1890 that 7Z was “a veritable mountain of amethyst with rich surface signs of silver” (ibid). This vein system is an - 20 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 example of a broad group of silver-bearing, carbonatequartz veins that typically occur in Animikie Group sedimentary rocks, often in close association with the Archean-Proterozoic unconformity and Midcontinent Rift-related diabase sills (Oja, 1967; Franklin et al., 1986; Kissin, 1992). They likely formed from metamorphically generated fluid from in the Midcontinent Rift and expelled along rift-bounding faults (Smyk and Frankin, 2007). largely based on a former outcrop exposure that was removed during highway reconstruction in 2011. “This is the only outcrop showing a complete ~ 3 m cross-section of the ejecta-bearing debrisite layer extending from Gunflint chert-carbonate up into the basal Rove Formation, which is overlain STOP 1-10: Terry Fox National Historic Monument 339836E / 5372406N This stop includes opportunities to view outcrop exposures near the Terry Fox National Historic Monument and lookout that commemorates Terry Fox’s 143-day, 5373-km Marathon of Hope run to raise money for cancer research in 1980, which continues to inspire global fundraising efforts. A number of rock types and features are exposed in the road cuts that flank the highway and access road near the monument (Figure 30). A prominent, columnarjointed Nipigon diabase sill (Terry Fox sill; Magnus, 2012; Magnus and Kissin, 2010) intrudes and caps these Rove and Gunflint formation sedimentary rocks. As a result, this site displays a complete stratigraphic section from the Gunflint Formation, through the Sudbury Impact Layer (SIL) and up into the overlying Rove Formation. Disconformities appear at both the base and top of the SIL (Addison and Brumpton, 2012). Figure 31. Rocks of the Sudbury Impact Layer (grey) and Rove formation (black) are visible in this photo from STOP 1-10. Geologist’s hand is located at the top of the Sudbury Impact Layer. The description of the SIL (Figures 31 and 32) at this location by Addison and Brumpton (2012) was Figure 30. This quarried rock face adjacent to the Terry Fox Lookout road displays a cross-section that includes (bottom left to top right) the Gunflint Formation, Sudbury Impact Layer (SIL), Rove Formation and Nipigon diabase. A Midcontinent Rift-related normal fault exhibiting approximately 4 to 5 metres of vertical displacement is visible near the left margin of the quarry face and is highlighted with a dashed line. Figure 32. The weathered outcrop (now gone) at STOP 1-10 in 2010 (Addison and Brumpton, 2012). Carbonatereplaced devitrified vesicular impact glass shapes and tektites were then visible on the weathered surface. The Ocean Transgression Sequence is composed of ankerite grainstones identical to those of the Gunflint and probably represents a limited transgression millions of years prior to the deposition of the Rove Formation (P. Fralick, personal communication, 2026). - 21 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 in turn by a diabase sill. An iron-rich alteration profile, heavily replaced by secondary pyrite, lies ~ 1 m below the base of the debrisite and a few metres northeast of the main outcrop. weave through this spherule-rich material but on a much finer scale than at Hillcrest Park. Red-brown agate 3-8 cm thick lies on top of the spherule-rich layer. Laterally discontinuous vertical digitate projections extend down from the top and project up from the base of this agate layer. They are similar in shape and size to the agate stalactites in vugs at Hillcrest Park, except that in this case the spaces between the projections were subsequently infilled by more agate. The red-brown colour is similar to that of the iron-rich alteration profile overlying it but it is a less saturated hue. The basal SIL is a recessively weathering, locally sheared, clastic layer about 0.5 m thick containing crushed spherule clusters, some of which are aligned subvertically instead of in the usual subhorizontal position. Several sets of subhorizontal slickensides, whose striae are aligned at a 140º azimuth, are found at various levels within this layer. Postdepositional anastomosing chert has replaced much of this basal sheared layer, obliterating considerable structural detail. Non-ejecta features include centimetre to millimetre-sized angular chert clasts and angular, subrounded to round Gunflint Formation iron carbonate clasts plus two rounded crystalline rocks with prominent alteration rinds. The presence of clasts with weathering rinds reinforces the idea that Gunflint clasts lacking such rinds were freshly fractured by impactgenerated earthquakes before being incorporated into the debrisite. An iron-rich alteration profile on top of the spherule-rich layer, consisting of hematite has been largely replaced by secondary pyrite. Prominent deformed spherule clusters are locally present. The total thickness of all these ejectabearing layers is 3 m. The top of this iron-rich layer marks a return to carbonate deposition. The basal 10-15 cm of this 80-100 cm thick carbonate zone is unstratified and shows dark, angular, commonly rectangular, millimetre-centimetre-sized rip-up mudstone clasts and probable Gunflint Formations clasts. This is followed by millimetre- to centimetre-scale layered carbonate strata topped by a zone with a few poorly defined, laterally discontinuous beds containing centimetre-scale, angular carbonate clasts. The main body of the 2.2 m thick debrisite lies in sharp contact over the basal sheared clastic unit. It is so heavily replaced by recrystallized dolomite that any possible ejecta features are only seen as vaguely outlined shapes on weathered surfaces or in thin section. Almost all detail, including any vesicles in possible DVIG- [devitrified vesicular impact glass] shaped clasts, has been destroyed. Tektites and microtektites may be present, based upon shape and rare faint devitrification textures. A single, polycrystalline, rounded quartz grain shows faint planar features. Both angular and rounded millimetre-scale chert clasts are also present, but not common. A 5-20 cm thick undulating, dark brown, recessively weathering, spherule-cluster-rich layer appears as a groove across the cliff face at the top of the dolomite-replaced debrisite. This mass of spherule clusters is much more concentrated than seen at any other location or than is suggested by faint shapes in the main dolomite-replaced layer immediately beneath it. These concentrated clusters seem to be the residuum of a thicker layer. Plentiful, thin anastomosing post-depositional chert strands The carbonate then makes an abrupt transition to 10-15 cm of gray siltstone and is overtopped by 10-15 cm of black, rusty weathering shale characteristic of the Rove Formation. The black shale is interrupted by 5 cm of chert before returning to 0.9-1.2 m of black, rusty weathering shale which is overlain in turn by a diabase sill more than 8 m thick. The shale is less friable than typical lower Rove shale, probably the result of low-grade metamorphism induced by the overlying sill.” A 050˚-060˚-striking normal fault, perhaps related to Midcontinent Rift-related extension, has displaced Animikie rocks and diabase 4 to 5 m. Koroscil (2013) identified thrust faults, mainly expressed as small discrete bedding plane faults with few kinematic indicators or piercing points to quantify the displacement. Thrust faults within the SIL were - 22 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 identified by slickenlines or slickenfibres on fault plane surfaces. The faults can be traced along strike until they are either covered by overburden or cut by a prominent normal fault which displaces all units in the hanging wall down to the south several metres (ibid). Reid-Sharp (2016) described faults, related damage zones and calcite vein breccias along the highway ~2 km northeast of STOP 1-10. Normal faults that transect Gunflint Formation + Archean basement rocks, strike east-northeast and dip to the southeast, were also ascribed to extension during Midcontinent rifting. The past-producing Thunder Bay Silver Mine is situated between the highway and the Terry Fox access road. Discovered in 1866 by P. McKellar, it was mined underground until 1874 via four shallow (8 to 21 m deep) shafts (Ontario Mineral Inventory, https://www. geologyontario.mines.gov.on.ca/mineral-inventory/ MDI52A06NE00005). Mineralization occurs in calcite-quartz veins that are hosted in chert-carbonates (Gunflint Formation) and shales (Rove Formation). The host rocks strike 034/22 southeast in the vicinity of the vein but subhorizontally 30.5 m to the northwest beneath a diabase sill 12.2 m thick. A 3 m wide composite vein or stockwork system consisting of 2.5 cm wide quartz-carbonate veinlets lies within and parallel to a fault that also strikes 034/65 northwest. Ore was mined locally over the total length of 182.9 m. Native silver and acanthite occurred in pockets 7.6 - 45.7 cm thick by 1.8 -12.2 m in length, the silver being in leaves and grains irregularly distributed in a gangue of quartz, with some calcite, galena, sphalerite, and pyrite. A second vein of calcite occurs in a parallel fault 6.1 m southeast of the composite vein (ibid; Sergiades, 1968). When first opened, two orebodies were found, one next to the north or hanging-wall and one in the middle (Tanton, 1931). The ore was brought to a stamp mill at the mouth of the Current River, 4 km south of the mine (Figure 33). Production totaled an estimated $20 000 (Bowen, 1911), or approximately 15 000 ounces of silver. Figure 33. Stamp mill of the Thunder Bay Silver Mine at the mouth of the Current River, ca. 1880. Shegelski (1982; Figure 34) and later described in the context of impact-related brecciation by Addison et al. (2010) and Addison and Brumpton (2012, Figure 35): “A bedrock exposure, about 5 m by 15 m, in a private yard in Thunder Bay contains a spectacular debrisite exposure composed mainly of Gunflint chert-carbonate breccia and ejecta, primarily DVIG [devitrified vesicular impact glass], which is surrounded and partially replaced by blocky calcite cement. The debrisite remnant preserved here is 0-0.5 m thick and unconformably overlies stromatolites and chloritic grainstone of the uppermost Gunflint Formation. An iron-rich alteration zone exists approximately 30 cm below the erosive contact between the debrisite and the Gunflint bedrock. DVIG clasts are up to 2 cm across. Vesicles range from round to ovoid to nearly flat. Angular quartz and feldspar grains, chert shards, and chloritic granules are also present. Quartz grains with PDFs have not been found here.” The SIL is also exposed nearby at Hillcrest Park and along Banning Street. STOP 1-11: Sudbury Impact Layer, Markland and Hill Streets 334163E / 5366301N (n.b. Private Property, ask for permission to access. Be very careful not to step on any plants. No hammers are allowed.) Another spectacular debrisite breccia of the Sudbury Impact Layer is exposed at the corner of Markland and Hill streets. This outcrop was mapped in detail by - 23 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 34. Detailed map of the debrisite breccia outcrop at STOP 1-11 by Shegelski (1982) Figure 35. (from Addison and Brumpton, 2012) A – Gunflint Formation stromatolites exposed on a glacially truncated surface, STOP 1-11. While it is not recognizable in the photo, debrisite lies over stromatolites at upper right of the photo. B – Angular to slightly subangular clast-supported Gunflint Formation breccia with a finer DVIG-rich and calcite-rich matrix, all of which lies directly on Gunflint stromatolites, STOP 1-11. The angular clasts suggest a short travel distance from their point of origin. C – DVIG clasts within a recrystallized calcite matrix, STOP 1-11. The silicate devitrification product supports growth of a black lichen, whereas calcite prevents lichen growth. The vesicles are calcite infilled. D – Orange, weathered accretionary lapilli in a recrystallized carbonate matrix, Hillcrest Park. - 24 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Field Trip Stop Descriptions - Day Two Day Two comprises stops to a variety of locations south and west of Thunder Bay (Figure 36). This area is underlain by a variety of rocks that record almost three billion years of local geologic history, spanning from the Neoarchean (ca. 2.7 Ga) to the Paleoproterozoic (ca. 1.8 Ga) and Mesoproterozoic (ca. 1.4 and 1.1 Ga). Unconsolidated glacial and post-glacial deposits and features attest to a long-lived, Pleistocene glaciation record. All GPS coordinates are NAD83, UTM Zone 16. Figure 36. Generalized geology of the Thunder Bay area, showing Day Two field trip stop locations. Geology from Map M2232 (Carter et al., 1973). STOP 2-1: Mount McKay Lookout (Anemki Wajiw) 0331126E / 5357384N (n.b. admission via a gate operated by Fort William First Nation) Our first stop provides not only a panoramic view of Thunder Bay and surrounding area, but also stacked Logan sills which have produced the iconic mesa topography of Mount McKay and other similar mesas to the south and west in the Animikie-underlain Logan basin, collectively known as the Nor’westers. This location had previously been described by Cundari et al. (2012). Mount McKay is also known as Anemki Wajiw (“Thunder Mountain”) in Ojibwe. The summit, at 482 m ASL, is approximately 300 m higher than Lake Superior. The stop is centered on the lookout area (Figure 37), which represents the top of the lower sill at approximately 337 m ASL. The upper, ~60 m thick, columnar-jointed sill and adjacent, hornfelsed Figure 37. View of the top of Mount McKay, capped by the ~60 m-thick, upper Logan diabase sill. The Lookout level is underlain by the top of the lower sill. - 25 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Rove Formation wacke can be accessed by way of a hiking trail which leads to the summit. Although stacked sills have been encountered in drilling, few examples exist in surface exposures. As many at 14 sills were reported, for example, in a 705 m-deep drill hole in central Pardee Township by Dumont Nickel Inc. (Assessment Files, Thunder Bay South Resident Geologist’s District, Thunder Bay). al., 2007). Resampling of Zr-enriched, pegmatoidal, upper portion of the Logan Sill capping Mount McKay and re-analysis of baddeleyite and magmatic zircon yielded an age of 1106.3 + 2.0 Ma (Bleeker et al., 2020). This, and similar ages elsewhere, led Bleeker et al. (2020) to favour a relatively sharp onset of high-volume mafic-ultramafic magmatism in the Midcontinent Rift at ca. 1110 to 1106 Ma. The rugged topography (Figure 38) has produced extensive colluvial deposits and talus slopes. Unconsolidated, sandy lacustrine and fluviolacustrine deposits occur below the bedrock- and colluvium-predominated slopes. Abandoned shoreline escarpments and beach bars, visible between the lookout and Lake Superior, reflect higher post-glacial lake levels (Burwasser, 1977). Feldspar-phyric patches, common near upper chilled sill contacts, are present in an exposure of the upper, chilled contact of the lower sill along a path to the west of the clearing (Figure 39). A tentative age of 1114.7 ± 1.1 Ma was determined from a Logan sill on Mount McKay, using a limited selection of very small baddeleyite grains (Heaman et Figure 39. Polygonal jointing in upper chilled surface of lower diabase sill, STOP 2-1. Figure 38. Shaded relief LiDAR image of area south of Thunder Bay, showing topographic relief (i.e. gently southdipping mesas/cuestas) resulting from erosion-resistant mafic sills and, to a lesser extent, siliceous wackes in the Rove Formation (data from https://www.arcgis.com/apps/ mapviewer/index.html?url=https://ws.geoservices.lrc.gov. on.ca/arcgis5/rest/services/Elevation/FRI_DTM_SPL/ ImageServer). STOPS 2-1 and 2-2 are also shown. Logan sills generally consist of fine- to coarsegrained, ophitic to intergranular, quartz tholeiitic diabase/gabbro (Smith and Sutcliffe, 1987; Geul, 1970, 1973). Coarse-grained, intergranular gabbro, locally rich in granophyric mesostasis, is common in the interior of the thicker sills. Geochemical data from sampling of the upper and lower sills by Hart and Magyarosi (2004) are provided in Figure 40. These sills represent the northernmost known extent of Logan diabase sills near Thunder Bay. Nipigon diabase sills occur within the city and extend northward to the Nipigon Embayment. - 26 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 40. Primitive mantle-normalized trace element plots for upper and lower sills at Mount McKay with Nipigon sill sample for comparison. Data from Hart and Magyarosi (2004) and Hollings et al. (2011). Normalizing values from Sun and McDonough (1989). Nipigon sills are characterized by generally lower incompatible trace element abundances, lower TiO2 content, and a distinct negative Nb–Ta anomaly. They typically have lower Gd/Ybcn ratios compared to Logan sills. Logan Sills are characterized by higher TiO2 and higher Gd/Ybcn ratios, indicating a greater degree of heavy rare earth element fractionation (cf. Hollings et al., 2010). Riverdale sills (STOP 2-2) are geochemically distinguishable from both Nipigon and Logan diabase (Figure 41). STOP 2-2: Riverdale Quarry 322418E / 5355233N (n.b. Private property; permission is required to access. Caution advised on site due to slip and fall risks associated with steep slopes and vertical rock faces) This former shale quarry exposes a ~20 m-thick section of the lower Rove Formation, overlain by a ~12 m-thick Riverdale, columnar-jointed, gabbronorite sill related to Midcontinent Rift magmatism (Figure 41). This location was previously described by Cundari et al. (2012): “Sampling by Smyk and Hollings (2007) identified this as a Riverdale gabbronorite Figure 41. Discrimination diagrams for mafic and ultramafic intrusions near Thunder Bay (from Cundari et al. 2012). Data are from Hollings et al. (2007a) and Puchalski (2010). Normalizing values from Sun and McDonough (1989). - 27 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 lie within the major element abundances; olivine gabbros are lower in SiO2 and elevated in MgO, Cr, Co, and Ni compared to the gabbronorite samples. The sill does not display any evidence for Figure 42. Riverdale gabbronorite sill capping section of Rove Formation clastic sedimentary rocks, Riverdale Quarry. (Photo taken in 2008. Arrow points to geologist for scale.) sill in Rove Formation shale, wacke and minor tuffaceous units. Subsequent detailed petrographic and geochemical analyses were carried out by Puchalski (2010). Samples were taken through stratigraphy at the quarry to investigate composition and contamination, as well as to test whether the sill had undergone differentiation. The following section provides a concise summary of those findings. differentiation as shown by the erratic trends of MgO, SiO2, TiO2, Cr and Ni through stratigraphy. An olivine gabbro in the center of the sill displays elevated MgO, Cr, and Ni values as well as a lower abundance of silica when compared to the surrounding samples. This is likely the result of a slightly more primitive magma intruding the center of the sill. The lack of chilled margins between the olivine gabbro and the gabbronorite suggest that the sill had not fully crystallized when the second pulse intruded. A sample of a 60 cm wide north-trending diabase dyke which intrudes the sill near the western end of the quarry is geochemically comparable to the surrounding Logan sills. Contamination by the Rove shale is evident in samples taken from close to the contact (<1 m above the contact). These samples display higher SiO2 values as well as lower Nb/Nb* and Gd/Ybn values than the rest of the unit. As the Rove shale displays significantly lower Nb/Nb* and Gd/Ybn values than that of the surrounding gabbronorite. The Rove shale is the likely source of this contamination signature. Two different pulses of magma are recognized within the Riverdale sill, based on contamination signatures denoted by negative niobium anomalies. The less-contaminated samples are typically found towards the core of the intrusion with rocks above and below displaying a greater degree of contamination. Samples taken within 60 cm of a shale xenolith do not display a distinct negative niobium anomaly. This shows that the source of contamination responsible for the negative niobium anomaly is not the Rove shale but is likely a crustal component from depth. εNd (T=1100 Ma) values of -1.6 to -1.9 for the Riverdale Sill are consistent with this model (Smyk and Hollings, 2009). The mafic intrusive rocks within the quarry are dominantly classified as gabbronorites with olivine gabbro present towards the center of the sill. The gabbronorites are generally fine-grained with plagioclase occurring as subhedral laths. Orthopyroxene is present in greater abundance than clinopyroxene, occurring as anhedral to subhedral crystals. Varying degrees of alteration are manifested as sericitization of plagioclase and chloritization of pyroxene. The olivine gabbro is texturally similar to the gabbronorite, albeit with a higher modal percentage of fine-grained, anhedral to euhedral olivine. In most samples, olivine is replaced by serpentine, producing secondary quartz and calcite, as well as minor magnetite. Alteration is significantly greater in the narrow, chilled margin at the contact. Pyrite occurs throughout the unit; minor chalcopyrite has also been noted. Sampling for whole rock major and trace element geochemistry was undertaken by Puchalski (2010) throughout the 10 m exposure at 1-m intervals. Olivine gabbro samples display broadly similar trace element characteristics to those of the gabbronorite samples. Differences Although the Riverdale sill is located near Logan sills, it remains petrographically and geochemically distinct from them [Figure 42]. Geochemical discrimination based on La/Smn (LREE) vs. Gd/Ybn (HREE) shows characteristics - 28 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 similar to those for the ultramafic units of the Nipigon Embayment (e.g., Disraeli, Kitto, Hele and Seagull), closely resembling the mafic to ultramafic Jackfish sill. The Jackfish sill is finergrained and displays a higher modal abundance of olivine than the Nipigon sills surrounding it (Hollings et al., 2007a). This suggests that the Riverdale sill may be genetically related to the ultramafic and mafic to ultramafic units of the Nipigon Embayment. This is consistent with the reversed polarity of the Riverdale sill (Hollings et al., 2010).” A number of features are visible at or near the exposed upper and lower sill contacts (Figure 43). Calcite-filled vesicles define a crudely developed Figure 43. Lower gabbronorite sill contact. (A) Delamination of Rove shales by injection of gabbronorite sill magma; (B) Chilled margin of gabbronorite sill against hornfelsed Rove shale. layer/joint filling(?) in medium-grained gabbronorite, ~1 m above the lower sill contact. Stoping and delamination of Rove shales is also evident here. Thin, parallel chilled margins, perhaps representing multiple influxes of magma, occur above the lower sill contact. A narrow (75 cm) diabase dyke with Logan sill-like geochemistry intrudes the Riverdale gabbronorite sill near the western end of the quarry exposure (Figure 44). Glacial striae are visible on exposed outcrop surfaces at 060˚ and 075˚. STOP 2-3: Sudbury Impact Layer, Highway 588 0307539E / 5357977N Figure 44. Narrow diabase dyke with Logan sill-like geochemistry intruding Riverdale gabbronorite sill, Riverdale Quarry. Scale card straddles the eastern dyke contact. This stop, while having lost much of the best exposure of the Sudbury Impact Layer (SIL) due to ongoing highway construction, still provides an opportunity to view some of the features associated with the SIL in the affected ankeritic, Gunflint Formation chertcarbonate rocks. The SIL was also intersected a few metres below surface in a shallow drill hole, collared in Rove Formation shales in an abandoned quarry approximately 300 m south-southwest of STOP 2-3. This location was previously described by Addison and Brumpton (2012): - 29 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 “When first observed in 2000, the Hwy 588 outcrop was a bedrock exposure in the ditch on the northwest side of the highway, 2.4 km southwest of the hamlet of Stanley. It was a glacially polished and striated surface showing erosively truncated stromatolites up to 0.5 m diameter, some of which were surrounded by accretionary lapilli 3-25 mm in diameter [Figure 45]. Ankeritic grainstone and chloritic grainstone surrounded other stromatolites. This exposure was subsequently blasted to deepen the ditch and the blasted rock now lines the ditch slopes, giving a highly fragmented cross-section and plan view of the exposure. Since then, we have exposed bedrock in the ditch about 50 m southwest of the first exposure. It shows a glacially striated surface of exposed stromatolites and shattered, but insitu black chert with an ankeritic grainstone filling in the cracks. The chert is assumed to have fractured during the compressional stage of impact-triggered earthquake waves with the fractures then opening during the dilational wave phase. Fine granular material then fell into the openings, preventing them from closing and subsequently the material was lithified. material show a variety of ejecta features, the most obvious being accretionary lapilli which have yielded quartz and feldspar grains showing planar deformation features (PDFs) and planar fractures. Planar features have not been found in larger subrounded and angular quartz and feldspar grains contained within the debrisite generally as opposed to within accretionary lapilli. This is the only site in which DVIG [devitrified vesicular impact glass], is not the most obvious ejecta feature within the debrisite. In fact, no DVIG has been observed, however carbonate and silica replaced clusters of spherules are present. Non-ejecta features include subrounded to round chert grains in carbonate cement, subcentimetre stromatolite fragments and mudstone and shale rip-ups. Chloritic, blotchy, black Gunflint Formation granules, similar in shape and size to microtektites, are present within the carbonate cement. Carbonate-replaced microtektite shapes are present but since they lack residual internal structure, it is impossible to determine if they were microtektites or carbonatereplaced Gunflint chlorite granules.” Thin sections prepared from the blasted STOP 2-3: Kakabeka Falls Provincial Park 0305738E / 5364400N; 0305178E / 5364663 (n.b. Entry/parking fee is required in Kakabeka Falls Provincial Park. Sample collecting and hammers are NOT permitted.) Two stops at Kakabeka Falls provide an opportunity to see both a thick section of Gunflint Formation rocks exposed in the gorge of the Kaministiquia River, and the basal, stromatolite-bearing units of the Gunflint unconformably overlying Neoarchean granitoid basement. This location was previously described by Pufahl et al. (2000) and Smyk (2012). Figure 45. Accretionary and armored lapilli draped unconformably over a stromatolite, composed of silicified carbonate, which was abraded to its present configuration likely by a base surge immediately preceding the deposition of the lapilli; STOP 2-3, polished surface. The gray component is primarily fine-grained, angular, fractured carbonate clasts whose individual crystals are usually <10 μm. These clasts are typically <50 μm but they may be as large as 500 μm. Quartz and feldspar grains are a minor component among the carbonate clasts within the lapilli. (caption modified from Addison and Brumpton, 2012). The park is dominated by a single, spectacular feature, Kakabeka Falls, which drops 39 m over sheer cliffs in Gunflint Formation sedimentary rocks (Figure 46). Kakabeka is an aboriginal word meaning “steep cliffs”. The age of the river gorge below the falls is still debated. If none of it existed prior to the glacial Lake Beaver Bay stage, then it is less than ca. 9700 years old. The portage around the falls contains artifacts ranging from the Paleoindian to the historic (fur trade) periods. - 30 - The falls owes its existence to the thin chert- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 are often attributed to deposition in intertidal or lagoonal subtidal environments (Pufahl et al., 2000). Figure 46. Kakabeka Falls and gorge, cut into flat-lying Gunflint Formation shales. Photo from https://hikebiketravel. com/a-trip-to-kakabeka-falls-near-thunder-bay/. carbonate bed which forms a resistant cap rock to the softer underlying shales. The river gorge is composed of a sequence of volcaniclastic shales (lessresistant, darker units) and tuffs (more-resistant, lighter coloured units). This sequence represents the major volcaniclastic horizon in the upper Gunflint Formation that is traceable to the south through the Mesabi Range. Note that shale is the predominant lithology in the Kaministiquia sections and this is, in fact, typical for the Gunflint Formation in general throughout the Thunder Bay region. Samples of lapilli tuff and reworked tuffs from the middle of the Gunflint Formation, collected by Fralick et al. (2002) at Kakabeka Falls yielded a euhedral zircon population with a U-Pb age of 1878.3 ± 1.3 Ma, believed to be nearly synchronous with the depositional age. The outcrop on the northern edge of the parking lot contains layers of banded/ribbon chert-carbonate within black, fissile shale. The alternating, dark grey chert and brown siderite-ankerite layers display slump and soft-sediment deformation features. Microscopic examination of banded chert-carbonates reveals delicate lamination in the chert which resembles the “ribbon texture” of algal mats. The interlayered carbonate bands contain complex, microspherical structures which likely resulted by nucleation from a gel state. Local thick beds of carbonaceous siderite (2-3 wt% carbon) form carbonate iron formation; contemporaneous deposition of carbon and carbonate suggests biological activity during iron deposition. Studies of the Gunflint Formation have described this type of sediment as forming in a deep, quiet water environment. However, similar carbonate sequences The rapids visible north of the highway bridge are formed by Archean granitoids. The slow-water area to the south is underlain by the Gunflint Formation. The basal conglomerate (Kakabeka member) is patchily preserved on Archean basement here. Silicified stromatolites are developed on the conglomerate or directly on the basement. This is the location from which samples collected from silicified stromatolites in the 1950’s yielded the first documented Gunflint cyanobacteria (Tyler and Barghoorn, 1954). The rock cut on Highway 590 immediately south of the intersection with Highway 11-17, west of the Kaministiquia River bridge, expose cherty carbonates at the base of the Gunflint Formation where it rests unconformably over Neorchean granitoid basement. Large-form stromatolites are developed at the unconformity (Figures 47 and 48). The stromatolitic, ribbon carbonates are abruptly overlain by a grainstone succession. Black anthraxolite veinlets and void fillings occur with vein quartz in the chert-carbonate rocks. Anthraxolite and pyrobitumen in the Gunflint Formation (Figure 49) has been noted and studied numerous researchers, including Tanton (1931), Ellesworth (1934), Goodwin (1956), Kwiatkowski (1975), Barghoorn et al. (1977), Hayatsu et al (1983), Mancuso et al. (1989), Rutter (2014), Rasmussen and Muhling (2019), and Rasmussen et al. (2021). Figure 47. Colloform stromatolite (left of hammer) in Ferich carbonate grainstones, Highway 590 exposure, STOP 2-4. The stromatolite was situated approximately 40 cm above the unconformity with Neoarchean basement granitoid rocks. Unfortunately, the stromatolite spalled from the outcrop face ca. 2016. Coin is 2.5 cm in diameter. - 31 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 49. Void-filling, conchoidal anthraxolite and quartz in sideritic Gunflint grainstone, Highway 590 exposure, STOP 2-4. of Superior lobe drift. The exposed sequence consists of approximately 6 to 7 m of well sorted, steeply-dipping sand and gravel of Superior provenance overlain by 2 to 3 m of silty Superior lobe till [Figure 50]. Clasts in the lower glaciofluvial unit consist primarily of Proterozoic metasedimentary rocks. Numerous cobbles and boulders belonging to the Gunflint Formation and Sibley Group are recognizable. Foreset beds dip steeply to the north and are likely of deltaic origin. The delta was probably built proglacially into an early phase of glacial Lake Kaministikwia. The feature therefore represents a location at which the advancing Superior Lobe stalled prior to reaching its maximum position at the Marks Moraine. Figure 48. Detailed view of the large, silicified, colloform stromatolite in Figure 47. Rasmussen et al. (2021) suggested that stromatolitic, black Gunflint cherts were saturated in syn-sedimentary oil. Thermally altered oil (pyrobitumen) occurs in the stromatolites and intercolumn sediments, fills pores and fractures, and coats detrital and diagenetic grain surfaces. Hayatsu et al. (1983) described two very distinct macromolecular materials in the Gunflint anthraxolite that suggested that the Thunder Bay area was once covered by Cretaceous or Jurassic marine sediments, similar to those documented in the Mesabi range of Minnesota. The delta is actually located within only 3 km of the Superior lobe limit and occurs at an elevation of about 375 m asl, 85 m below the maximum elevation of Lake Kaministikwia. The delta was STOP 2-5: Briggs Drive Gravel Pit 0304790E / 5369390N (n.b. Private Property, contact Township of Conmee for access permission) This stop not only highlights some interesting glacial sediments, but also provides an opportunity to examine large boulders of a variety of local rock types that have been transported by glacial ice and meltwater. This location was described by Bajc (2000): “At this stop, we will be looking at a section Figure 50. View, looking west, of steeply dipping, gravelly foreset beds of a delta constructed along the margin of the advancing Superior lobe, Briggs Drive gravel pit, STOP 2-5. Silty Superior lobe till caps the sequence. - 32 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 overridden by the Superior lobe resulting in the truncation of the foreset unit and removal of the topset beds. Several metres of silty, Superior Lobe subglacial till was deposited on top of the sands and gravels. “seven lenticular masses of brecciated, banded iron formation, in which pyrite has replaced a considerable part of the rock” (Carter 1990). The largest of these masses has a maximum width of 23 m and is 244 m long. Other discoveries include a 21-m wide body of pyrite containing magnetite and pyrrhotite and a 9 m wide by 15 m long zone of magnetite-pyrite-jasper ironstone. It is possible that the boulders found within the gravel pit [Figure 51] were derived from this area and Of particular significance is the occurrence of large angular to rounded boulders on the pit floor. The boulders were extracted from the lower glaciofluvial unit and, in some cases, do not appear to have been transported very far. Some of the larger boulders measure several metres across and still display striated surfaces. The boulders are derived from both Archean and Proterozoic source rocks. Several boulders of sulphidized iron formation and massive pyrite of Archean age were discovered in the boulder piles. One of the boulders measured over 1 m in diameter and consisted of massive pyrite and magnetite with 10 to 15% sphalerite disseminated in pyriterich sections. Sphalerite was also concentrated along fractures and adjacent to quartz veinlets throughout the rock. Two samples from the pyriterich zones returned values of: 1) 5.13% Zn, 18 ppm Cu, 19 ppm Pb, 260 ppb Au and 0.5 ppm Ag; and 2) 2.85% Zn, 16 ppm Cu, 20 ppm Pb, 245 ppb Au and 0.5 ppm Ag. A sample from the magnetiterich zone returned values of 850 ppm Zn, 25 ppm Cu, 5 ppm Pb, 25 ppb Au and <0.2 ppm Ag. A second sulphidized iron formation boulder measuring approximately 0.5 m in diameter and consisting almost exclusively of pyrite, returned values of 140 ppm Zn, 8 ppm Cu, 11 ppm Pb, 710 ppb Au and <0.2 ppm Ag. There are two possible source areas for the boulders. Superior lobe striae in the immediate vicinity of the pit are oriented at 320 to 330° Az. If the boulders were eroded and transported by Superior ice, then there is a 5 km window towards the southeast from which they could have been derived. Proterozoic metasedimentary rocks outcrop beyond the 5 km limit. Alternatively, the boulders could have initially been eroded by northern ice from a source to the north-northeast of the pit then remobilized by the Superior lobe. Exploration work during the early 1900s along the lower reaches of Brule Creek, 4 to 5 km northnortheast of the gravel pit, by B.L. Morrison, the Davis Sulphur Company and General Chemical Company resulted in the discovery of Figure 51. Pile of oversized boulders of a variety of local Archean and Proterozoic rock types, Briggs Drive gravel pit, STOP 2-5. Reddish silty Superior lobe till is visible at the top of the pit wall. Photo taken ca. 1999. that sphalerite was not recognized in the rock. It is not yet clear whether the sulphides indicate proximity to a VMS style zone of mineralization. Further work is required to assess the mineral potential of this area.” STOP 2-6: Temiskaming sedimentary rocks, Finmark 293709E / 5383903N This stop, having long been a “must-see” for local geology students, has benefited from new exposures - 33 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 on the south side of Highway 11-17 that were created during highway expansion and ballast quarry development ca. 2019. The outcrops on both sides of the highway expose excellent examples of clastic Neoarchean “Temiskaming-type” metasedimentary rocks (ca. 2690-2695 Ma) of the Shebandowan assemblage that still display many primary sedimentary features that provide clues as to the depositional environment. The Timiskaming-type successions of the SGB are interpreted to have been deposited in subaerial to shallow marine environments (Shegelski, 1980). Highway 11-17 by Koebernick and Fralick (1995) and Koebernick (1996) documented sedimentary structures and bed sequences consistent with shallow water, coastal sedimentation in three major depositional environments: tidal strandline, the shoreface, and the offshore (e.g. Figure 52). Koebernick (1996) noted: “The three environments and associated sub-environments record processes reflective of differing current activity which controlled and The ballast quarry immediately to the south has been developed in mafic, Neoarchean metavolcanic rocks of the Greenwater assemblage (ca. 2720 Ma) which presumably underlie the clastic rocks unconformably or are in fault contact with them. Parker (1980) noted that reversals of top directions and the presence of both easterly and westerly plunging minor folds, suggest that one or more episodes of folding have occurred. Detrital zircon geochronology by Corfu and Stott (1998) confirmed that the metasedimentary rocks in the Finmark area (<2691+3 Ma) and in the southern part of Adrian Township (<2700+4 Ma) are younger than the Greenwater assemblage. Because of the remarkable preservation of primary sedimentary features, this stop has been the focus of study for many years, including theses by Parker (1980) and Koebernick (1996). The metasedimentary sequence here comprises interbedded sandstonesiltstone-mudstone sequences which alternate with thick deposits of cross-stratified sandstones (Parker, 1980). The interlayered sequences contain many of the primary sedimentary structures characteristic of tidal flat deposits, such as flaser bedding, lenticular bedding, herringbone cross-bedding, mud cracks, mud drapes, and bipolar paleocurrent indicators. Parker (1980) noted that the clastic sedimentary rocks are composed of feldspar, rock fragments, quartz, and some mafic minerals. Modal analysis revealed that most of the sandstones in the area are arkosic arenites. Lithic fragments are felsic to intermediate and predominantly calc-alkalic volcanics, with lesser amounts of other igneous grains and sedimentary rock fragments. This led Parker (1980) to suggest that the clastic rocks probably represent immature detritus from proximal volcanic centers. A detailed study of the “Temiskaming-type” clastic rocks at this location and along this section of Figure 52. Stacked, trough cross-bedded sandstone beds with tangential foreset laminae, south side of Highway 1117, STOP 2-6. Ripples are preserved on bedding surfaces (lower photo). - 34 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 53. Dark alteration holes defining herringbone cross bedding in sandstone, north side of Highway 11-17, STOP 2-6. influenced deposition. The tidal environment was dominated by bidirectional tidal currents [Figure 53]. Deposition In the shoreface was predominated by unidirectional wave-produced currents which overprinted prevailing tidal current activity, in the distal portions of the shoreface environment though, deposition was once again controlled by tidal currents. In the offshore, deposition was controlled by storm currents which generated distinctive beds of hummocky cross-stratification. The tidal environment is composed of many sedimentary structures similar to those present in Phanerozoic and present-day tidal sequences. In the tidal flat sub-environment, vertical sequences of flaser, lenticular, wavy and coarsely interlayered bedding reflect current velocity fluctuations Intimately tied to spring - neap tidal cycles. The tidal channel sub-environment lacks many of the features characteristic of tidal channels described in the literature; such as extensive point bar development. Instead, the tidal channels of the study area appear to represent sequences deposited in relatively straight channels. Migration of sand waves and dune fields deposited the cross-stratified lithofacies of the shoreface environment. Similar to a highenergy, non-barred coastline, the proximal portion of the shoreface lacks any evidence of beach development. Instead, the shoreface records a rapid and discontinuous transition from the tidal strandline environment. Hummocky cross-stratification (HCS) [Figure 54], parallel- Figure 54. Hummocky cross-stratification, which is only formed and preserved by storm waves in depths between fair weather wave base and storm wave base. laminated and massive sandstone beds as well as siltstone and mudstone beds typify the offshore environment [Figure 55]. The HCS differs greatly in thickness and internal structure from HCS described in the literature. The HCS in the study area reflects restricted and/or variable sediment supply and flow conditions. A paleotidal range was determined from the sediments of the tidal environment. The range indicated a mesotidal Figure 55. Thinning- and fining-upward sequence, showing transition from medium-bedded sandstones to a mudstone/ siltstone-dominated package with thin sandstone interbeds, south side of highway, STOP 2-6. This likely represents deepening of the water, going from nearshore coarse-grained sands moved around on the bottom as dunes by fair-weather waves and currents, to deeper water deposits below fairweather wave base, representing tempestites (i.e. hummocky cross-stratified storm deposits and graded beds formed below storm wave base by the same geostrophic flows that formed hummocky cross-stratification in shallower water; P. Fralick, personal communication, 2026). - 35 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 56. Stratigraphic column of outcrops on north side of Highway 11-17, STOP 2-6, showing primary sedimentary features and paleocurrent measurements (P. Fralick, personal communication, 2026). environment and is comparable to Precambrian tidal ranges reported in the literature. Tidal rhythmites, present on the tidal flats, suggest a length of 26 days for the Neoarchean lunar month. Currents which deposited the tidal rhythmites produced both semi-diurnal and diurnal sediment sequences [Figure 56].” reversals. Evidence of shearing and brittle deformation, including quartz-carbonate veining, can also be observed, especially in the easternmost portions of the outcrop exposure on the south side of the highway (Figures 57 and 58). As noted above, in spite of the remarkable preservation of sedimentary structures, the Shebandowan assemblage sedimentary rocks in this area have experienced tectonic deformation and display features that include minor folds and younging Bedding-cleavage relationships indicative of folding are visible in the outcrops at this location. Bedding orientations vary from approximately 325/70 northeast on the north side of the highway to 110/70 south on the south side. The cleavage-bedding relationship is most easily observed in the thin-bedded mudstones and siltstones south of the highway, where the cleavage - 36 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 57. Quartz-carbonate veining in altered and deformed Figure 58. Photo illustrating small-scale folds, shears and rocks at the east end of the outcrop area on the south side of sigmoidal tension fractures in thin-bedded siltstone and mudstone at STOP 2-6. Highway 11-17 at STOP 2-6. orientation is approximately 085/85 south. The change in bedding orientation relative to cleavage, when combined with northward-younging indicators (e.g., graded bedding), indicate the probable presence of an anticline axis a short distance to the north. This interpretation is consistent with previous geological mapping completed by Carter (1985). The structures observed here may have developed during the same tectonic events that gave rise to orogenic gold mineralization at the nearby Eureka Gold Deposit, which is currently being explored by Delta Resources Limited. Eureka is located approximately 4 km to the west-northwest of here, and the deposit occurs within a structural corridor known as the “Shebandowan structural zone.” Gold mineralization at Eureka also has a close spatial association with the unconformity between the Greenwater and Shebandowan assemblages. Delta Resources has outlined the Eureka Gold Deposit over a 2.5-kilometre strike length, and to a vertical depth of 300 metres. Mineralization occurs over true widths ranging from 10 to 100 metres, and the deposit remains open in all directions. Gold is hosted by multiple generations of quartzankerite-pyrite veinlets that generally range from 1 mm to 10 cm wide and cross-cut multiple lithologies. Wider quartz veins up to 4.5 metres wide, and goldbearing silica-flooding zones are also found within the deposit. Host rock alteration is characterized by intense, texture-destructive ankeritization, silicification, albitization and sericitization combined with trace to 2% disseminated pyrite and trace arsenopyrite. The altered rocks typically contain anomalous gold. Feldspar-phyric monzonite to diorite dikes also have a close spatial association with the mineralization and are locally altered (https://www.deltaresources.ca/ delta-1-gold-project/). STOP 2-7: Pillowed Basalt, Mud Lake 315029E / 5376770N No trip would be complete without pillowed basalt! These roadside exposures along Highway 102 near Mud Lake display tholeiitic, mafic to intermediate volcanic Figure 59. Pillowed basalt flow, STOP 2-7, showing wellpreserved, close-packed pillows and hyaloclastite-filled inter-pillow spaces. - 37 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 rocks of the Neoarchean Greenwater assemblage, situated near the Quetico–Wawa subprovince boundary (Brown, 1995; Brown and Fogal, 1995). In this area, the degree of pillow preservation varies considerably. Figure 60. Close-up of pillow, STOP 2-7, showing contact of chilled upper selvage (large dashed line). However, well-formed, close-packed pillows, ranging from 10 by 15 cm to 30 by 60 cm in size, are locally preserved (Figure 59). Where discernible, younging directions within this unit are consistently to the north. Carbonate-filled amygdules, generally less than or equal to 1 mm in size and constituting up to 10% of the pillows by volume, are commonly present, radiating outward from the core of the pillows (ibid; Figure 60). North-younging, pillowed flows exposed at STOP 2-7 display well-preserved primary features, including autoclastic breccias (e.g. pillow breccia, inter-pillow hyaloclastite; Figure 61), close packing and pillow cusps, and calcite-filled amygdules. Larger, ovoid amygdules occur sparingly in the cores of pillows, while smaller, more numerous, pipe-like amygdules tend to be concentrated near pillow selvages. Pillows typically range between 25 cm and ~1m in size. The Mud Lake Cu-Zn occurrence (Ontario Mineral Inventory, https://www.geologyontario.mines.gov. on.ca/mineral-inventory/MDI000000002310) can be observed in a roadside outcrop located a few hundred metres northwest of STOP 2-7 along the highway. Pyrite, minor chalcopyrite and rare sphalerite are finely disseminated throughout, and adjacent to, a sericitized northeast-trending zone of shearing hosted within chemical metasedimentary rocks interbedded with felsic and intermediate pyroclastic metavolcanic rocks (Brown, 1995; Brown and Fogal, 1995). The mineralization was first uncovered during construction along Highway 102 in the mid-1970s. A grab sample collected in 1975 by staff of the Resident Geologist’s office, Thunder Bay, yielded values of 0.24% Cu, 0.87% Zn, 0.12 ounces Ag per ton and 0.005 ounces Au per ton (Fenwick and Scott, 1976). The felsic metavolcanic rock unit adjacent to the copper- and zinc-mineralized horizon yielded a U-Pb age of 2718+3 Ma (Corfu and Stott, 1998). A magnetic lamprophyre dyke, < 2m wide, crosscuts the pillowed flows at 065˚-080˚ and dips steeply north. Some of these late Neoarchean intrusions in this area were classified as kersantites (i.e. calc-alkaline, biotiteplagioclase-bearing lamprophyre) by Brown (1995). ACKNOWLEDGEMENTS Figure 61. Isolated-pillow breccia, STOP 2-7, showing amoeboid to angular pillow fragments in hyaloclastite-rich matrix. The authors would like to acknowledge the support and guidance of many former and present colleagues at the Ontario Geological Survey, Lakehead University and the Geological Survey Canada over the past several decades. This field guide has benefitted greatly from the comments, information and suggestions provided by Dr. Phil Fralick (Lakehead University), Riku Metsaranta (Ontario Geological Survey) and Dr. Wouter Bleeker (Geological Survey of Canada). Pete - 38 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Hollings assembled the final manuscript. We would also like to thank property owners who have provided permission to access several sites for the purposes of this field trip. REFERENCES Addison, W.D. and Brumpton, G.R. 2012. Field trips 1 & 13 Sudbury impactoclastic debrisites at Thunder Bay; In; Hollings, P., MacTavish, A. and Addison, W. (Eds.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 2 - Field trip guidebook, v.58, part 2, 2-26. 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. 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? Geological Society of America Special Papers, 2010, 465, p. 245-268. Barghoorn, E.S., Knoll, A.H., Dembicki, H., Jr., and Meinschein, W.G. 1977. Variation in stable carbon isotopes in organic matter from the Gunflint Iron Formation; Geochimica et Cosmochimica Acta 41, pp.425–430. Bleeker, W., Smith, J., Hamilton, M., Kamo, S., Liikane, D., Hollings, P., Cundari, R., Easton, M., and Davis, D., 2020, The Midcontinent Rift and its mineral systems: Overview and temporal constraints of NiCu-PGE mineralised intrusions, in Bleeker, W., and Houle, M.G., eds., Targeted Geoscience Initiative 5: Advances in the Understanding of Canadian NiCu-PGE and Cr Ore Systems—Examples from the Midcontinent Rift, the Circum-Superior Belt, the Archean Superior Province, and Cordilleran AlaskanType Intrusions: Geological Survey of Canada Open File 8722, p. 7–35, https://doi.org /10 .4095 /326880. Bowen, N.L. 1911. Silver in Thunder Bay District; Ontario Bureau of Mines Annual Report, v. 20, Part 1, pp. 119–132. Brown, G.H. 1995. Precambrian geology, Oliver and Ware townships; Ontario Geological Survey, Report 294, 48p. Brown, G.H. and Fogal, R.I. 1995. 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Geology of Forbes and Conmee townships; Ontario Geological Survey, Open File Report 5726, 188p. Carter, M.W., McIlwaine, W.H. and Wisbey, P.A. 1973: Nipigon-Schreiber, Geological Compilation Series, Thunder Bay District; Ontario Division of Mines Map Number 2232, at a scale of l inch to 4 miles or1:253,440. Cheadle, B.A. 1986. Alluvial-playa sedimentation in the lower Keweenawan Sibley Group, Thunder Bay District, Ontario; Canadian Journal of Earth Sciences, 23, pp.527–542. Corfu, F. and Stott, G.M. 1998. Shebandowan greenstone belt, western Superior Province: U-Pb ges, tectonic implications, and correlations; GSA Bulletin 110, pp.1467-1484. 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, pp.476-488. Davis, D.W. and Sutcliffe, R.H. 1985. 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(Eds.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 2 Field trip guidebook, v.58, part 2, 27-55. River–Disraeli Lake area, Nipigon Embayment, northwestern Ontario: lithogeochemical, assay and compilation data; Ontario Geological Survey, Miscellaneous Release—Data 133. Franklin, J.M. 1978. The Sibley Group, Ontario. In Rubidium–strontium isotopic age studies, Report 2. Edited by R.K. Wanless and W.D. Loveridge. Geological Survey of Canada, Paper 77-14, pp. 31– 34. Hayatsu, R., Winans, R. E., Newman, D. S., Mancuso, J.J. and Seavoy, R.E. 1983: Correlations between the chemical and the geologic origins of anthraxolite from the Gunflint Formation, Thunder Bay, Ontario; Economic Geology, Vol. 78, 1983, pp. 175-180. Franklin, J.M., Kissin, S.A., Smyk, M.C. and Scott, S.D. 1986. Silver deposits associated with the Proterozoic rocks of the Thunder Bay District, Ontario; Canadian Journal of Earth Sciences, v.23, pp.1576-1591. 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Penokean fold-andthrust deformation of the Paleoproterozoic Gunflint Formation near Thunder Bay, Ontario; 51st Institute on Lake Superior Geology, Annual Meeting, Nipigon, Ontario, May 2005, Proceedings Volume 51, Part 1, p.26. Fenwick, K.G. and Scott, J.F., 1976. 1975 Report of the North Central Regional Geologist; p. 39-54 in Annual report of the Regional and Resident Geologists, 1975, edited by C.R. Kustra, Ontario Division of Mines, Miscellaneous Paper 64, 146p. Geul, J.J.C. 1970. Geology of Devon and Pardee Townships and the Stuart Location; Ontario Department of Mines, Geological Report 87, 52 p. Geul, J.J.C 1973. Geology of Crooks Township, Jarvis and Prince Locations, and Offshore Islands, District of Thunder Bay; Ontario Department of Mines, Geological Report 102, 46 p. Goodwin, A.M. 1956. Facies relations in the Gunflint Iron Formation; Economic Geology, v.51, pp.565–595. Hamilton, J. S. 1996. Pleistocene landscape features and Plano archaeological sites upon the Kaministiquia delta, Thunder Bay District; Lakehead University Monograph in Anthropology #1, 112p. 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, no.8, p.1021-1040. 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 Magyarosi, Z. 2004. Northern Black Sturgeon Hollings, P., Cundari, R., Pulchalski, R. and Smyk, M.C. 2011. Geochemistry of Midcontinent Rift- related mafic intrusions, Thunder Bay area; Ontario Geological Survey, Miscellaneous Release—Data 261 – Revised. Hollings, P. and Smyk, M.C. 2008. Whatever happened to the Logan sills? Ongoing research into the geochemistry of Midcontinent Rift-related mafic intrusive rocks south of Thunder Bay: 54th Institute on Lake Superior Geology, Annual Meeting, Marquette, Michigan, May 2008, Proceedings Volume 54, Part 1, p.36-37. Hollings, P., Hart, T., Richardson, A., and MacDonald, C.A. 2007a. 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, no.8, p.1087-1110. Hollings, P.N., Smyk, M.C. and Hart. T. 2007b. Geochemistry of Midcontinent Rift-related mafic dykes and sills near Thunder Bay: New insights into geographic distribution and the geochemical affinities of Nipigon and Logan sills and Pigeon River and other dykes; 53rd Institute on Lake Superior Geology, Annual Meeting, Lutsen, Minnesota, May 2007, Proceedings Volume 53, Part 1, p.40-41. Hollings, P.N., Smyk, M.C. 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 - 40 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 thesis, Lakehead University, Thunder Bay, ON, 103 p. (2010) 553–571. Hughes, A. 2016. Petrology and geochemistry of the McKenzie Granite, northwestern Ontario; unpublished H.B.Sc. thesis, Lakehead University, Thunder Bay, Ontario, 78p. Hughes, A., Hollings, P., Smyk, M.C., Puumala, M.A. and Campbell, D.A. 2017. Petrology and geochemistry of the Neoarchean McKenzie Granite, northwestern Ontario; Ontario Geological Survey, Miscellaneous Release – Data 349, 14p. Jones R., Marcelissen, R., and Fralick, P., 2022. Sedimentology and stratigraphy of a large prevegetation deltaic complex; Frontiers in Earth Science 10, 27p., 875838. Kissin, S.A. 1992. Five-element (Ni-Co-As-Ag-Bi) veins; Geoscience Canada, v.19, pp.113-124. Koebernick, C.A. 1996. Neoarchean coastal sedimentation in the Shebandowan Group, Northwestern Ontario’ unpublished M.Sc. thesis, Lakehead University, Thunder Bay, ON, 248p. Koebernick, C.F. and Fralick, P. 1995. Neoarchean coastal sedimentation in the Shebandowan Group, northeastern Ontario; in 41st Annual Meeting, Institute on Lake Superior Geology, Proceedings Volume 41, Part 1 - Program and Abstracts, pp. 3132. Koroscil, J.P.J. 2013. Deformation of the Animikie Group north of Lake Superior; unpublished H.B.Sc. thesis; Lakehead University, Thunder Bay, Ontario. Krogh, T.E., Davis, D.W., Corfu, F., 1984. Precise U-Pb zircon and baddeleyite ages for the Sudbury area. In: Pye, E.G. (Ed.), The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, pp. 431–446. Kronberg, B.I. and Fralick, P.W. 1992. Geochemical alteration of felsic Archean rocks by Gunflint Formation-derived fluids, Quetico - Superior region, northwest Ontario; Canadian Journal of Earth Sciences, v. 29, pp.2610-2616. Kup, S., Tsujita, C.J., Jin, J., Shuster, J., Metsaranta, R.T. and Kurcinka, C.E. 2025. Microfossil preservation in the Paleoproterozoic Gunflint Formation: Comparative study of the Pass Lake and the classic Schreiber Beach localities, northern Ontario; in Summary of Field Work and Other Activities, 2025, Ontario Geological Survey, Open File Report 6421, p.9-1 to 9-14. Kustra, C.R., McIlwaine, W.H., Fenwick, K.G. and Scott, J.F. 1977. Proterozoic rocks of the Thunder Bay area, northwestern Ontario; Field Trip Guidebook, 23rd Annual I.L.S.G. Meeting, Thunder Bay, 47p. Kwiatkowski, D. 1975. Geology and geochemistry of the Kakabeka Falls anthraxolite; unpublished H.B.Sc. Landman, M. 2021. Archean orogenesis to Proterozoic rifting: A structural history of Pass Lake, Thunder Bay, Ontario; unpublished H.B.Sc. thesis, Lakehead University, Thunder Bay, Ontario. Machado, A.B., 1987. On the origin and age of the Steep Rock buckshot, Ontario, Canada; Chemical Geology, v. 60, pp. 337–349. MacNeish, R.S. 1952. A possible early site in the Thunder Bay district, Ontario, Annual Report of the National Museum of Canada for the Fiscal year 1950–51. National Museum of Canada Bull 126, p 25, pp. 27–28. Magnus, S. 2012. An investigation of the assimilation hypothesis in the Navilus sill, Thunder Bay, Ontario’ unpublished H.B.Sc. thesis, Lakehead University, Thunder Bay ON. Magnus, S., and Kissin, S., 2010. Assimilation and petrogenesis in the Navilus and Terry Fox sills, Thunder Bay, Ontario; in Institute on Lake Superior Geology, Proceedings and Abstracts, v. 56, part 1, p. 36-37. Mancuso, J.J., Kneller, W.A., and Quick, J.C. 1989. Precambrian vein pyrobitumen: evidence for petroleum generation and migration 2 Ga ago. Precambrian Research, v. 44, pp.137–146. Maric, M. 2006. Sedimentology and sequence stratigraphy of the Paleoproterozoic Rove and Virginia formations, southwest Superior Province; unpublished M.Sc. thesis, Lakehead University, Thunder Bay ON. McIlwaine, W.H. 1975. McTavish Township (southern half), District of Thunder Bay; Ontario Division of Mines, Preliminary Map P.990, scale 1:15 840. Metsaranta, R.T. 2022. Highlights of bedrock geology mapping in the Quetico Subprovince, north of Thunder Bay, northwestern Ontario; in Summary of Field Work and Other Activities, 2022, Ontario Geological Survey, Open File Report 6390, p.9-1 to 9-9. Metsaranta, R.T and Kurcinka, C.E. 2022. Reconnaissance bedrock geology mapping of the Animikie Basin and spatially associated Midcontinent Rift–related igneous rocks, Thunder Bay area, northwestern Ontario: A project introduction; Summary of Field Work and Other Activities, 2022, Ontario Geological Survey, Open File Report 6390, p.11-1 to 11-13. Metsaranta, R.T. and Walker, J.A. 2019. Precambrian geology of western McGregor Township and adjacent areas, northeast of Thunder Bay; in Summary of Field Work and Other Activities, 2019, Ontario Geological Survey, Open File Report 6360, p.11-1 to 11-10. Miller, J.D., Green, J.C. and Severson, M.J. 2002. - 41 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Terminology, nomenclature and classification of Keweenawan igneous rocks of northeastern Minnesota; in Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota; Minnesota Geological Survey, Report of nvestigations Moorhouse, W.W. 1960. Gunflint Iron Range in the vicinity of Port Arthur; Ontario Department of Mines, v.69(7): pp.1-40. Oja, R.V. 1967. Geochemical investigations of the Thunder Bay silver area; in Proceedings, Symposium on Geochemical Prospecting, Ottawa Ontario, 1967; Geological Survey of Canada, Paper 66-54, pp.211221. Parker, J.R. 1980. The Structure and Environment of Deposition of the Finmark Metasediments, Thunder Bay, Ontario; unpublished H.B.Sc. thesis, Lakehead University, Thunder Bay, ON Phillips, B. 2004. Of moraines, lake floors, deltas and shorelines: A brief summary of the deglaciation of the Kaministiquia embayment, Thunder Bay, Ontario; unpublished report, World Wide Website, http:// www.lakeheadu.ca/~geogwww/phillips/FOP%20 page_4.htm (accessed 2004). Phillips, B., Hill, C., Fralick, P. and Ross, B. 1994. Postglacial shorelines and Paleoindian migration along the northwestern shore of Lake Superior; Field Trip Proceedings of the 58th ILSG Annual Meeting -Part 2 Guidebook, 13th Biennial meeting of AMQUA, Minneapolis, MN. Phillips, B., Stewart, J., Hamilton, S., Julig, P. and Ross, B. 2000. Geoarchaeology of the Thunder Bay area; 46th Institute on Lake Superior Geology, Thunder Bay, Ontario, Field Trip Guidebook, 40p. Puchalski, R. 2010. The petrology and geochemistry of the Riverdale sill. Unpublished H.B.Sc. thesis, Lakehead University, Thunder Bay, ON. Pufahl, P., Fralick, P. and Scott, J.F. 2000. Geology of the Paleoproterozoic Gunflint Formation; in Institute on Lake Superior Geology, 46th Annual Meeting, Field Trip Guidebook. Purucker, M. 1983: Time of Formation of Soft Iron Ore on the Gunflint and Mesabi Ranges (Ontario, Canada and Minnesota, U.S.A.); Economic Geology, Vol. 78, 1983, pp. 502-506. 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. Pye, E.G. 1969. Geology and scenery, north shore of Lake Superior; Ontario Department of Mines, Geological Guidebook No.2, 148p. Rasmussen, B., and Muhling, J.R. 2019. Evidence for widespread oil migration in 1.88 Ga Gunflint Formation; Geology, v.4, pp.899–903. Rasmussen, B., Muhling, J.R. and Fischer, W.W. 2021. Ancient oil as a source of carbonaceous matter in 1.88-billion-year-old Gunflint stromatolites and microfossils; Astrobiology, v.21, Number 6, 2021; DOI: 10.1089/ast.2020.2376. Reid-Sharp, R. 2016. Characterizing Deformation of Gunflint Formation in Contact with the Archean Basement; unpublished H.B.Sc. thesis, Lakehead University, Thunder Bay, Ontario. Rezka, R. 1987. Shallow marine, storm-dominated deposits of the Shebandowan–Wawa greenstone belt; unpublished BSc thesis, Lakehead University, Thunder Bay, Ontario, 79p. Rogala, B. 2003. The Sibley Group: a lithostratigraphic, geochemical, and paleomagnetic study. Unpublished M.Sc. thesis, Lakehead University, Thunder Bay, Ontario, 254 p. Rogala, B., Fralick, P.W., and Metsaranta, R. 2005. Stratigraphy and sedimentology of then Mesoproterozoic Sibley Group and related igneous intrusions, northwestern Ontario: Lake Nipigon Region Geoscience Initiative. Ontario Geological Survey, Open File Report 6174, 87 p. Rogala, B., Fralick, P.W., Heaman, L.M., and Metsaranta, R. 2007. Lithostratigraphy and chemostratigraphy of the Mesoproterozoic Sibley Group, northwestern Ontario. Canadian Journal of Earth Sciences, v.44. Rutter, A.M. 2014. Analysis of anthraxolite and Precambrian carbonates of Kakabeka Falls, Ontario, Canada; American Geophysical Union, Fall Meeting 2014, abstract id. B41B-0040, Pub Date: December 2014 Bibcode: 2014AGUFM.B41B0040R. Scott, J. 1990. Geology of MacGregor Township; Ontario Geological Survey, Open File Report 5719, 82p. Sergiades, A.O. 1968. Silver cobalt calcite vein deposits of Ontario; Ontario Department of Mines, Mineral Resources Circular 10, 498p. Shegelski, R.J., 1980. Archean cratonization, emergence and red bed development, Lake Shebandowan area, Canada. Precambrian Research 12, pp.331-347. Shegelski, R.J. 1982. The Gunflint Formation in the Thunder Bay area; in Franklin, J.M. ed, Field Trip Guidebook 4: Winnipeg, Manitoba; Geological Association of Canada, p.14-31. Sindol, G.P., Babechuk, M.G., Petrus, J.A. and Kamber, B.S. 2020. New insights into Paleoproterozoic surficial conditions revealed by 1.85 Ga corestonerich saprolith; Chemical Geology, v.545, 5 - 42 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 July 2020, 119621, chemgeo.2020.119621. https://doi.org/10.1016/j. Smith, J.W., Bleeker, W. and Hamilton, M. 2025. The 1093 Ma Crystal Lake Intrusion: A nickel-copper mineralized intrusion emplaced during the younger southwest–northeast rift phase of the Midcontinent Rift (North America); GSA Bulletin, http://pubs. geoscienceworld.org/gsa/gsabulletin/article-pdf/ doi/10.1130/B37649.1/7366228/b37649.pdf. Smith, A.R. and Sutcliffe, R.H. 1989. Precambrian geology of Keweenawan intrusive rocks in the Crystal LakePigeon River area: Ontario Geological Survey, Map P.3139, scale 1:50 000. Smyk, M.C. 2012. Field Trip 5 – Guide to the Thunder Bay area; In; Hollings, P., MacTavish, A. and Addison, W. (Eds.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 2 - Field trip guidebook, v.58, part 2, 74-81. 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, pp.10411053, doi:10.1139/E07-013. Smyk, M. and Hollings, P., 2007. Midcontinent rift-related mafic intrusion north of the international border; Proceedings of the Institute on Lake Superior Geology v. 53: pp.53-80. Smyk, M.C. and Hollings, P. 2009. Geochemistry of Midcontinent Rift-related mafic intrusions, Thunder Bay area; Ontario Geological Survey, Miscellaneous Release-Data 261. 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. Geological Society, Special Publication No.42, 313-345. Sutcliffe, R.H. 1989. Mineral variation in Proterozoic diabase sills and dykes at Lake Nipigon, Ontario; Canadian Mineralogist, v.27, p.67-79. 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. Tyler, S.A. and Barghoorn, E.S., 1954. Occurrence of structurally preserved plants in Precambrian rocks of the Canadian Shield; Science v. 199, p.606-608. Williams, H.R. 1991. Quetico Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, pt.1, p.383-403. 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, p.485-539. Yip, C.I., 2016. Sedimentology and geochemistry of regressive and transgressive surfaces in the Gunflint Formation, northwestern Ontario; unpublished M.Sc. thesis, Lakehead University, Thunder Bay ON, 321 p. Zaniewski, K., Phillips, B. and Dean, F. 2020. Northwestern Ontario: The Thunder Bay region; February 2020; DOI:10.1007/978-3-030-35137-3_6; In book: Landscapes and Landforms of Eastern Canada, pp.159-177. Smyk, M.C., Hollings P. and Heaman, L.M. 2006. Preliminary investigations of the petrology, geochemistry and geochronology of the St. Ignace Island Complex, Midcontinent Rift, northern Lake Superior, Ontario; Institute on Lake Superior Geology, 52nd Annual Meeting, Sault Ste. Marie, ON, Program with Abstracts, v. 52, p.61-62. - 43 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Trip 2 - Geology of the Quetico Subprovince and Shebandowan greenstone belt north of Thunder Bay Riku Metsaranta and Gaetan Launay Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, Sudbury, Ontario, P3E 6B5 Introduction This field trip examines the geology of the southern Quetico Subprovince (QS) and its tectonically intercalated contact with the Shebandowan greenstone belt (SGB) north and west of the City of Thunder Bay. Much of the content of this guidebook is informed by a multiyear, 1:50 000 scale bedrock mapping project that is being carried out in the area by the Ontario Geological Survey. The area encompassed by this guidebook represents the southern half of the multiyear bedrock mapping project area (see Figure 1). At the time of this field trip, a new bedrock geology map of the southern half of the project area and associated data are in preparation. Fieldwork on the northern half of the larger project area should be completed during the summer of 2026. In total, the new mapping will cover an area of approximately 4200 km2 of which approximately 70% had never been mapped at the 1:50 000 scale prior to this work. Some of the results of this bedrock mapping are summarized in interim publications (Metsaranta 2015; Metsaranta and Walker 2019; Metsaranta and Hamilton 2020, Metsaranta and Kamo 2021, Metsaranta 2022, Launay and Metsaranta 2023, Launay and Metsaranta 2024). The field trip will focus on the Archean geology of the area depicted on Figure 1. Although this is a oneday field trip, we have included 15 stops dispersed over a large area. We will not be able to visit all stops in one day. The order of the stops is organized from west to east in the SGB, followed by a south to north traverse along Highway 527 across the QS. Stops are labelled as “Optional” or “Planned”. “Planned” outcrops are the stops we will endeavour to visit during the field trip. Optional stops are included to put into perspective many of the “Planned” stops. As they are easily accessible, participants can visit these “Optional” outcrops on their own. As we are attempting to visit a high number of outcrops over a large area in one day, time spent on each outcrop may be limited. UTM coordinates used throughout the guidebook are NAD 83 Zone 16. Background Regional Geological Context The Quetico Subprovince is a vast geological entity that extends, at minimum, from central Minnesota to western Quebec. In “subprovince-style” subdivisions of the Superior Province (e.g. Card and Cielieski 1986, Williams 1991) the QS is bounded to the north by the Wabigoon Subprovince and to the south by the Wawa Subprovince. In more recent subdivisions (e.g. Percival et al. 2006, 2012; Stott et al. 2010) of the Superior Province into “terranes” and “domains” the Quetico Subprovince is commonly referred to as the Quetico basin or Quetico terrane and it is bounded to the south by the Wawa-Abitibi terrane and to the north by the Western and Eastern Wabigoon terranes and the Marmion terrane. In this guidebook, we will refer to the “Quetico” as the Quetico Subprovince (QS) to avoid any interpretive tectonic implications. Similarly, rather than discussing subprovinces or terranes bounding the QS to the south, we will simply refer to the Shebandowan greenstone belt (SGB). The detailed geology of the QS (as a whole) is somewhat poorly understood. Systematic OGS mapping of large portions of the QS has not been carried out previously at 1:50 000 or 1:20 000 scale. Consequently, accurate bedrock geology maps of much of the QS do not exist, nor do large scale regional geochronology or geochemistry datasets that are tied to geological mapping. General regional geological syntheses of the QS are provided by Percival (1989) and Williams (1991). Additional synoptic descriptions of the QS are included in Percival et al. (2006, 2012) and these include a summary of existing geochronological constraints. Additional influential studies on the metamorphic history of the QS include Pan, Fleet and Heaman (1996); Valli et al. (2004) and a recent PhD study by Rehm (2025) among others. The QS has historically been interpreted to have been deposited in a fore-arc setting (e.g., Percival 1989, Williams 1991). In reality, the tectonic setting is likely more complex. Geochronology indicates most of the QS was deposited after circa 2700 Ma. However, - 44 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 1 (A) Total magnetic field image of the map area (Ontario Geological Survey 2017), underlain with lidar imagery (Ministry of Natural Resources and Forestry 2023). (B) Geological map (modified from Launay and Metsaranta, 2023) of the field trip area showing the location of stops presented in this guidebook. Note that Stops 13-15 are just to the north of the area portrayed by this map. Geological abbreviations: BLI, Barnum Lake intrusion; CCF, Crayfish Creek fault; CLI; HLG, Hilma Lake granite; HLI, Hadwen Lake intrusion; HLIC, Hades Lake intrusive complex; KF, Kingfisher fault; MFP, Moving Post fault PLIC, Penassen Lakes intrusive complex; QDZ, Quetico deformation zone; RLIC, Roll Lake intrusive complex; SFIC, Silver Falls intrusive complex; SIC, Shabaqua intrusive complex; TBLLF, Thunder Bay–Loon Lake fault. constraints vary by location (see discussion in Percival et al. 2006; and references therein) with some authors indicating deposition between approximately 2698 Ma and 2696 Ma and others indicating deposition after approximately 2692 Ma. Maximum depositional ages are poorly constrained because of the limited availability of representative and consistent detrital zircon datasets across the Quetico Subprovince, making stratigraphic and tectonic interpretations challenging. A framework for deformation and metamorphism summarized in - 45 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Percival et al. (2006 and references therein) suggests polyphase deformation and metamorphism spanned from close to the time of deposition to approximately 2650 Ma. Williams (1991) described four discrete deformation episodes affecting the QS. Although work on the metamorphic history of the Quetico differs regarding details, most work converges on the prevalence of subprovince-wide high temperature, low pressure metamorphism, which was likely preceded in some places by early medium-pressure and temperature metamorphism. Metamorphic grade increases in the eastern part of the QS where it reaches granulite-facies conditions (see Pan et al. 1998). Intrusive rocks are abundant in the QS, their characteristics are explored in the field guidebook and general characteristics are summarized by Williams (1991). Much attention has been paid to the boundary between the QS and the Wabigoon subprovince (e.g. the Beardmore-Geraldton greenstone belt), however, much less has been paid to the geology of the southern boundary. The SGB (and correlative greenstone belts in Minnesota) is a relatively narrow, arcuate greenstone belt, that extends from the Pass Lake area (east of Thunder Bay) to Northern Minnesota. The SGB is described in detail in Williams et al. (1991). A more recent geochronology based tectonostratigraphic framework for the SGB was proposed by Corfu and Stott (1998) and this remains in common usage as a stratigraphic and structural framework. Minor modifications to the Corfu and Stott (1998) framework have been added by Lodge (2016). In contrast to the QS, much of the SGB has been mapped by the OGS at 1:20 000 scale. However, much of this mapping was carried out prior to technological advancements like the widespread use of U-Pb geochronology, routine high precision trace element geochemistry, access to lidar imagery and high-resolution airborne magnetic data. That said, the OGS also has an on-going multiyear bedrock mapping project in progress to map much of the eastern part of the SGB at 1:20 000 scale (e.g., Lodge 2014; Ratcliffe 2016, 2017, 2019). The general greenstone belt-wide tectonostratigraphic framework for the SGB described by Corfu and Stott (1998) includes circa 2720 Ma aged rocks of the Greenwater assemblage (mainly tholeiitic mafic metavolcanic rocks and lesser ultramafic metavolcanic rocks, mafic-ultramafic intrusive rocks and metasedimentary rocks), circa 2718 Ma age rocks of the Burchell assemblage (mainly calc-alkalic felsic to intermediate metavolcanic rocks), circa 2695 Ma aged rocks of the Kashabowie assemblage (mainly calc-alkalic intermediate metavolcanic and metasedimentary rocks), circa 2690Ma aged rocks of the Shebandowan assemblage (calc-alkalic intermediate metavolcanic rocks, shallow marine and fluvial metasedimentary rocks) and younger than circa 2682 Ma aged rocks of the Auto Road assemblage (conglomerate). Corfu and Stott (1998) envisaged a structural history of D1 thrusting that interleaved the Greenwater assemblage along with the Burchell assemblage with the Kashabowie assemblage followed by a regional unconformity overlain by younger rocks of the Shebandowan and Auto Road assemblages deposited during D2 transpression. D2 transpression ceased by about 2680 Ma. According to the framework of Corfu and Stott (1998) intrusive rocks in the SGB include older gneissic tonalitic rocks to the south of the belt with ages as old as approximately 2750Ma, syn-Greenwater assemblage mafic-ultramafic intrusions, syn-Kashabowie assemblage tonalite, syn-Shebandowan assemblage monzodiorite-granite (Tower stock) and post-tectonic circa 2680 Ma aged intrusions of biotite-hornblende bearing diorite to granodiorite. Locally, evidence for magmatic rocks with ages around 2710 Ma are present in some parts of the SGB, e.g. Kabaigon porphyry (Corfu and Stott 1998). Geology of field trip area This fieldtrip will examine the geology of four distinct geological domains (see Figure 1): 1) the northeastern part of the Shebandowan greenstone belt, 2) the Lappe domain, 3) the southern Quetico domain and 4) the Dog Lake injection complex (Figure 1). Figure 2 is a geological timeline summarizing important events affecting the different geological units in the QS-SGB boundary zone and southern QS compiled from unpublished OGS data and various other sources. Synoptic reviews of these domains are given below. Additional details are provided in the field trip stop descriptions and will be augmented by discussions in the field. - 46 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 2 Geological timeline summarizing the main volcanic, sedimentary, intrusive, and structural events affecting the Shebandowan greenstone belt and the Quetico Subprovince. Ages are compiled from Corfu and Stott, 1998; Corfu, 2000; Kamo, 2013; Wang et al., 2020 and preliminary OGS data. Age bars include analytical uncertainties. Shebandowan greenstone belt (Stops 1, 2, 3, 5 and 6) This field trip guidebook only examines a narrow portion of the northern margin of the SGB that was covered by our mapping (Figure 1). In this area, the northern boundary of the SGB is generally eaststriking dextral shear zone interpreted to be the eastward extension of the Crayfish Creek fault. (Figure 1). Within this area of the SGB, we recognize four mappable units at the 1:50 000 scale which are subdivided into two informal groups, each comprising two informal formations. The older Greenwater group (a less repetitive Group-level name should be devised) consists of the Greenwater and Mud Lake formations and the younger Shebandowan group consists of the Strawberry Hill and Auto Road formations. Although not specifically identified in our area of mapping, we would include the “Kashabowie assemblage” (e.g. Corfu and Stott 1998) with the Shebandowan group. These units correspond with the “older” and “younger” portions of the SGB as described by previous workers (e.g., Corfu and Stott 1998, Lodge 2016) in most respects. However, we feel that moving towards a “sub-assemblage level” nomenclature is warranted to begin a framework for more detailed characterization of supracrustal rock variability at the regional scale. This is particularly true for rocks of the Shebandowan assemblage (in the sense defined by previous workers) which is lithologically heterogeneous. Greenwater group The Greenwater formation (Stops 1 and 6) consists mainly of tholeiitic mafic metavolcanic rocks, minor ultramafic metavolcanic rocks, synvolcanic gabbroic intrusions and minor clastic and chemical metasedimentary rocks. We do not have specific age constraints on the Greenwater formation; however, we infer that it is part of the “older” SGB based on lithological similarities with rocks of the Greenwater assemblage sensu stricto. At more detailed mapping scales, the Greenwater formation could likely be further subdivided (e.g. mafic dominated vs ultramafic dominated portions). The Mud Lake formation (Stop 1) consists mainly of calc-alkalic fragmental volcaniclastic rocks and locally coherent flows of felsic to intermediate composition. In the area, Corfu and Stott (1998) determined the age of this unit to be 2718 +/- 3 Ma. These are likely equivalent to Burchell assemblage of Lodge (2016). - 47 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Shebandowan Group The Strawberry Hill formation (SHF, Stop 2) typically consists of coarse-grained polymictic breccias characterized by hornblende phenocrysts in a dark matrix. The SHF is typically mafic to intermediate and calc-alkalic. It typically has elevated magnetic susceptibility and displays pink to red hematite alteration and/or pale green epidote alteration. Locally, it also occurs as more massive, hornblende-phyric mafic flows or shallow intrusions. At some localities, breccias or flows are associated with thinly layered tuffs, and reworked tuffs with clear pyroclastic textures such as bombs that deform layering. Geochronology by Corfu and Stott (1998) and our own work indicate deposition/eruption of the SHF at circa 2690 Ma. Although the bulk of SHF rocks are undeformed, they are locally cut by narrow ductile shear zones, and it is in sheared contact with older rocks. The SHF forms a thin, but important marker unit that can be correlated across much of the central SGB in the area depicted by Figure 1. Breccias of the SHF are compositionally similar to, and of the same age as the Tower stock (just west of the area depicted in Figure 1) which hosts low-grade disseminated, intrusion related gold mineralization and includes marginal breccias similar to the SHF (e.g. Carter 1992). The Auto Road formation (ARF, Stops 3 and 5) comprises mainly polymictic, matrix supported conglomerate. Local occurrences of trough crossstratified, likely fluvial sandstones, are also considered part of the ARF. Calc-alkalic mafic metavolcanic rocks, including apparently pillowed flows are locally intercalated with ARF sedimentary rocks. The ARF was deposited after 2682 ± 3 Ma based on geochronology in Corfu and Stott (1998) and as such, it clearly postdates the SHF. Dextral shear zones are a common feature of the SGB. Although it is not clear in Figure 1, geophysical patterns in adjacent areas like the LD and the SGB outside of our mapping area suggest that D2 shear zones post date D1 thrust faults. Sinistral northeast-trending shear zones in the SGB such as the Kingfisher and Thunder Bay-Loon Lake faults are relatively younger than the dextral shear zones based on mapping inferred off-sets. Potassium-rich calc-alkalic suite intrusions (PRCAS) form a minor component of the SGB as shown on Figure 1. These include massive to weakly foliated hornblende-biotite-magnetite quartz monzonite to monzogranite dominated intrusions of unknown age. These may be related to similar intrusions in the SGB like the Kekekaub pluton (circa 2680 Ma) or perhaps correlate with the Tower stock (circa 2690Ma). Lappe Domain (Stops 4, 7 and 9) The Lappe domain comprises mainly metasedimentary rocks (wacke and siltstone) similar to those of the southern Quetico subprovince to the north. Its southern boundary is the Crayfish creek fault whereas it is bounded to the north by the Moving Post fault. The LD is characterized by thin fault bounded panels of mafic metavolcanic and mafic intrusive rocks comparable to the Greenwater formation intercalated with the metasedimentary rocks. Where best preserved, these mafic panels contain pillowed mafic flows, local banded iron formations, local thin ultramafic schists (sheared flows or thin sills) and local sulfidic mudstones. The margins of the mafic panels are commonly sheared and often preserve well developed steeply plunging stretching lineations indicating dip-slip, probable thrust motion. Locally LD rocks are folded, however we do not have a sufficient coverage of detailed younging data or outcrop scale fold observations to determine the nature of folding in the LD. Although we do not have direct age constraints on metavolcanic rocks in the mafic panels, preliminary data suggests some thin gabbro bodies in LD metasedimentary rocks are intrusive (i.e. not structurally interleaved). These provide minimum age constraints for LD sedimentation; combined with preliminary detrital zircon data, and considering analytical uncertainties, LD sedimentation is bracketed between about 2698 and 2689 Ma. At another locality, Corfu (2000) determined and age of circa 2718 Ma for a gabbro within one of the mafic panels in Ware township. If this age is reliable, then some of the mafic rocks in the LD correlate with the Greenwater formation. Another hypothesis to consider is that this age could reflect zircon inheritance. Regardless of the age of the mafic panels and whether they represent tectonic slivers of older rocks, or if they are part of the “stratigraphy”, observed thrusts faults indicate that the boundary between the SGB and the QS likely represents a zone of fold-thrust deformation. The timing of LD sedimentation corresponds with the timing of deposition of the Kashabowie assemblage in the SGB (see Corfu and Stott 1998). Observed thrust - 48 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 faults in the LD may therefore correspond to “D1” of Corfu and Stott (1998) and this interpretation links the timing of early QS deformation with early deformation in the SGB. East of the Kingfisher fault, large, multiphase, PRCAS intrusive complexes, (the Penassen Lakes intrusive complex and the Roll Lake intrusive complex) were emplaced into the LD. These intrusions consist of an early mafic phase (hornblendite to monzogabbro) roughly coeval with intermediate a hornblende-biotitemagnetite monzodiorite to quartz monzonite phase and a late phase that is volumetrically dominant and composed of biotite monzogranite. Compositionally similar rocks are found to the north in the QS and the Dog Lake injection complex and ages determined for the different phases are consistent regionally as summarized in Figure 2. Locally, peraluminous granitic pegmatites occur in the northern LD (e.g. Walkinshaw pegmatites) but are not associated with any obvious parental granite. LD metasedimentary rocks are typically low metamorphic grade, dominated by biotite or chlorite, quartz, plagioclase assemblages. However, near intrusions they locally contain contact metamorphic porphyroblasts of andalusite and/or cordierite and in some areas experienced partial melting. Southern Quetico domain (Stops 10, 11 and 12) The southern Quetico domain comprises mainly metasedimentary rocks (wacke and minor siltstone) with rare intermediate tuffaceous horizons, rare mafic tuff (possibly boninitic). Bedding and foliations in the southern QS are typically east to northeast trending and “D2” folds typically have east to northeast trending axial surfaces. Southern QS metasedimentary rocks gradually become more recrystallized towards the north. Metamorphic assemblages are typically dominated by biotite and locally biotite-garnet. Andalusite- and cordierite-bearing metamorphic assemblages are also relatively common and may related to proximity to intrusions. Staurolite-bearing assemblages are present locally but rare. The southern QS was intruded by numerous potassium-rich calc-alkalic suite intrusions (PRCAS). These are depicted on Figure 1 and include from west to east, the Shabaqua intrusive complex, the Silver Falls intrusive complex, the Trout Lake intrusion, the Barnum Lake intrusion, the Whitelily Lake intrusive complex and the Hades Lake intrusive complex. These intrusions and intrusive complexes are variably complex mixtures of mafic (hornblendite-monzogabbro), intermediate (monzodiorite-quartz monzonite) and felsic (monzogranite) phases. S-type granites are also common in the southern QS, these include the Hilma Lake granite, the Voutilainen intrusion and the Hadwen Lake intrusion. Peraluminous granitic pegmatites are also abundant and spatially related to S-type granite bodies. Dog Lake injection complex and Quetico deformation zone (Stops 12, 13, 14 and 15) Strain and degree of metasedimentary rock recrystallization increase abruptly in the vicinity of the Quetico deformation zone (QDZ) in the northern part of the area. In this domain, QS metasedimentary rocks are recrystallized and comprised mainly of biotite-quartzfeldspar+/- magnetite paragneiss and locally also sillimanite-cordierite-garnet bearing paragneiss. The Dog Lake injection complex (DLIC) is characterized by paragneiss intruded by a high volume of both peraluminous and HPCAS intrusions (injections) that were emplaced synchronously with intense dextral transpression along the QDZ. Intrusions of both suites are commonly schlieric with strong fabrics defined by schlieren and magmatic minerals. At the contact with HPCAS intrusions, paragneisses are commonly strongly magnetic which likely resulted from their oxidation by fluids exsolved from these HPCAS intrusion triggering the crystallisation of magnetite. These zones are also particularly rich in biotite which locally give the paragneisses the appearance of melanosome. Migmatitic rocks are a volumetrically minor component of the DLIC and comprise mainly patchy metatexite likely related to heating by the high volume of intrusive rocks in the area. Syn-tectonic fabrics are ubiquitous in intrusive rocks of the DLIC. Strong fabrics are generally steep and east striking to east-northeast striking. Dextral shear bands with strikes of approximately N290-300 are common as are approximately N40 striking and N15-20 striking subvertical sinistral shear bands. C-S fabrics are common and suggest syn-emplacement dextral strike slip and locally north side up thrust components of motion. Shearing related fabrics in the paragneisses have the same orientations and kinematics and syn-tectonic emplacement fabrics in the granitoid rocks. Folding of gneissosity is common and folding - 49 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 tends to be gently plunging. Intense, steeply dipping, east striking mylonitic zones are present locally. Late brittle-ductile strike slip and thrust motion occurred along the QDZ locally overprinting ductile fabrics. This pattern is repeated at map scale where the sinistral north to northeast-trending shear zones branches onto dextral east-trending QDZ in the Dog Lake area (see Figure 1). These observations suggest a transition from an earlier ductile transpressive deformation to later more brittle-ductile deformation. Field trip stop descriptions Stop 1 (Optional) - Mafic metavolcanic rocks of the Greenwater formation and felsic to intermediate metavolcanic rocks of the Mud Lake formation 315073E 5376695N (Greenwater formation) 314610E 5377220N (Mud Lake formation) Park on the shoulder of Mud Lake Road. Wear reflective vests, stay on shoulder or in the ditch. Traffic is heavy on Hwy 102 so be very cautious crossing the highway. A series of outcrops in this area shows mafic metavolcanic rocks considered to be part of the Greenwater formation and felsic metavolcanic rocks of the Mud Lake formation. These two formations are typical of the older part of the Shebandowan greenstone belt. From the junction of Hwy 102 and Mud Lake Road, outcrops immediately to the east and west are mainly mafic metavolcanic rocks of the Greenwater formation. These rocks display well preserved volcanic features such as prominent pillows (Figure 3A) and local pillow breccias. The pillowed flows are subvertical and striking to N75. Based on pillow cusps, they appear to young northward. Quartzepidote veining, local red-pinkish alteration and east striking brittle-ductile shear zones can also be seen at this outcrop. Younger, mica-phyric mafic lamprophyre dikes are also present. Farther west, closer to the north end of Mokomon Lake (314610E, 5377220N) felsic metavolcanic rocks of the Mud Lake formation comprising tuff breccias, lapilli tuffs and locally flows are present in outcrops located on the north side of the highway. These felsic metavolcanic rocks host a thin, semi-massive sulfide (sphalerite, pyrite, pyrrhotite) horizon known as the Mud Lake VMS occurrence (Figure 3B). The rocks Figure 3. Metavolcanic rocks of the Greenwater and Mud Lake formations. A) Large pillows in mafic flow, Greenwater formation. B) Sulfide mineralization in felsic metavolcanic rocks, Mud Lake formation. here are strongly foliated (260/75) with local shearing (234/80) and locally cut by rusty shallow dipping faults. Felsic rocks from this outcrop have an age of 2718 +/- 3 Ma based on Corfu and Stott (1998). The age of Greenwater formation (Greenwater assemblage) rocks throughout the greenstone belt is mainly inferred from adjacent felsic to intermediate units and maficultramafic intrusions. Although the contact between these two units appears sharp at this location, mafic rocks of the Greenwater formation are intercalated with felsic to intermediate rocks of the Mud Lake formation in other locations, suggesting that the contact may have originally been gradational. - 50 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Stop 2 (planned) - Strawberry Hill formation, Shebandowan group 312462E 5378109N Park along the access road to the quarry on the southern side of Dawson Road. Participants must always wear reflective vests and remain on the road shoulder, as traffic along this section is relatively dense and driver visibility is poor. This stop exposes a representative outcrop of the Strawberry Hill formation of the Shebandowan group. It consists of a massive, undeformed, mafic to intermediate calcalkaline, hornblende-phyric, matrix-supported breccia (Figure 4A). The breccia is polymictic, containing angular to subrounded clasts composed predominantly of plagioclase-phyric, medium- to coarse-grained pink monzonite, along with subordinate clasts of mafic and intermediate volcanic rocks, all set within a fine-grained, dark-green matrix. Clasts of monzonitic composition locally preserve an internal magmatic foliation, highlighted by the alignment of plagioclase phenocrysts. This magmatic fabric suggests that the intrusion from which the clasts were derived was likely emplaced syn-tectonically. The absence of visible bedding, combined with the generally angular to sub-rounded nature of the clasts, suggests formation in a high-energy volcanic environment, possibly associated with a cryptodome, and likely proximal to the volcanic source. Alternatively, this unit may represent a subvolcanic magmatic breccia, comparable to the breccias around the Tower stock described by Carter (1992). Based on TIMS U–Pb ages reported by Corfu and Stott (1998), magmatism related to this breccia likely occurred at approximately 2692±6 Ma. Thus, the SHF is likely contemporaneous with the 2690 ± 3 Ma Tower stock (Corfu and Stott 1998) Although the unit is largely massive and lacks visible foliation, discrete mylonitic shear bands are locally present (Figure 4B). These shear bands are commonly associated with carbonate alteration and quartz–carbonate veining. The shear zones generally strike northeast and locally preserve kinematic indicators consistent with a thrust motion, indicating a northsideup sense of shearing (Figure 4B). The breccia is also cut by multiple generations of quartz–carbonate veins, indicating postdepositional brittle deformation coeval with hydrothermal activity. These veins locally offset both clasts and matrix but do not significantly disrupt the overall massive character of the breccia. Stop 3 (Planned) - Auto Road formation, Shebandowan group 316851E 5378641N Park at the entrance of the private dirt road south of Korpela Road. As Korpela Road is narrow, please ensure that your vehicle does not obstruct traffic or restrict access along the road. Figure 4 Representative photographs of the Strawberry Hill formation outcrop (Stop 2). (A) Massive poorly sorted matrix supported polymictic intermediate breccia. (B) Discrete mylonitic shear band with asymmetric kinematic indicators (C-S fabric) indicating a north-side-up sense of shearing. This stop exposes a representative outcrop of the Auto Road formation of the Shebandowan group. It consists of a strongly foliated, poorly sorted, matrixsupported polymictic conglomerate. The conglomerate contains predominantly pebble to boulder-sized clasts of plagioclasephyric, medium to coarse-grained pink monzonite and monzogranite, many of which are - 51 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Road formation deformation event to approximately 2688 ± 0.8 Ma, whereas TIMS U–Pb ages reported by Corfu and Stott (1998) indicate a maximum depositional age of 2682 ± 2 Ma for the conglomerate. Together, these ages suggest that the intrusions from which the clasts were derived were deformed after 2688 Ma, prior to erosion and deposition of the conglomerate sometime after 2682 Ma. The conglomerate was subsequently deformed and affected by dextral shearing likely related to the Crayfish Creek fault. Stop 4a (Planned) - Sheared mafic rocks, mediumbedded wackes and thrust deformation at the Northern boundary of Lappe domain. 326238E/ 5383267N Park along Moving Post Road or in the parking lot of the old Lappe Store if permission is obtained. Wear reflective vests, be mindful of traffic, shoulders of the road are narrow and visibility is poor. Figure 5 Representative photographs of the Auto Road formation outcrop (Stop 3). (A) Strongly foliated polymictic conglomerate. (B) Dextral asymmetrical C-S fabric wrapping around a clast of pink monzogranite. Note the discordant internal foliation within the clast. strongly flattened parallel to the foliation (Figure 5A). The matrix is sandy, medium to coarse-grained, and characterized by a darkgreen color. Several monzonitic and monzogranitic clasts preserve an internal foliation that is discordant with the matrix foliation (Figure 5B), indicating that these intrusive rocks were deformed prior to erosion and deposition. This relationship suggests a preAuto Road formation deformation event affecting the source intrusions. The conglomerate exhibits a strong east–west trending foliation and is overprinted by a northwesttrending (approximately N300°) dextral shearing, highlighted by the development of asymmetric C–S fabrics wrapping around the clasts (Figure 5B). Preliminary ages obtained from a foliated intrusive clast constrain the maximum age of the preAuto This outcrop illustrates strongly deformed, tholeiitic mafic metavolcanic rocks (Figure 6A) of the northernmost mafic metavolcanic unit of the Lappe domain. Mafic metavolcanic rocks at this exposure are characterized by N250 striking north-dipping strong foliation and well-developed northward plunging lineation. The apparent dip-slip motion along the shear zone is north-side-down. In its present geometry, the true kinematics of the dip-slip motion along this structure is equivocal (Figure 6B). However, an interesting and commonly repeating pattern along strike is that metamorphosed north dipping wackes located north and south of this mafic metavolcanic unit consistently young southward indicating that the whole stratigraphic succession is overturned. This pattern was also observed in the past OGS mapping campaign of MacDonald (1939, observe printed version of old map). If our interpretation of the kinematics of this structure and the facing direction in the bounding metagreywacke units are correct, a possible interpretation of this structure is that it represents an originally northward verging thrust that was subsequently steepened and overturned. The stratigraphic relationship between the mafic metavolcanic rocks at this locality and surrounding wackes is not well constrained. Commonly both contacts of the mafic unit are sheared and there is a paucity of rocks suitable for geochronology in the exposures that we have mapped. Reliable younging - 52 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Stop 4b (Optional) - Southward younging Lappe domain metasedimentary rocks 326207E 5382509N This optional outcrop is located approximately 750 m south of Stop 4a along Dog Lake Road. It exposes mediumbedded wackes displaying a southward younging direction, as indicated by graded bedding and load casts (Figure 7A and 7B). The southward younging of these clastic metasedimentary rocks occurs near the mafic metavolcanic rocks and associated thrust fault observed at Stop 4a, providing important constraints on local stratigraphic facing and structural relationships. Figure 6 Representative photographs of sheared mafic volcanic flow from the Lappe Domain (Stop 4). (A) Strongly foliated and sheared mafic volcanic flow. (B) C-S fabric wrapping around quartz eyes indicating north-side down sense of shearing indicators within the mafic metavolcanic unit are almost nonexistent. At one locality, farher to the east, southward younging pillows were observed along a similar structure in a similar setting. We have not observed strong evidence of stratigraphic continuity between the surrounding wackes and the mafic unit. In this case, observed younging directions do not argue against stratigraphic continuity, however the sheared nature of contacts makes interpretation difficult. A thin, mica-phyric lamprophyre dike is also present at this locality. Mafic lamprophyre dikes are common throughout the area in the SGB, LD and QS. They have not been successfully dated and their contact relationships and relationships to structures and other intrusive suites is difficult to interpret as in many places relative age relationships are contradictory. This may Figure 7 Clastic sedimentary rocks of the Lappe Domain indicate multiple generations of lamprophyric mafic (Stop 4b). (A) Medium bedded wackes with graded beds. (B) intrusion are present regionally. Flame structure showing a southward younging direction. - 53 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Stop 5 (Optional) - Cross bedded sandstone, Auto Road formation, Shebandowan group 329269E 5379186N Park along the south side of Peterson Road, near the entrance to the private residential access road located on the curve. This outcrop is situated on private property. Authorization from the landowner is required prior to accessing the outcrop, and participants must ensure that permission has been obtained before entering on the property. This large outcrop exposes a well-preserved section of sandstone in sheared contact with a calc-alkaline mafic volcanic flow (Figure 8A). The sandstone displays a variety of well-developed sedimentary structures, including crossbedding and channelized geometries, indicative of a fluvial depositional environment (Figure 8B). The sandstone locally contains plagioclase phenocrysts, suggesting a potential juvenile volcaniclastic component. The mafic lava flow is locally pillowed and occurs in sheared contact with the sandstone (Figure 8C). The shear zone is characterized by a penetrative east– west-striking foliation that dips steeply to the south. A well-developed stretching lineation, plunging steeply (~55°) to the east, is observed on foliation planes. Locally, kinematic indicators are preserved and indicate a dextral sense of shearing. Together with the steeply plunging lineation, these observations suggest a dextral transpressional deformation with a top-to-the northwest thrust component. Younging directions determined from cross bedding indicate a consistent southward younging across the outcrop. The sandstone is also crosscut by late, narrow mafic dikes, indicating post-depositional magmatic activity (Figure 8D). Although samples collected for U–Pb geochronology Figure 8 Outcrop of cross-bedded sandstone of the Auto Road formation (Stop 5). (A) Aerial drone photograph showing crossbedding and channel structures. (B) Close up on cross-stratified sandstone with southward younging direction. (C) Strongly sheared and foliated mafic volcanic flow occurring within the sandstone. (D) Late mafic dikes crosscutting sandstone. - 54 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 did not yield datable mineral phases, the sandstone is interpreted to be part of the Shebandowan group, most likely the Auto Road formation, based on its characteristic fluvial depositional environment. Stop 6 (Planned) - Greenwater assemblage mafic metavolcanic rocks, eastern extension of Crayfish Creek deformation zone 340868E 5376210N Park on the shoulder of Mount Baldy Road at its junction with Hwy 527. Wear high visibility vests. The shoulder here is narrow and logging truck traffic can be heavy. Be careful if crossing the highway. Footing is uneven and there are commonly garbage and glass in the ditches. This outcrop represents the sheared contact between the Shebandowan greenstone belt (to the south) and the Lappe domain (to the north) and may represent the eastward extension of the Crayfish Creek deformation zone. The northern part of the outcrop consists mainly of east-striking, steeply south-dipping, sheared, ankerite altered, magnesium-rich mafic rocks (Figure 9A). During mapping, kinematics of the shearing at this locality were not determined confidently. Precise identification of protoliths in the northern part of the outcrop is problematic as most primary features were obliterated. However, towards the south, rock types are well preserved and include massive, fine- to mediumgrained mafic volcanic flows (Figure 9B), mafic pillowed flows, a thin pyrite-bearing nodule black mudstone, and quartz-feldspar porphyritic intermediate dikes. Enigmatic weakly boudinaged dikes of mafic to ultramafic composition locally cut the main shearing fabric in the northern part of the outcrop. Minor Proterozoic calcite veins are also present. Lappe domain metasedimentary rocks to the north were deposited after circa 2700 Ma and perhaps after circa 2690 Ma, depending on the interpretation of preliminary detrital zircon data. Based on geochronology performed elsewhere in the SGB, the Greenwater formation is inferred here to have an age of circa 2720 Ma. Therefore, this shear zone juxtaposes rocks that differ in age by at least 20 million years. At this locality, shearing could represent a transpressive dextral reactivation of an earlier shear zone that interleaved units of disparate age. Note that there does not appear to be a rhyolitic unit equivalent to the Mud Lake formation north of the Greenwater formation as seen in Stop 1. This feature could be attributable to either a fault-related subtraction or a lateral stratigraphic discontinuity. The “QFP” dikes cutting the Greenwater formation here have the same appearance as dikes commonly observed in the Lappe domain and in the southern Quetico. Unfortunately, these dikes have proven difficult to date due to the lack of mineral phases amenable to U-Pb geochronology. Pyritic black shales like those at this locality occur locally in the Greenwater and Mud Lake formations. Figure 9. Greenwater formation metavolcanic rocks (Stop 6). (A) ankerite altered mafic-ultrmafic schist. (B) Massive, plagioclase-phyric mafic flow. - 55 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Stop 7 (Optional) - Lappe domain metasedimentary rocks 341256E 5378262N Park at the “weigh-scale” on the west side of Highway 527. Wear high visibility vests and use a high level of caution crossing the highway to the outcrop. Ditches also have uneven footing and garbage. This outcrop illustrates low metamorphic grade rocks of the southern Lappe domain. Here ~ 250° striking, north dipping but southward younging metasedimentary rocks include thin- to mediumbedded intercalated siltstone and sandstone (Figure 10A) overlain by very thickly bedded poorly sorted volcaniclastic sandstone (Figure 10B). Near the base of the thick sandstone bed, there is a folded horizon of thinly interbedded sandstone and siltstone. This may represent syn-depositional soft deformation. Locally, sparse carbonate nodules are present in the outcrop; these have been interpreted as early diagenetic features. Farther north, calc-silicate nodules (amphibole, epidote, locally garnet) are common in the QS and probably represent more metamorphosed equivalents of these carbonate nodules. As alluded to in the description of Stop 6, Laserablation-ICP-MS zircon geochronology was carried out on samples from this outcrop. This data suggest deposition of these rocks about the same time as much of the Shebandowan group. However, analytical precision does not permit precise chronostratigraphic correlation with the Strawberry Hill or Auto Road formations. Maximum depositional ages for tidally influenced shallow marine sediments in the Finmark area have maximum depositional ages of about 2691 Ma based on limited population, single crystal IDTIMS geochronology reported by Corfu and Stott (1998). The Finmark metasedimentary rocks occur in a similar structural setting based on geophysical interpretation. Stop 8 (Optional) - Northern margin of the Penassen Lakes intrusive complex 345318E 5385843N This stop requires parking on the shoulder of Hwy 527. Use hazard lights and traffic cones to increase visibility. Wear reflective vests and be mindful again of the traffic on Hwy 527. Be also mindful of soft shoulders when parking and walking. This outcrop represents part of the northern contact of the Penassen Lakes intrusive complex. At this outcrop early hornblende monzodiorite is crosscut by intermediate aged hornblende quartz monzonite, which is cut by late pink leucocratic biotite monzogranite dikes (Figure 11A and 11B). Also visible are narrow mica-phyric mafic lamprophyre dikes and xenoliths of metawacke and possibly mafic metavolcanic rocks. Figure 10. Lappe domain metasedimentary rocks at Stop 7. (A) steeply dipping, southward younging, thin- to mediumbedded, low metamorphic grade siltstone and sandstone. (B) Poorly sorted, lapilli and intraformational-sedimentary-clast bearing thick-bedded volcaniclastic sandstone. Multiphase, oxidized (magnetite bearing) intrusive complexes are a common feature of the SGB, LD and QS (see Figures 1 and 2). Early mafic phases of these intrusive complexes were emplaced between about 2677 and 2670 Ma, whereas later pink monzogranite phases appear to be about 5 million years younger (Figure 2). Although relative and absolute age differences between different phases of these intrusions are observed, we refer to them collectively as the potassium-rich calcalkalic suite. - 56 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Stop 9 (Planned) - Lappe Domain metasedimentary rocks and Walkinshaw peraluminous granitic pegmatites 346697E 5388265N Turn to the east on side road located just south of Stop 9 coordinate (Seagris Rd; no sign). If this side road is too rough for vehicles, use the road leading to summer camps on the northern end of Walkinshaw Lake, farther north. Wear high visibility vests, stay on shoulder or in ditch. Use extreme caution when crossing the road. Outcrops are relatively tall at this locality, be mindful of the potential for falling rocks in places. Figure 11. Northern margin of the Penassen Lakes intrusive complex at Stop 8. (A) Metasedimentary country rock fragments intruded by grey monzodiorite, crosscut by pink quartz monzonite to monzogranite dikes. (B) Amphibolephyric quartz monzonite with xenoliths of metamorphosed wacke. We consider this area part of the Lappe domain as we are south of the northernmost mapped mafic metavolcanic panel and the inferred eastward extension of the Moving Post fault. The Penassen Lakes intrusive complex and the similar Roll Lake intrusive complex to the north appear to post date the Moving Post fault based on geophysical interpretation and the presence of amphibolitic mafic pillowed flows occurring as large, strongly deformed inliers within the RLC and the emergence of a thin shear zone bounded mafic panel on the east side of the RLIC (Figure 1). This provides a clear relative age constraint on the Moving Post fault. These rocks are not, however, post-tectonic as we will see in later stops. This long outcrop contains relatively undeformed metasedimentary rocks (Figure 12A and 12B) in the northern part of the Lappe domain along with peraluminous granitic pegmatite dikes locally containing green mica, black tourmaline (Figure 12C and 12D) and disseminated molybdenite. These pegmatites we refer to as the Walkinshaw pegmatites. This area of metasedimentary rocks is surrounded by several large intrusive complexes (Potasssium-rich calc-alkalic suite), the Whitelilly Lake, Roll Lake and Penassen intrusive complexes, as well as several minor, sub-concordant, pink monzogranite bodies that are too small to map at 1:50 000 scale. At this locality sedimentary features are well preserved (see Figure 12A) and the overall strain appears low. Foliation-bedding orientation relationships and measured intersection lineations suggest that folds in this area likely plunge steeply. Locally, thin shear zones have steep lineation plunges. Finer-grained beds locally have well developed porphyroblasts of andalusite (Figure 12B) and some cordierite, both are commonly replaced by muscovite. These assemblages are consistent with high temperature-low pressure metamorphism. At this locality we interpret that the observed metamorphic assemblage results from proximity to the many large intrusions in the area. The Walkinshaw pegmatites are somewhat enigmatic. They occur in relatively low metamorphic grade rocks and there is no clear peraluminous “parent” granite nearby. A speculative explanation could be that small volume melts were locally produced by partially melting adjacent to contacts of nearby intrusions (PRCAS). There is local evidence for such melts, but only in very small volumes. - 57 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 12. Lappe domain metasedimentary rocks and the “Walkinshaw” peraluminous granitic pegmatites (Stop 9). (A) northward younging wacke bed based on scouring and normal grading. (B) Andalusite porphyroblast-rich bed. (C) and (D) Black-tourmaline-rich subconcordant muscovite pegmatite dike. Stop 10 (Planned) - Onion Lake pegmatites and folded Quetico metawacke 346817E 5398027N Turn east on Cliff Rd (no sign) and drive about 250m down the dirt logging road and park to one side. Be aware of potential for logging traffic or other road users. Wood ticks are common at this site in spring and early summer. Use caution while walking around on the uneven ground. This outcrop shows a clean exposure of a peraluminous granitic pegmatite typical of pegmatites we refer to as the Onion Lake pegmatites. At this stop, a biotite-muscovite-garnet bearing pegmatite-aplite dike (Figure 13A and 13B) strikes roughly northeast and appears to post-date a strong foliation affecting typical Quetico Subprovince metasedimentary rocks. The pegmatite displays prominent interlayering of pegmatitic and aplitic rock and unidirectional solidification textures (Figure 13A). This pegmatite is not far south of the contact of a peraluminous granite we refer to as the Voutilainen intrusion. Several large “whalebacks” of pegmatite are present in this area. Although not entirely clear at this locality, the Onion Lake pegmatites are boudinaged and commonly have internal fabrics defined by magmatic phases implying a syn-tectonic emplacement. Regionally, the northeastward strike of pegmatite contacts is parallel to sinistral shear zones which are interpreted to be antithetic structures related to the overall dextral shearing related to the Quetico deformation zone. Granites in this area occur near the southern margin of prominent deformation related to the QDZ. At least two distinct generations of peraluminous granitic pegmatites are present in the southern Quetico. A second generation of pegmatites, younger than the one at this locality crosscut at a high angle the foliation related to the Quetico deformation zone and therefore are post-tectonic. These later pegmatites are much less abundant, and we will not be able to examine the younger generation of pegmatites on this trip. Metasedimentary rocks at this locality are strongly deformed. Relatively steeply plunging, east-northeast striking z- to m-folding is revealed by prominent quartz-feldspar veins (Figure 13C). These quartzfeldspar veins are very common in the southern QS. - 58 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 13. Onion Lake peraluminous granitic pegmatite and folded Quetico metasedimentary rocks (Stop 10). (A) Pegmatite-aplite layering, unidirectional solidification textures, biotite-muscovite pegmatite and garnet aplite (B) Close-up of abundant garnet in aplitic phase, (C) boudinaged S-type granite dike sub-parallel to axial plane of “D2” folds in Quetico metawacke. Based on cross-cutting relationships they appear to predate most of the mapped intrusive suites. Note the narrow, boudinaged peraluminous granitic dike emplaced sub parallel to the axial plane of the folds appearing to crosscut the veining (Figure 13C). Outcrop scale folds of sedimentary layering are relatively uncommon in the southern Quetico making overall understanding of fold-geometries somewhat difficult. These folds are likely “D2” in the nomenclature of Williams (1991) and limbs of similarly oriented folds are elsewhere postdated by dextral shearing (D3). A prominent linear magnetic anomaly, caused by a relatively magnetic wacke unit, shows a clear regional z-folding pattern and this is likely the general geometry of “D2” folding in the southern QS. In this area, many traverses across-strike documented well preserved younging indicators that show multiple reversals over relatively short distances. These reversals likely represent parasitic folds in hinge zones of larger z-fold enveloping surfaces. Towards the north, in the Quetico deformation zone, fold orientations change and tend to be east-striking, upright or slightly inclined with gentle plunges. These folds have been interpreted as being part of the D3 event. Stop 11 (Optional) - Mylonite and late brittle-ductile deformation Quetico deformation zone 348059E 5401951N Pull vehicles over near the north end of long outcrop and park on the shoulder of Hwy 527. Keep the duration of stop relatively short. If longer stop required park on logging road just to the south. Wear reflective vests, use hazard lights. Traffic can be heavy. At this locality, highly deformed metasedimentary rocks and sheared and boudinaged muscovite pegmatites are present. Locally, north dipping brittle structures offset pegmatite dike contacts with a north over south sense of displacement (Figure 14A). Pale - 59 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 QS domain to the south and the Dog Lake injection complex to the north. Grey amphibole-biotite-magnetite quartz monzonite at this outcrop is crosscut by later pink, moderately magnetic biotite monzogranite (Figure 15A and 15B). The quartz monzonite is moderately foliated and syn-tectonic. The quartz monzonite and the biotite monzogranite are compositionally similar to bigger, typically less strained intrusions located farther south e. Conversely, these types of intrusions are highly strained to the north in the Dog Lake injection complex. Preliminary geochronology suggests that despite highly variable degrees of strain, PRCAS intrusions have comparable ages in the Lappe domain, southern Quetico Subprovince and in the Dog Lake injection complex. Weakly deformed intrusions like the Trout Lake and Barnum Lake intrusions are essentially contemporaneous with highly sheared syn-tectonic equivalents in the DLIC. Figure 14. Sheared Quetico metasedimentary rocks and peraluminous granitic pegmatites, Quetico deformation zone (Stop 11). (A) brittle minor off-set thrust fault post-dates dextral shearing related to the ductile phase of the QDZ. (B) Pale green, siliceous mylonite related to the Quetico deformation zone. green siliceous rocks exposed at the north end of the outcrop are likely a mylonitic band (Figure 14B) within the larger Quetico deformation zone. To the south, dark grey rocks are strongly deformed wacke. Stop 12 (Planned) - Intermediate and felsic phases of potassium-rich calc-alkalic intrusive suite and syn-tectonic schlieric S-type granite 348541E 5403237N Pull over on the shoulder of Hwy 527. Use hazard lights, wear reflective vests, use extreme caution when crossing the highway. This large outcrop illustrates contact relationships between intermediate and felsic phases of the potassium-rich calc-alkalic intrusive suite (PRCAS) and a schlieric biotite-rich peraluminous granite. This outcrop represents the transition from the southern At the north end of the outcrop, syn-tectonic, schlieric, low magnetic susceptibility peraluminous leucogranites are in contact with rocks of the PRCAS intrusions. C-S fabrics in the granite, and narrow sheared bands (Figure 15C and 15D) indicate north side up thrusting with a dextral strike-slip horizontal component during emplacement. East-northeast striking foliations locally bear shallowly plunging stretching and mineral lineations indicating strike slip motion. These lineations may have formed under brittleductile conditions after the main phase of ductile syngranite shearing. The relative age of the peraluminous intrusion and the intermediate PRCAS phase is not immediately clear at this exposure. Regionally, the intermediate phase of the PRCAS slightly predates the bulk of S-type granite intrusions, and pink biotite monzogranites are typically younger. In some areas, hybridization of magmas have been documented. Stop 13 (Planned) - Shear-hosted leucogranitic injections, Dog Lake injection complex 346928E/ 5407112N (this stop is not on Figure 1) From Highway 527, turn right onto Hiccup Road (a logging road). Park on the curve near the stack of logs. This road is not active at the time of writing, and traffic is expected to be minimal. This outcrop exposes paragneiss of the Dog Lake complex intruded and variably digested by large volumes of syntectonic leucogranitic injections (Figure - 60 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 15. Potassium rich calc-alkalic suite intrusions and schlieric peraluminous granite at Stop 12. (A) Grey, foliated, feldspar porphyritic, hornblende-biotite quartz monzonite cross but by dikes of pink biotite monzogranite; (B) Porphyritic texture in quartz monzonite. (C) and D) syn-emplacement C-S fabrics defined by biotitic schlieren in biotite-rich peraluminous granite indicate a north-side up thrust component of motion (C) and dextral strike slip component of motion (D) during peraluminous granite emplacement. 16A), producing raft and schollen textures that can be easily confused with diatexite, and may therefore lead to misleading interpretations. The outcrop displays variable degrees of assimilation and digestion of the sedimentary host rocks by granitic melts, locally producing biotite schlieren within the leucogranite. Leucogranite injections are heterogeneous, ranging from coarse-grained to porphyritic textures (Figure 16B). They occur as sheeted to irregular intrusions that clearly exploit east-west trending foliation planes and northwest-trending dextral shear bands within the paragneiss host rock. This strong structural control indicates syntectonic emplacement facilitated by regional dextral zones related to the Quetico deformation zone. Locally, antithetic sinistral shear bands are also observed. Wider and more homogeneous granite injections locally contain garnet and pegmatitic segregations (Figure 16C and 16D). The presence of these pegmatitic segregations indicate a high degree of melt fractionation, which is incompatible with in situ partial melting. This interpretation is supported by chondritenormalized REE patterns, which display a strongly fractionated geochemical signature and variably developed negative Eu anomalies (Figure 17), comparable to those observed in the S-type granite intrusions of the Quetico subprovince. These features indicate that granitic melts had already undergone significant plagioclase fractionation at depth forming cumulates before extraction of the residual melt along shear zones. Although some textures locally resemble those seen in migmatite, their interpretation as true anatectic melts is not supported by the mineralogy. - 61 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 16 Photographs of shear-hosted leucogranitic injections of the Dog Lake injection complex (Stop 13). (A) Dense network of leucogranitic injections emplaced along foliation planes and dextral shear bands. Note the presence of paragneiss rafts within wider injections (B) Close-up view of dextral shear bands. Note the presence of biotite-rich schlieren in some leucogranitic injections. (C) Schlieric garnet-bearing dyke of leucogranite. Note the presence of pegmatitic pods indicating fluids segregation. (D) Close up on garnets within the dyke of leucogranitic. Melanosome-looking parts are mostly only composed of biotite and lack typical peritectic mineral phases (e.g., garnet, cordierite, sillimanite) expected from insitu partial melting of metasedimentary protoliths. Instead, this outcrop is interpreted as a migration zone for granitic melts generated deeper in the crust, which were channeled upward along regional dextral shear zones. During ascent, these melts assimilated sedimentary host rocks, resulting in the development of characteristic schollen and schlieric textures observed at this outcrop. The large volume of ascending granitic melt likely induced high-temperature, low-pressure metamorphic conditions in the surrounding paragneiss, as evidenced by the development of sillimanite–cordierite assemblages. Locally, the paragneiss also experienced limited partial melting, interpreted to have been induced by thermal input associated with the emplacement of this large volume of granitic melts. Stop 14 (Planned) - Patchy metatexites, Dog Lake injection complex 346711E 5406946N (this stop is not on Figure 1) From the previous stop, return to Highway 527 and cross the highway onto Doodie Road. Park at the entrance of the road. The outcrop is located on the north side of the road. This outcrop exposes an interlayered biotite–quartz– feldspar and garnet–biotite–quartz–feldspar migmatitic paragneiss, interpreted as a patchy metatexites (Figure 18A). Peritectic garnet is present within leucosome patches, providing clear evidence for insitu partial melting of fertile pelitic layers. The estimated melt proportion is relatively low (approximately 5–10%). - 62 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 17 Chondrite normalized rare earth element (REE) patterns of leucogranitic injections from the Dog Lake injection Complex compared with Stype granite intrusions occurring in the Quetico Subprovince. Bedding is largely preserved, indicating low melt connectivity and limited melt extraction (Figure 18B). This migmatitic paragneiss is gently folded, with fold axes plunging eastward. The east–westtrending foliation dips steeply to the north. The paragneiss is cut by boudinaged dikes of white, coarse-grained leucogranite, similar to the leucogranitic injections observed at the previous stop. This outcrop highlights a significant volumetric contrast between the limited amount of melt generated in situ within the paragneiss and the much larger volume of granitic injections observed at the previous stop (Stop 13), despite that the two localities are separated by only ~300 m. This contrast is a strong indication that the granitic melts observed elsewhere in the Dog Lake injection complex were not produced in situ but instead represent migrated and fractionated melts generated at deeper crustal levels, which were subsequently channeled upward along regional shear zones. The spatially restricted partial melting and migmatitization observed at this outcrop therefore do not represent the source of the Stype granites but rather reflect a thermal response to advected heat associated with the emplacement of large volumes of granitic melt in nearby shear corridors, rather than a widespread regional anatexis. Stop 15 (Optional) - Syn-tectonic schlieric pink monzogranite, Dog Lake injection complex 347031E 5405626N (this stop is not on Figure 1) Continue driving south along Doodie Road for approximately 1.5 km. This road is not active at the time of writing; however, road shoulders may be narrow or unstable, so vehicles should be parked directly on the road where it is safe to do so. This final stop exposes a representative outcrop of syntectonic, magnetitebearing pink monzogranite (Figure 19). The intrusion displays a heterogeneous texture, ranging from porphyritic to locally pegmatitic, highlighting the high fluid content of the granitic melt during its emplacement. The granite contains biotite schlieren, resulting of the complete digestion of the paragneiss host rock. Locally, relict structures of the - 63 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 19 Photograph of syntectonic, schlieric biotite pink monzogranite (Stop 15). Note the presence of dextral and antithetic sinistral shear bands. shear corridors. Melt migration was facilitated and channelized by regional transpressional deformation acting in the Quetico Subprovince and its adjacent greenstone belts. REFERENCES Card, K.D., and Ciesielski, A. 1986. DNAG No. 1: subdivisions of the Superior Province of the Canadian Shield. Geoscience Canada, 13: 5–13. Figure 18 Representative photographs of garnet bearing, patchy metatexites from the Dog Lake Injection Complex (Stop 14). (A) In situ partial melting localized within fertile pelitic layers of the paragneiss. Note that primary bedding structures are largely preserved. (B) Close up view of photo A. The upper layer shows a clear lack of segregation between leucosome and melanosome indicating low melt connectivity. original sedimentary bedding can still be inferred within these schlierenrich domains. The intrusion is strongly foliated and affected by welldeveloped dextral shear bands and antithetic sinistral shear bands, along which pegmatitic pods are locally emplaced. This outcrop illustrates how regional dextral shear zones within the Dog Lake complex have acted as efficient pathways for multiple types of granitic melts, which derived from different sources in the lower crust. The pink monzogranite suite is interpreted as the final, most fractionated product of the potassiumrich calcalkaline intrusive suite. These relationships reinforce the interpretation that the Dog Lake complex represents a major migration zone, where granitic magmas produced at depth were focused, transported, and emplaced along Carter, M. W., 1992, Geology and mineral potential of the Tower syenite stock, Conmee Township, District of Thunder Bay, in Dressler, B. O., Baker, C. L., and Blackwell, B., eds., Summary of field work and other activities 1992: Ontario Geological Survey Miscellaneous Paper 160, p. 60–63. Corfu, F., 2000. Extraction of Pb with artificially too-old ages during stepwise dissolution experiments on Archean zircon. Lithos, 53, nos. 3–4, p. 279–291. 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. Kamo, S.L. 2013. Report on U-Pb geochronology (CA-IDTIMS and LA-ICPMS) of rocks from the Grenville and Superior provinces of Ontario; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 50p. Launay, G.A. and Metsaranta, R.T. 2023. Precambrian bedrock geology mapping in the Onion Lake and Sunshine areas, Quetico and Wawa Subprovinces, northwestern Ontario; in Summary of Field Work and Other Activities, 2023, Ontario Geological Survey, Open File Report 6405, p.11-1 to 11-12. - 64 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Launay, G.A. and Metsaranta, R.T. 2024. Mapping regional fractionation patterns in S-type peraluminous granite and pegmatite intrusions in the southern Quetico Subprovince; in Summary of Field Work and Other Activities, 2024, Ontario Geological Survey, Open File Report 6413, p.9-1 to 9-11. Lodge, R.W.D. 2014. Precambrian geology of Aldina Township; Ontario Geological Survey, Preliminary Map P.3776, scale 1:20 000. Lodge, R.W.D., 2016. Petrogenesis of intermediate volcanic assemblages from the Shebandowan Greenstone Belt, Superior Province: evidence for subduction during the Neoarchean. Precambrian Research, 272, p. 150–167. MacDonald, R.D. 1939. Gorham Township and vicinity, District of Thunder Bay, Ontario; Ontario Department of Mines, Map 48C, scale 1:63 360. Ministry of Natural Resources and Forestry 2023. Forest Resources Inventory leaf-on LiDAR; Ministry of Natural Resources and Forestry, Science and Research Branch, Forest Resource Information Unit, online data, April 10, 2022 update, https://geohub.lio. gov.on.ca/maps/lio::forest-resources-inventory-leafon-lidar/about. [accessed April 27, 2023] Metsaranta, R.T. 2015. Preliminary results from geological mapping of the Quetico Subprovince, the Shebandowan greenstone belt and Proterozoic rocks north of Thunder Bay; in Summary of Field Work and Other Activities, 2015, Ontario Geological Survey, Open File Report 6313, p.15-1 to 15-20. Metsaranta, R.T. 2022. Highlights of bedrock geology mapping in the Quetico Subprovince, north of Thunder Bay, northwestern Ontario; in Summary of Field Work and Other Activities, 2022, Ontario Geological Survey, Open File Report 6380, p.9-1 to 9-9. Metsaranta, R.T. and Walker, J.A. 2019. Precambrian geology of western McGregor Township and adjacent areas, northeast of Thunder Bay; in Summary of Field Work and Other Activities, 2019, Ontario Geological Survey, Open File Report 6360, p.11-1 to 11-10. Metsaranta, R.T. and Hamilton, M.A. 2020. A precise U/ Pb age for a north-trending mafic dike from the western flank of the Marathon swarm, East Bay area, northwestern Ontario; in Summary of Field Work and Other Activities, 2020, Ontario Geological Survey, Open File Report 6370, p.7-1 to 7-9. Metsaranta, R.T. and Kamo, S.L. 2021. A uranium–lead baddeleyite age for the Midcontinent Rift–related Lone Island Lake intrusion, northwestern Ontario; in Summary of Field Work and Other Activities, 2021, Ontario Geological Survey, Open File Report 6380, p.12-1 to 12-8. Ontario Geological Survey 2017. Ontario airborne geophysical surveys, magnetic data, grid data (ASCII and Geosoft® formats), magnetic supergrids; Ontario Geological Survey, Geophysical Data Set 1037— Revised. Pan, Y., Fleet, M.E., and Heaman, L. 1998. Thermo‑tectonic evolution of an Archean accretionary complex: U–Pb geochronological constraints on granulites from the Quetico Subprovince, Ontario, Canada. Precambrian Research, 92: 117-128. Percival, J.A. 1989. Late Archean Quetico accretionary complex, Superior Province, Canada. Geology, 17: 23–25. Percival, J.A., Sanborn‑Barrie, M., Skulski, T., Stott, G.M., Leclair, A.D., and Corkery, M.T. 2006. Tectonic evolution of the western Superior Province from NATMAP and Lithoprobe studies. Canadian Journal of Earth Sciences, 43: 1085–1115. 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. In: Tectonic styles in Canada: the Lithoprobe perspective. Geological Association of Canada, Special Paper 49, p. 321–378 Ratcliffe, L.M. 2016. Precambrian geology of Sackville Township, Shebandowan greenstone belt, Wawa– Abitibi terrane; Ontario Geological Survey, Preliminary Map P.3802, scale 1:20 000.Ratcliffe, L.M. 2017. Precambrian geology of Adrian Township, Shebandowan greenstone belt, Wawa–Abitibi terrane; Ontario Geological Survey, Preliminary Map P.3813, scale 1:20 000. Ratcliffe L.M. 2019. Precambrian geology of Marks Township, Shebandowan greenstone belt, Wawa– Abitibi terrane, northwestern Ontario; Ontario Geological Survey, Preliminary Map P.3830, scale 1:20 000. Rehm, A. G. 2025. “Tectonometamorphic Evolution, Fluid Production, and Evaluation of Gold Liberation in the Quetico Metasedimentary Belt, Canada.” Ph.D., Laurentian University Sudbury, Ontario. 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. Valli, F., Guillot, S., and Hattori, K.H. 2004. Source and tectono‑metamorphic evolution of mafic and pelitic metasedimentary rocks from the central Quetico metasedimentary belt, Archean Superior Province of Canada. Precambrian Research, 132: 53–72. Wang, S., Kuzmich, B., Hollings, P., Zhou, T. and Wang, F. 2020. Petrogenesis of the Dog Lake Granite Chain, Quetico Basin, Superior Province, Canada: Implications for Neoarchean crustal growth. - 65 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Precambrian Research, 346: 105828. 4, Part 1, p.485-541. Williams, H.R. 1991. Quetico Subprovince; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.383-403. 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 - 66 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Trip 3 - Geological assemblages, regional structural framework and tectonic evolution of the Neoarchean Shebandowan greenstone belt Dorothy Campbell, P.Geo and Justin Jonsson P.Geo Resident Geologist Program, Ontario Geological Survey, Ministry of Energy and Mines, Thunder Bay, Ontario, P7E 6S7, Canada Introduction This trip provides an overview of the geological assemblages, regional structural framework, and tectonic evolution of the Neoarchean Shebandowan Greenstone Belt (SGB) and their relationship to gold and base metal mineralization. The SGB is situated in the western Wawa Subprovince (Superior Province) and extends 150 km from the Ontario–Minnesota border in the west to northeast of Thunder Bay in the east (Figure 1). The SGB is locally in fault contact with the Quetico Subprovince to the north and bounded by the older (2750 Ma) Northern Light–Perching Gull Lakes batholith (tonalitic gneiss) and younger granitic intrusions to the south (Lodge 2016). The SGB is characterized by a complex history of early rifting, subduction-driven volcanism, tectonic accretion, and later transpressional deformation. The SGB comprises three main assemblages and two primary deformation events (Williams et al. 1991; Stott and Corfu 1991; Corfu and Stott 1998; Percival 2006; Lodge 2016; Reynolds et al. 2023; Dorval et al. 2026): BLF=Burchell Lake fault; USSZ=Upper Shebandowan Lake shear zone; SGFZ=Squeers Lake-Greenwater Lake fault zone; TLFZ=Tinto Lake fault zone; CCF=Crayfish Creek fault; LSSZ=Lower Shebandowan Lake shear zone; MLS=Moss Lake stock; BLS=Burchell Lake stock; HGC=Haines gabbroic complex; HS=Hermia stock; HLS=Hood Lake stock; GLS=Greenwater Lake stock; LGP=Little Greenwater Lake pluton; PCS=Pinecone stock; KS=Kekekuab stock; PS=Peewatai stock; SS=Shebandowan stock; TS=Tower Stock Figure 1. Regional Geology of the Shebandowan greenstone belt (modified from Kuster, Lesher and Houlé, 2022; modified from Sotiriou et al 2019; Lodge 2016; Osmani 1997a; Corfu and Stott, 1998). - 67 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Greenwater assemblage (2722-2719 Ma): volcanic suites characterized by thick sequences of tholeiitic mafic volcanic rocks, ultramafic flows (komatiites) and sills, iron formations, mafic intrusions, and minor FII- and FIII-type felsic volcanic rocks. Burchell assemblage (2719-2716 Ma): calcalkalic, dominantly intermediate volcanic rocks and lesser FI-type felsic volcanic rocks and with no known ultramafic sills or intrusions. This subdivision was first defined by Williams (1991) on the basis of younging directions but rejected by Corfu and Stott (1998) due to lack of chronological distinction and re-interpretation of structural architecture. Lodge (2016) interpreted more recent higher-precision geochronology as supporting a similar subdivision to that of Williams (1991). Kashabowie assemblage (2695 Ma): syn-D1, represents renewed activity on the SGB after a long hiatus. It is less voluminous and more evolved than Greenwater assemblage. Calc-alkaline to alkalic intermediate/felsic volcanic rocks, associated diorites, tonalites (e.g., Shebandowan Lake Pluton), tectonically interleaved with older 2720 Ma volcanic suites. This subdivision was first introduced by Corfu and Stott (1998) as part of their re-interpretation of older assemblage classifications. D1 Compressional Deformation event (2695 and 2690 Ma): associated with calc-alkaline magmatism and intra-arc deformation (thruststacking and interleaving). Shebandowan assemblage (2690-2680 Ma): syn-D2 Timiskaming-type assemblage, unconformably overlies older Greenwater assemblage, composed of calc-alkalic to alkalic volcanic rocks and associated coarse clastic Timiskaming-type sedimentary rocks, iron formation and late sanukitoid plutons. D2 Transpressional Deformation event (2685– 2680 Ma): marked the final accretionary phase of the Wawa subprovince evolution of the Superior Craton, termed the Shebandowanian phase of the Kenoran Orogeny (Stott and Corfu 1991). This late-stage dextral transpression and obliqueslip deformation represents the development of Timiskaming-type pull-apart basins and regional Timiskaming-aged structures. Auto Road assemblage (<2682 Ma): distinctly younger sedimentary assemblage in the SGB, dominated by conglomerate-sandstone units (with clasts of volcanic and granitoid origin). Corfu and Stott (1998) describe the assemblage as a small sedimentary basin, informally termed the “Auto Road assemblage”. The western limb of the SGB is often divided from the central and eastern portions of the belt by an informal north-south boundary roughly, at the town of Kashabowie (Figure 1). There are differences in the distribution of assemblages between the west and east sides: the Kashabowie assemblage is situated mostly along the western limb, the Shebandowan assemblage is situated on the central-eastern side, and the Auto Road assemblage is restricted to a small area on the eastern side. The Greenwater/Burchell assemblage(s) make up the large majority of the preserved supracrustal rocks, despite comprising just ~6 million years of the >40 million-year evolution of the SGB (Figure 1). The volcanism recorded by these assemblages appears to have been two-stage: an extensional plume-rift setting recorded by the Greenwater assemblage followed by a compressional subduction-arc setting recorded by the Burchell assemblage (Figures 2, 3; Lodge 2016). The deposition of chemically distinct Kashabowie assemblage volcanic rocks occurred after a ~21 million-year hiatus, recording a later compressional subduction-arc setting (Figure 3). These rocks are contemporaneous with the D1 structural event, which involved the interleaving and thrust-stacking of the Kashabowie and Greenwater units (Reynolds et al. 2023). Subsequently, the Shebandowan assemblage, represents the final stages of the Shebandowan accretionary event. These “Timiskaming-type,” deposits unconformably overlie the Greenwater assemblage. They are characterized by a mix of clastic sediments (conglomerate, sandstone), iron formation, and calc-alkalic to alkalic volcanic rocks (Figure 3), interpreted to have formed in transtensional, pullapart basins along the flanks of transpressional uplifts (Reynolds et al. 2023). Due to their tectonic setting, these rocks are strongly associated with structurally controlled, late-orogenic gold mineralization (Figure 4). - 68 - Based on the spatial distribution of southwest- �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 2. Schematic illustration of possible tectonic evolution of the Shebandowan greenstone belt in both plan view and crustal cross section (from Lodge 2016). Colors of units correspond to legends in Figure 1. Note sketch is not to scale. (A) Initial plume-dominated tectonic setting forming Greenwater assemblage. (B) Subduction-dominated tectonic setting forming Burchell Assemblage. A change in plate motion results in the initiation of subduction and formation of a calc-alkalic arc dominated by andesitic strata. Subduction of ridge results in high geothermal gradient and melting of slab to produce adakitic melts. Hybridization of mantle and slab derived melts results in magnesian andesites (Mg# > 50) from Lodge 2016. trending metavolcanic rocks on the western limb, Osmani (1997a) defined three distinct geological units that remain in use by current explorers (Figure 6): • Central Felsic Belt (CFB): a <5 km-wide core of the Burchell/Kashabowie assemblage. • Northern Mafic Belt (NMB) and Southern Mafic Belt (SMB): mafic metavolcanic rocks of the Greenwater assemblage, flanking the CFB to the north and south respectively. There are three past-producing mines in the SGB: the North Coldstream copper mine (1957-1967) and the Ardeen gold mine (1932-1936, 1942) in the western part of the belt, and the Shebandowan nickel-copperPGE-cobalt mine (1971-1998) in the eastern part of the belt (Figure 5). Tectonic associations provide spatial context for mineral prospectivity in the SGB (e.g. Lodge et al. 2015, Lodge 2016, Reynolds et al. 2023). Magmatic Ni-Cu-PGE mineralization occurs in the Greenwater assemblage mafic-ultramafic intrusive rocks, notably the sill-hosted deposit comprising the past-producing Shebandowan mine. The mine operated for most of 1971-1998, producing 9.29 Mt at 1.75% Ni, 0.88% Cu, 0.06% Co and 1.83 g/t PGEs. Clusters of magmatic sulfide occurrences also occur in the Haines gabbro (~7 km northwest of the mine) and in the Bateman Lake area (~40 km east of the mine). Although no economic deposits that are definitively of volcanogenic massive sulfide (VMS) affinity exist in the SGB, several prospects exist in spatial association with Greenwater/Burchell assemblage felsic metavolcanic rocks. The North Coldstream copper-gold-silver deposit, located 10 km southwest of Kashabowie on the western arm of the SGB, is a - 69 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 past-producing (2.48 Mt at 1.87% Cu, 0.28 g/t Au, and 5.53 g/t Ag from 1957-1967) atypical deposit variably interpreted to be intrusion-related (e.g. Farrow 1994) or volcanogenic (Reynolds et al. 2023); it is currently being explored by Gold X2 Mining Inc., who tentatively interpret the deposit as a sheared, remobilized VMS system. In more recent years, orogenic gold has become the main focus of mineral exploration in the SGB. Gold is primarily controlled by late tectonic D2 structural zones, in contrast to lithologically controlled magmatic and VMS mineralization associated with Greenwater/ Burchell assemblages. Gold mineralization is generally hosted within ductile-brittle shear zones, particularly near regional fault zones (Figure 4) or adjacent to “Timiskaming-type” unconformities (Figure 14). Gold is typically hosted by quartz-carbonate-pyrite veins and veinlet networks cross-cutting all lithologies. On this field trip, we will look at some specific examples these structural zones: • Moss Gold Deposit (Gold X2 Mining Inc.) and the 111 Zone (Bold Ventures Inc.): gold mineralization occurs near regional fault zones, within sheared diorites, felsic dykes/sills and mafic to intermediate metavolcanic rocks. • I-Zone (Delta Resources Limited): gold-bearing quartz ladder veins within a felsic dyke (brittle, extensional), intruding Timiskaming iron formation. • Eureka Zone (Delta Resources Limited): a key target for gold exploration at the unconformity Figure 3. Evolution of Greenwater/Burchell, Kashabowie, and Shebandowan assemblages (from Reynolds et al. 2023). Figure 4. Schematic section of western Shebandowan greenstone belt (from Reynolds et al. 2023). - 70 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 5. Geology map of the Shebandowan greenstone belt showing location of field trip stops. NCM: North Coldstream Mine, See Figure 1 for all other abbreviations. - 71 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 between the Greenwater and Shebandowan assemblages, marking a “Timiskaming-type” unconformity. Stops 1 to 2 - Moss Gold Deposit (Gold X2 Mining Inc.) Permission is required from company to access sites Gold X2 Mining Inc. (Gold X2) is exploring the Moss gold deposit, a high-tonnage low-grade deposit (Figure 6), located 100 km west of Thunder Bay on the western limb of the SGB. Gold X2 recently completed a Preliminary Economic Assessment, releasing an updated resource estimate as of January 16, 2026, (Dorval et al. 2026) for the deposit as follows: • Indicated: 2.125 Moz Au at 1.03 g/t with 3.160 Moz Ag at 1.53 g/t • Inferred: 3.910 Moz Au at 0.97 g/t Au with 6.273 Moz Ag at 1.55 g/t Engineering trade-off studies & design work is underway and a feasibility study is anticipated for Q3 2027 (Gold X2 Mining Inc., Corporate Presentation, April 12, 2026). The Moss deposit is structurally controlled and situated within intermediate to felsic metavolcanic rocks of the Central Felsic Belt (CFB; Figure 6). Primarily hosted by sheared diorite (Figure 7), the deposit developed during and after intense ductile deformation, with 2 distinct tectonic-hydrothermal events identified (Reynolds et al 2023; Dorval et al. 2026). Alteration occurs in different styles and intensities but is generally composed of albite, biotite, sericite, chlorite, carbonate, epidote and pyrite (typically 2-10% of the rock; locally up to 15%). Gold mineralization occurs in complex arrays of smallscale quartz-carbonate-pyrite veinlets, breccias, and stockworks with higher grades within more intense, narrow shear zones (Nwakanma 2024; Dorval et al. 2026). The sulfide assemblage is dominated by pyrite, with minor chalcopyrite, sphalerite, and molybdenite. Rare, high-grade tellurides are associated with the high-grade gold mineralization (Reynolds et al. 2023; Dorval et al. 2026). Stop 1. Moss Gold Deposit (Portal) N83 Z15 U 668730E 5379177N At this stop, highly sheared diorite and feldspar porphyry has been variably silicified, chloritized, hematized, sericitized and sulphidized (Figure 7). The outcrop is highly fractured and exhibits a network of narrow quartz-carbonate-pyrite veinlets. The now closed-off portal, developed in the mid-1980s by Tandem Resources Limited and Storimin Exploration Limited, lead to historical underground workings and gold zones at the 230-foot (70 m) level (Figure 8). Figure 6. Geology map shear hosted Moss Lake Deposit (in red) modified from Dorval et al. (2026). - 72 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 8. Underground plan of the 230-foot (70 m) level, showing gold-bearing zones of the Moss deposit (from Osmani 1997a; modified from an underground plan of Tandem Resources Limited - Storimin Exploration Limited, 1989). Figure 7. Gold-mineralized diorite at the Moss deposit that has been variably silicified, chloritized, hematized, sericitized and sulphidized. Stop 2. Discovery Outcrop The Moss property has a long history of exploration dating back to 1936, when Mining Corporation of Canada completed 5 trenches that exposed a zone of mineralization later known as the Main Zone (often referred to in historical records as the Snodgrass showing). Gold was initially identified in a mineralized zone hosted by sheared dacite and felspar porphyry near the northern contact with diorite. The zone measured approximately 25 feet (7.6 m) in width and 600 feet (180 m) in length. In 1945, Lobanor Gold Mines Limited followed up with 12 diamond drill holes which ultimately led to the development of the Moss deposit (Figure 9). Subsequently, more than 30 companies explored various smaller sections of the property that were consolidated in 2014-2016 by Wesdome Gold Mines Ltd. In May 2021, Gold X2’s predecessor (Goldshore Resources Inc.) acquired the Moss Gold property from Figure 9. Map showing location of initial trenches, drill holes and gold assay results by Lobanor Gold Mines (1945) (from Harris 1970). - 73 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Wesdome. In 2024-2025, the property was further expanded by: i) purchasing the “Coldstream claims” and acquiring Kesselrun Resources Ltd., whose Huronian project claims include the past-producing Ardeen mine, ii) staking the Hillcrest property (Crayfish Creek Fault extension) and claims covering the Squeers-Greenwater Fault Zone extension, to the north and south of the Moss deposit, respectively, and iii) optioning Sky Gold’s Star Lake property, based on OGS gold-in-till anomalies (Figure 10). These expanded land holdings are strategic and a testament to the importance of regional structures for gold exploration. Stop 3. 111 Au Zone Trench - Burchell Lake Au-Cu Property (Bold Ventures Inc.) N83 Z15 U 676840E 5380320N Permission is required from company to access site The Burchell Lake Au-Cu property, located 95 km west of Thunder Bay, is adjacent to Gold X2’s Moss property to the west. While the property hosts multiple Au and Au-Cu showings, this stop focuses on Bold’s newly discovered 111 Au Zone (Figure 11 and 12). Initial grab samples returned 59.9 g/t and 68 g/t Au (Figure 11). Sampling at the 111 Au Zone by the Regional Resident Geologist (2025) returned up to 61.2 g/t Au and >1.2% Cu. Early assay results from 2026 drilling at the 111 Au Zone, BL-26-001 returned 0.42 g/t Au over 19 m, including 1.1 g/t Au over 5.0 m, and 2.7 g/t Au over 1 m. At this stop, silicified mafic to intermediate metavolcanic rocks are crosscut by a northeasttrending anastomosing shear zone (Figures 11, 12). A 14 m-wide halo of anomalous gold (see red dotted outline on Figure 12) has been outlined, flanked with zinc and copper mineralization. Gold mineralization is associated with disseminated pyrite and stringers of chalcopyrite, hosted in strongly silica‑ and sericite‑altered metavolcanic rocks. Locally, the rock is characterized by intense shearing and alteration obscuring the protolith, potentially a sheared and highly silicified metavolcanic rock or diorite. A narrow, relatively undeformed felspar porphyry occurs adjacent to the shear zone (Figure 12). Osmani (1997b) mapped this location as a felsic metavolcanic-dominated portion of the Southern Mafic Belt, though there was no outcrop exposure at the 111 Au Zone at the time of his mapping. Corfu and Stott (1998) reported a U–Pb zircon age of 2721 ± 1 Ma from a felsic metavolcanic flow less than 2 km to the northeast, interpreted to represent its eruption age. The mafic metavolcanic Figure 10. Map showing Gold X2’s 2025-2026 land acquisitions: Kesselrun’s Huronian project with the past-producing Ardeen Mine (red oval), Hillcrest and Squeers-Greenwater projects (yellow ovals), and Sky Gold’s Star Lake property (blue) with a cluster of gold in-till anomalies (orange and yellow dots). - 74 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 11. Map of land position, major showings, and 111 Au Zone trench highlighted with red oval (from Bold Ventures Inc., news release, October 20, 2025). Figure 12. Geology map showing the 111 Au Zone with channel sample results for gold, copper and zinc (from Bold Ventures Inc, news release, October 20, 2025). - 75 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 rocks here may either be tectonically interleaved with or conformable with the felsic metavolcanic rocks mapped by Osmani (1997b). Stop 4. Pillowed vesicular basalt at Swamp River N83 Z15 U 714018E 5390896N This outcrop is an example of typical Greenwater assemblage tholeiitic mafic metavolcanic rocks. The outcrop is glacially polished with well-defined striations that trend 25°. Glassy pillow selvages and abundant vesicles are well preserved at this location. Pillows are deformed (~10:1 aspect ratio) in the same orientation as foliation, striking 100° and dipping steeply south. Original mineralogy is replaced by a typical greenschist facies assemblage of chlorite, hornblende, sericite, saussurite, and carbonate (Morin, 1973). Pillow selvages appear to have been loci for fluid movement, as evidence by localization of pyrite and carbonate to the selvages. Morin (1973) mapped these pillows as younging to the north-northeast – can you see this? Stop 5. Timiskaming-type conglomerate N83 Z15 U 715392E 5387505N From Aubet and Campbell (2012). At this location two facies of the epiclastic suite of Timiskaming-type rocks are exposed. The dominant rock type is poorly sorted, highly foliated conglomerate (Figure 13). Note the heterolithic nature of the fragments, including minor Keewatin-type red jasper fragments. This particular outcrop is highly deformed with the clasts being stretched, forming a welldeveloped lineation plunging steeply to the southeast. Note the abundant iron carbonate alteration within the sandy matrix. In fault contact with the conglomerate to the west are mudstone and siltstone. Here we have near vertical mineral lineations normal to rolls on the bedding planes. This unit is finely bedded with grading, although present, obscured by the deformation. Stop 6. Autoclastite N83 Z15 U 715698E 5385810N This location is an example of ultramafic metavolcanic rocks of the Greenwater assemblage, featuring a flow-top breccia with a mixture of transported sub-angular, blocky clasts exhibiting some nice examples of random spinifex-textures and Figure 13. Timiskaming Conglomerate variolitic textures. The clasts range in size from 0.5 cm to 20 cm in diameter. Although this particular outcrop was not mapped by Rogers (1995), Rogers and Berger (1995) reported other nearby ultramafic metavolcanic units to be generally narrow (<50 m thick) and discontinuous (<1 km along-strike). Both olivine and pyroxene spinifex have been reported in the eastern SGB (e.g. Hinz 2018). Stops 7 to 10. Delta-1 Au Property (Delta Resources Limited) Permission required from company to access sites The Delta-1 Gold property (formerly Shabaqua Gold Project) is located near Shabaqua, 50 km west of Thunder Bay. The area has a long history of exploration dating back to 1930s, where numerous companies and prospectors carried out prospecting, geological, geochemical, and geophysical surveys, trenching, sampling and diamond drilling programs. While the Eureka deposit is the company’s flagship, Delta Resources significantly expanded the Delta-1 property in 2024 by acquiring more than a dozen properties from numerous companies and prospectors. The Delta-1 property now has multiple gold prospects and occurrences covering a 35-km strike extent of several regional-scale structural zones, near or at the unconformity between Shebandowan (Temiskaming- - 76 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 type) metasedimentary rocks and Greenwater metavolcanic rocks (see black dotted lines in Figure 14). (Portofino Resources Inc., news release, November 17, 2020). Stop 7. I-Zone gold-bearing quartz ladder veins N83 Z15 U 714705E 5382490N Modified from Aubet and Campbell (2012) The I-Zone (and associated gold showings) is an exploration target situated proximal to the Crayfish Creek Fault (Figure 14), a major regional structure currently presenting as brittle but likely with a protracted brittle-ductile history. The I-Zone gold occurrence consists of felsic dikes intruding Timiskaming oxide facies iron formation intercalated with argillite. The felsic dikes are host to gold-bearing quartz-tension/ ladder veins with 3%-5% pyrite and localized visible gold (Figure 15). Fractures opened up in the dike due to the ductility contrast of the enclosing ironrich argillites and the felsic dike. Later hydrothermal fluids, likely carrying gold reacted with the iron oxides resulting in the formation of pyrite and precipitation of native gold. Historical findings at the site include Landore Resources’ 1995 drill program, which intersected 4.32 g/t Au over 41 m, 4.53 g/t Au over 14.4 m, and 4.36 g/t Au over 20.4 m. Additionally, a 2008 mini-bulk sample conducted by Mengold Resources yielded an average grade of 9.9 g/t Au. Portofino Resources Inc. reported 2020 sampling at the I-Zone returned up to 45.9 g/t Au with 6 of 14 samples returning more that 5 g/t Au Figure 15. Simplified geology at the I-Zone (modified from Aubut et al., 1990). Stop 8 Eureka Zone (2024 Delta-1 Eureka Trench above drill hole D1-23-60) N83 Z16 U 290200E 5385348N In 2017, Doug Parker and Barbara D’Silva generated renewed interest in gold exploration in the Shabaqua area, on the eastern limb of the SGB. The ParkerD’Silva team followed up on historical data and OGS gold-in-till anomalies with prospecting, mechanical stripping, and rock sampling, which successfully led to the discovery of the Eureka Gold Zone (Figure 16). In 2019, Mr. Parker optioned the property to Delta Resources Inc. (Delta). Since then, Delta has advanced the project with 140 diamond drill holes (totaling more 40 000 m), confirming a mineralized zone for more than 2.5 km along strike and to a depth of 400 m (Figure 17). The Eureka Zone is situated adjacent to the unconformity between Shebandowan (<2690 Ma) and Greenwater (2720 Ma) assemblages. The Figure 14. Geology map showing Delta 1 Gold property (black outline), regional structural zones (black dashed lines) and gold showings (red stars) situated at or near the uniformity between Greenwater-Shebandowan assemblages (from Delta Resources Inc. website, Projects; Delta-1; Regional and Property Geology, April 2026). - 77 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 ankerite-pyrite veinlets. The quartz-ankerite-pyrite gold veinlets crosscut all lithologies and are hosted within a 300-400 m wide corridor of ankerite-silicasericite altered rocks. The Greenwater assemblage host rocks at this stop are comprised of mafic metavolcanic and ultramafic flows, weathered to a dark rusty brown with rock textures nearly or completely obliterated (Figure 18). Stop 9. Bylund Trench N83 Z15 290490E 5385211N Figure 16. Doug Parker’s 2017-18 prospecting and mechanical trenching programs generated renewed interest in gold exploration in the Shabaqua area. unconformity between the Greenwater (ultramafic and mafic to intermediate metavolcanic rocks) and Shebandowan (Temiskaming-type metavolcanic and metasedimentary rocks) assemblages has a close spatial association with gold occurrences, widely known as prospective for gold exploration (Figure 14). At this stop, Delta’s 2024 trenching program exposed an 11 m surface section of the Eureka Gold Zone, directly above drill hole D1-23-60. This drill hole returned an intersection of 1.79 g/t Au over 128.5 m (including 2.16 g/t Au over 97.5 m), while channel sampling from the surface trench returned an average grade of 1.23 g/t Au over 11 m (Delta Resources Inc., news releases, September 12, 2023, and September 25, 2024). Gold mineralization at the Eureka Zone is hosted by a stockwork of 1 mm to 10 cm wide quartz- At this stop, stockworks of gold-bearing quartzankerite-pyrite veinlets are situated within a broader 300-400 m carbonate-sericite-silica-altered halo that hosts anomalous/low-grade gold mineralization. The mineralized trend strikes ~110° and dips approximately 50-55° north. Mineralization is hosted within Greenwater assemblage rocks – most commonly, a feldspar-phyric tholeiitic basalt. That unit is not seen at this trench; what we see here a silica-rich rock that has historically been interpreted as chert but is being re-evaluated by the company as at least partially comprising highly silicified metavolcanic rocks. On the northeastern end of the trench, silicified komatiite or komatiitic basalt is present, displaying beautiful spinifex texture. Intense pyrite-ankerite alteration is widespread, but gold grades from this trench are relatively low. Three generations of quartz-carbonate veins are visible at this trench: i) NE-trending, steeply NW- Figure 17. Longitudinal section of the Eureka Zone - 78 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 18. Eureka Zone stockwork of 1 mm to 10 cm wide gold-bearing quartz-ankerite-pyrite veinlets hosted by Greenwater assemblage rocks are weathered to dark rusty brown with rock textures nearly or completely obliterated. dipping, ii) NW-trending, steeply dipping, and iii) NE-trending, shallowly dipping. Vein sets 1 and 2 are conjugate and are post-dated by vein set 3. Vein set 1 is the main gold-bearing set. Stop 10. Finmark Metasedimentary Rocks N83 Z16 U 293525E 5383950N From Puumala and Cundari (2023) At this stop we will have an opportunity to view a remarkably well-preserved roadside exposure of Shebandowan assemblage clastic metasedimentary rocks. The following description of these rocks is provided by Carter (1990). The rocks are mainly thinly bedded, the beds ranging in thickness from 5 cm to 12 cm. Primary sedimentary structures comprising load casts and flame structures, small-scale ripple structures, and cross bedding, are well developed in these rocks in the road exposures along Highway 11- 17 about 2.5 km west of the eastern boundary of Horne Township, and in the outcrops immediately southeast of these. Parker (1980) indicates that the “Finmark metasedimentary belt” consists of sandstone-siltstonemudstone sequences that alternate with thick units of cross-stratified sandstone. These sequences display many of the primary sedimentary structures that are characteristic of tidal flat (e.g., rhythmic layering, lenticular, wavy and flaser bedding) and tidal channel (e.g., herringbone cross stratification, large scale crossstratification) depositional environments respectively. Petrology of these rocks indicates that the primary sediment source was a felsic to intermediate volcanic terrain (Parker 1980). Characteristics of the clasts and rock fragments are consistent with a proximal Shebandowan assemblage sediment source. - 79 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Stop 11. Auto Road Assemblage (Optional) mafic clasts to nearly undeformed. N83 Z16 U 313838E 5377726N The Auto Road assemblage comprises a small sedimentary basin in south-central Ware township. It was first provisionally subdivided by Corfu and Stott (1998) on the basis of a U-Pb in youngest detrital zircon age of 2682±3 Ma – this is 9 m.y. younger (and outside of error provisions) than the youngest detrital minerals (zircon and titanite) in the Shebandowan assemblage. The assemblage is affected by D2 deformation and therefore provides a lower constraint on both sedimentation and regional transpression in the SGB. Corfu and Stott (1998) comment: The map pattern suggests that this conglomerate-sandstone unit is interbedded with Greenwater assemblage basaltic units (Brown, 1995), yet the polymictic conglomerate includes feldspar-hornblende-phyric volcanic clasts typically found within the Shebandowan assemblage. Also common are coarse granitoid cobbles as well as clasts of various volcanic lithologies. The results for sandstone sample Au presented below demonstrate that this is indeed one of the youngest supracrustal units of the Shebandowan greenstone belt as well as of the neighboring Quetico Subprovince, justifying its separate designation. Acknowledgements We would like to thank Gold X2 Mining Inc., Bold Ventures Inc., and Delta Resources Limited for permission to access parts of their properties and for their generous time, knowledge and support in preparing for this field trip. REFERENCES Aubet, A. and Campbell, D. 2012. Field trip 4 - Shebandowan Mine Area In; Hollings, P., MacTavish, A. and Addison, W. (Eds.), Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 2 - Field trip guidebook, v.58, part 2, 2-26. Aubut, A., Lavigne Jr., M.J., Scott, J. And Kita, J. 1990. Metallogeny, Stratigraphy and Structure of the Shebandowan Greenstone Belt; Field Trip 3 Guidebook, Mineral Deposits of Central Canada, CIM Thunder Bay Branch. Campbell, D.A. and Rainsford, D.R.B. 2020. Nickelcopper-cobalt-PGE potential in the Shebandowan greenstone belt; in Ontario Geological Survey, Resident Geologist Program, Recommendations for Exploration 2019–2020, p.69-74 At this location, felsic intrusive, chert, and felsic to mafic intrusive clasts of up to 40 cm in size are deformed (up to ~5:1 aspect ratio) by D2 transpression. Mafic clasts are deformed to a roughly uniform degree, while felsic clasts vary from similarly deformed as Carter, M.W. 1990. Geology of Goldie and Horne townships; Ontario Geological Survey, Open File Report 5720, 189p. Corfu, F. and Stott, G.M. 1998. Shebandowan greenstone belt, western SuperiorProvince: U–Pb ages, tectonic implications, and correlations. GSA Bulletin 110,1467–1484. Dorval, A., Lussier, D., Michaud, C., Taschereau, C., Vanier-Larrivée, N., Shankie, S. 2026. Preliminary Economic Assessment NI 43-101 Technical Report, Moss Gold Project, Ontario Canada, prepared for Gold X Mining Inc. by G. Mining Services Inc. Farrow, C.E.G. 1994. Base metal mineralization, Shebandowan greenstone belt, District of Thunder Bay in Summary of Field Work and Other Activities 1994, Ontario Geological Survey, Miscellaneous Paper 163, p. 22-97 to 22-104 Harris, F.R. 1970. Geology of the Moss Lake area, Ontario Geological Survey, Geological R085, 89p. Hinz, S.L.K. 2018. Geochemistry and petrography of the ultramafic metavolcanic rocks in the eastern portion of the Shebandowan greenstone belt, northwestern Ontario; Lakehead University, unpublished MSc thesis, 157p. Figure 19. Polymictic conglomerate of the Auto Road assemblage. Inco Limited Ontario Division 2001. Shebandowan Mine closure plan Part I of II: unpublished report; Ministry of Energy and Mines, Thunder Bay Mining Division; - 80 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Thunder Bay District, 84p. Survey, Special Vol. 4, Part 1, pp.145-238. Kuster, K., Lesher C.M., and Houlé, M.G. 2022. Geology and geochemistry of mafic and ultramafic bodies in the Shebandowan mine area, Wawa-Abitibi terrane: implications for Ni-Cu-(PGE) and Cr-(PGE) mineralization, Ontario and Quebec, Geological Survey of Canada Scientific Presentation 130, 25p. Lodge, R.W.D., Gibson, H.L., Stott, G.M., Franklin, J.M. and Hudak, G. 2016. 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, vol. 52, p. 196-214. Lodge, R.W.D. 2016. Petrogenesis of intermediate volcanic assemblages from the Shebandowan greenstone belt, Superior Province: Evidence for subduction during the Neoarchean: Precambrian Research, v.272, p.150–167. Morin, J.A. 1973. Geology of the Lower Shebandowan Lake area, District of Thunder Bay. Ontario Geological Survey, Report 110, 45p. Nwakanma, M.U. 2024. Characterization of alteration and mineralization of the Moss gold deposit, Shebandowan greenstone belt, Northwestern Ontario, Lakehead University, Department of Geology, Masters Thesis, 173p. Osmani, I.A., 1997a. Geology and mineral potential: Greenwater Lake area, west-central Shebandowan greenstone belt; Ontario Geological Survey, Report 296, 135p. Osmani, I.A. 1997b. Precambrian Geology, BurchellGreenwater Lakes area, west half; Ontario Geological Survey, Map 2622, 1: 20 000. Parker J. R. 1980: The Structure and Environment of Deposition of the Finmark metasediments, Thunder Bay, Ontario. Unpublished Hon.B.Sc. Thesis, Lakehead University, Thunder Bay, Ontario, 90 p. Percival, J.A., Sanborn-Barrie, Skulski, T., M., Stott, G.M., Helmstaedt, H., and White, D.J. 2006. Tectonic evolution of the western Superior Province from NATMAP and Lithoprobe studies. Geological Survey of Canada, NRC Research Press Web site at http:// cjes.nrc.ca.on 4 September 2006. Puumala, M. and Cundari, R. 2023. Geological highlights of the Thunder Bay area, Thunder Bay South Resident Geologist’s Office, unpublished field trip guide, 23p. Reynolds, N., Field, M., Fung, N., Peruse, C., Raponi, R., Ugarte, E., Gupta, N. 2023. NI 43-101 Technical report mineral resource estimates for the Moss Gold and East Coldstream deposit, Ontario, Canada, prepared for: Goldshore Resources Inc., 285p. Rogers, M.C. 1995. Precambrian geology, Duckworth township; Ontario Geological Survey, Map 2621, 1:20 000. Rogers, M.C. and Berger, B.R. 1995. Precambrian geology, Adrian, Marks, Sackville, Aldina and Duckworth townships. Ontario Geological Survey, Geological Report 295, 66p. 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. Sotiriou, P., Polat, A., Frei, R. 2019. Petrogenesis and geodynamic setting of the Neoarchean Haines Gabbroic Complex and Shebandowan greenstone belt, southwestern Superior Province, Ontario, Canada: Lithos, v.324-325, p.1–19. Stott, G.M. and Corfu, F.1991.Uchi subprovince. In Geology of Ontario. Edited by P.C. Thurston, H.R. Williams, R.H. Sutcliffe, and G.M. Stott. Ontario Geological - 81 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Trip 5 - Archean Geology and Metallogeny of the Rainy Lake Wrench Zone K. Howard Poulsen Geological Consultant USA. It includes approximately 2600 km of rocky shoreline plus more than 1600 islands and covers an area of approximately 930 square kilometers. Rainy Lake is fed from the east by the Seine River waterway and is drained westward by the Rainy River which leads to the even larger Lake of the Woods and the Winnipeg River system (Fig. 1). For centuries it has been part of the historic water link between the Atlantic and Arctic watersheds: it was known as Tekamaniwen to the indigenous inhabitants of the region and as Lac a la Pluie to the French voyageurs and fur traders. Rainy Lake and Lake of the Woods are remnants of the vast glacial Lake Agassiz which formed by melting of the Wisconsin continental ice sheet approximately 13,000 years ago. The predominately Archean bedrock in the Rainy Lake region (Figs. 1, 2) is now exposed in arched, glacially-sculpted outcrops within areas of generally thin and discontinuous surficial cover overgrown by boreal forest. A little learning is a dangerous thing; Drink deep, or taste not the Pierian spring: There shallow draughts intoxicate the brain, And drinking largely sobers us again. Fired at first sight with what the Muse imparts, In fearless youth we tempt the heights of Arts, While from the bounded level of our mind Short views we take, nor see the lengths behind; But more advanced, behold with strange surprise New distant scenes of endless science rise! So pleased at first the towering Alps we try, Mount o’er the vales, and seem to tread the sky, The eternal snows appear already past, And the first clouds and mountains seem the last; But, those attained, we tremble to survey The growing labors of the lengthened way, The increasing prospects tire our wandering eyes, Hills peep o’er hills, and Alps on Alps arise! Alexander Pope, 1711 Foreword Rainy Lake is a body of fresh water which straddles the border between Ontario, Canada and Minnesota, My first visit to Rainy Lake was in summer 1966 when I helped a geophysical operator evaluate the longwire electromagnetic survey method for our employer Dr. Ray Oja, a geological consultant working out of Thunder Bay. The equipment test was focused on the Grassy Portage Bay property of Noranda Mines Ltd. which included copper mineralization on their Halkirk Figure 1: Southwestern Superior Province with locations of selected mineral deposits. The area of Figure 2 is outlined by the dashed rectangle. - 82 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 2: Rainy Lake Geology – Watten (Northrock) prospect: C.J. Hodgson who later became one of my thesis supervisors at Queen’s had completed his MSc thesis on this deposit in 1959. I returned to the property in the early 1970’s with Dr. Mel Bartley who was then consulting for Northrock Mines and I helped him log a section of diamond drill core from the deposit which is located on the south flank of a feature known as the Rice Bay Dome (Fig. 2). Mel, who was a well-regarded geologist and one of the founders of Lakehead University, also consulted around that time for George Armstrong of Fort Frances. George was a successful highway construction contractor who, along with Mike Hupchuk, was also an avid part-time prospector. They had discovered Zn mineralization in 1971 near Pocket Pond east of Rice Bay. Mel and I made a site visit to Pocket Pond in fall 1972 and I prepared a report on the geophysical data for the property (Poulsen, 1973). At that time, I noted the existence of abundant outcrops along the drill roads near Armstrong’s trenches which were extremely large for the time - they had been excavated by his road construction crew! Jim Franklin, for whom I had been a research assistant at Lakehead University, left in July 1975 to join the GSC in Ottawa while I became a full-time technician in the geology department. I also began an independent look at roadside outcrops around Thunder Bay with a view toward identifying a possible thesis topic with Dick Ojakangas, the well-regarded sedimentologist at Duluth who was also famous for his Finn jokes. Around the same time, however, I told Jim about the interesting geology and the massive sulfide mineralization at Rainy Lake and we decided to investigate further. Armed with Andrew C. Lawson’s classic GSC Memoir 40 as well as Fred Harris’ more recent Ontario geology maps as guides, we visited several outcrops on Rainy Lake in summer 1976. In particular we revisited the Pocket Pond property where we confirmed my original observation that, based on pillow shapes, the strata appeared to be overturned. We also visited Lawson’s classic outcrops at Bear’s Passage using a Zodiak boat only to realize when we got there that they now are located at a boat launch site which is easily accessible by road! We nonetheless had also found the key outcrops and agreed that the staurolite-bearing metasedimentary rocks are also overturned. We later met John Wood of the Ontario Geological Survey who was mapping to the east at Mine Centre and he showed us outcrops of Lawson’s Seine conglomerate and of gold-bearing quartz veins near Bad Vermilion Lake. Once Lakehead University gained approval for its own M.Sc. program in geology, I applied to become the first (part-time) graduate geology student at Lakehead to study the problem of the apparently - 83 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 overturned stratigraphy around the Rice Bay Dome and Bear’s Passage. The thesis work began in 1978 under the supervision of Manfred Kehlenbeck and the newly hired structural geology professor, Graham Borradaile. The resulting structural analysis showed that, contrary to the long-standing interpretation of the Rice Bay Dome as originally proposed by Lawson, the evidence was clearly in favour of downward facing folds and extensive stratigraphic inversion but it was not easy to convince others of this: I had, somewhat unwittingly, stumbled into a larger problem that had already dominated the discussion of the Archean rocks northwest of Lake Superior for many decades. The Institute on Lake Superior Geology was initiated in 1955 as an annual meeting, most times with companion field trips, to discuss developments in geological understanding of both the U.S. and Canadian side of Lake Superior. Early meetings emphasized iron ore deposits which were of the greatest economic importance at that time but evolved into an exploration of a much more eclectic range of topics. Mel Bartley and Ed Pye of Port Arthur were among the early participants and Samuel S. Goldich, a pioneer of geochronology, was a founder and frequent contributor. I attended my first ILSG meeting at Madison, Wisconsin in 1973. Among the speakers were Paul Sims and Klaus Schultz of the U.S. Geological Survey and Don Davidson and John Green of the University of Minnesota at Duluth. A memorable and perhaps prophetic moment came during a presentation on Proterozoic stratigraphy of the Lake Superior area by the mild-mannered John Green who suggested that the Puckwunge Formation should be excluded from the Keweenawan Group. The proposal brought loud and angry condemnation by the short, red-faced Sam Goldich even before the talk was completed. As it turns out, although Goldich was a painfully shy and quiet individual in social situations, he was equally fierce and combative in professional settings. This proved to be the case again in 1976 at the ILSG meeting at St. Paul Minnesota. I attended Goldich’s excellent field trip to the Archean gneisses of the Minnesota River Valley, including a visit to a small outcrop in a swamp where he believed he had sampled and analysed the oldest rock on Earth as reported by that time. When one of his former graduate students questioned the statistical validity of Goldich’s data regression, the offender was told in no uncertain terms that, if he didn’t like the method, he could just leave the field trip immediately! At that time Goldich was a member in high standing of an international group of geologists with a strong interest in Precambrian geochemistry and geochronology. The Canadian leader within this group was Alan M. Goodwin of the University of Toronto who had conceived of and organized the multidisciplinary Superior Geotraverse Project which ran from 1970 to 1978. Near the project’s end Goodwin organized the Archean Geochemistry Conference in summer 1978 to highlight the significant results. This meeting involved the “who’s who” of Precambrian geochemistry at the time and, after a series of conference presentations at the Quetico Centre, the group headed west to Fort Frances-International Falls. Sam Goldich and Zell Peterman led a one-day field trip on route to illustrate aspects of the geology at Rainy Lake. Peterman had completed his MS thesis with Goldich in 1959 on the metasedimentary rocks of the Rice Bay Dome and now was a geochemist with the US Geological Survey in Denver. With prior arrangement by Jim Franklin who was a formal participant, I was able to tag along unofficially and silently as a beginning graduate student. The emphasis at each stop was placed on the chemical composition of rock units as recorded on hand-written file cards which Goldich drew from a deck at each outcrop. At one exposure there was considerable debate about whether a xenolith-rich lamprophyre dike might actually be a new locality of the Seine conglomerate but the important localities at Bear Passage and Pocket Pond were not part of the field trip. The classical stratigraphic interpretation of the eminent geologist Andrew C. Lawson was adhered to and I was not in any position to offer an objection. I made my first formal presentation on “Polyphase Deformation of Archean Rocks at Rainy Lake, Ontario” on May 10th of the following year at the Institute of Lake Superior Geology meeting at Duluth where I made the case for overturned strata around the Rice Bay Dome, contrary to Lawson’s original interpretations. Sam Goldich, who often referred reverently to “Professor Lawson”, was upset by this, so much so that he was unable to speak to me about it in person: he sent a delegation of Zell Peterman and Paul Sims instead to voice his displeasure. Both were apologetic and conciliatory and asked if it would be possible for me to arrange a field trip to visit the outcrops in question and this was set for later in the summer. In addition to Goldich, Sims and Peterman, the field trip participants included Dick Ojakangas - 84 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 and David Southwick from Minnesota as well as the Ontario geologists John Wood, Charlie Blackburn and Dick Beard. The event did not start well with Goldich clearly muttering something to the effect of “young punks don’t know anything” but the situation improved somewhat with successive stops. The second to last was at Lawson’s famous Bear’s Passage exposure: the group hadn’t fully had a chance to look closely at the outcrop before Dick Ojakangas, with a pointing of his thumb, indicated the southwestward direction of younging of the northwestward-dipping graded beds. This prompted Goldich to become agitated and declare that he didn’t believe graded bedding was a reliable criterion: Ojakangas, who had completed his PhD on the Cretaceous turbidites of the Great Valley Sequence in California, responded calmly that he had measured at least a thousand similar beds there without conflict. The final stop was at the exposure of pillow basalt near Pocket Pond where the field relationships proved to be even more convincing. Paul Sims, a nononsense individual whose standard dinner included a martini, a rare steak and a salad, followed by a footlong cigar, had remained quiet throughout the day but, when confronted with the outcrop, he turned to Goldich and said “there’s no question about this Sam, the section is overturned”. This prompted Goldich to smile, walk over and shake my hand, saying “well young man, you showed me something important today that I didn’t know before – let’s go back to Fort Frances and drink some of that “Canadian” beer”. He was always cordial to me from that point onward and made a point of connecting again in the field the following season. That one-day field trip was essential in demonstrating the credibility of the field observations and a second presentation on overturned Archean successions at the 1980 ILSG meeting at Eau Claire met with little resistance. A companion journal paper which previously had been rejected by the editor was ultimately accepted for publication with minor revisions by the Canadian Journal of Earth Sciences on June 17, 1980. The involvement of representatives from the Ontario and Minnesota geological surveys also proved to be important. Charlie Blackburn and John Wood later asked me to incorporate many of the field stops into one leg of a multi-day OGS-led excursion on Western Wabigoon Geology for the May 1982 GAC Meeting at Winnipeg. Dick Beard also lobbied hard for the funding of my subsequent work for Sandy Colvine of the Mineral Deposits Section of the OGS on the mineral deposits of the Mine Centre-Fort Frances area. Much of the field work for the OGS had been completed by 1981 and formed the basis of a third ILSG presentation at International Falls in 1982. Dave Southwick of the Minnesota survey was the organizer of the meeting and late in 1981 asked me if I would lead a related field trip focussed on the mineral deposits of the area. I agreed and we set a limit of 25 participants but this was a period of renewed in interest in gold exploration so registration quickly filled up. Dave contacted me again in the New Year and asked if we could double the limit to 50 participants and I reluctantly agreed but that limit was also reached in a short time so Dave developed a waiting list which grew to more than 20 requests. He contacted me once more to ask whether I would accommodate additional participants if he could find a Greyhound bus and provide shuttle vans and drivers to speed up the logistics at some field stops. I once again agreed and an exhausting, but gratifying, one-day, 10-stop field trip was delivered to 77 participants on May 5, 1982. Two weeks later, John Wood and I also led a field trip with a structural-stratigraphic focus as part of the larger excursion organized by Charlie Blackburn for the Winnipeg meeting. What follows is an attempt to illustrate the historical development of ideas about Rainy Lake geology using outcrops in the Mine Centre – Fort Frances corridor (Fig. 2). It includes a concise historical overview and a summary account of the highlights of the regional geology followed by updated descriptions of representative field localities. The motivation for doing this involves three main considerations. The first is purely practical and involves the precision and accuracy of outcrop locations. After the passage of more than forty years, some of the sites described in the guidebooks of 1982 are difficult to re-locate, especially for someone without prior knowledge of the area. Most mapping at that time was largely carried out on nonrectified aerial photographs with considerable local distortion and the modern digital tools of geographic positioning were not available. Furthermore, destruction of vegetation in some areas and new growth in others has rendered past bush trails and landmarks to be obscure, if not impossible, to recognize, even for the author of the guidebooks. Abandonment of old bush roads has also given way to new gravel access roads and the widening of highways has compromised some outcrops while exposing new ones. A second consideration is the currency of information and ideas. - 85 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 The work in the late 1970’s and early 1980’s was done within a limited context without much consideration of comparable situations globally: this is somewhat ironic because one of Lawson’s goals had been to use this region as a global type example of an Archean granitegreenstone belt. It is also important to recognize that new contributions have been made to the understanding of the geology of this area since the original guidebooks were written. The final consideration is historical. In its time the “Seine-Coutchiching problem”, rooted in the simple but laborious task of geological field mapping, influenced discussions among geologists world-wide but many of the details about the background to the debate have not been adequately recorded, especially in the context of the outcrops themselves. Every effort has been made to avoid duplication of points which are adequately covered in the past guidebooks for this region and the material presented below is meant to serve mainly as a source of information for field trip leaders and new researchers to draw on to supplement the existing documents. The Seine-Coutchiching Problem The discovery of gold at Lake of the Woods in 1878 followed by the building of the C.P.R. line prompted Arthur Selwyn, director of the Geological Survey of Canada, to instruct one of its senior mappers to begin a survey of the geology of this area. Part of the task given to Robert Bell and his young assistants, A.C. Lawson and J.W. Tyrell, was to have the geology carefully worked out as a type locality for the “socalled Huronian system” (Zaslow, 1975, p. 184). Bell left Lawson and Tyrell at Bigstone Bay in spring 1883 to map the shoreline geology and topography while he Figure 3: Simplified geology of the Coutchiching Rapids area using the colour scheme of Figure 2. surveyed a line northward toward Red Lake. When he returned, he checked their results and directed them to work separately, Lawson on geology and Tyrell on topography, before they all reconvened at Rat Portage (now Kenora) at the end of the season. By the beginning of the 1884 field season, Andrew Cowper Lawson had graduated from the University of Toronto with the gold medal in natural science and was put in charge of the project. By then, at age 23, he had been taken on staff at the Geological Survey of Canada and, during that field season, he continued the geological work at Lake of the Woods. A.E. Barlow and W.H.C. Smith independently mapped the topography southward toward Rainy Lake. Lawson’s geological report and maps resulting from the work conducted at Lake of the Woods were published in 1885, questioning the approach the Survey had taken in the mapping Precambrian rocks up to that time (Zaslow, 1975). Lawson (1885) argued that the greenstone which he termed “Keewatin” is clearly intruded by foliated granitoid rocks. He termed these “Laurentian” in keeping with the original GSC terminology introduced by its first director W.E. Logan to denote quartzo-feldspathic basement gneiss upon which all supracrustal rocks had been deposited. Although Lawson’s productivity was admired, his geological interpretations were doubted by more senior geologists. Perhaps because of the attention he gained and the fact that he had by now received an M.A. from the University of Toronto, Lawson was able to prevail on Selwyn to support his further academic advancement. He was allowed to attend courses during the winter in the U.S.A.: he was the first of many GSC geologists to follow this course of action for decades to come. Lawson, with Smith as topographer, began by mapping the canoe routes between Lake of the Woods and Rainy Lake in 1885 and in 1886 focussed on systematic mapping of the Rainy Lake area (Fig. 2). The following season, Lawson filled in the details in the Rainy Lake district while Smith moved eastward along the Seine River Route and southward to the Canada-U.S. border. Lawson recognized a series of metasedimentary rocks exposed along the Coutchiching rapids at the outlet of Rainy Lake into Rainy River at modern-day Fort Frances and International Falls (Fig. 3). He traced these rocks farther northeastward from the type locality into the Rice Bay and Bear’s Passage areas (Figs. 4, 5) where the evidence suggested that rocks of the Coutchiching series are even older than the Keewatin (Fig. 6a). This - 86 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 4: Simplified geology of the Rice Bay area only added further to his dispute with GSC management leading to the heavy editing of his map and first report on Rainy Lake geology (Lawson, 1888). He also provided additional evidence that, rather than being a fundamental basement gneiss, the Laurentian rocks are actually deformed and metamorphosed intrusions that show evidence of cross-cutting the supracrustal rocks. A hand-written version of the report was also submitted for his PhD thesis at John’s Hopkins University (Lawson (1888) where he applied the relatively new technique of optical petrographic description to thin sections from his field samples. Although the accomplishments of Lawson and his colleague W.H.C. Smith were significant, it was the resulting geological interpretation that met continued Figure 5: Simplified geology of the Bear’s Passage area. Note that “Bear’s Passage” refers to the strait linking Swell Bay to Redgut Bay but the terms “Bear Pass” or “Bear Passage” have also been used over time to describe the nearby area. resistance from management and the 1887 report was heavily edited (Saslow, 1975). Lawson, however, largely prevailed and showcased his results at the Fourth International Geological Congress at London in 1888 and the American Association for the Advancement of Science meeting at Toronto in 1889. During the 1889 field season Lawson completed the mapping of the Hunter Island Sheet southeastward of Rainy Lake with Smith but resigned from the Geological Survey of Canada in spring 1890. He worked for a while as a geological consultant in Vancouver but soon accepted a faculty position at the University of California at Berkley, where he pursued an illustrious career for the Figure 6: Portrayals of stratigraphic order at Rainy Lake (1887-1999). - 87 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 next sixty years. Although increasingly accepted overall, some aspects of Lawson’s interpretation of Rainy Lake geology continued to be questioned by his peers. In particular, Coleman (1898) noted that the sedimentary sequence at Shoal Lake (Fig. 7), portrayed by Lawson as belonging to the Coutchiching, includes conglomerate with abundant rounded clasts of both Keewatin greenstone and Laurentian granitoid rocks which are also exposed nearby. The U.S. Survey, which was responsible for mapping the southward extensions of the area covered by Lawson and Smit, took particular exception to Lawson’s interpretations. A special committee on stratigraphic nomenclature for the Lake Superior region was therefore convened by the U.S. Geological Survey and the Geological Survey of Canada and the resulting report was published in the Journal of Geology (Adams et al., 1905). It was critical of Lawson’s interpretation and suggested that there is evidence for the Coutchiching rocks to be younger than the Keewatin, a point that was later re-affirmed Minnesota by Van Hise and Leith (1909). Lawson was irate over the findings of the special committee and R.W. Brock, who was by then the director of the Geological Survey of Canada, invited Lawson to re-study the key parts of his original Rainy Lake map sheet in 1911. Lawson was also given the mandate to examine the rocks farther east along the Seine River toward Steeprock Lake and Sapawe. Several practical developments had ensued since the first mapping, including a gold rush to Mine Centre in the 1890’s, construction of the CNR south line through the area circa 1906 and major forest fires in the region in 1910, generating much new bedrock exposure. By then Lawson was in mid-career and had gained pre-eminence in many aspects of geological science in the western U.S. so that, when he produced his famous Geological Survey of Canada Memoir 40 in 1913 and an accompanying map in 1914, they were accepted without revision. In the memoir he reaffirmed his interpretation of the field relationships in the Rice Bay and Bear Passage areas where he observed the Coutchiching metasedimentary rocks to dip at moderate angles below Keewatin strata. He also issued a bitter challenge to the members of the special committee (Memoir 40, p.13-14): “The facts here recited in regard to this line of contact, particularly near the railway on the shores of Bear Passage and the south end of Redgut Bay, taken in connexion with the relations of the Coutchiching to the granite, appear to me to prove conclusively the superposition of the Keewatin upon the rocks mapped by me as Coutchiching in the report of 1887. I invite the attention of the International Committee and of the U.S. Geological Survey to this section and challenge them in view of the facts there Figure 7: Bad Vermilion Lake area - 88 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 apparent and easily accessible, to deny the relations of the Keewatin and Coutchiching as I mapped and described them a quarter of a century ago. The fact that these eminent authorities have denied in toto the existence of the Coutchiching series as a constituent member of the Archean below the Keewatin, without any attempt to verify the very explicit statement of the evidence in regard to this section contained in the report of 1887 places them in a curious light from the point of view of scientific method.” While forcefully retaining his contention that the Coutchiching rocks at Rice Bay (Fig. 4) and Bear’s Passage (Fig. 5) are positioned below the Keewatin, Lawson also admitted to an error in the Bad Vermilion Lake area (Fig. 7). There he proposed the new term “Seine Series” for the conglomeratic metasedimentary rocks that he had previously included as a basal part of the Keewatin. He now confidently placed the basal Seine conglomerate unconformably above both the Laurentian granitoid rocks and the Keewatin metavolcanic rocks in that area (Fig. 6b). He also noted the presence of trough crossbeds in the sandy portions of the Seine series near Old Mine Centre (Fig. 7) and used the newly recognized criterion of determining the directions of stratigraphic younging using their shapes. This allowed him to define a synclinal fold within the Seine sedimentary sequence in a narrow belt extending eastward along the Seine River (Fig. 2). He also recognized that the Seine Series locally extended farther eastward beyond the Rainy Lake area. In so doing, he mistakenly classified sedimentary rocks at Sapawe (Fig. 1, then known as Iron Spur) as part of the Seine Series. The granitoid Blalock stock cuts metasedimentary rocks discordantly at that locality so Lawson introduced the new term “Algoman” for such intrusions which he believed to be generally younger than the Seine (Fig. 6b). Although subsequent studies at Sapawe support Lawson’s contention of a late-tectonic intrusion, they also have consistently portrayed the intruded sedimentary rocks there as part of the Coutchiching rather than the Seine. Nonetheless, Lawson offered other acceptable field and petrographic distinctions that argue for the existence of a younger set of Algoman granitoid intrusions in the Rainy Lake area proper. They tend to contain higher proportions of K-feldspar than the dominantly sodic tonalitic rocks which comprise the Laurentian. He also mapped a narrow band of conglomerate near Hopkins Bay, west of Rice Bay (Fig. 2), and tentatively correlated it with the Seine sequence: at that locality he also presented strong evidence that the conglomerate is cut by younger Algoman granitoid rocks. As he had in the 1880’s Lawson used the International Geological Congress, this time at Ottawa in 1913, to promote his revised view of Rainy Lake geology and Precambrian stratigraphy in general. A debate was staged between Lawson and C.K. Leith to present arguments for and against the findings of the special committee: as later recalled by Leith, Lawson had a “slashing style” and “while I came out feeling I had presented the facts, I also felt Lawson had chewed me up and thrown me to the wolves” (Dott, 2001, p.1007). Lawson’s views were further solidified during the 1913 International Congress field trip that he led on Sunday, August 17th for approximately 90 participants who had traveled by C.N.R. to the Mine Centre and Bear’s Passage train stations after similar visits at Iron Spur and Steep Rock Lake the day before. At Mine Centre, participants were given the option of riding in horsedrawn wagons from the station to the Golden Star Mine along the Shoal Lake Road or taking a short boat ride across Bad Vermilion Lake to a walking trail leading to the mine (Fig. 7). At Bear’s Passage, one group was assigned to a boat trip which visited lakeshore outcrops along Redgut Bay and Bear’s Passage (Fig. 5) while others made a traverse though a similar geological section exposed relatively new rock cuts along the C.N.R. railway line. A carefully prepared itinerary and field guide for both of the historic localities (Uglow, 1913) allowed Lawson to illustrate the nature of each of his five stratigraphic units and his observations on the relationships among them. Lawson’s revised interpretation of the geology of the Lake of the Woods and Rainy Lake regions prevailed for another decade (Bruce, 1925) before the idea that there was still a “problem” was revived by F. F. Grout of the Minnesota Geological Survey. Grout (1925) reviewed the field relationships at the type locality of the Coutchiching near International Falls (Fig. 3) and at outcrops which Lawson had assigned to the Seine in the area of Neil Point farther to the east. Grout affirmed Lawson’s use of cross-bedding as an indicator of stratigraphic younging at Neil Point but offered an alternative overall interpretation which placed both the Seine and Coutchiching above the Keewatin. He then moved northeastward across the international boundary to a locality south of Bear’s Passage known as Morton Island (Fig. 5). There he used the newly recognized - 89 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 field criterion of graded bedding to deduce that the direction of stratigraphic younging in the Coutchiching is northward and away from the Keewatin volcanic rocks. He also visited the key localities at Rice Bay, Bear Passage, Shoal Lake and also at Jackfish Lake southwest of Steeprock Lake. In all cases he raised objections to Lawson’s positioning of the Coutchiching and placed it above both the Keewatin and the Seine (Fig. 6c). The problem expanded in scope when, during a subsequent field examination with T.L. Tanton of the Geological Survey of Canada, it became apparent that Grout had made a significant observational error on the Minnesota side of the boundary by misinterpreting an intrusion breccia to be a conglomerate of sedimentary origin. Tanton (1927) took great pains to publicly point this out at a Geological Society of America Precambrian Symposium, noting that the error was made by “a Minnesota geologist”. Grout, in a discussion of Tanton’s paper, duly acknowledged his own mistake but also stated combatively that a “structure section sketched in the field by Tanton shows more errors than any before”. Rather than resolving the problem, these exchanges only served to accentuate it. Some years later, Tanton (1936) made a comparable error by misinterpreting the Seine conglomerate at Shoal Lake to have been intruded by the Laurentian granitoid rocks rather than being deposited above it (Fig. 6d). Apart from his local mistake, Grout’s overall arguments found some traction and provided the impetus for additional field work. J.E. Hawley, a graduate of the University of Wisconsin and a professor at Queen’s University, was well versed in stratigraphy and structural geology. Along with a review of the Shoal Lake area, he conducted a study of the Seine and Coutchiching eastward though Jackfish Lake and past Sapawe. He concluded (Hawley, 1930) that part of the problem was, in some localities at least, that the contacts between the Coutchiching and Keewatin are arguably occupied by faults so that attitudes of strata alone provide inconclusive evidence of stratigraphic order. F.F. Grout also remained influential at that time and recommended the Seine-Coutchiching Problem to P.L. Merritt who conducted a study of the entire corridor from Rainy Lake eastward along the Seine River watershed for his Ph.D. thesis at Columbia University (Merritt, 1934). His conclusions concerning the two metasedimentary sequences supported Grout’s interpretation (Fig. 6c) and he suggested that, with the notable exception the clastic rocks at Rice Bay and Bear’s Passage, the term Coutchiching should be abandoned altogether and that all other sedimentary units should be included in the post-Keewatin, Seine Series above the basal conglomerate. Like Hawley, Merritt also provided detailed documentation of a fault contact between the Keewatin and sedimentary units at various localities and traced a continuous fault from Calm Lake eastward through Sapawe as far as Dog Lake, 60 km north of Thunder Bay: this is known today as the Quetico Fault. He also proposed (Merritt, 1934, p. 371) that “the fault movement along the contact is believed to combine a horizontal shear with an associated overthrust to the south”. Grout had a further influence on the expanding Seine-Coutchiching problem in that he inspired Francis Pettijohn, his field assistant during the work at Rainy Lake, to take on pioneering work in the study of Archean sedimentary rocks in general. Pettijohn did his undergraduate and graduate work at the University of Minnesota and, in accepting the Penrose Medal for 1975 from the Geological Society of America, he acknowledged the importance of Grout’s tutelage as well as the short time that he spent studying with A.C. Lawson at Berkley in 1927-28 to learn more about the alternative view. Pettijohn’s Ph. D. thesis documented the Abram Lake conglomerate in the Minnitaki Lake area in part because it resembled both the Seine conglomerate at Rainy Lake and the Ogishke conglomerate at Knife Lake Minnesota (Fig. 1). He later summarized his findings at all three of these localities as well as at several others in the northern Lake Superior region (Pettijohn, 1937) to also conclude that the majority of sedimentary units are arguably younger than the Keewatin. He also raised the possibility, however, that not all of the units included in the Keewatin need be of the same age and that local intercalation between volcanic and sedimentary rocks may locally be possible. As important as these insights ultimately proved to be, Pettijohn’s resolution of the Seine-Coutchiching problem also called for the abandonment of both of Lawson’s local units (Seine and Coutchiching) in favour of an overarching “Knife Lake Series” composed of similar rocks which had precedence of definition in the geological literature in Minnesota. This also reinforced Grout’s views and a summary paper on the topic (Grout et al., 1951) notably included a section entitled “No Coutchiching Recognized in Minnesota”. Nonetheless, the “SeineCoutchiching problem” was kept alive intermittently well into the 1970’s even to the extreme point of academic speculations that no unconformities existed - 90 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 at all and that there was simply a proximal to distal lateral facies equivalence (Fig. 6e) among broadly age-equivalent Keewatin, Seine and Coutchiching rocks (Bass, 1961; Ayres, 1971; Mackasey et al., 1974; Goodwin, 1977). Sam Goldich who was a graduate of the University of Minnesota rejoined that institution in 1948 as a professor and director of the Rock Analysis Laboratory where he and Alfred Nier gained international reputations as pioneers in isotope geochemistry and geochronology. An outcome of that work was the application of geochronological methods to the resolution of stratigraphic problems in the Lake Superior region (Goldich et al., 1961; Goldich, 1968). Goldich compromised on the question of the SeineCoutchiching problem by favouring the term “Knife Lake Group” over “Seine Group” above the Keewatin but also allowed for the possibility (using a question mark for emphasis) of the existence of Coutchiching metasedimentary rocks below it. Goldich tackled the geology of the Rainy Lake area head on by supporting three field-based M.S. theses at the University of Minnesota (Alt, 1959; Frye, 1959; Peterman, 1959) and the resulting maps and samples became the basis for on-going geochronology and geochemical studies (Peterman et al, 1972; Goldich and Peterman, 1980). With time, Lawson’s original terminology and interpretation of stratigraphic order was largely supported by the data but with the added implication that all of the constituent rock-forming events took place in less than 100 million years with only local evidence for younger post-metamorphic retrogression. At the time of the studies by Goldich and his colleagues an important fact remained: apart from Grout’s work in Minnesota, no geologist other than A.C. Lawson had mapped systematically in the Rainy Lake area. This had been undertaken by him at a scale of one inch to four miles in 1885-87 and, with the assistance of H.C. Cooke and R.C. Wallace, at one inch to one mile in 1911. New mapping in greater detail was therefore ultimately undertaken by the Ontario Division of Mines beginning in the 1970’s (Davies, 1973; Blackburn, 1973; Harris, 1974; Wood et al., 1980 a, b; Fumerton, 1985). The outcrop mapping of Fred Harris is perhaps the most notable because it provided an advanced level of lithostratigraphic detail, at a scale of one inch to ½ mile, while covering the historically controversial Rice Bay and Bear’s Passage areas. He also provided new local evidence for stratigraphic younging in the Keewatin strata including the first recorded use of the shapes to pillows in basaltic flows in this area. Harris (1974) avoided the use of the historical stratigraphic terms but his table of formations tends to support Lawson’s original interpretation of metasedimentary biotite schists at the base of the sequence. John Wood provided a comparable level of mapping at Mine Centre (Wood et al., 1980 a, b) with a focus on the Seine and Coutchiching metasedimentary rocks (Wood, 1980). Companion studies of the geology in the Minnesota portion of the Rainy Lake area were conducted under the auspices of the Minnesota Geological Survey and the US Geological Survey (Ojakangas, 1972; Southwick, 1972; Southwick and Ojakangas, 1979; Southwick and Sims, 1980). Poulsen (1980) made extensive use of the report and maps of Harris (1974) as a foundation for structural and metamorphic studies in the Rice Bay and Bear’s Passage areas. It eventually became clear, however, that one of the flaws in Lawson’s original interpretation was that it relied on the assumption that structural superposition of the Keewatin above the Coutchiching equates to stratigraphic superposition as well in rock packages that are arguably overturned (Poulsen et al., 1981). This led to a revised interpretation of stratigraphic order (Fig. 6f) but one without geochronological constraint. The full essence of the Seine-Coutchiching problem was ultimately clarified by application progressively improved methods of U-Pb analysis of zircons (Davis, 2023). Strategic sampling of each of Lawson’s five lithostratigraphic units across several sites where field relationships had been established (Davis et al., 1989; Davis et al., 1990; Fralick and Davis,1999) provided the results that form the basis of the current chronostratigraphic chart for the area (Fig. 6f). Keewatin metavolcanic rocks and Laurentian metaplutonic rocks were shown to be of similar age (circa 2727 Ma) whereas detrital zircons from the Coutchiching suggested a younger age (circa 2700 Ma) and a granitoid clast and detrital zircons from the Seine conglomerate even younger (<2693 Ma). The age of crystallization of the Algoman intrusions was estimated to be in the range of 2693 to 2684 Ma. In total, the geochronological constraints have shown that most of the historical interpretations, including those of Lawson, had both merits as well as flaws whereas the lateral facies concepts which were so widely and uncritically accepted in the 1970’s have proven to be entirely untenable. The net result, however, is that - 91 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Lawson’s placement of the Coutchiching beneath the Keewatin on the grounds of the dip of strata alone was the main source of geological error. It was not that he did not understand the difference because, in Memoir 40, he appears to have been the first geologist to have compare directions of dip to the directions of younging (way-up or bedding top) in cross-bedded arenite of the Seine sedimentary unit. Lawson (1913) did acknowledge that he had learned about the utility of truncated cross-bedding from his field assistant J.D. Trueman, then a graduate student who in turn had been taught this by W.O. Hotchkiss at the University of Wisconsin (Dott, 2001). Hotchkiss was also familiar with upward-fining grain size variation in sandstone mudstone sequences but this was not yet in common use and therefore the significance of graded bedding, as preserved at Bear Passage area, was not yet appreciated by Lawson in 1913. It took the work of F.F. Grout (1925) to demonstrate that the graded beds at Morton Island indicate that the Coutchiching beds there stratigraphically overlie the Keewatin. Lawson also certainly would have been aware of the stratigraphic use of pillowed volcanic flows, as advocated by Morley E. Wilson (1913) in Memoir 39 of the Geological Survey of Canada, but was of the opinion that this method was unsound because he believed that pillows, then referred to as ellipsoidal structures, were of intrusive origin (Lawson, 1912). As one looks back, the Timiskaming-Keewatin problem evolved along similar paths as another great geological debate that played out during much the same time frame, the Highlands controversy of Scotland. That problem also involved many observers who were focussed on small, geographically separated parts of a bigger problem and it has been said that, in many cases, they did not know what they did not know (Oldroyd, 1990). A case in point is the famous anecdote concerning T.L. Tanton of the Geological Survey of Canada and E.B. Bailey of the British Geological Survey (Dott, 2001). Tanton, who had graduated from the University of Wisconsin in 1915 under the supervision of C.K. Leith, led a group of Princeton geologists on a tour of Rainy Lake as part of their geological trip across Canada by rail in 1927. The group included two eminent guests from overseas, L.W. Collett from Switzerland and E.B. Bailey of Britain (Bailey, 1927). Tanton demonstrated the utility of cross-bedding and graded bedding to determine wayup in metasedimentary rocks, using Lawson’s examples from the Seine Group at Shoal Lake and Grout’s Coutchiching outcrops at Morton Island respectively. This resulted in Tanton’s inclusion as a participant on a reciprocal visit to Scotland where he convinced Bailey that the Dalradian strata at Ballachulish are overturned (Tanton, 1930; Bailley, 1930, Dott, 2001). Another point of communality between the two controversies is that the reputations of the observers, especially Sir R.I. Murchison in the Highlands and A.C. Lawson at Rainy Lake, tended to get in the way of the geological facts. This should not overshadow the reality, however, that in his first, youthful burst of mapping from 1882 to 1889 and in his mid-career re-study from 1911 to1913, Lawson identified the five lithological building blocks which are representative of the architecture of virtually every Archean greenstone belt in the world. As Oldroyd (1990) has pointed out for the Scottish Highlands controversy, it is not only about who was right and who was wrong but it is also about the process of narrowing in on a consensus view based on the facts at hand. Similar sentiments were expressed by Lawson himself in the introduction to his 1913 report which fuelled the Seine-Coutchiching problem in the first place. The overall lesson of the problem seems to be that: “Science is never ‘settled’ but evolves by the accumulation of facts, new ideas and vigorous open discussion and debate. Consensus is irrelevant in science; only truth matters.” (Dewey and Ryan, 2022, p.1834). Rainy Lake Wrench Zone Poulsen (1986b) introduced the term “Rainy Lake wrench zone” to distinguish rocks between the E-W Quetico Fault and the ENE Rainy Lake – Seine River Fault from the Quetico metasedimentary belt to the south and the main mass of the Wabigoon granite-greenstone belt to the north (Fig. 8). The rationale for highlighting the wrench zone involved many different geological aspects (Poulsen, 1986b) but the most prominent are the distinctive lenticular. s-shaped lithostratigraphic domains which merge with the discordant boundary faults. Broadly similar patterns are also evident in the steep metamorphic foliation which affects the Seine as well all of the older lithostratigraphic units. This is also the area in which the Seine – Coutchiching problem mainly played out and where generations of geologists contributed to the understanding of diverse aspects of its geology. It is also the focus of this field guide which can be used to illustrate the major lithostratigraphic - 92 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 8: Simplified geological map of the Rainy Lake Wrench Zone. units that comprise the wrench zone as well as the related topics of deformation, metamorphism and metallogeny. Keewatin Lawson applied the term “Keewatin Series” to all of the Archean metavolcanic rocks in the Rainy Lake area mainly to distinguish them from foliated quartzo-feldspathic rocks of probable plutonic origin. He initially did this in a descriptive way (Lawson, 1885) but his Rainy Lake reports (Lawson, 1887; Lawson, 1913) also provided petrographic detail and genetic interpretation. The Keewatin rocks within the wrench zone include lithofacies which are common to Archean greenstone belts in general. Mafic volcanic rocks predominate in the northwestern part of the zone, particularly at Windy Point, Nickel Lake and Pocket Pond. The rocks at these locations were metamorphosed to amphibolite facies assemblages so that primary features are difficult to document in the resulting foliated mafic tectonites. In places where strain is moderate it is relatively easy to identify pillows and varioles but the level of distortion in many places (Fig. 10a) makes it difficult to confidently use the shapes of pillow to confidently define directions of younging (Borradaile and Poulsen, 1981). A notable exception is at Pocket Pond (Fig. 9a) where adequate evidence of stratigraphic polarity is preserved (Fig. 10b). Felsic metavolcanic rocks predominate in the southeastern part of the wrench zone where they are commonly intercalated with rocks of andesitic composition (Fig.10c), The rhyolitic rocks are commonly quartzphyric and included both coherent (Fig. 10d) and volcaniclastic (Fig. 10e) facies. Figure 9: Pocket Pond locality - 93 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B v C D E F Figure 10: a) westward plunging shape lineation defined by deformed pillows in metabasalt, Nickel Lake area; a) Pillow basalt, Pocket Pond; c) Amygdaloidal basalt, Port Arthur Copper; d) spherulitic rhyolite, Ottertail east; e) Rhyolite Breccia, Wind Bay; f) volcaniclastic ferropicrite, Belacoma area. One outstanding unit within the Keewatin sequence is composed of a distinctive ultramafic volcaniclastic rock (Fig. 10f) which is exposed in the Grassy Portage Bay area (Fig. 4). The unit was first recognized by Harris (1974) who classified it as an intermediate volcanic rock, mainly because of its common bright green, chloritic appearance along with volcaniclastic textures. Poulsen (1980) prosaically termed it “magnetic green rock” which is composed mainly of Mg-chlorite plus actinolite and magnetite. He provided lithogeochemical analyses to show that the rock has an ultramafic bulk composition but incorrectly classified it as a komatiite, a rock type with which it shares only some chemical similarities. He also compared the unit, both chemically and texturally to the betterknown picritic Steep Rock Ashrock approximately 100 km to the east and suggested that their separation might be due to dextral displacement on the Quetico Fault (Fig. 1). Steve Schaefer conducted a study of the ultramafic units at both localities and confirmed their volcanic origins (Schaefer and Morton, 1991). He also provided the acronym GUP (Grassy Portage - 94 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Ultramafic Pyroclastic) for the rocks at the Rainy Lake locality. Goldstein and Francis (2008) pointed out the differences in the chemical composition of this unit compared to komatiites: the GUP shows higher FeO, TiO2 and incompatible elements (e.g., Nb) as well as displaying fractionated rather than flat rare earth element patterns. Goldstein and Francis (2008) reclassified the rocks as pyroclastic ferropicrites, noting that they are examples of relatively rare Ferich volcanic varieties that were likely derived from partial melting a mantle source that was enriched in Ti and rare earth elements. A further characteristic of the GUP is that it contains microdiamonds which were discovered in 2008 by MetalCORP Limited at the Beaver Pond Occurrence (Hinz et al., 2010). All of these observations have significantly improved the understanding of this unusual volcanic unit but questions remain regarding its stratigraphic position and regional significance. Despite the remarkable similarity to the Dismal Ashrock at Steeprock, the notion of a strike-slip separation of the same unit is still feasible but not fully demonstrated. Tomlinson et al., (2003) reported a maximum age of 2780.4 +/-1.4 Ma for the Dismal Ashrock based on analyses of inherited zircons and argued that it is feasible for it and overlying basalts (Witch Bay formation) to be as young as other sequences in the Western Wabigoon Subprovince: by extension, this would include the mafic-ultramafic volcanic units in the Grassy Portage Bay area. If the ages of the ultramafic volcanic rocks at the two distant localities prove to be different, however, it would mean that an alternative explanation for their similarity would have to involve operation of similar processes at different times. In that case the communality might be sought in the mantle composition and depth that led to the formation and deposition of these unusual rocks. Interflow metasedimentary rocks comprise a common but volumetrically small component of the Keewatin sequence. Although in places these rocks could be mistaken as providing evidence for interdigitation with Coutchiching biotite schists or with volcaniclastic rocks of intermediate composition, in most cases, they are arguably metalliferous, synvolcanic sedimentary units which range from pyritic mudstone and minor sandstone, to chert-magnetite banded iron-formation (Fig. 11a) and pyritic massive sulfide deposits (also termed sulfide facies ironformation). The sulfide-bearing varieties were targets for possible sulfur production in the period around World War I when, particularly at Nickel Lake, they were noted to contain anomalous concentrations of CoNi-Zn-Cu. In places zinc is also a locally anomalous component and, at Pocket Pond, the small sphalerite lenses discovered in the 1970’s are associated with iron-formation intercalated with metabasalt (Fig. 9). From a strictly geological perspective the presence of laterally extensive interflow units proves valuable for establishing a sense of stratification within the Keewatin because they not only can be mapped discontinuously in outcrop and drill core but cane be easily traced accurately by magnetic and electromagnetic surveys. A.C. Lawson’s 1914 geological map of Rainy Lake also includes two mafic plutonic rock types which he regarded to be part of the Keewatin sequence: extensive units of what he termed hornblende gabbro in the Grassy Portage Bay area (Fig. 4) and anorthosite in the Bad Vermilion Lake area (Fig. 5). Subsequent mapping has shown that he underestimated the total volumes of mafic plutonic rock in both cases and this was with good reason. It is now well-understood that thick, mafic submarine lava flows are capable of slow cooling rates to produce what can easily be accepted as a “gabbro-textured” facies that grades vertically and laterally over short distances into finer grained basaltic rocks, making their visual distinction from plutonic equivalents difficult. Furthermore, where amphibolite facies metamorphism has affected mafic volcanic rocks, recrystallization tends to coarsen the texture and obscure primary features: this is certainly the case in the northwestern western part of the Rainy Lake Wrench Zone. Finally, considerable local variations in textural detail are common in layered mafic intrusions (Fig. 11b, c, d) that that are difficult to map at a reconnaissance scale. What Lawson did map, however, were two of most extensive, distinctive and homogeneous plutonic phases, leucogabbro in the Grassy Portage Intrusion (Fig. 11e) and coarse anorthosite in the Bad Vermilion intrusion (Fig. 11 f). Detailed mapping by Hodgson (1959) at Grassy Portage Bay and by Harris (1974) at both localities provided much better definition of the full extents of these intrusions. Ashwal et al. (1983) undertook a more advanced petrological and geochemical study of the Bad Vermilion anorthosite and concluded that it represents the remnants of a subvolcanic magma chamber from which aliquots of magma had been extracted as extrusive lava flows. Poulsen and Hodgson (1985) reviewed the disposition of the different phases of both intrusions and the sulfide - 95 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B D C D F E Figure 11: a) chert-magnetite iron-formation, Pocket Pond; b) thin-layered gabbro, Grassy Portage intrusion, west side of Redgut Bay; c) thick-layered gabbro-melagabbro, Northrock East trenches, Grassy Portage intrusion; d), Pegmatitic Gabbro, Northrock E trenches, Grassy Portage intrusion; e) leucogabbro, Grassy Portage intrusion; f) coarse anorthosite, Bad Vermilion intrusion, Scott Islands (the edge of the compass in the bottom left measures 10 cm) and oxide mineralization within them, providing support to the idea that they are examples of synvolcanic layered intrusions resulting from cumulus growth and magma fractionation. Both intrusions also received attention for their economic potential during exploration programs for Cu-Ni-PGE sulfide and Ti-V oxide mineralization (Hinz et al., 2010). The Bad Vermilion Intrusive complex and the surrounding metavolcanic rocks have recently been described as an arc-related “ophiolite” sequence (Wu et al., 2016) but this is highly unlikely given the dominance of rhyolite in the volcanic section and the absence of both peridotite and sheeted dikes. The ages of the Keewatin units were largely unknown until the mid-1970’s despite many attempts to apply modern geochronological methods (Goldich, 1968; Tilton and Grunenfelder, 1968; Hart and Davis, 1969; Peterman et al., 1972). At that point improvement in analytical precision and accuracy allowed U-Pb geochronology on carefully constrained samples to impact stratigraphic interpretations (Davis, 2023). Davis et al. (1988) applied these methods in the Rainy Lake Wrench zone to show that the units which were historically classified as Keewatin formed around 2727 - 96 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Ma in a remarkably short interval of five to six million years (Fig. 5f). This included direct analysis of rhyolitic rocks from both the northwestern and southeastern part of the zone as well as the indirect constraint of mafic (Grassy Portage Gabbro) and felsic (Mud Lake trondhjemite) that cut the volcanic rocks. More recent attempts to provide additional ages of volcanic rocks near the Bad Vermilion Intrusion have proven to be unsuccessful in light of the lower analytical precision and accuracy of the methods employed (Wu et al. 2016). Coutchiching Lawson’s 1914 map of Rainy Lake outlines three areas of Coutchiching rocks labeled as “mica schist, paragneiss and phyllite”. The most extensive area occurs south of the Seine River Fault in Quetico Subprovince where the term Quetico metasedimentary rocks also applies (Fig. 12a). The other two major localities are located within the Rainy Lake Wrench Zone: a southern belt extending from Fort FrancesInternational Falls northeastward along Swell Bay and a northern one as a partially annular zone within the Rice Bay Dome (Figs. 2, 3, 4). Grout (1925) A B C D E F Figure 12: a) Quetico metasedimentary rocks, Bleak Bay area; b) thick-bedded wacke, Sandpoint Island c) graded beds cut by ENE cleavage, Morton Island (N to top of photo); d) knotty biotite schist containing retrograded staurolite and garnet, Great River Road; e) graded beds, Bear’s Passage boat launch; f) graded beds in greenschist facies turbidites, Old Station Road (diagonal lines are glacial striae). - 97 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 who was the first to recognize graded bedding in the metasedimentary rocks in the Swell Bay belt (Fig. 12b, c) and to apply it to establish local stratigraphic polarity at Morton Island (Fig. 13). This was confirmed by Merritt (1934) who also notably interpreted the colour banding of graded rocks in the Swell Bay as “varves”. R.W. Ojakangas who often lamented that the Canadian glaciers had been cruel to Minnesota re-examined the sparse exposures of the Coutchiching rocks at Ranier, Minnesota near Lawson’s type locality (Fig. 3). He described the rocks there as metagreywacke noting that they originally consisted of alternating beds of greywacke sandstone (or simply wacke) and mudstone deposited, not by glacial processes but by turbidity currents on submarine fans (Ojakangas, 1972; 1982). Most authors who have studied the belt of Coutchiching rocks along Swell Bay have also noted that they been clearly intruded by younger granitoid rocks (Algoman) and that the rocks on the northern shore of Swell Bay display amphibolite facies, pelitic metamorphic assemblages (Fig. 12d) involving biotite, muscovite, garnet, cordierite and staurolite (Ojakangas, 1982; Poulsen, 1980). The higher metamorphic grade has also been implicated by many authors for obscuring primary features such as graded bedding in the metasedimentary rocks. The most contentious interpretations of the stratigraphic significance the Coutchiching rocks in the Swell Bay corridor result from observations in the Bear’s Passage area (Figs. 4, 14). The Keewatin at this locality consists of a northwestward-dipping section composed of the upper part of the southeastwardyounging Grassy Portage layered intrusion overlain by a thin unit of what are arguably mafic metavolcanic Figure 13: Morton Island locality (adapted from Poulsen and Wood, 1982) rocks. The staurolite-bearing metasedimentary rocks, although locally folded, also dip to the northwest and are cut discordantly by granodiorite of the Bear’s Passage intrusion (Fig. 4). Although minor reversals in polarity of grading suggest local folding within the Coutchiching rocks in the Bear Passage area, a good quality exposure at their contact with Keewatin strata (Fig. 12e) demonstrates that the metasedimentary rocks are overturned (Poulsen, 1980). This plus the observations at Pocket Pond (Fig. 9) and Morton Island (Fig. 13), provides the geological evidence in favour of the Coutchiching being younger than the Keewatin. The most conclusive evidence, however, ultimately came from U-Pb analyses of detrital zircon in biotite schist near Tunnel Bay and in well preserved greenschist grade metagreywacke (Fig. 12f) northeast of Shelter Cove (Fig. 6) which represents the northeastward extension of the exposures at Morton Island (Davis et al. 1989). The age of the Coutchiching is constrained by the youngest detrital zircon grains at approximately 2704 Ma and by the circa 2692 Ma age of across-cutting felsic dike (Fig. 5 f). The presence of much older detrital grains (2930, 2940 and 3060 Ma) also suggested a potential contribution of detritus to the Coutchiching from a source area comparable to the Marmion domain north of the Quetico Fault extending in the Steeprock Lake area (Fig. 1). Similar conclusions were reached for the Quetico metasedimentary rocks by Davis et al. (1990). Laurentian Lawson (1887) used the term Laurentian for variably foliated granitoid rocks in general but by 1913 he only applied it at only three localities, Bad Vermilion lake, Figure 14: Bear’s Passage locality - 98 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Rice Bay and Grassy Island near Neil Point (Fig. 2). The Laurentian granitoid rocks at Bad Vermilion Lake occupy three sinuous bodies that are broadly co-spatial with the Bad Vermillion layered intrusion (Fig. 7). The sodic granitoid rocks range from tonalite (Fig. 12a, b) to trondhjemite (Fig. 12c) in composition (Goldich and Peterman, 1980) and likely occupy the remnants of sills that are broadly concordant with stratigraphic layering in the surrounding northward-younging Keewatin volcanic rocks (Fig. 7). It is noteworthy that Lawson (1887) was the first to suggest that they might be syn-volcanic, subvolcanic intrusions. This point was verified by Davis et al. (1989) who established nearly identical ages of 2728 Ma for the intrusive Mud Lake trondhjemite near the Stellar gold deposit and an overlying rhyolite west of the Port Arthur copper deposit. The Laurentian rocks at Grassy Island likely represent an isolated remnant of the same stratigraphic section to the southwest (Fig. 2). A noteworthy characteristic of some outcrops of tonalite, particularly near gold-bearing quartz veins, is a quartz-rich sericitic rock (Fig. 12b) that was termed “protogene” by the early gold explorers in the region and results from plagioclase-destructive metamomatism related to carbonatization associated with brittle-ductile shear zones in the tonalite (Diamond and Marshall, 1990). The Laurentian rocks exposed in the core of the Rice Bay Dome (Fig. 4) are much more difficult to interpret, in part due to overprinting deformation and amphibolite facies metamorphism. Lawson’s 1914 map classified them to include as an inner body of granite and granite gneiss with an intrusive relationship with an outer annulus of Coutchiching biotite schist. Subsequent petrographic studies documented the distinctions among the lithologies (Frye, 1959; Peterman, 1959) and the term “paragneiss” was ultimately given to the innermost rocks. The existence of a large pluton was questioned and both the paragneiss and biotite schist were considered to be different components of the Coutchiching (Peterman, 1959; Peterman et al. 1972). By the same token, however, a small volume of deformed quartz-feldspar dikes and sills were shown to cut the paragneiss within the dome (Peterman et al. 1972). Harris (1974) took much the same approach and, apart from areas where the minor granitoid dikes and sills were particularly abundant, he mapped most of the interior of the Rice Bay Dome as being composed mainly of “biotite-feldspar-quartz schist” which he also assigned to the lower metasedimentary unit (i.e. the Coutchiching). Goldich and Peterman (1980) continued to view the rocks in the interior of the Rice Bay Dome as being composed of paragneiss derived from epiclastic sedimentary rocks but they also presented chemical data to show that they are different from the Coutchiching biotite schists and metagreywackes. Poulsen (1980) used the non-genetic term “grey gneiss” for rocks in the interior of the Rice Bay Dome (Fig. 15d, e) and also showed that they are fundamentally different in chemical composition from the annulus of biotite schist that envelopes them (Fig. 4). The minor deformed quartz-feldspar porphyry dikes (Fig. 15 e, f) are in, turn, different in chemical composition from both the grey gneiss and biotite schists (Poulsen. 1980; Goldich and Peterman, 1980). Dick Ojakangas (personal communication, circa 1980) provided the novel suggestion that some of the grey gneisses actually may have been felsic volcanic rocks rather than felsic intrusions. This prompted Poulsen (1984) to opt for the uninspiring descriptive term quartzo-feldspathic gneiss to distinguish the Laurentian rocks from the Coutchiching biotite schists. Davis et al., (1989) reported a U-Pb zircon age of 2725+/-2 Ma from a sample of the quartzofeldspathic gneiss near Moran’s Bay (Fig. 4) to demonstrate its probable chronological equivalence with both the Keewatin rhyolite and the Laurentian Mud Lake trondhjemite in the Bad Vermilion Lake area. One of the notable lithogeochemical attributes of the biotite-rich Laurentian gneisses within the Rice Bay dome is their local deficiency in Na and Ca and their excess in Mg and Fe relative to their high silica and low Ti contents (Goldich and Peterman, 1980; Poulsen, 1980). One explanation for this is that they were locally subjected to plagioclase-destructive metasomatism which would also explain the presence of staurolite, andalusite and/or cordierite within them at specific sites. Such alteration in well-known in the environments of volcanic-associated massive sulfide deposits. Beakhouse (1984) evaluated this possibility in the western part of the Rice Bay dome where he identified the metamorphic assemblage quartzchlorite-garnet-anthophyllite-staurolite with possible large relict grains of cordierite at one locality and common garnet over a larger area. Teck Corporation subsequently verified these mineralogical anomalies with further mapping and lithogeochemical surveys to conclude that the alteration is likely related to pyritic massive sulfide mineralization within an iron- - 99 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B C D E F Figure 15: a) Bad Vermilion tonalite, Mine Centre; b) Sericitized tonalite (“protogene”), Mine Centre c) Mud Lake trondhjemite, Stellar gold property; d) grey quartzo-feldspathic gneiss cut by leucocratic dikes, Rice Bay, e) quartz-phyric grey gneiss cut by quartz-feldspar-phyric dike, Laurentian gneiss unit, Moran’s Bay; f) deformed quartz (dark) and feldspar phenocrysts in qfp dike, Moran’s Bay formation unit near the outer part of the Rice Bay dome (Alderman, 1988). In summary, despite incremental advances in establishing the geological facts concerning the Laurentian gneiss of the Rice Bay dome, considerable uncertainty remains about its origin. It has been established to be age equivalent and compositionally similar to both the Keewatin and Laurentian rocks in the Bad Vermilion Lake area but much study is required to establish its stratigraphic significance with respect to the Coutchiching and Keewatin rocks which structurally overlie them. The weights of evidence suggest, however, that the definitively intrusive aspects of the dome are attributable to the minor volume dikes and sills for no absolute ages have been established. On lithogeochemical grounds they may represent a phase on the younger Algoman intrusive suite (Goldich and Peterman, 1980). - 100 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Seine Of all of Lawson’s many achievements at Rainy Lake, it was arguably the recognition of the sedimentary rocks of the Seine Series, the interpretation of their probable depositional paleoenvironment and the demonstration of a high-angle unconformity beneath them that have best withstood the test of time. Although the overall level of exposure is uneven, the critical localities where this is best illustrated are located between Shoal Lake and Bad Vermilion Lake in the Mine Centre area (Fig. 7). In particular, exposures of the basal conglomerate (which Lawson termed “fanglomerate”) near the Golden Star Mine (Fig. 16) and the overlying arenite facies exposed on islands in Shoal Lake to the south provided the diagnostic evidence for Lawson’s arguments. Figure 16: Simplified geology of the Golden Star locality. Rocks of the Seine series occupy the area to the southeast of the trace of the unconformity and the underlying rocks of the Laurentian and Keewatin are located to the northwest. The S-shaped configuration of the unconformity trace is likely meaningful, not only because it mimics larger patterns in the wrench zone as a whole (Fig. 8) but also because the northsouth segment reflects lower than average intensity of superimposed strain. Lawson was the first to note that this is in part responsible for the convincing preservation of contact relationships. The basal Seine conglomerate dips shallowly southeastward whereas Lawson showed that an interflow chert-carbonate unit within the Keewatin dips moderately northward. A relatively minor refinement (Pouslen and Wood, 1982) is that pillowed metabasalt overlies the chertcarbonate marker and indicates a northwestwardyounging for the Keewatin rocks. In other words, there is evidence for back-to-back younging across a highangle unconformity. Although weakly aligned due to overprinting strain, a critical point of observation is that clasts in the basal conglomerate show no evidence of a prior metamorphic foliation (Fig. 17a). The derivation of the coarse gritty matrix of the basal conglomerate from the underlying Laurentian tonalite is also clearly evident when compared to the intrusive rocks below the nonconformity. The shallow dipping basal conglomerate (Fig. 17b) to which Lawson ascribed an alluvial origin has been mapped along a persistent ridge of fair outcrop (Fig. 16) but topographically recessive arenite which overlies it to the east is poorly exposed. The Seine arenite unit is well-exposed at Shoal Lake where cross-bedded sandstone (Fig. 17c) provides stratigraphic polarity as well as supporting the common interpretation of a fluviatile origin. Cross-bedded sandstone (Fig. 17d) also can be traced farther eastward along the Seine River (Fig. 2) where it can be demonstrated to be overlain by an upper unit of coarse, polymictic conglomerate (Fig. 17e) and, in some cases, intercalated with it (Fig. 17f). The uppermost conglomerate unit is notable for an abundance of granitoid clasts and Davis et al. (1989) reported an age of 2696.1+5/-3 demonstrating that it was sourced in a granitoid body that was much younger than the Laurentian which provided detritus for the basal Seine Conglomerate. Davis (1990) further constrained the depositional age of the sandsized fraction from arenite at Horsecollar Junction (Fig. 2) by noting the presence of abundant detrital zircons with a U-Pb age of approximately 2693 Ma., effectively the same age as the Bear Pass pluton. This fact contradicted Lawson’s original contention that all of the Algoman intrusions could be defined on the basis of the fact that they are younger than the Seine (see below). Nonetheless, Lawson’s original interpretation of the Seine Series mainly as a product of Archean alluvial and fluvial sedimentary processes has been reinforced and elaborated upon by several authors (Ojakangas, 1972; Wood, 1980; Fralick and Davis, 1999; Czech and Fralick, 2002). Although his language was somewhat dense, the overall message of Lawson’s paleoenvironmental interpretation is paraphrased as follows: “it seems a fair inference that the conglomerate represents a gravelly flood plain… The distribution of the conglomerate … indicates the course of a river (following) the dominant structural lines … at a time which antedates the intense complication which folded - 101 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B C D E F Figure 17: a) basal Seine conglomerate inter-clast quartz grit derived from underlying tonalite, Golden Star arear; b) stratified basal conglomerate, Golden Star area; c) trough cross-bedded Seine arenite in plan view, Shoal Lake showing younging toward the top of the photo; d) deformed cross-bedded Seine arenite and pebble conglomerate in cross-section view, Seine River Bridge; e) deformed polymictic conglomerate in cross section view, east of Mine Centre; f) sandstone interbed in coarse upper Seine conglomerate in plan view west of Wild Potato Lake. Note the angle between bedding (arrows) and foliation. and deformed the conglomerate” (Lawson, 1913, p.62). In other words, he envisioned the location of Seine conglomerate and arenite to have been controlled by syn-sedimentary faults to account in part for it’s elongate map pattern (Fig. 2). Algoman Lawson’s 1914 map portrays five different varieties of intrusive rocks at Rainy Lake in decreasing order of perceived age which he classified with the term Algoman: basic facies of syenite, syenite gneiss, granite and granite gneiss, banded and streaked gneiss and porphyroid gneiss. Harris (1974) made similar distinctions which allowed the least deformed Algoman rocks to be discussed in terms of three distinct spatial and compositional suites: the Rocky Islet Bay complex, the Swell Bay intrusions and the large and conspicuous Ottertail Lake Intrusion (Fig. 3). The first of these are dominated by quartz monzonite syenite and mafic syenite and are commonly feldspar-phyric (Fig. 18a, - 102 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B C D E F Figure 18: a) feldspar-phyric quartz monzonite, Raspberry Island; b) porphyritic quartz monzonite, Rocky Islet Bay; c) granodiorite cut by vertical sheeted quatz-pyrite veins, Bear’s Passage; d) xenolithic monzodiorite, Ottertail Lake intrusion; e) Intrusion breccia with granitoid matrix, western Ottertail Lake intrusion; f) incipient brecciation and granitoid infilling of metamorphic tectonite, Ottertail Lake intrusion. b), The Swell Bay intrusions, exemplified by the Bear Pass Pluton (Fig. 18c) are composed mainly of quartz monzonite and granodiorite (Goldich and Peterman; 1980). Some of these intrusions are compositionally zoned with mafic to intermediate margins and felsic interiors (Cram, 1923; Harris, 1974). The Ottertail Lake intrusion is also compositionally zoned from marginal hornblende-biotite quartz monzonite to interior leucocratic quartz monzonite in the interior (Goldich and Peterman, 1980): wallrock xenoliths are common in the marginal phase (Fig. 18d) and internal magmatic breccias (Fig. 18e, f) are well developed in what Lawson interpreted to be roof pendants of deformed and metamorphosed Keewatin rocks. Goldich and Peterman (1980) demonstrated that the Algoman intrusive rocks commonly contain abundant K-feldspar and have much higher Sr contents than Laurentian tonalite and trondhjemite. The Ottertail Lake intrusion, also with high overall Sr content, displays a fractionation trend of increasing Rb:Sr - 103 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 ratio toward its interior. Shirey and Hanson (1984, 1986) and Stern et al. (1989) further defined specific lithogeochemical characteristics of the Algoman rocks at Rainy Lake to show that they are also distinctive from other granitoid rocks at a global scale. Relative to their intermediate silica content (55-60%), they contain anomalous Mg, Sr, Ba, Ni, Cr and are strongly enriched in light rare earth elements. Stern et al. (1989) proposed their formation from hydrous melting of mantle that had been enriched large ion lithophile elements though prior metasomatism. Davis (1990) provided an estimate of 2693 +/- 2 Ma age for the Bear Pass pluton and coupled with the 2686+2/-1 Ma age of the Ottertail Lake intrusion (Davis et al. 1989), demonstrated that the sanukitoid magmatism spanned the time bracket for inferred for deposition of the Seine conglomerate and arenite above a profound angular unconformity. A recent study by Bjorkman et al. (2024) has demonstrated the widespread distribution of the sanukitoid suite of rocks across the Wabigoon Subprovince, including the Ottertail Lake Intrusion. This has been interpreted to represent a significant shift in magmatism at approximately 2690 Ma that can be explained by metasomatism and magmatism in a suprasubduction setting leading to collisional deformation and metamorphism that is commonly attributable to the Kenoran Orogeny. Deformation and Metamorphism The emphasis on protoliths and stratigraphic relationships that has historically dominated the discussion of the geology does not outweigh the fact that most of the rocks are clearly metamorphic tectonites as well. Lawson (1913) recognized this and attributed commonly observed foliation and lineation (“pencilling” in his terminology) to compressive deformation related temporally to the Algoman granitoid suite. Rocks in the southeastern part of the wrench zone have been metamorphosed to greenschist facies mineral assemblages and rocks of the amphibolite facies are dominant in the northwest (Fig. 8). Significant areas of retrograde metamorphism have also been noted (Peterman et al. 1972; Poulsen, 1984) and this has been taken to be the explanation why most geochronological approaches have yielded unreliable protolith ages. It is likely that the overall distribution of preserved prograde assemblages is the result a combination of both local contact and regional dynamothermal metamorphism. The common existence of minor structures of dynamothermal metamorphic origin such as foliation (Fig. 12c, 17f, minor folds (Fig. 19a, b) and lineation (Fig. 10a) are reflections of local strain. Rheological contrasts within and among lithological units have also been well established to be important in controlling the local strain intensity in the Rainy Lake area, particularly in the Seine conglomerate (Hsu, 1971; Jackson, 1982; Czeck et al., 2009). The highest strains are also common in features which are arguably shear zones in which strong foliation is accompanied by asymmetric distribution of foliation (Fig. 19c, d, e, f) that mimics the overall structural pattern in the wrench zone as a whole (Fig. 8). Following the lead of Peter Hudleston (1986) in the Vermilion district of Minnesota, dynamic interpretations invoking dextral transpression have been invoked by several authors to explain the overall structural style of the Rainy Lake wrench zone (Poulsen, 1986b; Borradaile et al., 1988; Poulsen et al., 1992; Czeck and Hudleston, 2003; Fernandez et al., 2013). Beyond the local importance of dynamothermal metamorphic fabrics, however, the larger structural features in the wrench zone also of considerable interest. Foremost among these is the angular unconformity at the base of the Seine sedimentary sequence in the southeastern part of the zone (Fig. 20a) and it also provides an ideal temporal reference point for understanding the deformational and metamorphic history of the area. As illustrated above, the fact that lithic clasts in the basal conglomerate above the unconformity show no evidence of pre-depositional metamorphic fabrics yet clasts throughout the Seine have been variably strained during post-depositional dynamothermal metamorphism is an important one. It illustrates the insufficiency of using the development of foliation alone as a means of tracking a protracted structural history. A second notable structural aspect at Rainy Lake is the stratigraphic evidence for significant overturning of beds in the northwestern part of the zone (Fig. 20b). Poulsen (1980) suggested that this might have resulted from the overprinting of early recumbent folds by younger upright ones but, given the observation that the first-formed foliation in these rocks is also folded in the Rice Bay dome, the possibility of late-overturning of what may have been at one time steep strata can’t be entirely ruled out. A third topic of importance is the fact that, since their recognition in the 1930’s, there also has been a great deal of attention paid to the major faults that define the boundaries of - 104 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B C D E F Figure 19: a) shortening and transposition of felsic dikes cutting Coutchiching biotite schist, north of Noden Causeway; b) folded felsic sills, Great River Road; c) asymmetric boudinage of the interior of a mafic dike relative to its foliated margins, Noden Causeway; d)) asymmetric boudinage in felsic metavolcanic rocks south of the Olive gold mine, e) asymmetric shapes of clasts in Seine meta-conglomerate adjacent to the Rainy Lake – Seine River Fault south of Seine River Bridge; f) tight asymmetric folds in mylonite, Little Turtle Lake landing. the wrench zone. The rocks that now help to define the Quetico Fault at Rainy Lake were originally mapped by Lawson (1913) as part of a narrow belt of “porphyroid gneiss” extending westward from Little Turtle Lake at Mine Centre to Cheery Island. He recognized that the red porphyroid gneiss “has a pronounced cataclastic structure and that the schistosity of the rock is referable to deformation involving shearing of the mass” (Lawson, 1913, p.94). He stopped short of relating the rocks to a fault, however, interpreting them instead to represent the deformed southern margin of a large granitoid batholith: this is somewhat ironic because he is the geologist who, by this time, had named the San Andreas Fault and had compiled the definitive technical report on the Great San Franciso Earthquake of 1906. By the time F.R. Harris remapped the area, however, it had been recognized that the rocks here belong to the greater than 350 km long Quetico Fault based on the interpretation linears on air photo mosaics (Parkinson, 1962). Harris (1974) went on to describe the rocks in the fault as crushed granite, augen gneiss and mylonite and, like Lawson before him, locally showed gradational contacts with adjacent banded gneissic rocks which he termed migmatite. Kennedy (1984) studied 14 sites along the entire Quetico Fault, including 3 in the Rainy Lake wrench zone, and concluded that the mylonitic foliation on average resulted, not strictly from cataclastic processes, but from ductile flattening - 105 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 A B Figure 20: Schematic cross-sections through a) Rice Bay – Bear’s Passage and b) the Bad Vermilion – Shoal Lake areas. See Figure 8 for the locations of the sections (adapted from Davis et al., 1989). The Quetico fault is located at the northern end of both sections. based on measured axial ratios of deformed mineral aggregates and object-object strain estimates. She also used quartz c-axis fabric measurements and analysis of brittle micro-faults and ductile shear zones to argue for overall dextral displacement on the fault. Kennedy (1984) showed that the microfaults and minor shear zones dominantly strike NW and have dextral shear sense. She further argued that transition from ductile behaviour (mylonite) to brittle is consistent with the current level of exposure representing deformation at a crustal depth of 10-15 km. Borrradaile and Kennedy (1982) also showed evidence of flow-banding in veins of pseudotachylite at Crowrock Inlet as evidence of frictional melting in the fault zone. Peterman and Day (1989) reported a Rb-Sr isochron age of 1947+/23 Ma to suggest that the pseudotachylite from both the Quetico and Seine River faults resulted from Proterozoic reactivation of the Archean faults. Metallogeny A commonly understated geological feature of the Rainy Lake wrench zone is the simple abundance of mineral occurrences within it in comparison to the adjacent areas on either side. Poulsen (2000b) enumerated 88 of them in total and demonstrated that they include examples that are representative of multiple deposit types (Figs. 8, 21) which, in turn, are thought to relate to multiple geological processes. Syngenetic deposits include stratabound metalliferous sediments in the mafic sections of the Keewatin including banded iron formation, pyritic massive sulfide deposits with locally anomalous zinc sulfides (Nickel Lake and Pocket Pond). Numeous Zn-Cu occurrences (Port Arthur Copper, Lochart Lake, Wind Bay, Gagne Lake, Pidgeon) demonstrably possess the descriptive of volcanic-associated massive sulfide deposits in general. Basal Cu+/-Ni sulfide mineralization (North Rock) and magnetite+/ilmenite mineralization (Seine Bay, Mironsky) is clearly associated with the Grassy Portage and Bad Vermilion Lake layered gabbroic intrusions (Poulsen and Hodgson, 1984). Quartzpyrite-molybdenite veins show a spatial association with Algoman granitoid rocks and sheeted veins of this type within the Bear Pass Pluton are similar in style to those in the deeper parts of granitoid-related - 106 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 main geological features. The western starting point of the road log Km 0.0 (92.4) is at the lookout tower on the waterfront in Fort Frances and the eastern ending point Km 92.4 (0.0) is at the highway bridge across the Seine River near Crilly. The highway distances are approximate and, although the stops are described from west to east, they can be visited in any order depending on topical interests. Figure 21: Metallogenic Summary of the Rainy Lake Wrench Zone Phanerozoic vein and stockwork deposits. The goldbearing quartz veins in the Mine Centre area which were the focus of a gold rush in the 1890’s (Coleman, 1894; Winchell and Grant, 1895), are readily classified in modern terms as “orogenic” deposits characterized by ribbon quartz, carbonate-sericite alteration and spatial control by minor shear zones (Poulsen, 1986a). A recurring question about the metallogeny of the Rainy Lake wrench zone concerns the apparent absence of economically viable mineral deposits compared to the numerous occurrences. While it is true the there is strong similarity between the make-up of the rocks in the Rainy Lake wrench zone and the central volcanic complexes at Chibougamau, Val d’Or and Noranda in the Abitibi subrprovince, the discrepancy in metal endowment may simply be explained in the context of the geological deposit types. For example, the metal endowment of syngenetic massive sulfide systems is thought to be negatively influenced by shallow water environments, the lack of a well-defined lithocap or by cooler upwelling fluids and this might apply to Rainy Lake. A notable characteristic of the orogenic Auquartz veins at Mine Centre the kinematic evidence for strike-slip stress conditions for vein formation at Rainy Lake in contrast to conditions for reverse faulting allowing for higher fluid pressure at Val d’Or in the Abitibi Subprovince (Poulsen et al., 1992). Road Log and Field Stops A traverse which follows Highway 11 along the Rainy Lake wrench zone provides an opportunity to examine representative outcrops which illustrate its The lookout tower at Fort Frances is located on the north shore of the Rainy River near its outlet from Rainy Lake (Fig. 22). The rock exposures which Lawson (1887) originally chose as a type locality of the Archean metasedimentary biotite schist at the Coutchiching Rapids were flooded upon construction of the power dam to the west of here circa 1906. Since then, representative outcrops that illustrate the Coutchiching Group have been described nearby at Ranier, Minnesota by Ojakangas et al. (1982, Stop 1) and Jirsa and Hemstad (2010, Stop 6-2). Drive east along Front Street and join Highway 11 and continuel eastbound from Fort Frances. Lake Road intersects the highway at Km 1.9 (90.5). Continue through the land of the Couchiching First Nation past Couchiching Drive at Km 3.3 (89.1). Note the discrepancy between the modern spelling compared that of the geological unit which was based on the version used topographically circa 1887. Continue past the C.N.R. Railway Crossing (Km 5.5 (86.9)) and over the crest of the Noden Causeway bridge and continue past the intersection with a side road to the north marked “Scenic Lookout”. This sideroad (Km 8.0, 84.4) leads to stop 13 of Czeck and Poulsen (2010). Continue eastward on Highway 11 and turn in to the next (unmarked) sideroad (Km 8.8, 83.6) which leads northward to a parking area beneath the hydro tower. This is STOP NC (Noden Causeway). This is an instructive stop (Fig. 23) in that this is one of the many islands in Rainy that would have been mapped both topographically by triangulation by W.H.C. Smith and geologically by Andrew C. Lawson in the 1880’s. The rocks here consist mainly of foliated quartz monzonite of which Lawson first assigned to the Laurentian but later revised to the Algoman intrusive suite which he described as “mica syenite” belonging to a larger Pukamo Island intrusion (Lawson, 1913). Harris (1974) correlated these rocks with the Rocky Islet Bay Complex west of Rice Bay which are comprised mainly of felsic to intermediate granitoid rocks of variable composition. The main unit - 107 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 22: Simplified geology of the Fort Frances segment. Field stops NC – Noden Causeway; GA – George Armstrong Drive Figure 23: Noden Causeway stop (NC) - 108 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 here is cut by a variety of dikes which offer contrasts in structural competence compared to the surrounding granitoid rock. Note the s-shaped asymmetric foliation pattern in one of the mafic dikes (site 1) that mimics the regional structural pattern of the Rainy Lake Wrench Zone as a whole (19c). Return to Highway 11 to resume the road log. Km 11.8 (80.6) – George Armstrong Drive intersects highway 11 from the east; turn in and park near the mailboxes to examine STOP GA (George Armstrong). This is also stop 11 of Czeck and Poulsen (2010) and Point of Interest 27 described in Pye (1968). This is the first area of significant exposure of the Coutchiching rocks northeast of their type locality at Fort Frances. Although the nature of the contact with the Keewatin rocks is obscure (Fig. 24), it is still a good place to examine the differences between the metasedimentary biotite schists which are cut by felsic intrusive rocks (site 1) and the metavolcanic Figure 24: George Amstrong Drive stop (GA) amphibole-biotite schists (site 3). Both units are now metamorphic tectonites which exhibit moderate to high strain but the variability of layer thickness in the metasedimentary units is consistent with their inferred origin as submarine turbidites (Ojakangas et al., 1982). Further evidence for the superimposed strain is evident at (site 2) where at least four generations of dikes cut the metasedimentary rocks and display the variable effects of folding and boudinage depending on their structural competence and pre-strain orientation with respect to bedding (see also Czeck and Poulsen (2010) and Druguet et al. (2008). Continue eastward along Highway 11 past Commissioners Bay which is the location of a zircon sample from a Keewatin felsic which yielded a U-Pb age of approximately 2727 Ma (Davis et al., 1989). Km 18.4 (74.0) – Windy Point Bridge Km 21.1 (71.3) – outcrops on both sides of Highway 11. This is STOP SM (Sims) and corresponds in part to the Windy Point locality described by Pye (1968). The outcrops here) display deformed pillowed and variolitic metabasalt which is a dominant lithology within the Keewatin volcanic sequence on the flanks of the Rice Bay Dome (Fig. 25). It is important to examine the exposure (site 1, Fig. 26)) carefully in three dimensions because primary pillow shapes which are inherently variable are further distorted by superposition of a moderate amount of tectonic strain. This result is log-shaped pillows with long axes that plunge moderately westward (Fig, 10a). The effects of the strain can be further appreciated by examining the cm-scale light-coloured patches that stand out against the darker amphibolitic background of the metabasalt (especially at site 2). They are varioles which predictably would have formed originally as spherical patches due to devitrification of glassy volcanic rock but here their shapes reflect their tectonic distortion with a flat aspect corresponding to a foliation and a long axis which plunges westward in the foliation. Note also that the dark pillow selvedges offer rheological contrasts with the rest of the basaltic material so that the down-plunge elongation is also expressed in places in the outcrops by boudinage of individual pillows. Pye (1968) described these outcrops without reference to their volcanic origins at all while still emphasizing the lineation and the sets of joints perpendicular to it. Even where the pillows are clearly defined the considerable strain makes it difficult to draw satisfactory conclusions about primary stratification and directions of younging. Harris (1974) and Poulsen (1980) suggested, albeit with some doubt, that the stratigraphic section in this area faces downward and eastward. Km 24.9 (67.5) – The Nickle [sic] Lake Shores Road which intersects Highway 11 from the south leads to STOP NL (Nickel Lake). This was stop 1 of Poulsen (1982). This area illustrates the fact that, although the term Keewatin is synonymous with metavolcanic protoliths, it also contains clastic and chemical interflow sedimentary units which include oxide, sulfide, carbonate and silicate facies of iron-formation. These rocks are important from at structural point of view in that they typically have sharp magnetic and - 109 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 25: Simplified geology of the Swell Bay segment. Field Stops: SM-Sims; NL-Nickel Lake; MB – Moran’s Bay; GR-Great River Rd.; PP-Pocket Pond; BC- Belacoma; GP-Grassy Portage; BL- Bear’s Passage boat launch; BB- Bear’s Passage bridge; TB-Tunnel Bay Figure 26: Sims stop (SM). - 110 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 electromagnetic geophysical responses which aids in the definition of their position in areas of poor exposure. The rock here (Fig. 27, site 1) is typically referred to as chert-magnetite, banded iron formation (BIF) and is a lithology that is commonly folded at all scales. At Nickel Lake the iron-formation defines a structural synform (historically the Nickel Lake Syncline) which plunges shallowly westward along axes coincident with those of the minor folds and with the axes of maximum elongation in the adjacent volcanic rocks. The curved traces of folds observed here in a downplunge view was originally interpreted by Poulsen (1980) to present a type 3 (coaxial) fold interference pattern. It is equally possible, however, that they result from a single deformation with a strong westward plunging linear component of strain (i.e. L-s tectonite). potentially represent coeval subvolcanic intrusions or younger sills Algoman which are responsible for the cross-cutting relationships. Both lithofacies display prominent polycrystalline quartz aggregates which are likely deformed phenocrysts which help define both the tectonic foliation and a prominent lineation which plunges shallowly eastward at this locality (Fig. 15f). Figure 28: Moran’s Bay stop (MB) Km 29.0 (63.4) intersection between Highway 11 and Highway 502 (Fig. 25). This is STOP GR (Great River Rd.) which corresponds to Stop 10 of Czeck and Poulsen (2010). Figure 27: Nickel Lake stop (NL). Km 26.9 (65.5) – outcrops on both sides of highway 11 but a particularly large one on the south side. This is STOP MB (Moran’s Bay) and is described as Stop D.1 in Poulsen and Wood (1982). The outcrop is located on the south limb of the prominent antiformal Rice Bay Dome (25). It provides ample illustration of the rocks Lawson (1914) mapped as Laurentian granite and granite gneiss in the interior of the dome (Fig. 28). Both Lawson (1913) and Harris (1974) interpreted the unit to be at least in part intrusive into the mantling Coutchiching metasedimentary rocks but the details remain in considerable doubt. R.W. Ojakangas was the first to suggest that the wispy banded, quartz-phyric, grey, foliated quartzofeldspathic can also be interpreted as a deformed rhyolite. This unit yielded a U-PB zircon age of 2725+/-2 Ma (Davis et al., 1989). It is cut by more competent sheets of coarser quartz-feldspar porphyry (Fig. 15e) which could Folded quartz-phyric intrusions on the north side of Highway 11 west of the intersection (site 1, Fig. 29) cut amphibole-biotite schists containing local ironformation which were included with the Coutchiching biotite schist on the maps of Lawson (1914) and Harris (1974) but which show greater similarity to Keewatin units elsewhere. The porphyritic felsic intrusions have been generally included in the suite of Laurentian intrusions but the molybdenite-bearing quartz veins exposed here are also a characteristic of Algoman intrusions elsewhere. Despite these uncertainties of interpretation and the somewhat transitional nature of the contacts, it is clear these rocks serve to separate the inner core of the Rice Bay dome from a structurally higher annular band of moderately southeastwarddipping Couchiching biotite schists which are well exposed approximately east of the intersection (site 2). It is also possible to make a short side-trip form this intersection northward along highway 502 for 2.2 km to its intersection with the Baseline Bay side road which enters from the east. This is STOP PP (Pocket Pond) and corresponds to Stop D.2 of Poulsen and Wood (1982). - 111 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Km 30.2 (62.2) – bush road and outcrops on north side of Highway 11. This is STOP BC (Belacoma) corresponding to stop 3 of Poulsen (1982, stop D.3 of Poulsen and Wood (1982) and stop 5 of Hinz (2010). Figure 29: Great River Road stop (GR) The critical outcrops (Fig. 30, site 1) that demonstrate the overturned stratigraphic section on the northern limb of the Rice Bay dome are now heavily overgrown and no longer instructive. A good sense of the nature of the northeastward-dipping contact between the Coutchiching metapelites and the distinctive green, magnetic ultramafic unit which here represents the Keewatin volcanic rocks can still be observed along Highway 502 (site 2). Continuity of the lithostratigraphic units and their moderate northeasterly dips in this area were established with the assistance of ground magnetic and electromagnetic surveys and by diamond drilling which targeted Cu-Zn mineralization associated with the interflow iron-formation units in the section. Although the contacts among the units are sharp and well defined there is no conclusive evidence for them to be erosional-depositional in origin but the evidence for an overturned volcanic sequence is sound (Fig. 30). This is a continuation of the Coutchiching-Keewatin contact which extends southward from Pocket Pond and westward to Nickel Lake and sharply defines the eastern closure of the Rice Bay Dome. The volcaniclastic ferropicrite unit here is exposed over a wider area than at Pocket Pond and the full nature of the contact is uncertain. The ultramafic rocks near the contact with the structurally underlying Coutchiching biotite schists (Fig. 31, site 1) are foliated as but appear to be progressively less deformed eastward (sites 2 and 3). Nonetheless, graded bedding of reasonable quality suggests the Coutchichiing strata are overturned in support of the observations at Pocket Pond. The cluster of outcrops near the beaver pond (site 3) have been documented by Schaefer and Morton (1991), Goldstein and Franceis (2008) and Hinz (2010) and the inference is that this unit is composed of relatively rare mantlederived ultramafic coherent and pyroclastic rocks that locally contain well-preserved accretionary lapilli (Fig. 10f). Return to Highway 11 and resume the road log. Figure 31: Belacoma stop (BC) Km 31.3 (61.1) C.N.R. overpass Km 31.9 (60.5) – numerous outcrops on both sides of Highway 11; safe parking is available beneath the powerline on the west side of the highway (Fig. 32). This is Stop GP (Grassy Portage) and corresponds to Stop D.4 of Poulsen and Wood (1982). Figure 30: Pocket Pond stop (PP) The gabbroic rocks exposed here are part of the metamorphosed Grassy Portage layered mafic intrusion and include plagioclase-rich leucogabbro (site 1) that gives way southward to garnet-bearing quartz- 112 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 ferrodiorite (site 2). The garnets are metamorphic porphyroblasts that likely crystallized owing to the favourable bulk composition of the diorite which has a higher Fe/Mg ratio and silica content than the leucogabbro. Lawson’s 1914 map of the area portrayed the leucogabbro as “hornblende gabbro” alone as an intrusion within the Keewatin while including the gabbro and melagabbro to the north and the garnetbearing quartz diorite to the south as Keewatin metavolcanic rocks. This inferred symmetry led to his interpretation of a synclinal axis centred on the leucogabbro but Harris (I974), Poulsen (1980) and Poulsen and Hodgson (1986) recognized all three lithofacies as distinctively different phases of a single layered mafic intrusion that shows progressive southward, upward in a stratigraphic sense, chemical and mineralogical fractionation. contact (site 3) is consistent with southward younging in the meta-turbidites and contradicts the structural order of the rocks based on dip alone. It is, however, consistent with the southward younging implied by the fractionation within the Grassy Portage layered mafic intrusion. Return to Highway 11 to resume the road log Figure 33: Bear’s Passage Boat Launch stop (BL) Km 36.4 (56.0) – Taylor’s Road intersects Highway 11 from the north Km 37.0 (55.4) – parking area and scenic view on South side of the highway (Fig. 34). This is STOP BB (Bear’s Passage Bridge) corresponding to Stop 7 of Poulsen (1982) and Point of Interest 2 of Pye (1968). Figure 32: Grassy Portage stop (GP) Km 33.6 (58.8) - the side road on the east side of Highway 11 leads to the boat launch at Bear’s Passage where parking is available at the lakeside (Fig. 33). This is STOP BL (Bear’s Passage Boat Launch) corresponding to Stop D.5 of Poulsen and Wood (1982) and locality 20 of Uglow (1913). This critical area of outcrop illustrates one of the most contentious points of the Seine-Coutchiching problem. The Keewatin rocks which are cut locally by a foliated lamprophyre dike structurally overlie gabbroic rocks of the Grassy Portage layered intrusion (site 1.) The Coutchiching rocks are staurolite-bearing biotite schists (site 2) and locally display evidence of primary graded bedding with is enhanced by the distribution of porphyroclasts in upper parts of individual beds. Graded bedding which can be observed directly adjacent to the relatively sharp Keewatin-Coutchiching The eastward dipping Coutchiching biotite schist, as exposed on the north side of the highway (site 2). is cut by granodiorite of the Bear Pass Pluton which contains sheeted quartz-pyrite-molybdenite veins which are exposed on both sides of the bridge (sites 1 and 3) The view southward from the lookout features Swell Bay and the belt of Keewatin volcanic rocks to the south of it. The Keewatin-Coutchiching contact is located on Morton Island to the southwest. Km 37.7 (54.7) – Bear Pass Road intersects Highway 11 from the north. From this location it is possible to make a side trip to STOP TB (Tunnel Bay) by driving northward for 1.3 km to the C.N.R. tracks and taking the first dirt road uphill to an exposure of Coutchiching metasedimentary rocks (Fig. 35). This area is near Tunnel Bay and localities 5 and 6 of Uglow (1913). These outcrops are located on the eastern limb of the antiformal culmination in the Bear’s Passage area. The demonstration of the existence of the antiform - 113 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 OS (Old Station Road) where the field relationships are comparable to those at Morton Island (Stop D.6 of Poulsen and Wood (1982)). Figure 34: Bear’s Passage bridge stop (BB) was essential to Lawson’s (1913) interpretation of the Coutchiching strata in the interior of this structure. It is also the location where D.W. Davis first demonstrated the effects of zircon inheritance from the Coutchiching metasedimentary rocks by felsic dikes related to the Bear Pass Pluton. One of the dikes near the stop of a steep outcrop can be viewed to the east at (site 2). An additional point of interest in these exposures (sites 1 and3) is that the main foliation is locally crenulated by a steep, northwest striking, transecting cleavage (S3 of Poulsen, 1980), which is particularly prominent in a 2 km-wide northwesterly trending corridor through this area. Although locally dominant at the mesoscopic and mircroscopic scales, where crenulation of the main biotite-rich foliation and rotation of metamorphic porphyroblasts are both evident, the effects of this deformation at the macroscopic scale are negligible. Return to Highway 11 to resume the road log Km 40.8 (51.6) – Old Station Road intersects Highway 11 from the north (Fig. 36). This is STOP Figure 35: Tunnel Bay stop (TB) Highway 11 at this locality (Fig. 37) is approximately parallel to the strike of stratification in the Coutchiching meta-sdedimentary rocks as well as to their mapped contact with Keewatin meta-volcanic rocks (Harris, 1974). The overall dip of bedding is steep to the southeast and in places a steep cleavage with a more northerly strike transects bedding to form a moderately eastward plunging intersection lineations. Exposure is plentiful but the clearest features of the Coutchiching beds are illustrated in flat outcrops on the south side of Highway 11 (site 1). The rocks here are metamorphosed to greenschist facies assemblages and primary features are reasonably well preserved: polarity in graded beds consistently indicate a northward direction of younging which is away from Keewatin volcanic rocks which are exposed at the shore of Rainy Lake south of here. Km 43.3 (47.9)) – Ottertail Landing Road intersects Highway 11 from the north. Km 46.2 (46.2) – a side road to a communications tower intersects Highway 11 from the north: turn in and park (Fig. 38). This is STOP OW (Ottertail West). The field relationships exposed in the outcrops east of the intersection on the north side of Highway 11 are comparable to those at stop 1 of Czeck and Poulsen (2010) which is located approximately 1 km to the west. This is an area in which a roof pendant composed of foliated metavolcanic and metasedimentary schists has been variably incorporated into granitoid rocks of the Ottertail Lake intrusion. The outcrops here provide a rare case where highway improvement has also resulted in outcrop improvement. A marginal phase of the Ottertail Lake intrusion (site 1) is composed of diorite containing abundant mafic xenoliths (Fig. 18d). Magmatic breccias (Fig. 18e) are well exposed along the highway to the east (site 2) and, at one location nearby, a narrow NNE-striking mylonitic zone cuts the intrusive rocks. The most critical point made by Lawson (1913) and most observers since is that there is abundant visual evidence for intrusion of felsic magma into previously foliated metamorphic tectonites. Km 47.5 (44.9) – the outcrop on the north side of the road was sampled by D.W. Davis to yield a U-Pb zircon age of 2686+/-3 Ma for this part of the Ottertail Lake intrusion. - 114 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 36: Simplified geology of the Ottertail Lake segment. Field Stops: OS- Old Station Rd; OW: Ottertail Lake West; OE-Ottertail East Figure 37: Old Station Road stop (OS) - 115 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Turtle River Road. Km 58.8 (33.6) Patten Park picnic area Km 62.0 (30.4) – Low outcrops are present on both sides of highway (Fig. 40) and a rusty waste dump is visible across a marshy area on the north side of Highway 11. This is STOP PA (Port Arthur Copper) which was stop 9 of Poulsen (1982) and Point of Interest 22 in Pye (1968). Figure 38: Ottertail West stop (OW) Km 48.2 (46.2) – Pearson’s Road intersects Highway 11 from the north Km 53.9 (38.5) – the outcrops of Ottertail Lake intrusive rocks exposed here are described as Stop 2 in Czeck and Poulsen (2010). Km 56.1 (36.3) – outcrops on both sides of the road expose the eastern margin of the Ottertail Lake Intrusion. This is STOP OE (Ottertail East) and corresponds in part to Stop D.7 of Poulsen and Wood (1982). The easternmost outcrop on the north side of Highway 11 (site 1, Fig. 39) exposes deformed spherulitic and flow-banded rhyolite that is common in the Keewatin volcanic section in this part of the belt. It is cut by granitoid phases or the Ottertail Lake Intrusion, including a distinctive feldspar-phryic variety containing xenoliths (site 2). The abundance of xenoliths decreases westward in these outcrops (site 3). Km 56.4 (36.0) intersection of Highway 11 and Figure 39: Ottertail East stop (OE) Access the rusty area from the west side of the water and cross a small Beaver Dam to reach the large area of exposure (Fig. 41). The main mineralized lithology is composed of foliated amygdaloidal andesite (Fig. 10c) containing disseminated and semi-massive lenses of pyrite, chalcopyrite and sphalerite (site 1). Stratified, rusty felsic volcanic rocks are exposed on the north side of the outcrop area (site 2). This is but one of several occurrences of syngenetic sulfide deposits hosted by the felsic portions of the Keewatin volcanic section extending more than 25 km southwestward beyond Wind Bay. It is also noteworthy that the base metal deposits are located up-section northward from the syn-volcanic Bad Vermilion Lake intrusive complex (Fig. 40). Km 63.4 (29.0 side road intersects Highway 11 from the north Km 67.0 (25.5) -the Mine Centre Road intersects Highway 11. This road can be followed north to Little Turtle Lake by travelling for 1.0 km to Government Road and continuing .5 km to the C.N.R. tracks. Bear right at the intersection with Queen St. and follow the dirt road to the public boat launch site. The outcrops near the shoreline constitute STOP LT (Little Turtle Landing) which corresponds to Stop D.9 of Poulsen and Wood, 1982). Lawson (1913) mapped the rocks that are exposed here as a distinctive lithological unit which he described as “porphyroid gneiss”. In doing so, he effectively defined a 60 km E-W segment of what is now known as the Quetico Fault without explicit reference to faults but certainly recognized the overall significance of the rock type in “that it has a pronounced cataclastic structure and that the schistosity of the rock is referable to deformation involving shearing of the mass” (Lawson, 1913. P.94). Today the lithologies which he described are regarded as variably deformed fault rocks which include protomylonite (site 1) which is exposed in the outcrop east of the parking area and mylonite (site 2) along the shore of Little Turtle Lake. - 116 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 40: Simpified geology of the Mine Centre segment. Field Trip Stops: PA-Port Arthur Copper; LT-Little Turtle landing; FG: Ferguson; GS-Golden Star; WC-Windy City Rd. Regrettably, a recently constructed dock partially obscures the best exposure of the folded mylonite (Fig. 19f) as described in Poulsen and Wood (1982). Km 68.2 (24.2) The Shoal Lake Road meets Highway 11 from the south (Fig. 40). This road leads to what is arguably the most significant geological feature in the entire belt – the angular unconformity at the base of the Seine Group metasedimentary rocks. Follow the (in places rough) Shoal Lake public road southward for 3.3 km to a point where it is met from the east by a recently constructed but as yet uncompleted sideroad. The is STOP FG (Ferguson) and outcrops in this area Figure 41: Port Arthur Copper stop (PA) Figure 42: Little Turtle Landing stop (LT) Return to Highway 11 and resume the road log - 117 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 offer a good view of the Bad Vermilion tonalite which Lawson assigned to his Laurentian suite of granitoid rocks. The recent excavation (site 1) has exposed the tonalite and small quartz veins with adjacent sericiteankerite alteration of the style exposed at the Ferguson gold prospect to the north (site 2). Lawson (1913) demonstrated conclusively that the tonalite cuts both the Bad Vermillion gabbro-anorthosite to the west and Keewatin volcanic rocks to the north which include moderately northward dipping interflow chertcarbonate units and a northward younging unit of pillow basalt. conglomerate (Fig. 17a, b) it is also clear that even the least competent lithic clasts possessed no tectonic fabric at the time of deposition across strata with a preexisting steep dip. Return northward to Highway 11 and continue eastward along it. Km 76.5 (15.9) an unmarked bush road meets Turn around and proceed back northward along the Shoal Lake Road for 2.2 km to a small rise with a low outcrop on the east side; pull to the right side of the road and park as safely as possible (Fig. 44). This is Figure 44: Golden Star stop (GS) Figure 43: Ferguson stop (FG) STOP GS (Golden Star) and the site of Stop D.10 of Poulsen and Wood (1982) and the contact described by Uglow (1913) as being marked by “brown flags” for the International Geological Congress Field Trip led by Lawson. The field relationships here have also been described and discussed more recently as Stop 1 of Czeck and Fralick (2020). The base of the Seine Group here dips gently eastward at high angle to stratification in the Keewatin rocks. Much of the outcrop (site 1) is now grown over but five small patches have been recently cleaned to clearly show the west to east transition from quartzbearing tonalite a), tonalite sand with rare clasts (b) to angular conglomerate (c) with interstitial sand (fanglomerate of Lawson) to polymictic pebble and cobble conglomerate (d, e). Although there is evidence of a weak tectonic foliation superimposed on the Highway 11 on the south side; pull in and park. This is STOP WC (Windy City road). The increasingly overgrown leads southward from here for approximately 500 metres to a sign which explains how a windstorm in 1988 flattened trees over a seven km2 area resulting in its nickname of “Windy City”. Reclamation of the area resulted in local removal of shallow glacial overburden to produce two-dimensional pavement exposures of cobble to boulder conglomerate which show the rheological effects of superimposed strain. These outcrops comprise the “Forest Tour” Stop 5 of Czeck and Poulsen (2010)) can be reached by continuing another 250 m southward beyond the sign and following the second sideroad to the southwest (approximate UTM NAD 83 Zone15 N: 536 800E, 5 398 500N). The outcrops exposed at the intersection along highway at its intersection with the Forest Tour Road, however, make for a good and easily accessible substitute stop. The outcrops occur along both sides of the highway and serve to illustrate three important aspects of Seine Group as a whole. The first is the distinction between the two main lithofacies: polymictic clast-supported conglomerate (site 1) versus thick-bedded, locally cross-bedded, arenaceous sandstone (site 2) which occupies the middle part of the Seine stratigraphic - 118 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 consistent with that of rocks which are regarded to be part of the Algoman suite (Davis et al., 1989). Km 78.4 (14.0) the Manion Lake Road meets Highway 11 from the north (Fig. 46). Km 82.2 (10.8) Horsecollar Junction – the road to the south leads to the Seine River village and the outcrops the deformed conglomeratic facies of the Seine Group on the north side of the highway constitute Stop 5 of Czeck and Fralick (2002). Km 92.1 (0.3) the Crilly Road meets Highway 11 from the north. Figure 45: Windy City Road stop. section – most evidence suggests that the strata young southward toward the polymictic conglomerate units. Second, a good three-dimensional view of the shape fabrics shows both elements of both foliation and eastward plunging lineation as well as the rheological differences in response to the bulk strain by clasts of different original composition and grain size. Third, a population of granitoid clasts is particularly noticeable in this part of the Seine stratigraphic section and these were commonly assumed to have been sourced in the Laurentian granitoid suite. A sample from this area (site 3) was collected and analysed by D.W. Davis to demonstrate that the age of a granitoid clast was actually Km 92.4 (0.0) Highway bridge across the Seine River (Fig. 47). This is STOP SR (Seine River Bridge) and is also described as Stop D12 of Poulsen and Wood (1982) and Stop 5 of Czeck and Fralick (2002). The outcrop southeast of the bridge provides an excellent visual representation in cross-section of the mixed arenite-conglomerate facies of the Seine Series. The overprinting steep foliation corresponds to pronounced shape fabrics in clasts at high angle to bedding in pebble conglomerate and the comparable shortening across the foliation is manifested by steepening and distortion of the foresets in the crossbedded sandstone units. The beds dip shallowly northward and this also corresponds to the inferred Figure 46: Seine River segment - 119 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Geological Magazine, v. 67, p. 77-92. Bailey, E.B., 1927, Across Canada with Princeton; Nature, v.120, p.673-675. Bass, M. N., 1961, Regional tectonics of part of the southern Canadian Shield; Journal of Geology, v. 69, p. 668702. Bauer, R.L., Czeck, D.M., Hudleston, P.J., Tikoff, B., 2011, Structural geology of the subprovince boundaries in the Archean Superior Province of northern Minnesota and adjacent Ontario; 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. 203–241. Figure 47: Seine River Bridge stop. direction of stratigraphic younging. Although not formally defined as a type locality for the Seine Series, the outcrops here are arguably a good reference locality. References Adams, F.D., Bell, R., Lane, A.C., Leith, C.K., Miller, W.G. and Van Hise, C.R., 1905, Report of the special committee for the Lake Superior Region; Journal of Geology, v.13, p.89-104. Alderman, D.J., 1988, Rice Bay geological mapping and lithogeochemical sampling; unpublished geological report, Falconbridge Limited, Winnipeg, 51 p. Alt, D.D., 1958, A review of the geology of Rainy Lake; unpublished M.S. thesis, University of Minnesota, 73p. Ashwal, L.D., Morrison, D.A., Phinney, W.C. and Wood, J., 1983, Origin of Archean anorthosites: evidence from the Bad Vermilion Lake anorthosite complex, Ontario; Contributions to Mineralogy and Petrology, v. 82, p. 259-273. Ayer, J.A. and Davis, D.W., 1997, Neoarchean 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, v. 81, p. 155-178. Ayres, L.D., 1970, Synthesis of Early Precmbrian stratigraphy north of Lake Superior; Institute on Lake Superior Geology Proceedings, Thunder Bay, Ontario, Abstract, p. 8. Ayres, L.D., 1969, Early Precambrian Stratigraphy of Part of Lake Superior Province Park, Ontario, Canada and its Implications for the Origin of the Superior Province; unpublished Ph. D. thesis, Princeton University. Bailey, E.B., 1930, New light on sedimentation and tectonics; Beakhouse, G.P., 1984, Reconnaissance investigations of granitoid and medium to high grade metasedimentary terrains: volcanic components and mineral potential; in Summary of Field Work, 1984; Ontario Geological Survey Miscellaneous Paper 119, p. 14-18. Bjorkman, K.E., Lu, Y., McCuaig, T.C., Kemp, A.I., Hollings, P., 2024, Linked evolution and in situ growth of the Wabigoon superterrane, Superior Craton: evidence from zircon U-Pb isotopes and whole-rock geochemistry. Precambrian. Research, v. 404, 107341. Blackburn, C.E., 1980, Towards a mobilist tectonic model for part of the Archean of Northwestern Ontario; Geoscience Canada, v. 7, p. 64-72. Blackburn, C.E., 1973, Geology of the Otukamamoan Lake area, Districts of Rainy River and Kenora; Ontario Division of Mines, Geological Report 109, 42p. Accompanied by Map 2266, scale 1 inch to 1 mile. 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, edited by L.D. Ayres, P.C. Thurston, K.D. Card and W. Weber; Geological Association of Canada, Special Paper 28. p. 89-119. Borradaile, G.J., 1982, Comparison of Archean structural styles in two belts of the Canadian Superior Province; Precambrian Research, v. 19, p. 179-189. Borradaile, G.J. and Kennedy, M.C., 1982, Pseudotachylite; in Atlas of Deformation and Metamorphic Rock Fabrics, edited by G.J. Borradaile, M.B. Bayly C. McA. Powell; Springer Verlag, Berlin, Heidelberg, New York, p. 366-367. Borradaile, G.J., and Poulsen, K.H., 1981, Tectonic deformation of pillow lava; Tectonophysics, v. 79, p. T17-T26. Borradaile G., Sarvas, P., Dutka, R., Stewart, R., and Stubley, - 120 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 M., 1988, Transpression in slates along the margin of an Archean gneiss belt, northern Ontario - magnetic fabrics and petrofabrics; Canadian Journal of Earth Sciences, v.25 , p. 1069-1077. Bruce, E.L., 1925, The Coutchiching rocks of the Bear’s Pass section, Rainy Lake; Royal Society of Canada Transactions, Section IV, v.19, p. 42-46. Bruce, E.L., 1927, Coutchiching delta; Geological Society of America Bulletin, v.39, p.771-782. Carreras, J., Czeck, D. M., Druguet, E., Hudleston, P. J., 2010, Structure and development of an anastomosing network of ductile shear zones; Journal of Structural Geology, v. 32, p. 656-666. Coleman, A. P., 1898, Clastic Huronian rocks of western Ontario; Geological Society of America Bulletin, v. 9, p. 223-238. Coleman, A.P., 1894, Gold in Ontario: it’s associated rocks and minerals; in Report of the Bureau of Mines, Ontario, v.4, p. 35-100. (printed 1985) Cooke, H.C., 1926, Between the Archean and Keweenawan is the Huronian; Royal Society of Canada Proceedings, series 3, v.19, sec. 4, p. 17-37. Cram, I.H., 1932, The Rest Island granite of Minnesota and Ontario; Journal of Geology, v.40, p.270-278. Czeck, D. M. and Hudleston, P. J., 2003. Testing models for obliquely plunging lineations in transpression: a natural example and theoretical discussion. Journal of Structural Geology 25, 959-982. Czeck, D. M., Maes. S. M., Sturm, C. L., and Fein, E. M., 2006. Assessment of the relationship between emplacement of the Algoman plutons and regional deformation in the Rainy Lake region, Ontario. Canadian Journal of Earth Sciences 43, 1653-1671. Czeck, D. M. and Poulsen, K. H., 2010, Deformation in the Rainy Lake Region: A Fabulous Display of Structures Controlled by Rheological Contrasts; Field Trip 3 Guide, 56th Annual Institute on Lake Superior Geology, International Falls, MN, v. 56, Part 2, p. 47-75. Czeck, D. M. and Fralick, P., 2002, Structure and Sedimentology of the Seine Conglomerate, Mine Centre Area, Ontario; Field Trip Guide 3, 48th Annual Institute on Lake Superior Geology, Kenora, Ontario, v. 48, Part 2, p. 37-68. 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, v. 31, 1365-1376. Davies, J.C., 1973, Geology of the Fort Frances area, Rainy River District; Ontario Division of Mines, Geological Report, v. 107, 35p. Accompanied by Map 2263, scale 1 inch to 1 mile. Davis, D.W., 2023, U–Pb geochronology: its development and importance in Canada; Canadian Journal of Earth Sciences. V. 60, p. 388-400. 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; 36th Annual Institute on Lake Superior Geology, Thunder Bay, Ontario, Proceedings and Abstracts, v. 36, pt. 1, p.20-21. 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., Poulsen, K.H. and Kamo, S.L., 1989, N e w insights into Archean crustal development from geochronology in the Rainy Lake Area, Superior Province, Canada; Journal of Geology, v. 97, p.379398. Diamond, L.W. and Marshall, D.D., 1990, Evaluation of the fluid inclusion crushing-stage as an aid in exploration for mesothermal gold-quartz deposits: Journal of Geochemical Exploration, v. 38, p. 285-297. Dewey, J.F. and Ryan, P.D., 2022, Discussion of Searle, ‘Tectonic evolution of the Caledonian orogeny in Scotland: a review based on the timing of magmatism, metamorphism and deformation’; Geological Magazine, v.159, p.1833-1836. Dewey, J. F., Dalziel, I.W.D, Reavy, R.J. and Strachan, R.A., 2015, The Neoproterozoic to Mid-Devonian evolution of Scotland: a review and unresolved issues; Scottish Journal of Geology, v. 51, p.5-30. Dott, R.H. Jr., 2001, Wisconsin roots of the modern revolution in structural geology; Geological Society of America Bulletin, v. 113, p. 996-1009. Druguet, E., Czeck, D.M., Carreras, J. and Castano, L.M., 2008, Emplacement and deformation features of syntectonic leucocratic veins from Rainy Lake zone (Western Superior Province, Canada); Precambrian Research, v. 163, p.384-400. Fernández, C., Czeck, D. M., and Díaz-Azpiroz, M., 2013, Testing the model of oblique transpression with oblique extrusion in two natural cases: steps and consequences; Journal of Structural Geology, v. 54, p. 85-102. Fralick, P. and Davis, D.W., 1999, The Seine-Coutchiching problem revisited: sedimentology, geochronology and geochemistry; in Western Superior Transect 5th Annual Workshop, edited by R.M. Harrap and H.H. Helmstaedt; Lithoprobe Secretariat, University of British Columbia, Lithoprobe Report 70, p. 66-75. Frye, J.K., 1959, Petrography of the ancient granites of the - 121 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Minnesota-Ontario boundary region, unpublished M.S. thesis, University Fumerton, S.L., 1985, Geology of the Calm Lake area, District of Rainy River; Ontario Geological Survey Report 226, 72p., accompanied by Map 2467, scale inch to ½ mile. Fumerton, S.L., 1982, Redefinition of the Quetico Fault near Atikokan, Ontario; Canadian Journal of Earth Sciences; v. 19, p. 222-224. Goldich, S.S., 1968, Geochronology of the Lake Superior region; Canadian Journal of Earth Sciences, v.5, p. 715-724. Goldich, S.S. and Fischer, L.B., 1986, Air abrasion experiments in U-Pb dating of zircon; Chemical Geology, v. 59, p. 195-215. Goldich, S.S. and Peterman, Z.E., 1980, Geology and geochemistry of the Rainy Lake Area; Precambrian Research, v. 11, p. 307-327. Goldstein, S.B. and Francis, D., 2008. The petrogenesis and mantle source of Archaean ferropicrites from the western Superior Province, Ontario, Canada; Journal of Petrology, v.49, p.1729-1753. Goodwin, A.M., 1977, Archean basin-craton complexes and the growth of Precambrian Shields; Canadian Journal of Earth Sciences, v. 14, p. 2737-2759. Grout, F.F., 1925, Coutchiching problem, Geological Society of America Bulletin, v. 36, p. 351-364; including discussion by A.C. Lawson and reply. Grout, F.F., Gruner, J.W., Schwartz, G.M. and Thiel, G.A., Precambrian Stratigraphy of Minnesota, Geological Society of America Bulletin, v.62, p. 1017-1078. Harris, F.R., 1974, 1974, Geology of the Rainy Lake area, District of Rainy River, Ontario Division of Mines, Geological Report 115, 94p. accompanied by Maps 2278 and 2279, scale inch to ½ mile. 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. Hawley, J.E., 1930, “Seine” or “Coutchiching”? Journal of Geology, v.38, p. 521-547. Discussion by J.E. Gill and Reply,1931, Journal of Geology, v. 39, p.655669). Hinz, P., White, C.R., Albers, P.B. and Tortosa, D., Mineral Deposits of the Mine Centre – Rainy River Area, Field Trip 7 Guide, 56th Annual Institute on Lake Superior Geology, International Falls, MN, v. 56, Part 2, p. 91-125. Hooper, P.R. and Ojakangas, R.W., 1971, Multiple deformation in Archean rocks of the Vermilion district, northeastern Minnesota, Canadian Journal Earth Sciences, v. 8, p. 423-434. Hsu, M.-Y., 1971. Analysis of strain, shape and orientation of the deformed pebbles in the Seine River area, Ontario; unpublished Ph.D. thesis, McMaster University, 167p. Hudleston, P.J., Schulz-Ela, Bauer, R.L. and Southwick, D.L., 1986, Transpression as the main deformational event in an Archean greenstone belt, Northeastern Minnesota: in Tectonic Evolution of Greenstone Belts, edited by M.J. deWit and L.D. 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G., Keay, J.P. and Panagapko, D., 1980b, Mine Centre Area (Eastern Half), District of Rainy River; Ontario Geological Survey, Preliminary Map P.2202, Geological Series, scale 1: 15 840 or 1 inch to ¼ mile, Geology 1976, 1977. Wu, T., Polat, A., Frei, R., Fryer, B.J., Yang, K.-G. and Kusky, K., 2016, Geochemistry, Nd, Pb and Sr isotope systematics and U-Pb zircon ages of the Neoarchean Bad Vermilion Lake greenstone belt and spatially associated granitic rocks, western Superior Province, Canada; Precambrian Research, v. 282, p. 21-51. Zaslow, M., 1975, Reading The Rocks: The story of the Geological Survey of Canada 1842-1972; MacMillan Company of Canada, Limited; 599p. Zhou, S., Polat, Ali, Longstaffe, F., Yang, K.G., Fryer, B.J. and Weisener, C., 2016, Formation of the Neoarchean Bad Vermillion Lake Anorthosite Complex and Spatially Associated Granitic Rocks at a Convergent Plate Margin, Superior Province, Western Ontario, - 124 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Canada; Earth Sciences Publications, v. 10 (https:// ir.lib.uwo.ca/earthpub/10) - 125 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Trip 6 - Amethyst Deposits of Thunder Bay Stephen Kissin Department of Geology, Lakehead University, Thunder Bay, Ontario, P7B 5E1, Canada and Greg Paju Resident Geologist Program, Ontario Geological Survey, Ministry of Energy and Mines, Thunder Bay, Ontario, P7E 6S7, Canada Canada Introduction Properties of Amethyst Amethyst, occurring in abundance in the Thunder Bay region, is purple gemstone variety of α-quartz. It has been known for some time that an iron impurity in quartz is the underlying source of amethyst coloration (Holden, 1925). However, incorporation of iron of alone cannot account for the formation of amethyst, as many varieties of quartz contain trace amounts of iron, yet amethyst is relatively rare, and large deposits of amethyst are very rare. interstitial sites. The color of amethyst is produced by absorptions of light in the visible region of the spectrum owing to the presence of Fe4+, as originally shown by Cox (1977). The proposed mechanism requires the coincidence of four geological conditions for the formation of amethyst: (1) The incorporation of Fe and Al, as well as Na or Li. This is not a limiting condition, as the small concentrations of these trace elements are readily available in hydrothermal solutions. (2) A source of ionizing radiation, either from U and Th or 40K in order to produce the defects in Fe and Al. (3) Deposition at generally rather shallow depth such that oxidizing conditions prevail and iron is in the form of Fe3+. (4) Deposition with a temperature range for the stability of Fe4+, the source of amethyst coloration. In a series of papers by Cohen and coworkers, culminating in a summary in Cohen (1989), a simultaneous sequence of reactions was proposed for the formation of amethyst. (1) (Al–O)- → (Al–O)° + eIonizing radiation forms a hole center from oxidizing the substitutional Al-O bond. (2) Na+ + e- → Na° Electron from step 1 is trapped by an interstitial alkali metal ion. (3) Fe3+int → Fe4+int + eInduced ionizing radiation forms a trapped hole center via oxidizing the interstitial Fe3+. (4) (Al–O)°+e- → (Al–O)Trapped hole center is satiated as [AlO°] is reduced via gaining the electron from step 3. The presence of iron is positions interstitial with respect to the SiO4 framework was established by Adekeye and Cohen (1986), in noting its correlation with pervasive Brazil law twinning in colored sectors of amethyst crystals. Data on incorporation of the alkalis Na, K and Li and trivalent Al and Fe in quartz were reported by Deer et al. (1963), who further noted that the incorporation of Al3+ (and presumably Fe3+), is compensated by the incorporation of Na+ or Li+ The mechanism proposed above is consistent with observed data and provides a logical mechanism for the formation of amethyst. However, Rossman (1994) noted that there are unestablished factors in the model such that its acceptance is tentative. Crystal forms expressed in amethyst are invariably simple, consisting only of combined positive {101} and negative {011} rhombohedra. The faces of one of the forms are generally largely and are designated as the major rhombohedron r, and the other form is designated as the minor rhombohedron z (Fig. 1). The only other form occasionally observed is the ditrigonal prism m (Frondel, 1962). Amethystine coloration is unevenly distributed in the crystal, generally with concentration in the major rhombohedral forms, in which Brazil law twins are also concentrated (Fig. 2; Frondel, 1962). The orientation of Brazil law twins in Figure 2, is typical - 126 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 2. Etched basal α-quartz illustrating the typical occurrence of Brazil law twinning in which alternate bands contain left- and right-handed α-quartz (after Frondel, 1962). Figure 1. A typical amethyst crystal viewed perpendicular to the c-axis, illustrating the combination of positive {101 ̅1} and negative {011 ̅1} rhombohedra. of their occurrence in α-quartz; however, in amethyst the twins are polysynthetic with a typical width of 0.1 mm (McLaren and Pitkethly 1982). The twin plane of the Brazil law is {101}, which separates right-handed and left-handed orientations of quartz. McLaren and Pitkethly (1982) demonstrated that the composition plane of the Brazil law twin provides space for incorporation of Fe3+ and that iron is preferentially concentrated along this composition plane in amethyst. Amethyst’s Name and Colour Origins The word amethyst has its origins from the ancient Greek word amethystos which may be translated as “not drunken”, from the Greek a-, “not” + methustos, “intoxicated”, as the gemstone was believed to prevent or lessen the effects of drinking alcohol. There is a common theme regarding the mythological origin of amethyst’s purple colouration. Bacchus (Dionysus to the Romans); the Greek god of winemaking, orchards, fruit, vegetation, fertility, festivity, insanity, ritual madness, religious ecstasy, and theatre, pursuing a maiden named Amethyste, who was refusing his affections. Amethyste prayed to the gods to remain chaste, a prayer answered by the goddess Artemis (Diana to the Romans), who transformed her into a white stone. Bacchus humbled by Amethyste’s desire to remain chaste, poured wine over the stone as an offering, dyeing the crystals purple. In another variation the god was insulted by a mortal, and vowing to slay the next mortal who crossed his path in retaliation created fierce tigers to carry out his wrath. The hapless mortal a young woman, Amethystos, was on her way to the shrine of the goddess Diana, when the tigers fell upon her. Her life was spared by the goddess, but the price was being transformed into a statue of pure quartz. Seeing what his anger had done, a remorseful Dionysus was so moved that tears of wine poured from his eyes onto Amethystos, staining her stature purple. Despite the belief in this origin story, there are no ancient texts supporting the myth, as compared to the ancient period that supposed birthed this story, it’s quite recent as it was written in 1569 by the French Renaissance poet Rémi Belleau (1528–1577), in the poem “L’Amethyste, ou les Amours de Bacchus et d’Amethyste” (Amethyst or the loves of Bacchus and Amethyste; Belleau, 1576). Amethyst Deposits in the Thunder Bay Area In his summary of the history of amethyst in the Thunder Bay area, Patterson (1985) reported that as early as 1642, Radisson described the use of - 127 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 “torquoise” as a gemstone by local indigenous peoples. Amethyst was an associated mineral in most of the lead-zinc and silver mines, and attracted the interest of a few prospectors. In the early 1860s, the McEachern brothers prospected for amethyst in the Thunder and Black Bay areas. In 1862, they mined two tons of amethyst crystals, which they barged to Toronto to sell in that city. About the same time, a shipment of amethyst from the Thunder Bay area was sold in Montreal. The success of the mineral as a valued item for sale even in an unprocessed state and the ease of mining encouraged other prospectors and developers to try searching for and producing amethyst (Garland 1994).In the 1880s, amethyst was mined northeast of Thunder Bay in a place now called Amethyst Harbour. Interest in Thunder Bay amethyst declined around the turn of the century with the development of deposits of high quality and inexpensive amethyst from Brazil. The deposit that became known as the Amethyst Mine Panorama was originally discovered in 1935. When the fire tower was built in the 1950s, near Elbow Lake in McTavish Township, the large amethyst veins were uncovered by the roadbuilders. In the early 1960s, the area was staked, and trenches exposed the veins in what is now the open pit for the mine. In large vugs near the surface of the vein deposit, amethyst crystal of spectacular size were obtained. The development of the deposit with wide-spread sales and distribution of specimens revitalized interest in amethyst in the region (Sinkankas, 1976; Garland, 1994). The interest and activity in amethyst deposits in the Thunder Bay area led to the proclamation in 1975 designating amethyst as Ontario’s provincial gemstone (Patterson 1985). A comprehensive report on amethyst deposits and mining activity in the Thunder Bay area was completed by Garland (1994). The interest and activity in amethyst deposits in the Thunder Bay area led to the Mineral Emblem Act in 1975 designating amethyst as Ontario’s provincial Mineral Emblem (Ontario, 1990; Patterson, 1985), with the 50th anniversary taking place in 2025. There are currently 15 amethyst quarries authorized to produce under the Ontario Ministry of Nature Resources Aggregate Resources Act within two areas northeast of Thunder Bay (Campbell et al. 2024). Twelve of these authorized amethyst extraction sites are in McTavish Township and are accessible from Highway 11-17. The other three authorized quarries are located in the Tartan Lake Area (north of MacGregor Township) in an area that is accessed via the Magone Lake Road from Highway 527. Four quarries operate as tourist attractions that are open to the public on a seasonal basis. A listing of these amethyst quarries, including information about their products and services (where available), is provided in Table 1 (Campbell et al., 2024). Geology Of Amethyst Mine Panorama (Thunder Bay Amethyst Mine) Geologic Setting The geological setting of the mine is complex, as an Archean and a Proterozoic record are preserved in the area. This record has been recently reviewed by Sutcliffe (1991) with an update by Addison et al. (2010) and will not be repeated in detail here. The Amethyst Mine Panorama (Thunder Bay Amethyst Mine) is hosted in the Archean Hilma Lake granite of McCrank et al. (1981). This pluton lies on the boundary of the Quetico Subprovince and the Wawa Subprovince, with typical greenstone lithologies on its southern margin and gneissic metasedimentary rocks on the northern margin. The Hilma Lake granite in the vicinity of the mine consists predominantly of monzonite, with compositional variation along the trend monzonitequartz monzonite-granite-granodiorite and pegmatite and pegmatitic textural variants (Jennings, 1985). Jennings’ study indicates that monzonite had been cut first by granodiorite, then by pegmatite, with some metasomatic alteration of early monzonite toward granodioritic composition. At the Greenwich Lake uranium occurrence, a vein-type occurrence located 10 km to the northwest, Franklin (1978) noted the presence of quartz monzonitic pegmatites containing 60-100 ppm U in the form of uraninite. As these pegmatites are apparently comagmatic with the Hilma Lake granite, its uranium-rich character is likely a general feature. The Proterozoic rocks were deposited on the eroded Archean surface; however, the Animike Group (Gunflint and Rove Formations) is missing in the vicinity of the amethyst mine. As indicated by Franklin et al. (1980), the Mesoproterozoic Sibley Group progressively onlaps Archean terrain in a northerly direction. The Sibley Group is presently absent in the vicinity of the Amethyst Mine Panorama, although its presence as abundant fragments in mineralized breccias within the vein system indicates that these sediments - 128 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Table 1. Amethyst quarries in the Thunder Bay Area authorized to produce under the Aggregate Resources Act (from Campbell et al. 2024) Deposit Name and Ownership Amethyst Mine Panorama Precious Purple Gemstones Ltd. Location (Licence) McTavish Township (622921) Products and Services Tourist attraction (pick-your-own and mine tours), specimens, decorative and landscaping stone, and tumbling stone, jewellery, giftware, carvings, faceted gemstones www.amethystmine.com/ Blue Points Amethyst Mine Jordan Vivian McTavish Township (624926) Tourist attraction (pick-your-own), specimens, decorative stone, aquarium stone www.tripadvisor.ca/Attraction_Reviewg155017-d3334892- ReviewsThe_Blue_Point_Amethyst_MineThunder_Bay_Thunder_Bay_District_Ontario.ht ml (lynswan@lakenet.com – email) Diamond Willow Amethyst Mine Big Pearl, Sward Lake B. Leroux and C. Fayle McTavish Township 3 permitted quarries, (626151, 625922, 626134) Tourist attraction (pick-your-own and mine tours), specimens, decorative and landscaping stone, slabs, tumbling stone, jewellery and giftware www.diamondwillowamethyst.com/ Keetch Quarry / Boulder Creek Amethyst Quarry L. Harasym McTavish Township (77956) Tourist attraction (pick-your-own), specimens https://mininglifeonline.net/company_page_487.html Assiniboia Amethyst Mine P. and T. Smitham McTavish Township (626091) Not open to the public, but may be visited by invitation only. Contact: https://assiniboiaamethystmine.weebly.com/ Bill’s Old Amethyst Mine K. Zytaruk McTavish Township (607322) Not advertised Canadian Shield Amethyst Mine K. Zytaruk McTavish Township (616261) Not advertised Tartan Lake Danbill Mine Auralite 23 Mine and Company Area (20227) Inc. Specimens, polished and tumbled stone, jewellery, tiles and countertop stone www.auralite23canada.com/home.html Gunnard Project M. Noyes and J.A. Gavin McTavish Township (625989) Not advertised Loon Lake Technical Services Quarry Loon Lake Technical Services McTavish Township (625067) Not advertised Tartan Lake Area Purple Haze Mine Auralite 23 Mine and Company (624879) Inc. Specimens, giftware, jewellery, decorative and landscaping stone from former owners at www.purplehazeamethyst.com/. Transferred to new ownership in late 2022. Roll Lake Amethyst Tartan Lake Area Auralite 23 Mine and Company (624838) Inc. McTavish Windy Ridge Amethyst L. Kowtuski Township (625831) Specimens, polished and tumbled stone, jewellery, tiles and countertop stone www.auralite23canada.com/home.html - 129 - Email: windyridge@live.ca �Proceedings of the 72nd ILSG Annual Meeting - Part 2 were present as basement cover during the forming of the deposit. The significance of the Sibley Group is unclear in the face of contradictory evidence concerning its age and depositional setting. Franklin et al. (1980) noted that the Sibley Group isdeposited at the location of a failed arm of an r-r-r triple junction, although they admitted to uncertainty as to the contemporaneity of sedimentation and rifting. Although some features of the Sibley Group are suggestive of a rift-filling deposit, the whole-rock Rb/Sr age of 1339±33 Ma (Franklin, 1978b) is approximately 200 Ma prior to the main stage of rifting of the Midcontinent (Keweenawan) Rift (Van Schmus et al., 1982). Cheadle (1986), however, concluded on the basis of sedimentological studies that the Sibley Group was not deposited in a classical aulocogen, but represents a deposit on a sagging crust preceding rifting. The Sibley Group was more recently dated by U/Pb geochronology in zircons in a basal rhyolite unit at 1537+10/-2 Ma (Davis and Sutcliffe, 1985). This timing makes a relationship with the Midcontinent Rift event unlikely, and Hollings et al. (2004) proposed that the Sibley Basin formed due to effects of a plume track that created an infracratonic basin. by breccias of granitic country rock and Sibley Group sedimentary rocks with large proportions of void space. The brecciated fault was subsequently mineralized by hydrothermal solutions. At least two periods of mineralization occurred, as an early generation of amethyst was clearly brecciated and subsequently coated by a second generation of amethyst. Figure 4 is an illustration of the state of the mine in 1987. At present, the main pit configuration is basically the same but has been deepened. In that year, an extension of the vein system to the east was developed, offset to the north by a few metres strike-slip fault. Jennings (1985) subdivided the mineralization patterns into three basic types: (i) open fracture fillings, (ii) breccias with tectonic and collapse subtypes, and (iii) “honeycomb” veins. Other deposits located at or near the Sibley -Archean unconformity include the Dorion lead -zinc -barite veins (Fig. 3). The ore-depositing solution was considered to be a basinal, connate brine by Franklin and Mitchell (1977), an interpretation supported by the fluid-inclusion studies of Haynes (1988). As illustrated in Figure 3, there is a close spatial relationship between the lead-zinc-barite veins and the amethyst, and both are spatially related to the Sibley-Archean unconformity. Geological features of the mine Amethyst Mine Panorama is located within a firstorder strike-slip fault, which strikes at 90- 100º and dips steeply to the south. This fault is roughly parallel to one 2.1 krn to the south, which strikes east-northeasterly (McIlwaine, 1971) and has a vertical displacement of at least 125 m (Jennings, 1985), forming a major boundary to the Sibley Group’s depositional basin. The strike-slip fault hosting the amethyst deposit is offset by seven first-order strike-slip faults, five of which are illustrated in Figure 4, which strike 162 - 150° and dip vertically producing en echelon, pull-apart structures in the main fault. These structures are filled Figure 3. Local geology and location map of amethyst deposits and lead-zinc-barite deposits, showing the relationship of the former to the margin of the Sibley Group outcrop and the Hilma Lake granite. The location of the producing Thunder Bay Amethyst Mine’s are indicated by stars. Bedrock geology and mineral occurrence locations modified from Ontario Geological Survey (2011; 2026). - 130 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 4. Diagram of the main pit of Amethyst Mine Panorama (Thunder Bay Amethyst Mine). The strike directions of these veins are strongly clustered in two groups, one slightly west of north and parallel to the second stage of strike-slip faulting, and one easterly, parallel to the principal directions of faulting. Open fracture fillings are common in the shallower zones of the deposit where low lithostatic pressure permitted the maintenance of open fissures following faulting. The veins are widest near the edges of collapsed breccias and at the intersections of oblique shears with the main fault zone. Fracture-fill mineralization occurred at the crystal-fluid interface as quartz crystals grew outward from the fracture walls. The crystals formed as parallel to radial growths with long crystal axes oriented perpendicular or subperpendicular to the growth surface. The crystal size invariably increases outward, and outward growth from opposite fractures resulted in an interlocking comb structure of euhedrally terminated crystals. This vein type may also contain vugs up to 2-3 m in diameter with large quartz crystals up to 10-15 cm in prism diameter. Tectonic breccias are here attributed to fault movement, as opposed to brecciation caused by collapse with variable degrees of fluid action. Some breccia fragments are surrounded only by a later portion of the paragenetic sequence, suggesting that multiple fault motion during the mineralizing event has occurred. Breccia fragments of this type are invariably angular and may consist of fragments of earlier deposited vein material, which may have been thermally bleached. Collapse brecciation is not always differentiated from tectonic brecciation, and some collapse breccias have undergone subsequent tectonic brecciation and vice versa. Evidence of collapse brecciation is seen in the occurrence of Sibley Group lithologies not present in the mine area now, together with granite and diabase as breccia fragments. Sibley Group fragments are particularly abundant within channel- or pipe-like structures in which fluid transport and abrasion have produced subangular to subrounded fragments, which have undergone an appreciable degree of sorting. Collapse-breccia fragments are typically coated with successive layers of chalcedony, colorless quartz, and amethyst, producing a cockade structure. Vugs have developed in open space produced in the breccia in which crystals with prism diameters of up to 10 cm have grown. Honeycomb veins are the result of quartz crystallization that has occurred in all directions from small nuclei, usually chalcedony, hematite, or silicified granite fragments, rather than from a fracture wall. The amethyst and quartz are more massive than in the other types of veins, but the growth is chaotic. - 131 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Mineralogy Amethyst and other varieties of quartz. Several varieties of quartz occur in Amethyst Mine Panorama, including colorless quartz; chalcedony; amethyst; the yellowish variety, citrine; and the greenish variety, prasiolite or “greened amethyst”. Smoky quartz has very limited development . The only variety of gemstone interest is amethyst, although the occurrence of the other varieties has aided in establishing the sequence of deposition. A grading system based on estimated intensity of coloration and clarity of specimens is in use at the mine, and this system has also aided in establishing the paragenetic sequence. Thus, the intensity of coloration may be from I (lightest) to IV (darkest) and clarity from a (clear) to f (opaque). Table 2 lists typical paragenetic sequences in an older sequence, which is present as breccia fragments in a younger sequence presently occupying the veins. The prasiolite in stages 4 and 5 of the older sequence appears to be thermally bleached amethyst on the basis of both its appearance and experimental evidence that heat-treated amethyst can be transformed to prasiolite (Lehmann and Bambauer, 1973). In the younger sequence, late-stage variations are noted, particularly as cappings to stage 5. A distinctive variety called “black gem”, a dark, brownish-black amethyst, is apparently characteristic of larger crystals grown in vugs in which iron-enriched, late-stage fluids were trapped. These frequently have final growth zone that contains abundant hematite inclusions, such that recent sales of such material has been called “Thunder Bay red”. It was this material, recovered in the early development of the deposit that led to the notorious statement by Sinkankas (1976, p. 204): “By far most of the amethyst is unsuited for lapidary purposes, with very little being free from flaws and hence useless for faceted gems or even baroques.” Figure 5 illustrating cut and faceted gemstone demonstrates the error in Sinkankas’ statement. The compositions of specimens of amethyst by neutron activation analysis for selected trace elements (Table 3) revealed the presence of subequal concentrations of Fe and Al. As well, low concentrations of Ge were sought based on absorption spectra that indicated its presence. The low Ti concentrations are perhaps related to the spotty occurrence of rutile needles in the amethyst, needles occurring when concentrations are relatively greater. Figure 5. Cut and faceted smoky quartz (top left) and amethyst from Amethyst Mine Panorama (Thunder Bay Amethyst Mine). Photo by S. Kissin. Table 2. Paragenetic sequences observed in the veins of Amethyst Mine Panorama Notes: Variations observed include (i) late-stage greenish and yellowish-amethyst; (ii) late-stage smoky quartz; (iii) discontinuous hematitic and milky quartz capping to crystal terminations; and (iv) development of black gem in crystals, deposited in vugs. - 132 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Table 3. Analyses of r-zones of amethyst for selected trace elements (in ppm; Kissin, 1997) Sample No. Fe Al Ge Ti DZS1* 217 447 0.5 n.d. LZS2 102 393 0.5 0.01 BSS3 273 369 1.0 n.d. TPS4 368 249 0.5 n.d. *DZS1 evidently contains solid inclusions, as high concentrations (in ppm) were noted; e.g. Ta 0.329, W 0.38, Eu 0.349, Sr 89.43, Zr 1.02, Nb 0.13, Ba 2985.26, La 3.85, Ce 0.35, U 0.387. All samples contain small, but detectable quantities of Co, Ni, Ga, Rb, Nb, Zr, Mo, Sn, Sb, Cs, La, Pr, Nd and U. Other non-sulfide minerals. Barite is rare in the veins at the Amethyst Mine Panorama, although it is abundant in other amethyst mines of the district, where it follows the final stage of quartz deposition. It was not observed in the course of the present study, but has been noted in the mine. Calcite is fairly common in thin, monomineralic veins, but was not observed within the amethyst-bearing veins. The genetic link between the calcite veins and amethyst veins, if any, is unclear. Hematite is abundant as minute- (< 0.1 mm diam.) solid inclusions in stage 5 of amethyst deposition and occurs sporadically at other stages of deposition as well. Hematite occasionally occurs as a daughter mineral in fluid inclusions, particularly in stage 5 of crystallization. Rutile occurs as needles that transect the growth zones of the quartz in scattered locations within the mine. The orientations of the needles are apparently random; however, the possibility of crystallographically controlled growth directions has not been considered in detail. Native copper occurs in association with copper and copperiron sulfides. Sulfides The common base-metal sulfides pyrite, chalcopyrite, galena, and sphalerite occur in small amounts throughout the vein succession and as veinlets and replacement bodies in altered granitic wall rock. Copper -iron sulfides, however, are predominant, and a sequence of the minerals cuprite-native copperchalcocite-covellite associated with hematite and pyrite was documented by McArthur et al. (1993). The copper -iron sulfides exhibit typical replacement textures (atoll structures, core-and-rim relationships) in occurrences both in amethyst growth stages and in wall rock. The assemblages bornite+pyrite and chalcopyrite+pyrite and chalcopyrite+pyrite occur in wall rock only; however, spatial relations of wallrock sulfides to the veins do not reveal any pattern, owing in part to their scarcity. Malachite is present as a supergene product derived from these hypogene copper minerals. Wall-rock alteration mineralogy. Hematitization, chloritization, and kaolinitization are prominent in envelopes surrounding the veins within zones of brecciated granite; however, the alteration extends only a few centimetres into the granites outside of the zone of brecciation. Intense hematitization occurs fairly generally in altered rock nearest the amethyst veins. The strongly hematitized zone is generally only a few centimetres thick, but weaker hematitization is notable throughout the altered zone. Outward from the hematized zone is an irregular zone of highly chloritized rock ranging from a few to a few tens of centimetres thick. Sometimes associated with the chloritization is diffuse epidotization, which produced a pistachio green tint over zones up to a metre wide. Kaolinitization is widespread and pervasive through the breccia zone, imparting a chalky, white appearance to relict feldspars. Other clay minerals, e.g., montmorillonite and illite, may also be present; however, they have not been sought in a detailed examination. The pervasive kaolinitization has allowed weathering to penetrate into the brecciated zone, resulting in a soft and loosely aggregated matrix in which the near-surface exposures of the amethyst are contained. The nature of this matrix has enabled a good deal of the amethyst to be mined with a minimum of blasting. The hematite-chlorite-epidote alteration assemblages in the presence of ubiquitous quartz are characteristic of the propylitic alteration typical in many hydrothermal ore deposits. The kaolinite and other clay minerals are characteristic of the argillic - 133 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 alteration of hydrothermal ore deposits. The two alteration types are analogous at least in their relative timing to early peripheral propylitic alteration, which is overprinted by argillic alteration stemming from downward-infiltrating meteoric water. Genesis of the deposit The genesis of the deposits of the Amethyst Mine Panorama were discussed in detail by McArthur et al. (1993) in the light of evidence obtained in their study. The conclusions of their study are given below; however, for details of the evidence, their paper should be consulted. Genetic speculations on the Amethyst Mine Panorama are hampered at the outset by questions as to the timing of amethyst deposition, as discussed in the Introduction. The spatial and geochemical affinities of the amethyst deposits with the Dorion lead-zinc-barite veins and the relationships of both to the depositional margin of the Sibley Group sediments suggest that all three are interrelated. Franklin and Mitchell (1977) proposed that the lead -zinc -barite veins formed when, during diagenesis and settling of the Sibley Group sediments, metal-bearing brines were formed when expelled connate waters mobilized metals from the Sibley Group sediments and(or) weathered granitic basement rocks below the Archean-Proterozoic unconformity. The solutions thus formed would have hypothetically migrated through the basal Pass Lake Formation aquifer to escape at basin-marginal faults. Precipitation of sulfide, carried in chloride- and sulfatebearing solution, occurred because of mixing of the relatively oxidized solution with H2S gas trapped at the Pass Lake Formation pinch-out. The amethyst deposits seem to be a variant of the lead-zinc-barite type of deposit in which the temperature was lower than that of the sulfide-rich lead-zinc-barite veins. The initially oxidizing to later reducing character of the solution is similar to that proposed for the lead-zinc-barite veins, but the relation to Pass Lake Formation pinch-outs is not present in most amethyst deposits. Rather, the amethyst deposits are generally hosted in granitic basement often with no Sibley Group sediments present. The amethyst deposits are richer in dissolved silica, having gained this component through the kaolinitization of feldspar during hydrothermal alteration of granitic country rock. As the amethyst deposits formed near the present or former unconformity with the Sibley Group, local reduction of the solution would have tended to occur as H2S was released during thermal breakdown of organic matter in the sediments. The quantity of sulfides precipitated would have been limited not only by the relatively small amount of H2S produced but also by the lower metal content of the solutions as compared with those depositing the Dorion lead-zinc-barite veins. The latter characteristic is inferred by a comparison of the results of this study with those of Haynes (1988) on the Dorion lead-zinc-barite veins. He found that fluid inclusions from these deposits are NaC1-CaC12H2O type on the basis of microthermometry and direct analysis of decrepitates. However, the fluid inclusions depositing sulfides are significantly more saline than those at the Amethyst Mine Panorama in that they contain daughter salts. The more saline and higher temperature (105-203°C) fluid inclusions indicate that solutions that they represent would have had a better metal carrying capacity as chloride complexes. The similarity of the solution components to those at Amethyst Mine Panorama lends support to the idea that the same event formed both types of deposits. The solutions depositing amethyst would have been cooler and less saline variants of those that formed the lead-zinc-barite veins. If the two types of deposit are genetically linked, both suffer from the problem of lack of knowledge of the timing of ore deposition. The maximum age of both is 1339 Ma, the whole rock Rb/Sr age of the Sibley Group (Franklin, 1978b), as both types of veins cut Sibley Group rocks and contain breccia fragments of them. Franklin and Mitchell (1977) did not suggest a specific timing for formation of the Dorion lead-zinc-barite veins; however, their suggested mechanisms for creation of the deposit favor a timing soon after the deposition of the Sibley Group sediments. The expulsion of pore water called upon would presumably occur during late diagenesis. However, as there is no evidence to suggest that the Sibley Group sediments have ever been deeply buried, the source of heat is a problem. If the timing of deposition were close to the formation of the Sibley depositional basin, it is possible that a thermal anomaly, perhaps augmented by seismic pumping, in the lower crust was responsible for both phenomena. Haynes (1988) suggested that the Dorion leadzinc-barite veins formed either in the environment of Keweenawan rifting or later, possibly in the Paleozoic. There is no geological evidence for activity in the Paleozoic in the western Lake Superior region, and the style of mineralization associated with Keweenawan - 134 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 events is different (silver deposits associated in part with Ni-Co arsenides; Franklin et al. 1986). Our preferred hypothesis is that the lead-zinc-barite veins and amethyst veins are associated with the timing of formation of and deposition in the Sibley basin. We, therefore, believe that these deposits are distinct from silver deposits in the Thunder Bay area and formed at a somewhat earlier time. The timing is, however, not at all certain. An attempt was made to directly date amethyst deposition by U/Pb age determinations on rutile needle inclusions in amethyst (Heaman and Easton, 2006). An age of 887±40 Ma with 68.6% discordance was determined on a very small sample with low uranium content. The authors indicated that these results should be viewed with caution as a large lead correction was needed. This age does not coincide with any known geological events in the area. As well, an attempt was made to date the cross-cutting diabase dikes, but was unable to yield any results. Summary Field and laboratory studies of the Amethyst Mine Panorama reveal the following: (3) Sulfide minerals including pyrite, chalcopyrite, galena, and sphalerite accompany amethyst deposition as small mineral inclusions and occur, as well, as veinlets and replacement bodies in altered granitic wall rock. Copper and copper-iron sulfides are most abundant and, together with native copper and cuprite, Eh-pH relationships indicate that the solutions forming the deposit were initially rather oxidizing and weakly acidic. In the course of crystallization, the solution became more reducing and slightly more acidic. (4) Fluid-inclusion studies indicate that in the younger sequence of quartz deposition, homogenization temperatures range from 146.5 to 114.7°C (mean 132.1°C) as contrasted with 91.2-40.9°C (mean 68.4°C) for amethyst. Eutectic temperatures of frozen inclusions indicate that the solution was of the NaClCaCl-H2O system, with possible concentration of an additional halide salt component in late-stage fluids. Few inclusions contain daughter minerals, and those found are hematite and sphalerite in late-stage fluids. Final melting temperatures indicate a trend of decreasing salinity in later growth stages. (5) Oxygen isotopic determinations on quartz indicate a range of δ180 outside that of juvenile (1) The vein system hosting amethyst deposits was formed by mineralization of an east-west-striking, waters and end-member basinal brines. Progressive steeply dipping strike-slip fault, opened into en mixing of basinal brine with local meteoric water echelon pull-apart structures by a series of later strike- is suggested. slip faults, also dipping steeply and intersecting the (6) Sulfur isotopic analyses of pyrite yield δ34S first-formed fault at high angles. Much open space of -0.4 to 0.6 ‰ and -1.4 ‰ in chalcopyrite. These with brecciated and vuggy textures resulted. Breccia volumes are consistent with derivation from H S 2 fragments include granitic host rock and Sibley Group gas liberated by thermal action protection on sedimentary rocks, implying that the latter were present organic material involving iron. The values are as a thin cover at the time of mineralization, although similar to those of the sulfur contained in sulfides they are erosionally removed from the mine area at in the Dorion lead-zinc-barite veins. present. At least one early generation of amethyst is included as breccia fragments, indicating that fault movement continued during mineralization. (2) At least two phases of amethyst crystallization separated by a period of brecciation are present. The older sequence contains five stages of quartz growth, the latter two of which were originally amethyst, but were thermally bleached to prasiolite by the influx of hot solutions that deposited the younger sequence of quartz. The younger sequence contains five and occasionally six stages of deposition, beginning with a stage of chalcedony and a stage of colorless quartz, followed by amethyst. Both sequences of deposition are traceable throughout the mine. (7) The presence Sibley breccia fragments cemented by quartz indicates that the veins cannot be older than 1339 Ma, the Rb/Sr age of the unit. However, a younger limit cannot be established at present. (8) On grounds of similarity in geological setting, proximity, composition of the ore-depositing solution, and sulfur isotopic composition, the amethyst veins are believed to be genetically related to the Dorion lead-zinc-barite veins. Both are believed to have been formed by solutions expelled and mobilized during diagenesis and compaction of the Sibley Group. The lead-zinc-barite veins formed in fractures at or near the margin of the Sibley depositional basin from solutions that were both hotter and more saline than those - 135 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 depositing amethyst. Amethyst-depositing solutions travelled longer distances in granitic basement, dissolving silica by alteration of feldspar. Although the amethyst-depositing solutions probably carried less metal as chloride complexes than did the solutions forming the lead-zinc-barite veins, less H2S at the site of deposition was probably the most significant factor causing a low sulfide content in the amethyst veins. (9) The temperature conditions under which amethyst forms appear to have a high temperature limit; at the Amethyst Mine Panorama this limit is no higher than approximately 115°C and may be as low as approximately 90°C. Temperatures as high as approximately 145°C but possibly as low as 115°C may be sufficient to thermally bleach earlier generations of amethyst in the influx of hot solutions. However, this theory of thermal bleaching has been recently criticized by Herbert and Rossman (2008), who attributed the development of greenish-grey to greenish quartz to the presence of H2O in the crystal. Our work (Klarner and Kissin, 201l) confirms the presence of water in IR absorption spectra; however, the water is largely contained in fluid inclusions, which are abundant and of secondary origin. Use of the highly focus beam of an FTIR microscope has shown that molecular water is of low and nearly identical concentration in both amethyst and “greened amethyst”. Experiments by Goetz (2014) demonstrated that heating at 250ºC for extended periods did not result in bleaching of amethyst, disproving that the 145ºC temperature caused bleaching of amethyst. This problem is unresolved at present. Geology of the Amethyst Mine Diamond Willow Unlike the years of extensive research undertaken at Amethyst Mine Panorama, the other known amethyst deposits in the Thunder Bay region are not well studied and with most information available being from Garland (1994) The original Diamond Willow Amethyst Mine was staked by Gunnard Noyes, in the 1960s with the mine initially operating in the 1970s. Following Gunnard’s passing in 1988, the mining leases were split into two parcels per inheritance and turning the original mine into the current Diamond Willow and Blue Points Amethyst Mines, owned by a son and daughter, respectively. The Diamond Willow Amethyst Mine operated until 2007 and was subsequently closed until 2015 (Garland, 1994). The currently producing Blue Points Amethyst Mine is the eastern extension of the original Diamond Willow mine and operates two and three-quarters of the four pits situated along the breccia zone (Fig. 6). The centre pit is the original and the largest, almost 60 m long and 4 m deep. A fence divides the pit between the two mines. The current Diamond Willow Amethyst Mine; the western extension of the original namesake mine site, operates one and one quarter of the four pits situated along this breccia zone (Fig. 7; Garland, 1994). This mineralized and well developed breccia zone occupies a vertically dipping fault zone, trending approximately 090° and extends for almost a kilometre. The fault separates Sibley Group conglomerates of the Pass Lake Formation from Sibley Group mudstone of the Rossport Formation (Garland, 1994). The current Diamond Willow Amethyst Mine is located at the western end of this fault/breccia zone, separating the Rossport Formation mudstones on the south from the Pass Lake Formation conglomerates on the north side. Both the mudstone and the conglomerate are well-layered, giving them a blocky appearance (Garland, 1994). The breccia zone varies from 1 to 5 m wide, and is characterized by a quartz­rich core and fragments of wall-rock material. In general, the fragment density increases away from the core, but is always matrix supported, the fragments are angular and representative of the wall rocks. Within the breccia, amethyst filled vugs can attain sizes of over 1 m and are lined with large, dark purple crystal points up to approximately 7.5 cm in diameter. The vugs also tend to be filled with a dense red clay; fault gouge, consisting of finely ground quartz, feldspar, chlorite, and biotite (Vos, 1982; Patterson, 1985; Garland, 1994). which must be removed in order to mine the amethyst. Light violet to a very dark, nearly black purple amethyst forms an extensive druse covering along the south wall, crystallizing between the mudstone and the breccia. The crystal points in this druse tend to be small, but are very well-formed, yielding excellent mineral specimens. Like Amethyst Mine Panorama, the amethyst crystals are sometimes coated with a layer of reddish brown hematite. Galena occurs as seams of crystals 1 cm in size, within the quartz at the west end of the exposed breccia zone, chalcopyrite-rich zones - 136 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Figure 6. Plan view of the Blue Points amethyst mine. Figure 7. Plan view of the Diamond Willow amethyst mine, refer to Figure 6 for legend. are associated with rusty stained or clear quartz crystals (Vos, 182; Garland, 1994). - 137 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Road Log Lakehead University to Amethyst Mine Panorama and Diamond Willow Amethyst Mine Leaving Lakehead University, we will follow the portion of the Trans-Canada Highway 11-17 which is the Thunder Bay Expressway. The flat terrain is the remnant of the bottom of the Nipissing stage of ancestral Lake Superior, and proceeding northeasterly, we pass upward through strandlines of the receding Pleistocene lake. Lakehead University itself is underlain by the Gunflint Formation at or near the top of the unit. Shaly rocks near the top of the formation are exposed in the bed of the McIntyre River that flows through the campus; however, in recent years blocks of rock containing the Sudbury ejecta debrisite were excavated during construction of new student residences. These placed in various places around the campus as ornamentation or barriers to vehicular traffic. Continuing, outcrops of a Logan sill diabase are exposed on the left side of the expressway. These sills form the caps of the prominent mesas south of town and underlie the high ground in the northern section of Thunder Bay, formerly the city of Port Arthur. Passing the junction of Red River Road (Highway 102), the expressway is on a level stretch marking the top of a Logan sill. The expressway then passes downhill to the Current River. In proceeding downhill outcrops of Logan sill diabase, Gunflint shale and Gunflint carbonate are successively exposed. The carbonate is ankeritic and is oxidized to yellowish orange. Climbing uphill from the Current River bridge, the highway is again cutting into diabase sill. A fault trends along the highway offsetting the sill on opposite sides of the highway. A few hundred metres farther along the highway, the sill is dropped downward by a fault trending perpendicularly to the highway. Recent work has shown that this sill, known locally as the Terry Fox sill, is a Nipigon sill (Magnus and Kissin 2010). Nipigon sills, which occur from here northeasterly to the Lake Nipigon area, are somewhat younger than Logan sills and can be distinguished on the basis of their trace element composition. Proceeding downhill and rounding a curve to the left, there is a high bluff on the left capped by a prominent diabase sill. The sill has intruded the top of the Gunflint Formation and the overlying Sudbury debrisite layer, which is capped by a thin remnant of Rove Formation shale. This is the only outcrop known in the Thunder Bay area that contains the complete debrisite layer. The east end of this outcrop is bounded by a fault that dropped down the section. Continuing onward, high ground on both sides of the highway are capped by sills; the sill on the right was extensively quarried for railway bed ballast and large stone for construction of the breakwall in the harbor. After the junction with Highway 527, the highway climbs the hill locally known as KOA hill. Prior to a widening of the highway about a decade ago, the angular unconformity between the Gunflint Formation and steeply dipping Archean metavolcanics was exposed on the left of the highway. The hill is formed by the outcrop of the Mackenzie granite, an unmetamorphosed and undeformed, late Archean pluton. The highway continues on top the of granite, which contains occasional roof pendants of Archean metavolcanics. After crossing the Mackenzie River sparse outcrops of granite are replaced by poorly exposed Gunflint Formation until just past the junction with Highway 587. Here, well-bedded red-stained carbonates of the Gunflint Formation crop out beside the highway. Passing onward to the East Loon Road, turn left onto the road, then right on Bass Lake Road. Continue to the turn off on the right to the private road to the mine. Proceeding along the mine road, it climbs steeply up from the Sibley basin onto the Archean Hilma Lake granite, ascending along a border fault surface. At the top of the grade, there is a chance to view Lake Superior with Black Bay, the Black Bay Peninsula and the Sibley Peninsula, clear weather permitting. A few more kilometres brings the road to the mine. To get to the Diamond Willow Mine, head back to Highway 11-17, and turn east towards Nipigon. Travel for approximately 13.4 km and turn left onto 5 Rd S, then make a right and drive to 5 Rd N,, crossing the railbed and make a left onto a private dirt road. Continue on this road staying right for 2.56 km, until a “Y” junction is reached and stay left until you reach the Diamond Willow Amethyst Mine parking area. - 138 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Amethyst Mine Tours amethystine color in quartz. Mineralogical Record, v. 20, p. 365-367. Note: Safety boots or shoes recommended. No sandals or open-toed shoes. Cox, R.T., 1977, Optical absorption of the d4 ion Fe4+ in pleochroic amethyst quartz, Journal of Physics C: Solid State Physics, v. 10, p. 4631-4643. The tours will pass through the operating mining areas, which is not available to ordinary tourists. No collecting is allowed in these areas After visiting the mining area, there will be an opportunity to look for specimens in a designated collecting area. The charge for specimens is by weight. Hammering or chiseling is not permitted in Amethyst Mine Panorama’s collecting area, hammering and chiseling are only permitted at Diamond Willow Amethyst Mine; however, only hammers up to 2 lb. max, are permitted and absolutely no sledge hammers and safety glasses must be worn when using hammers or tools while collecting. Specimens are also for sale in the shops. 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. References 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 W.U.Reimold, and R.L. Gibson, eds., Large Meteorite Impacts and Planetary Evolution IV: Geological Society of America Special Paper 465, p. 245-268. Deer, W.A., Howie, R.A., and Zussman, J. 1963, RockForming Minerals,Vol. 4 Framework Silicates. John Wiley and Sons, Inc., New York, 435 p. Franklin, J.M. 1978a, Uranium mineralization in the Nipigon area,Thunder Bay District, Ontario. in Current Research, Part A. Geological Survey of Canada, Paper 78-lA, pp. 275-282. Franklin, J.M., 1978b, The Sibley Group, Ontario, in Rubidium strontrium isochron age studies report 2. Edited by R.K. Wanless and W.D. Loveridge. Geological Survey of Canada, Paper 77-14, p. 31-34. Franklin, J.M., and Mitchell, R.H., 1977, Lead-zinc -barite veins of the Dorion area, Thunder Bay District, Ontario. Canadian Journal of Earth Sciences, v. 14, p. 1963-1979. 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 Science, v. 17, p. 633-651. Adekeye, J.I.D., and Cohen, A.J., 1986, Correlation of Fe4+ optical anisotropy, Brazil twinning and channels in the basal plane of amethyst quartz, Applied Geochemistry, v. 1, p.153-160. Franklin, J.M., Kissin, S.A., Smyk, M.C., and Scott, S.D., 1986, Silver deposits associated with the Proterozoic rocks of the Thunder Bay District, Ontario. Canadian Journal of Earth Sciences, v. 23, p. 1576-1591. Belleau, Rémi., 1576. Les amours et nouveaux eschanges des pierres précieuses : vertus et proprietez d’icelles ; Discours de la vanité, pris de l’Ecclesiaste ; Eclogues sacrees, prises du Cantique des Cantiques ([Reprod.]) / par Remy Belleau. Published by M. PatissonM. Patisson (Paris). Accessed from BnF Gallica: https:// gallica.bnf.fr/ark:/12148/bpt6k522648/f21.image. Last Accessed on November 14, 2025 Frondel, C. 1962, The System of Mineralogy, 7th edition, Vol. III Silica Minerals. John Wiley & Sons, New York and London, 334 p. Campbell, D.A., Jonsson, J.R.B., Kurcinka, C.E., Hinz, S.L.K., Sabiri, N., Meyer, G., McEachern, A.D. and Smith, A.M. 2024. Report of Activities 2024, Resident Geologist Program, Thunder Bay South Regional Resident Geologist Report: Thunder Bay South District; Ontario Geological Survey, Open File Report 6417, 128p. Cheadle, B.A., 1986, Alluvial-playa sedimentation in the Lower Keweenawan Sibley Group, Thunder Bay District, Ontario. Canadian Journal of Earth Science, v. 23, p. 527-541. Cohen, A.J., 1989, New data on the cause of smoky and Garland, M.I., 1994. Amethyst in the Thunder Bay area. Ontario Geological Survey, Open-file Report 5891, 197 p. Goetz, M.M. 2014. Heating Experiments of Amthyst from Thunder Bay Amethyst Mine. HBSc thesis, Lakehead University, 63 p. Haynes, F.M., 1988, Fluid-inclusion evidence of basinal brines in Archean basement, Thunder Bay Pb-Zn-Ba district, Ontario, Canada. Canadian Journal of Earth Sciences, v. 25, p. 1884-1894. Heaman, L.M., and Easton, R.M., 2006, Preliminary U/ Pb geochronology results: Lake Nipigon Region Geoscience Initiative. Ontario Geological Survey Miscellaneous Release – Data 191, 79 p. Hebert, L.B., and Rossman, G.R., 2008, Greenish quartz from the Thunder Bay Amethyst Mine Panorama, Thunder - 139 - �Proceedings of the 72nd ILSG Annual Meeting - Part 2 Bay, Ontario, Canada. Canadian Mineralogist, v. 46, p. 111-124. Holden, E.F., 1925, The cause of color in smoky quartz and amethyst, American Mineralogist, v. 10, p. 203-252. 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 Science, v. 41, p. 1329-1338. Jennings, E.A., 1985, Geology of the Thunder Bay Amethyst Mine and Precious Purple Gemstone claims. Report to Precious Purple Gemstones Ltd., Thunder Bay, Ont., 65 p. Kissin, S.A., 1997, Comprehensive research to colour enhance Canadian amethyst by heat treatment and irradiation. Final Report, Amsearch Colour Project, Northern Ontario Development Agreement SSC File #015SQ-2223440-2-9243, 36 p. and appendix. Klarner, J.M., and Kissin, S.A., 2011, Hydrothermal bleaching of amethyst at the Thunder Bay Amethyst Mine, Ontario. Geological Society of America Annual Meeting, Minneapolis, Paper No. 44-11. Lehmann, G., and Bambauer, H.U., 1973, Quartz crystals and their colors. Angewendtede Chemie International Edition, v. 12, p. 283-291. Magnus, S., and Kissin, S., 2010, Assimilation and petrogenesis in the Navilus and Terry Fox sills, Thunder Bay, Ontario; in Institute on Lake Superior Geology, Proceedings and Abstracts, v. 56, part 1, p. 36-37. McArthur, J.R., Jennings, E.A., Kissin, S.A., and Sherlock, R.L., 1993, Stable-isotope, fluid-inclusion, and mineralogical studies relating to the genesis of amethyst, Thunder Bay Amethyst Mine, Ontario. Canadian Journal of Earth Science, v. 30, p. 19551969. McLaren, A.C., and Pitkethly, D.R., 1982, The twinning microstructure and growth of amethyst quartz. Physics and Chemistry of Minerals, v. 8, p. 128-135. Ontario 1990. Mineral Emblem Act, 1990, c. M.13, s. 1 Ontario Geological Survey. 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release— Data 126 – Revision 1. Ontario Geological Survey, 2026. Ontario Mineral Inventory; Ontario Geological Survey, Ontario Mineral Inventory, online database (March 2026 update). Patterson, G.C., 1985, Amethyst in the Thunder Bay area of Ontario, Canadian Gemologist, V. 6, p. 104-116. Rossman, G.R., 1994, Colored varieties of the silica minerals, in P.J. Heaney, C.T. Prewitt, and G.V. Gibbs, eds., Silica: Physical Behavior, Geochemistry and Materials Applications, Mineralogical Society of America, Reviews in Mineralogy, v. 29, p. 433-467. Sinkankas, J., 1976, Gemstones of North America, Vol, II. D. Van Nostrand Company, Inc., New York, 494 p. Sutcliffe, R.H., 1991, Proterozoic geology of the Lake Superior area in P.C. Thurston, H.R. Williams, R.H. Sutcliffe, and G.M.Stott eds., Geology of Ontario. Ontario Geological Survey, Special Volume 4, Part 1, p. 627-660. Van Schmus, W.R., Green, J.C., and Halls, H.C., 1982, Geochronology of Keweenawan rocks of the Lake Superior region: A summary, in R.J. Wold and W.H. Hinze, eds., Geology and Tectonics of the Lake Superior Basin, Geological Society of America, Memoir 156, p. 165-171. Vos, M.A., Abolins, T., and Smith, V. 1982: Industrial. Minerals of Northern Ontario- Supplement 1, Ontario Geological Survey Open File Report 5388, 344 p. McCrank, G.F.D., Misiura, J.D., and Brown, P.A., 1981, Plutonic rocks in Ontario. Geological Survey of Canada, Paper 80-23, 171 p. - 140 - � 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, 2026 Description An account of the resource Institute on Lake Superior Geology. Thunder Bay, Ontario. May 21-22, 2026. 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 2026-05-21 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/99c083bfbf57b97590b6050acae4bd4e.pdf f1d247daa107d03ceffc4ef05679054d PDF Text Text 71st Annual Meeting Proceedings Volume 71 Part 1 – Program and abstracts Mountain Iron, Minnesota, May 14-17, 2025 �71st Annual Meeting Institute on Lake Superior Geology Mountain Iron, Minnesota May 14-17, 2025 Meeting Co-Chairs Amy Radakovich, Allison Severson, Eric Nowariak, Stacy Saari, Aaron Hirsch Proceedings Volume 71 Part 1: Program and Abstracts Edited by Co-Chairs i �71st Institute on Lake Superior Geology Volume 71 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trip 1: Transect of the Quetico subprovince Trip 2: Drill Core from three Cu-Ni deposits of the Duluth Complex Trip 3: How do you make iron and/or manganese in Proterozoic Iron Formation? Trip 4: New Geological Insights into the genesis of iron ores at Lake Vermillion – Soudan Underground Mine State Park Trip 5: Neoarchean alkalic intrusions in the Wawa and Quetico subprovinces Trip 6: Unique Keweenawan inclusion (Colvin Creek) in the Duluth Complex Trip 7: Classic outcrops of Northeastern Minnesota Trip 8: Glacial Lake Norwood and the Koochiching Lobe Reference to material in Part 1 & 2 should follow the examples below: Authors, 2025, Title in Institute on Lake Superior Geology, 71st Annual Meeting, Mountain Iron, Minnesota, Part 1 - Abstracts and Program, v. 71, part 1, p. xx-xx. Authors, 2025, Field Trip title in Institute on Lake Superior Geology, 71st Annual Meetings, Mountain Iron, Minnesota, Part 2 – Field Trip Guidebook, v. 71, part 2, p. xx-xx. Proceedings Volume 71, Part 1: Program and Abstracts and Part 2: Field Trip Guidebook are published by the 71st 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 black and white. 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 �Table of Contents Table of Contents ..................................................................................................................................... iii Institutes on Lake Superior Geology, 1955-2025 .................................................................................... iv Sam Goldich and the Goldich Medal ...................................................................................................... vii Goldich Medal Guidelines ....................................................................................................................... ix Goldich Medalists .................................................................................................................................... xi Citation for the 2025 Goldich Medal Recipient ...................................................................................... xii Honoring the Pioneers of Lake Superior Geology ................................................................................. xiv Pioneers of Lake Superior Geology ....................................................................................................... xiv 2025 Citation for Robert Bell (1841-1917)............................................................................................. xv Eisenbrey Student Travel Awards ........................................................................................................... xx Joe Mancuso Student Research Awards ................................................................................................. xxi Doug Duskin Student Paper Awards ..................................................................................................... xxii Board of Directors ................................................................................................................................ xxiii 2025 ILSG Meeting Volunteers ........................................................................................................... xxiv 2025 ILSG Meeting Session Chairs ..................................................................................................... xxiv Field Trip Leaders and Guidebook Authors .......................................................................................... xxv Mine to Mountain Bike Mecca: ........................................................................................................... xxvi Report of the chairs of the 70th annual meeting .................................................................................. xxvii Donations to Support the Annual Meeting ........................................................................................... xxxi TECHNICAL PROGRAM ................................................................................................................ xxxiii ABSTRACTS........................................................................................................................................ xliii iii �Institutes on Lake Superior Geology, 1955-2025 # 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 Date 1955 1956 1957 1958 1959 1960 1961 1962 1963 1964 1965 1966 1967 1968 1969 1970 1971 1972 1973 1974 1975 Place 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 iv Chairs 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 �# 22 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 56 Date 1976 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 2010 Place St. Paul, Minnesota 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 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 Chairs M. Walton 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.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 �# 63 Date 2017 Place Wawa, Ontario 64 2018 Iron Mountain, Michigan 65 66 2019 2020 Terrace Bay, Ontario Meeting cancelled 67 68 69 2021 2022 2023 Virtual meeting Sudbury, Ontario Eau Claire, Wisconsin 70 2024 Houghton, Michigan 71 2025 Mountain Iron, Minnesota vi Chairs 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 T.J. Bornhorst, E. Vye, P. Cobin, & J. Degraff A. Radakovich, A. Severson, E. Nowariak, S. Saari, A.C. Hirsch �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. Kalliokoski, 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 �Institute on Lake Superior Geology Goldich Medal viii �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 medalists 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 medalist, 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. Nominating Procedures ix �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 vitas, 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 �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 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 2024 Suzanne W. Nicholson 1988 Walter S. White 2007 Joseph Mancuso 2025 Robert Michael Easton 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 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 Gregory R. Brumpton 2025 GOLDICH MEDAL RECIPIENT Robert Micheal Easton Goldich Medal Committee Serving through the meeting year shown in parentheses. Dean Peterson (2022-2025) Big Rock Exploration, Industry Member (Committee Chair) Marcia Bjornerud (2023-2026) Lawrence University, Academic Member Robert Cundari (2025 - 2028) OGS, Government Member xi �Citation for the 2025 Goldich Medal Recipient Robert Michael Easton It is a great pleasure and honor to present the 2025 Goldich Medal to Dr. Robert Michael Easton, a highly respected senior scientist at the Ontario Geological Survey, in Sudbury, Ontario. Michael Easton, or ‘Mike’ as we know him, has been and is, without any doubt, among the leading and most productive geoscientists at the Geological Survey of Ontario (OGS) where he has spent much of his geological career (1982–2025). His curriculum vitae and publication list provide evidence for >600 publications and significant contributions — way too many to cite here. Even a short list of publications most relevant to the interests of the Institute and the geology of the Midcontinent Rift (MCR) spans four pages. The highlights include: • • • • a large number of peer-reviewed papers and reports; numerous extended abstracts in ILSG Proceedings volumes spanning the years from 1985 to 2023; meticulous editing of various ILSG Proceedings volumes; and the writing and editing of several comprehensive ILSG field trip guidebooks. In 2022, Mike co-lead and co-organized the 68th ILSG meeting in Sudbury, the first post-“peak COVID” meeting. We had proposed organizing this Sudbury meeting years earlier, an idea cooked up at another ILSG meeting in Terrace Bay, … but then COVID hit! It was a pleasure to organize this highly successful meeting with Mike, as one can always be 100% sure Mike will come through with everything. Although it was a joint effort, Mike took care of all the editing of both Proceedings volumes (Part I and II), and a fair bit of the local logistics. Born and raised in Ontario, Mike started his geology career with a BSc Honours degree (1976) at the University of Western Ontario, London, with a thesis titled "Geobotanical Studies in the Back River Volcanic Complex, NWT." He then moved on to the University of Hawaii, Honolulu, where he graduated (1978) with an MSc thesis on the "Stratigraphy and Petrology of the Hilina Formation: The oldest exposed lavas of Kilauea Volcano, Hawaii". In the late 1980s, I remember studying a treatise and guidebook on volcanology that was influential at the time, and this was authored by Mike Easton and his wife Monica (Easton & Easton, 1985)! Mike completed his graduate studies with a PhD from Memorial University (1982), in Newfoundland, with a thesis titled "Tectonic Significance of the Akaitcho Group, Wopmay Orogen, NWT." These studies brought him to the Slave craton of northern Canada, and its western active margin, the Paleoproterozoic Great Bear Magmatic Zone. His studies of these ancient terranes, sponsored in part by the Geological Survey of Canada, prepared him well for the complex geology of the Canadian Shield in Ontario. He then joined the OGS, where over the years he has taken on more and more senior roles but never gotten away from doing fieldwork. At the OGS, Mike has mentored and supervised numerous students and junior colleagues, including a good number of them working in areas along the northern xii �shore of Lake Superior. He has also taken on more and more editorial roles for various OGS publications, maps, and datasets. From 2002 to 2007, Mike was one of the scientific leads of the Lake Nipigon Geoscience Initiative (LNGI), and provided oversight on OGS mapping projects in the region. He was directly involved in some of the mapping, and particularly the geochronology sampling. He handled numerous publications for this large project and was a guest-editor on the final volume that published many of the LNGI results (Easton et al., 2007). Since then, Mike has been a frequent collaborator on other projects either directly or indirectly relevant to Lake Superior area geology. In the early 1990s, together with Terry Carter, Mike investigated the basement geology beneath the Paleozoic cover in SW Ontario (e.g., Easton & Carter, 1991, 1994, 1995), using geophysical data, and drill cores and cuttings, to locate the Grenville Front and the extension of the MCR in Ontario and into Michigan. Notably, they found that the Grenville Front was located some 100 km to the east of where previous interpretations had located it, and that metamorphosed equivalents of MCR rocks were likely present in the Grenville Front tectonic zone in Essex County. They were among the first to hypothesize that the final stages of MCR rifting and inversion were connected to the main tectonic phases of the Grenville orogeny. From 2002 to 2010, Mike was involved with the MCR digital data and publication collaboration between the OGS, the Minnesota Geological Survey, and the United States Geological Survey, specifically the compilation of Ontario geological, mineral deposit, geochemical, and geochronological data in GIS-compatible formats to allow incorporation into the USGS-led cross-border compilation for the Midcontinent Rift. In addition, also in 2010, he was a co-organizer and editor of four guidebooks for the 11th International Platinum Symposium (June 2010, Sudbury), a meeting that had a strong focus on the MCR, including a week-long field trip visiting deposits around Lake Superior. None of this would have ever happened without Mike’s efforts and contributions. Among his many other contributions to ILSG over the years, Mike also served (and still serves) as a board member for the Institute (2022-2025). After spending the last 53 summers doing fieldwork and research in the Grenville, the Southern Province, the Lake Superior area, or elsewhere in Ontario, Mike retired in March 2025. Given his outstanding accomplishments and amazing productivity over the years, either for the OGS or for various extra-curricular projects such as ILSG meetings, leading field trips, time-consuming editorial jobs, teaching as an adjunct professor, or supervising and mentoring many students (and never missing a beat!), it is a great honour to present Mike with the Goldich Medal Citation by: Wouter Bleeker, Senior Research Scientist, Geological Survey of Canada, Ottawa xiii �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 the inception of the Institute in 1955. Biographical sketches of those pioneers will be presented at future annual meetings so that all may 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 summarize the contribution of the nominee. 2) The Organizing 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) 2024 Roland Duer Irving (1847-1888) 2025 Robert Bell (1841-1917) xiv �2025 Citation for Robert Bell (1841-1917) Pioneer of Lake Superior Geology Robert Bell had a decided taste for the natural sciences, especially for geology. In 1856, at the age of 15, he secured a temporary position with the Geological Survey of Canada. He assisted Sir William Logan, the Survey’s Director, beginning an illustrious career with the GSC that would span half a century. While working summers for the Survey, Bell graduated in 1861 from McGill College and received the Governor General’s Medal. Two years later, after study at the University of Edinburgh, he joined the faculty of Queen’s College. All the while, he spent summers with the GSC and was made a permanent officer in 1869, named Assistant Director in 1877, Chief Geologist in 1890, and, finally, Acting Director in 1901. He also earned a medical degree in 1878 so that he was prepared for any mishap in the field. Bell is best-remembered for his extensive explorations in northern Quebec, Ontario, Manitoba and the eastern Arctic in the 1870s and 1880s. He mapped the rivers between Hudson Bay and Lake Superior and reconnoitred part of the route that would be adopted for the National Transcontinental Railway. In 1859, Bell assisted in mapping the north shore of Lake Huron and first visited the Lake Superior region in 1860, west of Sault Ste. Marie. His Report on the Geology of the Northwest Side of Lake Superior and of the Nipigon District was published in 1870. In 1870 and 1871, he continued to work north of Lake Superior. In 1872 and 1873, he assisted GSC Director Alfred Selwyn on a preliminary exploration westward from Lake Superior to Fort Garry (now Winnipeg). In 1876, Bell examined the eastern shore of Lake Superior, as well as the Garden River and Echo Lake areas and the northeastern shore of Georgian Bay. A reconnaissance survey was undertaken between Parry Sound and the Ottawa River. In 1881, Bell carried out additional surveys in the Hudson Bay basin and in the Lake Superior region. During 1883 and 1884, Bell continued work near Lake of the Woods. In 1887, he continued a survey, started in 1886, between the Montreal River and Lake Huron to clarify the nature of the Huronian, especially in connection with its mineral deposits. He served as a member of the Royal Commission on the Mineral Resources of Ontario from 1888 to 1889. Between 1888 and 1892, Bell mapped the Sudbury and French River areas. His 1890 paper, On Glacial Phenomena in Canada, was regarded as the most significant advance in Canadian glaciology since Logan’s first acceptance of glacial action in Canada in 1847. Bell was the first to recognize ice streaming in the Laurentide ice sheet and also noted the occurrence of diamonds in glacial drift from Ohio through Indiana, Michigan and Wisconsin, suggesting a possible provenance in Ontario. Of particular interest to the ILSG is Bell’s involvement on a Special Committee created in 1903 on the nomenclature and correlation of the Lake Superior region geology of the United States and Canada. Its findings led to the first joint report by geologists of the two countries. The Committee comprised C.R. Van Hise and C.K. Leith of the United States Geological Survey, A.O. Lane, State Geologist of Michigan; Robert Bell and Frank D. Adams of the GSC, and W.G. Miller, Provincial Geologist of Ontario. In August, 1904, the committee met in the Marquette district, and, during the six weeks following, visited the Gogebic, Mesabi, Vermilion, Rainy Lake, Lake of the Woods, Animikie, and xv �Huronian districts. As a result, both countries adopted a common stratigraphy and nomenclature that served as a basis for later iterations and our current stratigraphic framework. Bell also made field notes on flora and fauna, forests, climate, soil, indigenous people, ethnology and resources. He performed much of his field work without maps and had to do topographical surveys as he went along. It is estimated that Bell named over 3000 geographical features, prompting colleagues to call him the “Father of Canadian Place-Names”. Bell authored 32 GSC reports, 111 journal papers, 17 solo-authored geological maps, 46 other geological, topographical, and cadastral maps as senior author, and 38 maps as junior author. He gave numerous lectures to natural history, historical, and charitable societies. In 1865, at the age of 23, he was elected a fellow of the Geological Society of London. A chartermember of the Royal Society of Canada (1882), he became a fellow of the Royal Society of London in 1897. In 1903, he was made a companion of the Imperial Service Order and in 1906 was awarded both the Patron’s Medal of the Royal Geographical Society of London and the Cullum Geographical Medal of the American Geographical Society of New York. Bell retired from the GSC in 1908. His career exemplified the wide-ranging reconnaissance work performed by the government geologist in the late 19th century. He was a generalist who valued field work over more detailed, specialized study. Few could match Bell’s travels, eclectic interests, and length of service. As Ami (1927) memorialized, “Bell was especially fond of investigating and exploring regions hitherto untraversed. Pioneer work of this nature can scarcely be appreciated today, when newer and more up-to-date methods of examining a hitherto-unknown territory are employed.” Citation by: Mark Smyk (Lakehead University / Ontario Geological Survey (retired)) References Adams, F.D., Bell, R., Lane, A.C., Leith. C.K., Miller, W.G. and Van Hise, C.R. 1905. Report of International Committee on Lake Superior Geology; Journal of Geology, February-March, 1905; in Precambrian nomenclature; Ontario Bureau of Mines, Report for 1905, v.4, part 1, 1905, pp.269-277. Ami, H.M. 1927. Memorial of Robert Bell. Bulletin of the Geological Society of America, v.38, pp. 18-33; PLS. 1- https://archive.org/details/sim_geological-society-of-americabulletin_1927_38/page/n41/mode/2up?view=theater Brookes, I.A. 2016. All that glitters… The Scientific and Financial Ambitions of Robert Bell at the Geological Survey of Canada; Geoscience Canada, v. 43, pp. 147–158; http://www.dx.doi.org/10.12789/geocanj.2016.43.098. Waiser, W.A. 1998. Robert Bell. Dictionary of Canadian Biography, v. XIV (1911-1920), https://www.biographi.ca/en/bio/bell_robert_1841_1917_14E.html. xvi �In Memoriam James M. Franklin (November 9, 1942 – June 19, 2024) We note the passing, on June 19, 2024, of long-term member and former SEG President (2000) James (Jim) M. Franklin. Jim had a long and productive career in academia, government, and industry. He made landmark scientific contributions to our understanding of volcanogenic massive sulfide (VMS) deposits, black smokers and sea-floor massive sulfide (SMS) deposits, and the metallogeny of the Precambrian orogenic belts. Jim’s work with the Ontario Department of Mines in 1966 along the north shore of Lake Superior led to his PhD study of the Proterozoic geology and metallogeny of the Thunder Bay area. He was the first Professor of Economic Geology at Lakehead University (1970–1976) and was later named as a Fellow of Lakehead University in 2017 in recognition of his many contributions to that institution and to its fledgling Geology Department. He then spent more than 20 years at the Geological Survey of Canada where he led the marine minerals program and ongoing work on VMS deposits on land, such as at Sturgeon Lake. Late in his career at the GSC, he was Chief Scientist and responsible for the day-to-day scientific direction of the organization and helping to inform and educate politicians and bureaucrats on the importance of science to the economy and well-being of Canada. In 1998, after retiring from the GSC, he established Franklin Geosciences and had a highly successful career as a consultant and contributed to the discovery of mineral resources globally, while serving as a director and advisor to numerous companies and scientific organizations, including SEG. Jim was very generous with his time and provided guidance and mentorship to students and professionals alike. He received numerous recognitions for his contributions, including Fellow of the Royal Society of Canada, member of the Canadian Mining Hall of Fame, and recipient of the Logan Medal recipient the Geological Association of Canada and the R.A.F. Penrose Gold Medal from SEG. He supported ILSG as field trip leader, Proceedings editor and banquet speaker. xvii �In Memoriam Jorma “Joe” Kalliokosk (November 23, 1923 – June 3, 2024) Jorma “Joe” Kalliokoski passed away on June 3, 2024, at age 100. He was born in Harma, Finland on Nov 23, 1923. In 1931 his family moved to Sudbury, ON, and later to Timmins, ON. He graduated from Western University in London, ON, and later received his PhD from Princeton University. Joe started as a geologist with the Geological Survey of Canada and later worked for Newmont Exploration. He was hired by Princeton’s Department of Geology in 1956. In 1968, he moved to Michigan Tech as a Professor and Head of the Department of Geology and Geological Engineering, where he remained until his retirement in 1988. He was very proud of the department’s growth in research papers and in research funding during his tenure. During his long geology career, he had many travel adventures from the wilds of Canada, to remote areas of South America, and various locations in Europe. He had the ability to make new friends everywhere he went. Joe was a Fellow of SEG for a noteworthy 60 years, from 1958 to 2018. He served the Society in a number of volunteer positions, including SEG Councillor (1972–1974) and SEG President (1980). He served as Trustee of SEG Foundation, Associate Editor for Economic Geology, and Business Editor (1971–1977) and Director for the Economic Geology Publishing Company (PUBCO), the company that was established to publish the journal and later merged with SEG. Joe was also an active member and supporter of ILSG, serving as its Secretary-Treasurer and Chair of the Goldich Medal Committee. He delivered papers at ILSG on various topics, including unconformitytype Proterozoic uranium deposit potential in northern Michigan; the Jacobsville sandstone and tectonic activity; and new Precambrian geology mapping of the Upper Peninsula. He Chaired the 1972 meeting in Houghton and was awarded the Goldich Medal in 1989. xviii �In Memoriam James Alexander Grant (October 3, 1935 — October 3, 2024) James Alexander Grant died on October 3, 2024 – coincidentally also his birthday – at the age of 89. James “Jim” Grant was born in Inverness, Scotland, in 1935. After graduating from the University of Aberdeen, he left Scotland for Canada where he earned his M.S. at Queens University and then his Ph.D. at the California Institute of Technology (Caltech). After graduating from Caltech, Jim took a job in Minneapolis as a geology professor with the University of Minnesota. Jim and his family moved to Duluth in 1969 and he joined the geology department at the University of Minnesota-Duluth, where he would work with his beloved colleagues and students for the next 35 years. In the early 1970s, he helped launch UMD’s still-running geology summer field camp in Park City, Utah, bringing undergrad students out to the mountains for many years. Jim’s groundbreaking work in the 1980s on the isocon diagram is now used by geologists the world over. Over the course of his career, Jim made substantial contributions to the geology of the Lake Superior region. His seminal mapping of the Minnesota River Valley subprovince is still referenced today by the dozens who have since worked in the region. Jim taught hundreds of students over the course of his career who have gone on to contribute in many ways to the geology of the Lake Superior region. Among the most memorable experiences for his students were Jim’s metamorphic petrology trips through Michigan’s Upper Peninsula. xix �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 end of the annual meeting. 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. xx �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 2024, the ILSG Board of Directors selected two students to be granted research funding of $500 each from the Joe Mancuso Student Research Fund. The awardees were: Zsuzsanna P. Allerton, University of Minnesota- Twin Cities Omar Khalil Droubi, University of Wisconsin - Madison xxi �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. 2025 Student Paper Awards Committee Aaron Hirsch – Minnesota Geological Survey (Committee Chair) Carsyn Ames – Wisconsin Geological and Natural History Survey Paula Leier-Englehardt – HydroGeo Solutions LLC, Wisconsin Ross Salerno – United States Geological Survey Esther Stewart – Wisconsin Geological and Natural History Survey Nick Swanson-Hysell – University of Minnesota xxii �Board of Directors Amy Radakovich, Chair (2025-2028) - Minnesota Geological Survey Peter Hollings, Secretary (2019-2027) — Lakehead University Mark A. Jirsa, Treasurer (2022-2025) — Minnesota Geological Survey Mike Easton (2022-2025) — Ontario Geological Survey Carysn Ames (2023-2026) — Wisconsin Geological and Natural History Survey Theodore J. Bornhorst, (2024-2027) — Michigan Technological University Board members serve through the close of the meeting year shown in parentheses. xxiii �2025 ILSG Meeting Volunteers Angela Sipila - Mesabi Range Geological Society Henry Djerlev - Mesabi Range Geological Society Kim Berry - Mesabi Range Geological Society William Daniels - Mesabi Range Geological Society Ann Marie Prue - MN Department of Natural Resources 2025 ILSG Meeting Session Chairs Aaron Hirsch, Minnesota Geological Survey Robert Lodge, University of Wisconsin, Eau Claire Eric Nowariak, Minnesota Geological Survey Amy Radakovich, Minnesota Geological Survey Stacy Saari, Minnesota Department of Natural Resources, Lands and Minerals Allison Severson, Minnesota Geological Survey xxiv �Field Trip Leaders and Guidebook Authors Field trips have been the mainstay of the ILSG since its inception 71 years ago. We give special thanks to the field trip leaders and guidebook authors who volunteered their time and talent in carrying that tradition forward. Trip 1: Transect through the Quetico subprovince of northern Minnesota – Eric Nowariak (Minnesota Geological Survey), Mark Jirsa (Minnesota Geological Survey, retired) Trip 2: Drill Core from three Cu-Ni deposits of the Duluth Complex - Mark Severson (Natural Resources Research Institute, Teck Retired), Cullen Phillips (New Range Copper Nickel), Kevin Boerst (Twin Metals Minnesota) Trip 3: How do you make iron and/or manganese in Proterozoic Iron Formation? - Alex Steiner (Big Rock Exploration), Latisha Brengman (University of Minnesota, Duluth), Dean Peterson (Big Rock Exploration) Trip 4: New Geological Insights into the genesis of iron ores at Lake Vermillion – Soudan Underground Mine State Park - George J. Hudak (University of Minnesota, George Hudak Geosciences P.L.L.C.), Zsuzsanna P. Allerton (University of Minnesota), Annia Fayon (University of Minnesota) Trip 5: Neoarchean alkalic intrusions in the Wawa and Quetico subprovinces - Terry Boerboom (Minnesota Geological Survey, retired), Amy Radakovich (Minnesota Geological Survey) Trip 6: Unique Keweenawan inclusion (Colvin Creek) in the Duluth Complex - Mark Severson (Natural Resources Research Institute, Teck, retired), Allison Severson (Minnesota Geological Survey), Lauri Severson (Earth Science teacher, retired) Trip 7: Classic outcrops of Northeastern Minnesota - Dean M. Peterson (Big Rock Exploration), George J. Hudak (University of Minnesota, George Hudak Geosciences P.L.L.C.) Trip 8: Glacial Lake Norwood and the Koochiching Lobe - Phillip Larson (Vesterheim Geoscience PLC), Andrew Breckinridge (University of Wisconsin-Superior), Howard Mooers (University of Minnesota, Duluth) xxv �Mine to Mountain Bike Mecca: The story of the Redhead Mountain Bike Park Pete Kero PE, Senior Environmental Engineer Barr Engineering Co. Pete Kero, PE, is an environmental engineer and Vice President with Barr Engineering Co. He has over 30 years of experience in mine permitting, water management, reclamation, and repurposing across the United States. He was the visionary behind the award-winning Redhead Mountain Bike Park in Chisholm, Minnesota which repurposed several former iron mine pits and stockpiles into a destination-quality regional park for mountain biking, hiking, water recreation and all-terrain vehicles. The project has been featured by Outside Magazine, the Sierra Club and the nation-wide documentary film Biketown. Pete’s book Minescapes: Reclaiming Minnesota’s Mined Lands, which was published by the Minnesota Historical Society Press, won a 2024 Minnesota Book Award. This talk will describe the transformation of ten idled open pit iron ore mines in northeastern Minnesota into a world-class recreation destination for mountain biking, hiking and paddling. In addition to describing how and why the trails were built, the presentation will include technical details on sustainable trail design, the concept of intermediate recreational use, changes to mine pit fencing laws that allow for government-sanctioned recreational use of mine lands and the early results and benefits from the first 5 years of the park’s operation. xxvi �Report of the Chairs of the 70th Annual Meeting Theodore J. Bornhorst, Erika C. Vye and Patrice F. Cobin Houghton, Michigan The 70th Institute on Lake Superior Geology (ILSG) was held May 15 to 18, 2024 in Houghton, Michigan, with the meeting headquartered at the Memorial Union Building on the campus of Michigan Technological University. The meeting was sponsored by the A. E. Seaman Mineral Museum, the Great Lakes Research Center, and the Department of Geological and Mining Engineering and Sciences - all units of Michigan Technological University. The meeting was co-chaired by Ted Bornhorst (principal cochair), Erika Vye, Patrice Cobin, and Jim DeGraff; all co-chairs are affiliated with Michigan Technological University. In addition to being a co-chair Patrice Cobin and Julie Stark served as registrars for the 70th annual meeting. The institute was attended by a total of 182 participants of which 40 were students. The meeting consisted of two full days of technical sessions from Thursday morning 16th of May through Friday afternoon 17th of May, and two days for field trips, pre-and post-meeting. A total of 57 presentations were subdivided into 8 technical sessions; 6 technical sessions for 30 oral presentations (of which 5 were presented by students), and 2 poster technical sessions with a total of 27 poster presentations (of which 16 were presented by students). Three presentations were withdrawn. Since past meetings have not included a dedicated technical session for poster presentations, the chairs opted to include two poster sessions for the 70th meeting. We believe this facilitated more time for attendees to review the posters and facilitated interaction between the authors of posters and attendees. The technical sessions of the 70th annual meeting of ILSG were published in 2024 as Part 1 of Proceedings Volume 70 (111 pages). As is customary with ILSG meetings, the field trips were a highlight of the 70th ILSG. The meeting offered 7 field trips with 3 pre-meeting on Wednesday May 15, and 4 post-meeting trips on Saturday May 18. Overall, the field trips were well attended. There were 145 registrants for the 5 field trips that were able to be run. Demand for 4 of the trips exceeded capacity resulting in wait lists. Pre-meeting trip 1 was led by Ted Bornhorst (Michigan Tech) and focused on Mesoproterozoic “Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan.” Pre-meeting trip 2 was led by Tom Wright (Quincy Mine Hoist Association) and Jim DeGraff and Katherine Langfield (Michigan Tech) and focused on the “Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan.” Pre-meeting trip 3 focusing on “Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty” was scheduled to be led by Erika Vye, Charlie Kerfoot (Michigan Tech), Stephanie Swart (Michigan Department of Environmental Quality), and Dione Price and Evelyn Ravindran (Keweenaw Bay Indian Community). However, the trip could not be run because of low water levels and shifting sediment impeding access to the harbor. xxvii �Post-meeting trip 4 was led by Jim DeGraff, Katherine Langfield, and Dan Lizzadro-McPherson (Michigan Tech) and focused on “Keweenaw Fault Geometry and Kinematics: Clues to Its Nature and Origin.” Post-meeting trip 5 was led by Matt Portfleet (Adventure Mining Company) and Ted Bornhorst (Michigan Tech) and focused on the “Adventure Mine, Ontonagon County, Michigan: Geology and History of a Native Copper Mine.” Post-meeting trip 6 led by Chad Deering (Michigan Tech) ventured outside of the Keweenaw rift to investigate “Southern Complex Granitoids, Gneisses, and Migmatites: New Data, Discoveries, and Perspectives.” Field trip 7 led by Stan Vitton and Mohammad Sadeghi (Michigan Technological University) was scheduled to investigate “Landslides in the Glacial Lake Ontonagon Sediments,” but had to be cancelled due to lack of registrations. Field trip guides were published in 2024 as Part 2 of the Proceedings Volume 70 (194 pages). Five Doug Duskin Best Student Paper Awards were given for student oral and poster presentations as judged by the 2024 Student Paper Awards Committee chaired by Stacy Saari (Minnesota Department of Natural Resources). Zsusanna Allerton was awarded the best oral presentation. The best graduate student poster presentation was awarded to Yirou Xu. The best undergraduate student poster presentation was awarded to Lyndsie Vickers. Alice Martin and Alexander Lawrence were awarded the runner-up for graduate student poster and for undergraduate student poster respectively. The 70th ILSG awarded 14 Student Travel and Participation Awards to help defray the cost of presentations of their research and participation in the ILSG professional meeting. The eligibility of costs, as designated by the Eisenbrey Award, were expanded for the 70th ILSG Student Travel and Participation Awards. We thank the donors for supporting the student awards. The awards were made possible by the generous financial support from our corporate sponsor Eagle Mine – Lundin Mining, the Geological Society of Minnesota, and 23 individual donors. The awardees were Zsuzsanna Allerton, Farhan Ahmed Bhuiyan, Andrea Paola Corredor Bravo, Kevin Mexia Duran, Trent Ediger, Alex Lawrence, Jordan Peterzon, Lucas Robarg, Daniel Shakked, Vlad Sheshnev, Demily Thibodeau-Bello, Adam Vanderkin, Lyndsie Vickers, and Yiruo Xu. There were 6 Michigan Tech students whose registration fees were waived because they volunteered with logistics for the meeting. The ILSG social and banquet was hosted at the Memorial Union Building on Thursday evening May 15. There were 120 people at the annual banquet. Ted Bornhorst served as master of ceremonies for the postbanquet program. After the introductions, Peter Hinz gave a short presentation about a geological excursion to Hawaii. Amy Radovich announced the location of the 2025 meeting as Mountain Iron. The program continued with ILSG awarding the prestigious Goldich Medal to Suzanne W. Nicholson (recently retired from the U.S. Geological Survey). Laurel Woodruff (U.S. Geological Survey and Goldich Medalist in 2014) provided the citation for Suzanne. The co-chairs and the A. E. Seaman Mineral Museum recognized Ted Bornhorst with a plaque for his distinguished service to ILSG. A highlight of the banquet was the keynote presentation by Robert Hazen (Carnegie Institution for Science), an internationally recognized and distinguished mineralogist. His thought-provoking presentation was on “Mineral Informatics: A New Frontier in Understanding Earth.” The keynote presentation ended the banquet program. Hazen’s presentation was made possible by joint funding between the 70th ILSG and the A. E. Seaman Mineral Museum of Michigan Tech. Hazen gave a second presentation on Friday evening for the general public and as a bonus for ILSG participants. This presentation was the A. E. xxviii �Seaman Mineral Museum’s 2024 Edith D. and E. Wm Heinrich Lecture titled “Mineral Evolution: A case study of a new natural law.” The first presentation of the technical sessions was given by Jim Miller (Goldich Medalist in 2012) who gave the citation for Roland Duer Irving as the 2024 Pioneer of Lake Superior Geology. Irving is the 5th person to be recognized for their contributions to Lake Superior Geology prior to the initiation of the ILSG. The Institute’s Board of Directors met on Thursday May 16, 2024 to discuss ILSG business and approve the 2025 meeting location. The meeting was attended by Ted Bornhorst (Board Chair), Carysn Ames, Mark Smyk, Peter Hollings (Secretary), and Mark Jirsa (Treasurer). Guests at the meeting were the meeting co-chairs Patrice Cobin, Erika Vye, and Jim DeGraff, and Amy Radakovich (Assistant Treasurer) and also the Chair of the proposed 2025 Mountain Iron meeting (approved by the board see below). Stacy Saari, Alli Severson, Eric Nowariak, and Aaron Hirsch were additional guests supporting the proposed Mountain Iron 71st ILSG. Institute’s Board of Directors meeting notes were taken by ILSG Secretary Hollings, which are as follows: Accepted report of the Chairs for the 69th ILSG, as published in the Proceedings volume, and minutes of last Board meeting, May, 2023 (Hollings). 2. Received, discussed, and accepted 2023-2024 ILSG Financial Summary (Jirsa). 3. Received, discussed, and accepted 2023-2024 report of the Secretary (Hollings). 4. Approved Ted Bornhorst as on-going ILSG Board member and Amy Radakovich as Chair. 5. Discussed and approved replacing Dorothy Campbell as the “member from government” on Goldich Committee (end of term 2024) with Robert Cundari. 6. Approved Mt Iron as the site for the 71st annual ILSG meeting. The meeting will be Chaired by Amy Radakovich and hosted by the Minnesota Geological Survey. 7. A number of future meeting locations were discussed. Peter Hinz has offered Kenora as a future site, while Mark Puumala has offered Thunder Bay. 8. The confusion over the appointment of the Board Chair was discussed and it was agreed we would follow the Constitution with the incoming Meeting Chair assuming the role of Board Chair. 9. It was agreed that the purchase of additional safety equipment would be postponed for now. 10. The Secretary agreed to revamp the boilerplate material for the volumes to make it easier for the organisers of subsequent meetings. Carsyn agreed to revamp the Eisenbrey and Mancuso award applications. Bornhorst agreed to rewrite the Eisenbrey award document for Board consideration. The allowable expenses will be broadened so the award will be more than travel. 11. Discussed and approved renewal of Pete Hollings as Institute Secretary (end of term 2027). This was later approved by a vote of the membership. 12. Hollings mentioned that the ILSG proceeding volumes standing order sales remain the same as the recent past with only 5 institutions receiving them plus one sent to GeoRef. 1. xxix �13. The co-Chairs would like to thank all those who helped make the 70th annual meeting a success such as judging student papers, chairing sessions, leading field trips, driving for field trips, staffing the registration desk, caring for the projectors, general logistics and more. A special thank you goes to Julie Stark, who played a key role in online and onsite registration. The 70th ILSG was a milestone for a professional organization, as noted by Pete Hollings in a recently published article on ILSG in the Lake Superior Magazine - “not a lot of groups hang around 70 years.” Forty years ago, Ted Bornhorst chaired the annual meeting and Board of Directors. At this time the board had serious concerns about the survival of the organization. We are happy to report that ILSG continues to thrive and has done so by being a small, but vibrant organization. We believe that the combination of collegial, friendly, and open discussion and exchange of ideas on geology of the Lake Superior region between government, industry, and academic geologists has played a major role in ILSG’s survival for 70 years. We strongly believe that field relations are the foundation of geologic interpretation. The depth, breadth, and quality of ILSG field trips is another reason ILSG continues to thrive. What makes ILSG field trips special is that trip leaders are open to debate on their interpretation of an outcrop. Open - but not competitive - discussion is a hallmark of both ILSG field trips and technical sessions. Lastly, meetings would not be possible without people willing to serve as chair or co-chair and people willing to organize the annual conference, to lead field trips, and to serve on local committees. Chairing an ILSG meeting involves personal time, extra work, and a bit of extra stress as attested to by anyone who has risen to this challenge in past years. One of us (Bornhorst) has been principal chair for 6 meetings over 41 years, from 1983 to 2024. He agreed to be Chair one last time to mentor Erika and Patty with the hope that one day, one or both of them, will chair a future annual meeting, contributing to the continuation of ILSG. We hope that ISLG survives for many decades and into the next century and beyond. We are gratified by the positive comments by participants and are happy to have served the Lake Superior geological community. We look forward to the 2025 Mountain Iron ILSG meeting when we can be much more relaxed! Respectfully submitted, Theodore (Ted) Bornhorst, Erika Vye, and Patrice (Patty) Cobin Co-chairs, 70th Institute on Lake Superior Geology xxx �Donations to Support the Annual Meeting of the Institute on Lake Superior Geology A special thank you to our individual contributors Roger Anderson Allan MacTavish Dave Dahl xxxi �Donations to Support Student Participation at the Annual Meeting of the Institute on Lake Superior Geology A special thank you to our individual contributors Kate Clover Jim and Isagel DeGraff Tom Erickson Tom Fitz Aaron Hirsch Paula Leier-Engelhardt Bob Mahin Vince and Susan Mathews Jim Miller Allison Severson Mark and Lauri Severson John Verhoeven Gerry White xxxii �TECHNICAL PROGRAM xxxiii �Wednesday May 14, 2025 All field trips begin and end at the Mountain Iron Community Center Pre-meeting Field Trips May 14, 2024 8:00 am - 5:00 pm PRE-MEETING FIELD TRIPS Trip 1: Transect through the Quetico subprovince of northern Minnesota – Eric Nowariak (Minnesota Geological Survey), Mark Jirsa (Minnesota Geological Survey, retired) Trip 2: Drill Core from three Cu-Ni deposits of the Duluth Complex - Mark Severson (Natural Resources Research Institute, Teck Retired), Cullen Phillips (New Range Copper Nickel), Kevin Boerst (Twin Metals Minnesota) Trip 3: How do you make iron and/or manganese in Proterozoic Iron Formation? - Alex Steiner (Big Rock Exploration), Latisha Brengman (University of Minnesota, Duluth), Dean Peterson (Big Rock Exploration) Trip 4: New Geological Insights into the genesis of iron ores at Lake Vermillion – Soudan Underground Mine State Park - George J. Hudak (University of Minnesota, George Hudak Geosciences P.L.L.C.), Zsuzsanna P. Allerton (University of Minnesota), Annia Fayon (University of Minnesota) Wednesday evening May 14, 2025 5:00 pm - 8:00 pm Registration (Mountain Iron Community Center) 6:00 pm - 8:00 pm Poster Setup and Viewing (Mountain Iron Community Center) 6:00 pm - 8:00 pm Welcoming Reception (Mountain Iron Community Center) * 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 the paper if different than the first author. Thursday - May 15, 2025 7:15 am – 12:00 pm 8:00 am. Registration (Mountain Iron Community Center) Opening remarks (Mountain Iron Community Center) Amy Radakovich, Allison Severson, Eric Nowariak, Stacy Saari, Aaron C. Hirsch Co-Chairs, 2025 ILSG xxxiv �TECHNICAL SESSION I – ORAL PRESENTATIONS Session Chair: Amy Radakovich 8:20 Mark SMYK Robert Bell - Pioneer of Lake Superior geology 8:40 William J. HINZE and Mark B. LONGACRE Revisiting Gravity and Magnetic Anomalies of the Baraboo Range 9:00 Huifang XU and Tianyu ZHOU Battle between the bands: competitive precipitations lead to bands in banded iron formations 9:20 Howard MOOERS, Mark SEVERSON, Peter JONGEWAARD, and Phillip LARSON US Steel Corporation / Ralph W. Marsden iron ore collection 9:40 Matt CARTER Updates on the Minnesota Department of Natural Resource’s Drill Core Library 10:00 END OF TECHNICAL SESSION I 10:00-10:20 COFFEE BREAK TECHNICAL SESSION II – ORAL PRESENTATIONS Session Chair: Stacy Saari 10:20 Alan AUBUT A Contrarian View: Thoughts on the Genesis of the Tamarack Ni-Cu Deposit 10:40Cory PALIEWICZ and Joyashish THAKURTA Lithogeochemical Characterization of Manganese Mineralization at the Cuyuna Range, Central Minnesota 11:00 Guy N. EVANS and William E. SEYFRIED JR. Experimental Reproduction of Acidic Mafic-Ultramafic Hydrothermal Fluids with Implications for Linking Seafloor Lithology to Ore Mineral Solubility and Novel Geochemical Trapping Mechanisms 11:20 Wyatt BAIN, James TOLLEY, and Peter HOLLINGS An overview of the geology, tectonic setting, and occurrence of sulphide mineralization in the Lac Des Iles Intrusive Suite 11:40 Thomas BUCHHOLZ, Alexander FALSTER, and William SIMMONS A complex F-rich alkalic pegmatite in the pyroxene syenites of the stettin complex, Wausau Complex, Marathon county, Wisconsin 12:00 END OF TECHNICAL SESSION II xxxv �12:00-1:30 LUNCH BREAK and ILSG BOARD OF DIRECTORS MEETING - Buffet lunch provided- TECHNICAL SESSION III- POSTER PRESENTATIONS Session Chair: Robert Lodge 1:30-3:00 AUTHORS PRESENT AT THEIR POSTERS 2:40-3:00 COFFEE BREAK 3:00 END OF TECHNICAL SESSION III TECHNICAL SESSION IV – ORAL PRESENTATIONS Session Chair: Allison Severson 3:00 James V. JONES, Ross SALERNO, William F. CANNON, and Pau O’SULLIVAN Geologic implications of detrital zircon U-Pb ages from Archean and Paleoproterozoic strata in central Minnesota and the Gogebic Range of Wisconsin and Michigan, USA 3:20 R. SALERNO, W.F. CANNON, A. SOUDERS, J.M. THOMPSON, and J. VERVOORT Constraining the timing of crustal exhumation following the Penokean orogeny using U-Pb, SmNd, and Lu-Hf geochronology and microstructural analysis 3:40 James DeGRAFF, Chad DEERING, and James JONES III The Archean Carney Lake gneiss complex in Michigan’s Upper Peninsula: Preliminary subdivisions with age constraints 4:00 *Omar Khalil DROUBI, Erik SCHOONOVER, Mona-Liza SIRBESCU, Joshua GARBER, and Chlo BONAMICI Geochronology of lithium mineralization in the Florence pegmatite field, WI, USA 4:20 END OF TECHNICAL SESSION IV xxxvi �Thursday evening May 15, 2024 5:30 pm RECEPTION AND CASH BAR (Mountain Iron Convention Center) 6:30 pm ANNUAL BANQUET (Mountain Iron Convention Center) 2025 Goldich Medal Recipient: Robert Michael Easton Banquet Speaker: Pete Kero, Mine to Mountain Bike Mecca: The story of the Redhead Mountain Bike Park Friday - May 16, 2025 8:15 INTRODUCTORY REMARKS AND UPDATES (Mountain Iron Community Center) Amy Radakovich, Allison Severson, Eric Nowariak, Stacy Saari, Aaron C. Hirsch; Co-Chairs, 2025 ILSG TECHNICAL SESSION V – ORAL PRESENTATIONS Session Chair: Aaron Hirsch 8:20 Wouter BLEEKER, Michael HAMILTON, and Sandra KAMO Paleoproterozoic mantle plume tracks shaping the southern margin of the Superior craton and the geology of the Lake Superior region 8:40 Max ROHRMAN Plume control on the initiation of Mid-Continent Rift breakup using Unconformities: Implications for the Tectono-magmatic evolution and mineral deposits 9:00 James TOLLEY, Pete HOLLINGS, Kevin MEXIA DURAN, and Myles HARDING Evaluating Ni in Olivine as a Prospectivity Indicator for Magmatic Ni-Cu-(PGE) Deposits: A Preliminary Study from the Midcontinent Rift System. 9:20 James TOLLEY, Jacob HANLEY, James CROWLEY, Sasha TSAY, Zoltan ZAJACZ, and Pete HOLLINGS A Porphyry in a Rift? Constraining the Petrogenesis of the Jogran Porphyry, Mamainse Point, Ontario, Canada: Insights from Zircon and Melt Inclusion Geochemistry 9:40 Nicholas SWANSON-HYSELL, Eben B. HODGIN, Tadesse ALEMU, Anthony FUENTES, Yiming ZHANG, Sarah SLOTZNICK, and Luke FAIRCHILD Midcontinent Rift extension ceased and the rift inverted due to the Grenvillian orogeny 10:00 END OF TECHNICAL SESSION V xxxvii �10:00-10:20 COFFEE BREAK – Sponsored by MRGS TECHNICAL SESSION VI – POSTER PRESENTATIONS Session Chair: Robert Lodge 10:00-11:30 AUTHORS PRESENT AT THEIR POSTERS 11:30 END OF TECHNICAL SESSION VI 11:30-1:00 LUNCH BREAK - Buffet lunch provided- TECHNICAL SESSION VII – ORAL PRESENTATIONS Session Chair: Eric Nowariak 1:00 Steven D.J. BAUMANN Pembine-Wausau Terrane as an Icelandic style island overthrust onto Archean basement, instead of an island arc or continental fragment accretion 1:20 Jiří1 ŽÁK, Filip TOMEK, Václav KACHLÍK, František VACEK, Martin SVOJTKA, and Lukáš ACKERMAN Broadly coeval but migrating deformation, plutonism and deposition in the northeastern Superior Province, Québec: evidence of hot accretionary orogeny and oroclinal folding in the late Archean? 1:40 Mark SMYK, Pete HOLLINGS, Riku METSARANTA, Robert CUNDARI, Stephen KISSIN, and Colleen KURCINKA Basaltic rocks of the Animikie Group in Ontario: Geochemical characteristics and tectonic significance 2:00 W. F. CANNON, M. Rebecca STOKES, Ross A. SALERNO Micromineralogy and textures in the Sudbury impact layer on the Mesabi Iron Range, Minnesota: record of processes in the proximal-distal ejecta transition zone 2:20 END OF TECHNICAL SESSION VII 2:20 COFFEE BREAK xxxviii �TECHNICAL SESSION VIII – ORAL PRESENTATIONS Session Chair: Amy Radakovich 2:40 J.D. VERHOEVEN and Tim ZOWADA Origin of magnetic black sand found on the south Shore of Lake Superior 3:00 Erika VYE and Daniel LIZZADRO-MCPHERSON Geospatial Learning Resources to Explore Relationships with Keweenaw Geology 3:20 Allan MACTAVISH, Peter HINZ, +George HUDAK, Phil LARSON, Allan AUBUT, Terry BOERBOOM, Vern CHILTON, Jim DeGRAFF, Tom ERICKSON, Barb FAULKNER, Isabel SERRANO, Larry and ZANKO An informal review of the ILSG field trip to Hawaii: January and February 2025 4:00 END OF TECHNICAL SESSION VIII 4:00 Presentation of Student Awards Best Student Paper Awards – Student award committee Student Travel/Participation Awards – Amy Radakovich MRGS Awards – Mark Severson 4:30 Concluding Remarks and Field Trips Amy Radakovich, Allison Severson, Eric Nowariak, Stacy Saari, Aaron C. Hirsch; Co-Chairs, 2025 ILSG END OF TECHNICAL SESSIONS OF THE 71st ANNUAL MEETING xxxix �Saturday May 17, 2025 Field trips begin and end at the Mountain Iron Community Center 8:00 am – 5:00 pm POST-MEETING FIELD TRIPS Trip 5: Neoarchean alkalic intrusions in the Wawa and Quetico subprovinces Terry Boerboom (Minnesota Geological Survey, retired); Amy Radakovich (Minnesota Geological Survey) Trip 6: Unique Keweenawan inclusion (Colvin Creek) in the Duluth Complex Mark Severson (Natural Resources Research Institute, Teck, retired); Allison Severson (Minnesota Geological Survey); Lauri Severson (Earth Science teacher, retired) Trip 7: Classic outcrops of Northeastern Minnesota Dean M. Peterson (Big Rock Exploration); George J. Hudak (University of Minnesota, George Hudak Geosciences P.L.L.C.) Trip 8: Glacial Lake Norwood and the Koochiching Lobe Phillip Larson (Vesterheim Geoscience PLC); Andrew Breckinridge (University of Wisconsin-Superior); Howard Mooers (University of Minnesota, Duluth) xl �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 + Denotes author that will present the paper if different than the first author. Numbered Posters and Abstracts are in sequential order 1. *Zsuzsanna ALLERTON, George HUDAK, Guy EVANS, Xinyuan ZHENG, and Christian TEYSSIER Geochemical analyses of banded iron formations and formerly mined iron ore in the Lake Vermilion-Soudan Underground Mine State Park, NE Minnesota 2. *Madelyn BANKS, Latisha BRENGMAN, Athena EYSTER Linking whole rock geochemical data with micro-scale mineral characterization of oxidation reactions in the Biwabik Iron Formation, MN, USA 3. Howard MOOERS, Mark SEVERSON, Peter JONGEWAARD, and Phillip LARSON US Steel Corporation / Ralph W. Marsden iron ore collection 4. *Sarah JAROZEWSKI, Paige DUFFY, Cole BARRÉ, Latisha BRENGMAN, and Athena EYSTER Mapping oxidation reactions in iron-rich rocks from northeast Minnesota, USA. 5. *Celia L. CORTOPASSI, Zsuzsanna P. ALLERTON, Joshua M. FEINBERG Alteration of magnetic mineralogy in the Giants Range Batholith by the Duluth Complex 6. *Samara GRIES, Robert W.D. LODGE, Sara HANEL, and Robert HOOPER Rare-element Geochemistry of the Eau Claire River Complex Pegmatites 7. *Linsey HULA and Dyanna CZECK Emplacement of the Mesoproterozoic Wausau Syenite Complex, Wisconsin 8. *Renee O. JEUTTER and Robert W.D. LODGE Geology and Geochemistry of the Mesoproterozoic Round Lake Intrusion and associated TiMineralization, Northern Wisconsin 9. *Bekah R. THOMPSON and Robert W.D. LODGE Ni-Cu-PGE Mineralization at the Mineral Lake Intrusive Complex, northern Wisconsin 10. *Lyndsie A. VICKERS and Robert W.D. LODGE Zircon Petrochronology of the Eau Claire Volcanic Complex in the Marshfield Terrane of the Penokean Orogen, Northcentral Wisconsin 11. *Andrew A. CASPER and Robert W.D. LODGE R Geology and Mineralization of the Plover Au Prospect, Marathon County, Wisconsin xli �12. William FITZPATRICK Textural and chemical analysis of sphalerite ores from the Highland Subdistrict, Upper Mississippi Valley Zinc-Lead District, Wisconsin 13. *Haley P. JOHANNESEN and Robert W.D. LODGE Geology and Geochemistry of the Ritche Creek Cu-Zn deposit, North central Wisconsin 14. *Aidan O. KWIATKOWSKI and Robert W.D. LODGE Zircon Petrochronology of Wisconsin’s Volcanogenic Massive Sulfide Deposits, Northcentral Wisconsin 15. Sara PEARSON, Nolan GAMET, Molly SHALIFOE, Ashley QUIGLEY, and Robert MAHIN Michigan Geological Survey’s Contributions to the USGS Earth MRI National Mine Waste Inventory Effort 16. Ashley K. QUIGLEY, Robert A. MAHIN, and Nolan G. GAMET Critical Mineral Potential of the Northern Margin of the Watersmeet Gneiss Dome, MI USA 17. *MaryElizabeth SHALIFOE and Peter VOICE Identifying Abandoned Mine Surficial Features Using Mask R-CNN, Upper Peninsula Michigan. 18. Sophie CHURCHLEY and Philip FRALICK Unusual early diagenetic structures in the Paleoproterozoic Gunflint Formation, Ontario, Canada 19. Gordon MEDARIS Jr. and Dave MALONE Post-Penokean and Pre-Yavapai Magmatism and Sedimentation in Central Wisconsin (Southern Lake Superior Region) 20. Esther K. STEWART, Michael TAPPA, Ann BAUER, Latisha BRENGMAN, and Anthony PRAVE Sedimentologic and geochemical evidence of marine incursion to the Oronto Group basin, southern Lake Superior region, at ca. 1.08 Ga 21. Carsyn AMES and Brad GOTTSCHALK High resolution thin-section scanning and metadata capture- WGNHS Data Preservation Project 2024 early efforts 22. Nate DANIELS, Grace MCELLISTREM, Raeann VOGEL, and Michael BRAUNAGEL Architecture of the Douglas Fault damage zone, northwest Wisconsin 23. Mark B. LONGACRE and William J. HINZE, William Geologic Interpretation of Filtered Gravity and Magnetic Anomalies of the Baraboo Range 24. Jack MALONE, David MALONE, Raymond ANDERSON, Ryan CLARK Refining the Age and Occurrence of Basement Rocks in Northwest Iowa: Implications for Precambrian Tectonics and Magmatic Evolution of the Laurentian Midcontinent xlii �ABSTRACTS xliii �Geochemical analyses of banded iron formations and formerly mined iron ore in the Lake Vermilion-Soudan Underground Mine State Park, NE Minnesota ALLERTON, Zsuzsanna1, HUDAK, George1,2,3, EVANS, Guy1, ZHENG, Xinyuan1, and TEYSSIER, Christian1 1 Earth & Environmental Sciences, University of Minnesota, Minneapolis, MN 55455, USA Earth and Environmental Sciences, University of Minnesota, Duluth, MN 55812, USA 3 George Hudak Geosciences P.L.L.C., Duluth, MN 55804, USA 2 The Lake Vermilion-Soudan Underground Mine State Park in northeastern Minnesota is known for its underground tours in the former iron mine that was operational between 1884-1962 (Klinger, 1960). The mine contains lenticular-shaped ore bodies enclosed in variably altered banded iron formations (BIFs) that were upgraded to massive hematite iron ore during replacement-style hydrothermal alteration (Gruner, 1926; Klinger, 1960; Thompson, 2015). The timing of ore mineralization is constrained to 1.8-1.6 Ga (Allerton, 2024b). The widely accepted simplified genetic model for these ore deposits involves hydrothermal fluids that leached silica from BIFs and concentrated iron as hematite. Here we utilize historic and recently acquired whole rock major, trace and rare earth element lithogeochemical analyses to perform mass balance evaluations via the isocon method (Grant, 2005) and iron stable isotope geochemistry to propose a new hydrothermal model to better constrain the transition from BIF to iron ore. Eight BIF and twelve ore samples from Thompson (2015) were utilized for this study. BIFs show varying degrees of alteration adjacent to the orebodies, whereas iron ore samples comprise massive hematite ± chlorite. Our data include eight additional samples; four least altered and two hematite-altered BIFs collected from surface outcrops, and two iron ore samples that are 1) high-grade hematite ore with primary phase microcrystalline hematite-martite (MCHMT) with minimal chlorite and 2) lower grade ore with abundant secondary quartz and microplaty hematite (MPH; Allerton, 2024b). The Fe isotope analysis incorporates variably deformed gabbroic rocks and chlorite schist adjacent to the ore bodies as well. Lithogeochemical analyses of 14 BIFs and 14 iron ore samples indicate inverse correlation between SiO2 and Fe2O3(total); BIF has high SiO2 and low Fe2O3(total) contents, whereas iron ore displays low in SiO2 and Fe2O3(total). Statistical evaluations suggest that high strength field elements (HFSE) are immobile and therefore have been selected for isocon analysis. Utilizing a HFSE ‘best fit’ isocon, the system shows almost complete SiO2-loss (99%) and 54% Fe2O3-loss from least altered BIF to high-grade ore (Fig. 1A), suggesting that greater loss of silica relative to iron has resulted in a net concentration of iron. Moreover, there is secondary quartz and MPH in the lower grade ore based on petrography, indicating the lower-grade ore postdates the high-grade ore. Isocon analysis shows SiO2-gain and continued Fe3O2-loss from high-grade to lower grade ore (Fig. 1B). The Fe stable isotope results attest to this by presenting higher δ56Fe values for BIFs that decrease from less to more altered BIF. MCH-MT in highgrade ore displays even lower δ56Fe values, and chlorite within fractures of high-grade ore shows similar values to secondary quartz and MPH in heterogeneous ore, suggestive of paragenesis of two different hematite phases and gangue minerals (Fig. 2). Our new hydrothermal model proposes continuous removal of Fe, entailing coeval Si mobilization and removal from BIF to high-grade MCH-MT ore and re-deposition into lower grade MPH ore. These hypotheses are supported by detailed analyses of lithogeochemistry, mineral textures, and Fe stable isotopes. 1 �Figure 1: A) Diagram displays major oxides of least altered BIF (x-axis) against MCH-MT ore (yaxis) and B) MCTMT ore (x-axis) against MPH ore (y-axis). Ratios show gains and losses are calculated based on the slope of HFSE best fit isocons. Figure 2: Diagram shows decreasing δ57Fe/ δ56Fe values of less and more BIFs, MCH-MT and MPH ore samples, and lithologies adjacent to the ore bodies in Soudan; foliated gabbro, gabbroderived schist, chlorite schist. Values are calibrated to BHVO-2 iron standard commonly used in Fe stable isotope geochemistry. REFERENCES Allerton, Z., Hudak, G., Teyssier, C., Fayon, A., Daniŝik, M., Courtney-Davies, L, and Larson, P., 2024b. Geochronology and geochemistry of hematite ore in northeastern Minnesota: Institute on Lake Superior Geology, Proceedings Volume 70, Part 1 – Program and Abstracts, p. 4-5. Grant, J.A., 2005. Isocon analysis: A brief review of the method and applications: Physics and Chemistry of the Earth, Parts A/B/C, v. 30, p. 997–1004, doi: 10.1016/j.pce.2004.11.003. Gruner, J. W., 1926. Hydrothermal alteration of iron ores of the Lake Superior type—a modified theory: Economic Geology, v. 32, p.121-130. Klinger, F.L., 1960. Geology and ore deposits of the Soudan mine, St. Louis County, Minnesota [thesis]. Thompson, A., 2015. A hydrothermal model for metasomatism of Neoarchean Algoma-Type banded iron formation to massive hematite ore at the Soudan Mine, NE Minnesota [thesis]. 2 �High resolution thin-section scanning and metadata capture- WGNHS Data Preservation Project 2024 early efforts AMES, Carsyn1 and GOTTSCHALK, Brad1 1 Wisconsin Geological and Natural History Survey, UW-Madison, 3817 Mineral Point Rd, Madison, WI 53704 USA The Wisconsin Geological and Natural History Survey (WGNHS) has recently undertaken an effort to scan approximately 3800 of the 4800 historical thin sections held in WGNHS collections as part of the USGS-National Geological and Geophysical Data Preservation Program (NGGDPP). This work builds upon a number of previous projects including: a 2011 NGGDPP project to inventory all thin sections in the WGNHS collections, an internal project to catalog fields notebooks and refine locations of recorded samples, a pilot study to develop a workflow for scanning and editing high resolution photos of thin sections, and a project to inventory and collect metadata from an extensive collection of samples donated to WGNHS by Gene LaBerge (UW-Oshkosh). Building on the lessons learned from these prior studies and methods outlined in Leung and Mcdonald (2023), we have developed a workflow to scan thin sections using a Plustek OpticFilm 8200i film scanner (Figures 1a and 1c) and SilverFastSE Plus software with settings shown in Figure 1b. Forty-eight-bit raw images are produced in both plane, non-polarized light and cross-polarized light (Figure 2). Images are edited post scanning in Adobe Lightroom to enhance the sharpness and exposure to better replicate what users see when viewing thin sections with a petrographic microscope. Scanning and editing images takes approximately 10 minutes per thin section. Photos are stored in TIFF format and are intended to be served on the WGNHS Dataviewer for public access. Thin sections included in this project capture a wide range of lithologies from several Wisconsin counties. Many samples represent some of the first efforts to survey the natural resources and map the geology of northern Wisconsin. The original data associated with the thin sections is archived in historic field notebooks archived at the WGNHS and includes documentation of geomorphology, bedrock and glacial geology, and magnetic susceptibilities of encountered bedrock units. Locations are recorded in Public Land Survey System (PLSS) notation. Samples with at least section level location information were included in this project; many of the locations given in the field notes can be narrowed down to quarter-section designation with certainty. In the initial phase of this ongoing project we have focused on scanning and entering metatdata for samples in and around Florence County, Wisconsin. Precambrian iron formation in this area was mined from 1880-1931 to produce some three million tons of hematite and limonite ore (Brown B., 2021). This project has focused on capturing lithological information from samples in this area, which is characterized by complex Precambrian stratigraphy and structure. Upon project completion, high resolution thin section images will be made publicly available online using the WGNHS Dataviewer. Additionally, all metadata will be uploaded to the USGS’s ReSciColl collection and the WGNHS internal database (Geobase). 3 �Figure 1: A) Plustek Optic Film 8200i scanner and acompyning film tray. B) Scanner settings to be used during the proposed project. Note the 600 ppi preset and further 7,200 ppi adjustable resolution. Thin sections are scanned in 48-bit HDR Raw C) Tray with card stock paper cut to better hold thin sections. Note the two slots on the right are fitted with linear polarizing screens that sandwich the thin sections. The polarized screens are oriented to cross polarize the light when scanning. Figure 2: High resolution thin section images scanned as part of the pilot project. A. and C. were scanned using plane, non-polarized light; B. and D. were scanned using cross polarized light. REFERENCES Brown, B., 2021. Florence Iron Mine: Historical Maps Showing Location of Surface Development, Regional Setting, and Underground Workings. Wisconsin Geological and Natural History Survey WOFR2018-03: 5. Leung, D. D.V., and Mcdonald, A.M., 2023. Picture-perfect petrography: affordable thin-section scanning for geoscientists in the digital era. The Canadian Journal of Mineralogy and Petrology, 61: 10451050. 4 �A Contrarian View: Thoughts on the Genesis of the Tamarack Ni-Cu Deposit. AUBUT, Alan1 1 Sibley Basin Group Ltd., PO Box 304, Nipigon, ON P0T 2J0.Canada The Tamarack Ni-Cu deposit has been attributed as being of intrusive origin (Goldner, 2021; Taranovic et al., 2018). There are many nickel deposits hosted by ultramafic bodies that display clear evidence of being the product of extrusive flows, often exhibiting the same key features used to invoke an intrusive origin (e.g. Arndt, 1975; Hill et al., 1995; Hubbert and Sparks, 1985; Marston et al.,1981). This includes the nickel deposits of the Kambalda district of Australia, Pechenga in the Kola Peninsula of western Russia, Raglan in northern Quebec and Thompson in northern Manitoba. All have been, or currently are, attributed to the intrusion of ultramafic sills (e.g. Bleeker, 1990; Marston et al., 1981; Melezhik et al., 1994). Key evidence in support of this model is that the ultramafic bodies typically exhibit at least some differentiation and are sub-concordant to the host sediments. This tendency to default to an intrusion model now includes the Tamarack deposit in Minnesota even though an extrusion model is more valid. The major komatiite hosted nickel deposits listed above share common features: 1) the nickel mineralisation is hosted by ultramafic rocks; 2) the sulphides are at the stratigraphic base of the host ultramafics; 3) the ultramafic rocks are hosted by, or in contact with, sulphidic and carbonaceous argillaceous rocks; 4) the ultramafic bodies are stratabound and generally conformable to the host lithology; and 5) they are hosted within extensional basins usually with a significant sedimentary component with Kambalda being the one exception. But there is a density “problem” in that ultramafic magmas are typically denser than the host rocks, especially when they are sedimentary. When rocks melt, they become about 10% less dense. In the case of ultramafic rocks, the average density is about 3.0 g/cc (Nisbet et al., 1993) while the crust has a density of 2.7 g/cc or less. To move upward from the mantle through the crust there must have been a mechanism other than buoyancy. “Overpressure” is a valid explanation (Sleep, 1974, 1992). Magma plumes in a mantle plume move upward due to buoyancy to the Mantle-Crust boundary. There it collects and then moves laterally thus creating extensional forces in the overlying crust. This accumulating magma would be constrained by the overlying lithostatic load and in doing so would build up overpressure. If the crust thins enough vertical fractures can form allowing the trapped magma to escape due to the built-up overpressure exceeding the lithostatic load. At surface the hot, dense ultramafic magma would then flow over, and into, deep water sediments where the magma would mechanically and thermally erode and assimilate sulphide rich sediments. Tamarack shows all the same characteristics as other Ni-Cu deposits associated with rift basins and features that are more easily explained by extrusive flow of komatiitic magma. As such the intrusive emplacement model currently favoured should be reviewed and serious consideration given to emplacement by extrusion of a high-density magma driven by overpressure. 5 �REFERENCES Arndt, N.T., 1975. Ultramafic rocks of Munro Township and their volcanic setting; Unpub. Ph.D. Thesis, Univ. Toronto. Bleeker, W., 1990. New Structural-Metamorphic constraints on Early Proterozoic oblique collision along the Thompson Nickel Belt, Manitoba, Canada; In Lewry, J.F. and Stauffer, M.R., eds., The Early Proterozoic Trans-Hudson Orogen of North America: Geological Association of Canada, Special Paper 37, p. 57-73. Goldner, B.D., 2011. Igneous Petrology of the Ni-Cu-PGE Mineralized Tamarack Intrusion; Unpub. M.Sc. Thesis, Univ. Minesota. Aitkin and Carlton Counties, Minnesota; Canadian Journal of Earth Sciences, 44, 1087-1110. Hill, R.E.T., Barnes, S.J., Gole, M.J. and Dowling, S.E., 1995. The volcanology of komatiites as deduced from field relationships in the Norseman-Wiluna greenstone belt, Western Australia; Lithos 34, p. 159-188. Huppert, H.E. and Sparks, R.S.J., 1985, Komatiites I: Eruption and Flow; Journal of Petrology, Vol. 26, Part 3, pp. 694-725. Marston, R.J., Groves, D.I., Hudson, D.R. and Ross, J.R., 1981, Nickel sulfide deposits in Western Australia: a review; Economic Geology, Vol. 76, pp. 1330-1363. Melezhik, V.A., Hudson-Edwards, K.A., Skuf'in, P.K and Nilsson, L.P., 1994a, Pechenga Area, Russia Part 1: geological setting and comparison with Pasvik, Norway; Transactions of Institution of Mining and Metallurgy (Sect. B: Applied Earth Science), Vol. 103, p B129-B145. Nisbet, E. G., Cheadle, M. J., Arndt, N. T., & Bickle, M. J. (1993). Constraining the potential temperature of the Archaean mantle: a review of the evidence from komatiites. Lithos, 30(3-4), 291-307. Sleep, N. H., 1974. Segregation of Magma in the Ascending Mantle. The Journal of Geology, 82(2), 131– 142. Sleep, N. H., 1992. Time Dependence of Kilauea Volcano Structure from Hotspots to Trench Due to Overpressure in the Asthenosphere. Journal of Geophysical Research: Solid Earth, 97(B8), 11773– 11782. Taranovic, V., Ripley, E.M., Li, C. and Shirey, S.B., 2018. S, O, and Re-Os Isotope Studies of the Tamarack Igneous Complex: Melt-Rock Interaction During the Early Stage of Midcontinent Rift Development; Economic Geology, v. 113, no. 5, pp. 1161-1179. 6 �An overview of the geology, tectonic setting, and occurrence of sulphide mineralization in the Lac Des Iles Intrusive Suite BAIN, Wyatt1, TOLLEY, James 2, and HOLLINGS, Peter 2 1 Department of Earth Sciences, Western University, 1151 Richmond St, London, ON N6A 5B7 Canada 2 Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada The Lac des Iles (LDI) mafic-ultramafic complex hosts a world-class platinum group element (PGE) deposit and is spatially associated with a suite of mafic-ultramafic satellite intrusions (i.e. the LDI-intrusive suite; LDI-IS). The intrusions are hosted in the crystalline rocks of the Wabigoon subprovince, along its eastern contact with the sedimentary and volcanic rocks of the Quetico subprovince. Previous work identified textural and geochemical similarities between the LDI-IS and the mineralized rocks of the LDI complex that likely reflect a temporal and genetic association, and perhaps a similar degree of prospectivity for PGE mineralization (Stone et al., 2003). Here, we present an overview of the geology and setting of the LDI-IS, as well as new geochronology, isotopic data, and parental melt modelling. The LDI-IS (Tib Lake, Legris Lake, Wakinoo Lake, Demars Lake, Dog River, Taman Lake, and Buck Lake; Fig. 1a) are mostly leucogabbro to gabbronorite in composition but commonly include hornblende gabbro, hornblendite, and minor peridotite and pyroxenite. Zircon U-Pb ages for mineralized gabbro from the Buck Lake (2698.1 ± 1.6 Ma), Wakinoo Lake (2696.6 ± 0.8 Ma), Demars Lake (2694.1 ± 1.5 Ma), Legris Lake (2690.6 ± 0.8 Ma), Dog River (2689.9 ± 0.7 Ma), and Tib Lake (2685.9 ± 1.6 Ma) intrusions show a spatial trend of younging to the north and demonstrate a temporal association with the Lac des Iles Mine Block intrusion (2689.0±1.0 Ma; Stone, 2010; Fig 1 b). Trace element profiles for modelled parental melts are similar across most of the LDI-IS and are consistent with an arc setting and a common parental magma source reservoir. However, modelled REE profiles for some cyclic units in the Tib lake intrusion were more evolved and enriched in light rare earth elements. Similar patterns are reported in modelled parental melts from North LDI and are consistent with mixing between primitive and more evolved, siliceous magmas (Djon et al., 2017). Though magma mixing influenced the geochemical evolution of the Tib lake intrusion, cyclic units with more evolved signatures were not significantly mineralized. Whole rock εNdT values of gabbroic rocks from the LDI-IS and the Lac des Iles complex overlap with the tonalitic rocks of the Wabigoon subprovince in older intrusions and trend toward increasingly negative values in younger intrusions (Fig. 1c). This suggests assimilation of Wabigoon tonalite by LDI-IS parental magmas early in the formation of this magmatic system, and greater degrees of contamination by Quetico metasediment over time. Magmatic sulphides from the Legris Lake intrusion have δ34S values that overlap the mantle range but trend toward the composition of Wabigoon tonalite (Bain et al., 2023). This suggests that external S or Si addition from the tantalite drove sulphide saturation during its formation. However, a comparison of whole rock S/Se and Cu/Pd ratios of mineralized lithologies across the LDI-IS suggest that sulphide melt retention during emplacement was a more crucial control on the occurrence of PGE-bearing sulphide mineralization than the source of S or the timing of sulphide saturation. 7 �Figure 1: a. Regional geologic map showing locations of Thunder Bay, the Lac des Iles mine (in red), and the Lac des Iles intrusive suite (in blue). b. U-Pb ages for individual intrusions in the Lac des Iles intrusive suite. c. Whole-rock εNdT values for the LDI intrusive suite and host rock lithologies. North LDI and South LDI data from Brügmann et al. 1997 REFERENCES Bain, W.M., Hollings, P.N., Djon, M.L., Brzozowski, M.J., Layton-Matthews, D., Dobosz, A., and Stern, R.A., 2024. Geochemical evolution and parental magma of the Lake Legris mafic-ultramafic complex, Ontario. Mineralium Deposita 59:85-108 Brügmann, 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: 223−239. Djon, M.L., Olivo, G.R., Miller, J.D., and Peck, D.C., 2017. Stratiform platinum-group element mineralization in the layered northern ultramafic center of the Lac des Iles Intrusive Complex, Ontario, Canada. Ore Geology Reviews, doi: 10.1016/j.oregeorev.2017.03.011. Stone, D., Lavigne, M.J., Schnieders, B., Scott, J., and Wagner, D., 2003. Regional geology of the Lac des Iles area. Ontario Geological Survey, Open File Report 6120: 15–25. Stone, D. 2010. Precambrian geology of the central Wabigoon Subprovince area, northwestern Ontario. Ontario Geological Survey, Open File Report 5422:1-130. 8 �Linking whole rock geochemical data with micro-scale mineral characterization of oxidation reactions in the Biwabik Iron Formation, MN, USA BANKS, Madelyn1, BRENGMAN, Latisha1, EYSTER, Athena2 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Heller Hall, 1114 Kirby Drive, Duluth, MN 55812, USA 2 Department of Earth and Climate Sciences, Tufts University, Lane Hall, 2 North Hill Road, Medford, MA 02155, USA Oxidation and hydration reactions in iron-rich chemical sedimentary rocks are of critical interest because they signify post-depositional changes often linked to later weathering and fluid alteration. Evaluating oxidation and hydration reactions present in iron formations is therefore required to separate out depositional signals in mineralogical and geochemical data from those that link to post-depositional mineral reactions and enrichment processes (Geymond et al., 2022). The Biwabik iron formation is a part of a well-preserved, sub-greenschist lithologic assemblage containing three major meta-sedimentary formations known together as the Animike Group (e.g. Severson, 2009 and references therein). Previous work (e.g. Duncanson et al., 2024 and references therein) demonstrated the preservation of numerous mineral reactions in the Biwabik iron formation, making it an ideal location to test how mineral reactions link directly to whole rock geochemical signals. To evaluate the relative timing of different oxidation and hydration reactions and how they link to whole rock geochemical data, we integrate core, petrographic, and scanning electron microscope observations with whole rock digestion ICP-MS geochemical datasets from core LWD-99-01 (n = 60) of the Biwabik iron formation. Two key oxidation reactions identified in this work include (1) magnetite to hematite and (2) carbonate to magnetite. Mineral reactions are documented by cross-cutting relationships (Figure 1A-D). The mineral reaction of magnetite to hematite (possibly via the recrystallization of metastable maghemite, 2(αFe3O4) + H2O ↔ 3(γFe2O3) + H2); Geymond et al., 2023) is present in all four informal lithologic subunits of the Biwabik iron formation, occurring in 48% (n = 23) of samples (n = 48) across these units. The mineral reaction of carbonate to magnetite (3FeCO3 +H2O → Fe3O4 + 3CO2 + H2, Duncanson et al., 2024) is also present in all four informal lithologic subunits of the Biwabik Iron Formation, occurring in 77% (n = 37) of samples (n = 48) across these units. Based on 64 EDS point analyses of 4 representative samples from each subunit of the Biwabik iron formation, dominant carbonate minerals range from siderite at the base of the stratigraphy, to ankerite, dolomite, and calcite towards the top. Combined, carbonate compositional variability and zonation indicate element exchange during multiple generations of post-depositional fluid alteration, and cross-cutting relationships between carbonate-magnetite, and magnetite-hematite indicate post-depositional oxidation via fluid interaction with pre-existing reduced iron phases. Dissolution of carbonate may have created porosity providing pathways for oxidizing fluids, and further oxidation. Despite these later oxidation reactions, whole rock geochemical data preserves lithology specific signals of oxic vs. anoxic conditions, independent of the presence of the post-formational reactions outlined above. Lower stratigraphic units preserve oxic signals even with ferrous iron phases like siderite and greenalite preserved, while upper stratigraphic units preserve anoxic signals, despite the presence of hematite. Overall, bulk geochemical data from lithologic subunits of the Biwabik Iron Formation do not preserve clear signals associated with post-depositional mineral assemblage modification and oxidation documented by detailed petrographic work. 9 �Figure 1: LWD-99-01 reflected light photomicrographs documenting cross-cutting relationships between mineral phases. A. Upper Slaty sample MIR-17-15 carbonate granule cross-cut by euhedral magnetite (mag) in 20x. B. Lower Cherty sample MIR-19-14 carbonate granule (carb) cross-cut by euhedral magnetite (mag) in 5x. C. Upper Cherty sample MIR-19-15 magnetite crystal (mag) cross-cut by platy hematite (hem) in 20x. D. Lower Slaty sample U-05 magnetite crystals (mag) crosscut by platy hematite (hem) at the edge of a silicate granule in 10x. REFERENCES Duncanson, S., Brengman, L., Johnson, J., Eyster, A., Fournelle, J., Moy, A., 2024. Reconstructing diagenetic mineral reactions from silicified horizons of the Paleoproterozoic Biwabik Iron Formation, Minnesota. American Mineralogist, 109, 339-358. Geymond, U., Briolet, T,. Combaudon, V., Sissmann, O., Martinez, I., Duttine, M., Moretti, I., 2023. Reassessing the role of magnetite during natural hydrogen generation. Front. Earth Sci., 11, 1169356. Geymond, U., Ramanaidou, E., Lévy, D., Ouaya, A., Moretti, I., 2022. Can Weathering of Banded Iron Formations Generate Natural Hydrogen? Evidence from Australia, Brazil and South Africa. Minerals, 12, 163. Severson, M., Heine, J., Patelke, M., 2009. Geologic and Stratigraphic Controls of the Biwabik Iron Formation and the Aggregate Potential of the Mesabi Iron Range, Minnesota. University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR- 2009/09, 173, 37 plates. 10 �Pembine-Wausau Terrane as an Icelandic style island overthrust onto Archean basement, instead of an island arc or continental fragment accretion BAUMANN, Steven D.J. Midwest Institute of Geosciences and Engineering Since at least the 1960s, we have thought of the Pembine-Wausau Terrane (PWT) as an island arc or continental fragment accretion, smashed between the Superior Craton to the north and the Marshfield Terrane to the south. We all have seen a fault zone appear on geologic maps of the border between the Upper Peninsula of Michigan and northeast Wisconsin called the Niagara Fault Zone (NFZ). There is only one major problem, no one has ever found the NFZ. It doesn’t outcrop anywhere, it does not appear in well records, nor clearly on gravity maps, or magnetic maps. Often where it is inferred it can be interpreted other ways. And the NFZ isn’t reflected in any smaller chronostratigraphically equivalent structures that do outcrop. I have found white unbaked quartzite pebbles (the Sturgeon Quartzite) north of the NFZ (fig. 1). The host rock of these pebbles according to maps, are metamorphic rocks that supposedly have an igneous protolith. I find that very hard to reconcile with present modeling. As I have looked at the highly deformed rocks of the Florence Wisconsin, Iron Mountain Michigan, and Norway Michigan areas, I have come to the conclusion that many rocks mapped as metaigneous, are in fact, metasedimentary. I have been working on a local cross section for several years with my observations. I am coming to the conclusion that the mafic rocks and metasediments to the north of where the NFZ has traditionally been mapped, are more or less continuous and correlative to the mafic and metasedimentary rocks to the south of it. Interpretation of the rocks is understandably very difficult as the rocks are highly metamorphosed and deformed. This work is preliminary. My interpretations could change. But this is where the evidence is leading me thus far. So, if the area that is mapped as the NFZ is not a fault zone, what is it? I see it as one of two possibilities. It could be more of a shear zone formed from a more lateral accretion of a volcanically active, partially rifted Archean sliver, similar in appearance to Baja California. Sheering would be hard to see expressed in the rocks, just as it is for other covered shear zones further north. The second possibility is that the PWT was originally an Icelandic style island on a spreading center that would eventually become subducted under the Superior Craton, similar to the East Pacific Rise, before subduction switched to the south as the Marshfield Terrane approached. Its suspected Archean basement could be explained by a thin skinned over thrusting of the PWT over a small sliver of Archean crust, while volcanism was ongoing. The age of the xenocryst zircons expected to be Archean are only 2,607+22 Ma (VanWyck and Johnson. 1997). This is similar to many Archean ages of the Superior Craton. It is still a possibility the Penokean was a continental fragment like the Marshfield Terrane, only far more incomplete and still covered with younger deposits, but this cannot be the default without understand the nature of the NFZ, if it even exists. The Archean basement of the PWT could also be some sort of an extension of nearly in situ Superior Craton, that hosted the PWT as it formed, or it was overridden by the PWT. I am currently favoring the second interpretation. In this case no NFZ is needed to explain anything observed, at least locally. Everything can be explained by dominantly ductile deformation, at least in the upper crust. This is really reflected in the rocks at Piers Gorge and in the local Michigamme Formation, which locally do not express any Penokean aged faults of any 11 �significance. It also would explain the contemporaneous crustal thinning to the east in the Sudbury area if we had a subducting rift. This is something that forearc extension and island arc accretion cannot explain on their own. This would also put the continental suture further south, at the Eau Claire Sheer Zone. Figure 1: Adapted from Baumann, 2021 REFERENCES Baumann, S.D.J., 2021. The Misunderstood Penokean Orogeny. Midwest Institute of Geosciences and Engineering, publication G-102021-1A VanWyck, 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. GSA Bulletin; July 1997; v. 109; no. 7; p. 799–808; 8 figures, 2 tables 12 �Paleoproterozoic mantle plume tracks shaping the southern margin of the Superior craton and the geology of the Lake Superior region BLEEKER, Wouter1, HAMILTON, Michael 2, and KAMO, Sandra 2 1 Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8, Canada; wouter.bleeker@canada.ca 2 Jack Satterly Geochronology Laboratory, Department of Earth Sciences, University of Toronto, 22 Ursula Franklin Street, Toronto, ON M5S 3B1, Canada All Archean cratons are fragments of late Archean “supercratons”, i.e. the larger landmasses to which these craton fragments trace their origin (Bleeker, 2023). At least two large independent supercratons, Superia and Sclavia, named after their well-preserved internal fragments, had formed by the late Archean and underwent progressive breakup during the early Paleoproterozoic, from ca. 2.2 Ga to 1.9 Ga. Based on well-populated apparent polar wander paths, these supercratons moved independently; hence, the mythical notion of a single, longlived, late Archean supercontinent “Kenorland” is incorrect, aside from being untestable. Craton fragments can be correlated and put back together again by matching distinctive basement geology, by correlating overlying pre-breakup basin stratigraphies, and by correlating remnants of pre- to syn-breakup large igneous provinces, particularly their dyke swarms (Bleeker and Ernst, 2006). Of all the dispersed Archean craton fragments, more than 10 trace their origin back to supercraton Superia, representing “nearest neighbour” fragments to the Superior: Karelia, Kola, Wyoming, Hearne, Kaapvaal, Pilbara, Yilgarn, Zimbabwe, North Atlantic craton, and possibly Dharwar. Kaapvaal-Pilbara joined a growing Superia late in the game, at ca. 2650 Ma, separating again ~600–700 Myr later, leaving the ancient Minnesota River Valley terrane behind. With a robust reconstruction of Superia, numerous other important insights follow, including that of ancient mantle plume tracks (Figure 1). Here we discuss evidence for two major plume tracks that shaped the southern margin of the Superior craton, the 2480-2440 Ma “Matachewan” plume track, and the 2125-2050 Ma “Marathon” plume track. Both these plume tracks started within Superia’s core, before crossing over to then-contiguous crust of “greater Karelia” and Kaapvaal, respectively. The Matachewan plume track was initiated in the Sudbury area with a suite of 2480-2472 Ma mafic layered intrusions emplaced at the base of the Huronian Supergroup. It then triggered the ca. 2461 Ma giant Matachewan dyke swarm, which converge to a magmatic centre well to the south. At ca. 2450 Ma, the plume crossed over to then-contiguous “greater Karelia” (Davey et al., 2020) where it spawned additional dyke swarms and a flare-up of large layered intrusions, some as young as 2440 Ma. The much younger Marathon plume was initiated at ca. 2125 Ma with a giant radiating mafic dyke swarm, the Marathon dykes, with a focal point in the eastern Lake Superior area. It then spawned progressively younger mafic dyke swarms to the southwest before crossing over to the contiguous Kaapvaal craton where it spawned carbonatites at 2060 Ma, and finally the emplacement of the Bushveld Complex at 2056 Ma, the long axis of which is aligned with the plume track (Figure 1). Both plume tracks show well-defined age progressions indicating plate velocities of ~1–5 cm/yr. SOME REFERENCES Bleeker, W., 2003. The late Archean record: a puzzle in ca. 35 pieces. Lithos 71(2-4): 99-134. 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: Hanski, E., Mertanen, S., 13 �Rämö, T., Vuollo, J. (Eds.) Dyke Swarms—Time Markers of Crustal Evolution, AA Balkema, Rotterdam, p. 3-26. Davey, S.C., Bleeker, W., Kamo, S.L., Vuollo, J., Ernst, R.E., and Cousens, B.L., 2020. Archean block rotation in Western Karelia: Resolving dyke swarm patterns in metacraton Karelia-Kola for a refined paleogeographic reconstruction of supercraton Superia. Lithos 368: 105553. Fiorentini, M.L., O’Neill, C., Giuliani, A., Choi, E., Maas, R., Pirajno, F., and Foley, S., 2020. Bushveld superplume drove Proterozoic magmatism and metallogenesis in Australia. Scientific Report 10(1): 19729. Figure 1. Paleogeographic reconstruction of late Archean–early Paleoproterozoic supercraton Superia, involving >10 of the better-known Archean craton fragments from around the world, with the wellpreserved Superior craton as its signature internal fragment. Vaalbara and several other cratons (e.g., Wyoming) formed a single, large, ancient superterrane that collided with the southern margin of growing Superia at ca. 2650 Ma. After a period of stasis, supercraton Superia underwent progressive rifting and breakup from ca. 2.2 Ga to 1.9 Ga. Selected Paleoproterozoic mafic magmatic events are shown, with a focus on two well-defined mantle plume tracks, the “Matachewan” plume track (bold grey arrow) and the “Marathon” plume track (bold purple arrow), both with clear age progression. The Marathon plume track, which initiated at 2125 Ma with a giant radiating dyke swarm, crossed over into the adjacent Kaapvaal craton where it culminated in the emplacement of the Bushveld Complex. The actively rifting Superia plate was likely at a stand-still at Bushveld time (ca. 2056 Ma), allowing the plume tail to erode and dramatically thin the Kaapvaal lithosphere and setting up the conditions for the emplacement of Earth’s largest mafic layered intrusive complex. Ponding of voluminous sublithospheric plume magma resulted in outflow to distal localities (dashed arrows), possibly as far as Karelia-Kola (e.g., Kevitsa, 2058 Ma) and the Yilgarn (e.g., Mount Weld, ca. 2060 Ma; cf. Fiorentini et al., 2020). 14 �A COMPLEX F-RICH ALKALIC PEGMATITE IN THE PYROXENE SYENITES OF THE STETTIN COMPLEX, WAUSAU COMPLEX, MARATHON COUNTY, WISCONSIN BUCHHOLZ, Thomas1, FALSTER, Alexander2, and SIMMONS2, William 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494, 2MP2 Research Group, Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217, USA 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 composed of various syenite phases. This abstract is an update to a study of this dike in ILSG 2024. The sub-horizontal pegmatite is weathered, mineralogically and texturally zoned, and includes numerous syenite screens. Thin 2-3 cm reaction zones are common at contacts, with small miaroles, scattered patches of abundant, tiny pink zircons, fergusonite-(Y), and other minerals. Small miarolitic cavities are common throughout the dike. Overall mineralogy is complex, typical for fractionated alkalic pegmatites. Pyroxenes are largely absent except for highly altered replacements and sparse unaltered remnants of hedenbergite. Early formed pyroxene(s) appear to have been destabilized by later oxidation, altering Fe2+-rich pyroxenes to quartz and smectite-group clays ± goethite with sparse remnants of hedenbergite, and allowing crystallization of more oxidized (Fe3+ rich) magnetite and arfvedsonite. Similar reactions may have altered early-crystallizing chevkinite(Ce) (or a similar LREE-Ti species) and possibly aeschynite-(Ce), to an unidentified Ti-Ce4+-Fe phase: relatively common soft, pale yellow to creamy to brown grains of varying morphologies typically containing high Ti-Ce-Fe contents with traces of other elements. Cerium is likely present as Ce4+ based on the absence of associated LREE3+ (La, Nd, Pr). Alteration under oxidizing conditions may have removed LREE3+, Si and other elements, leaving immobile Ti, Ce4+, minor Fe3+ and trace amounts of other elements. Fluorapatite occurs as abundant hexagonal prisms in intermediate zones of the dike; generally highly altered with elevated to very high LREE (Ce-dominant) and Si contents, while similar reddish crystals in pegmatite units near the lower contact show more typical very low LREE contents. This may be the result of alteration/partial replacement of fluorapatite by fluorbritholite as discussed by Betkowski et al (2016). Work also continues on a rare unidentified Ba-silicate mineral, where lack of Al precludes Ba-feldspars. Several small-volume units and isolated occurrences contain minerals not normally found in alkalic pegmatites, including cassiterite, Hf-enriched zircons (up to 5.5 wt.% HfO2, vs 1.48 wt. % HfO2 in pegmatite margin zircons), fluorcalciomicrolite (D-site occupancy Ta 1.05, Nb 0.65, Ti 0.30; Σ 2), tantalite-(Mn), and barite. Sphalerite in unweathered lower portions of the dike is notable in containing about 0.7 wt. % Indium. Later oxidizing conditions are evident in late crystallization of siderite (now goethite), and LREE fluocarbonates. Crystallization of fergusonite-(Y) (to date Nb-dominant, ≈Nb 1.96, Ta 0.04; Σ 2), being rich in MREE and HREE and lacking redox sensitive Ce, appears to have continued throughout dike crystallization. 15 �REFERENCES Betkowski, Wladyslaw B., Harlov, Daniel E. and Rakovan, John F., 2016. Hydrothermal mineral replacement reactions for an apatite-monazite assemblage in alkali-rich fluids at 300-600° C and 100 MPa, American Mineralogist 101, 2620-2637. Van Wyck, N. 1994. The Wolf River A-type magmatic event in Wisconsin: U/Pb and Sm/Nd constraints on timing and petrogenesis (abstract): Institute on Lake Superior Geology, 40th Annual Meeting, Part 1, Program and Abstracts, 81-82. Aeschynite-(Ce) Albite Anorthoclase Barite Bavenite(?) Bertrandite Calcite Cassiterite Columbite-(Fe) Fayalite Fergusonite-(Y) Fluoro-arfvedsonite Fluorannite Fluorapatite Fluorite Fluorcalciomicrolite Fluorcalciopyrochlore Graphite Hedenbergite Ilmenite K-feldspar Kainosite-(Y) Magnetite Molybdenite Monazite-(Ce) Niocalite? Phenacite Quartz Siderite Sphalerite Thorite Titanite Zinnwaldite Zircon Zircon (metamict) (Ce,Ca,Fe,Th)(Ti,Nb)2(O,OH)6 NaAlSi3O8 (Na,K)AlSi3O8 BaSO4 Ca4Be2Al2Si9O26(OH)2 Be4(Si2O7)(OH)2 CaCO3 SnO2 Fe2+Nb2O6 Fe2+2SiO4 YNbO4 [Na][Na2][Fe2+4Fe3+]Si8O22F2 KFe2+3(Si3Al)O10F2 Ca5(PO4)3F CaF2 (Ca,Na)2(Ta,Nb)2O6F (Ca,Na)2(Nb, Ti)2O6F C CaFe2+Si2O6 Fe2+TiO3 KAlSi3O8 Ca2(Y,Ce)2(Si4O12)(CO3) · H2O Fe2+Fe3+2O4 MoS2 Ce(PO4) (Ca,Nb)4(Si2O7)(O,OH,F)2 Be2SiO4 SiO2 FeCO3 ZnS Th(SiO4) CaTi(SiO4)O KFe22+Al(Al2Si2O10)(OH)2 to KLi2Al(Si4O10)(F,OH)2 Zr(SiO4) Zr(SiO4) Table 1. Dike Mineralogy 16 Common Rock-forming Rock-forming Rare Rare Rare Common Uncommon Rare Rare Common Rock-forming Rock-forming Common Common Rare Rare Rare Uncommon Common Uncommon Rare Common Uncommon Common Rare Rare Rock forming Common Rare Rare Uncommon Uncommon Very common Common �Micromineralogy and textures in the Sudbury impact layer on the Mesabi Iron Range, Minnesota: record of processes in the proximal-distal ejecta transition zone CANNON, W. F., STOKES, M. Rebecca, SALERNO, Ross A. U.S. Geological Survey, Geology, Energy & Minerals Science Center, Mail Stop 954, Reston, VA 20192 The Sudbury Impact Layer (SIL) (1849 Ma), deposited here within hours of the giant meteor impact at Sudbury, Ontario, is known from drill core at four locations on the Mesabi Iron Range (Fig. 1) along a trajectory distance as great as 980 kilometers from the impact point. It records an instant of high energy deposition of about one meter of mixed ejecta and local bedrock within an otherwise quiescent sequence of siltstone and iron formation. Optical, scanning electron microscope, and Raman spectroscopy data provide details of the SIL that reveal some of the complexities of ejecta transport and deposition. Data presented here are from the Nashwauk occurrence where four drill holes provide continuous samples across the layer. Figure 1. Geologic map of the Mesabi Iron Range showing the four locations where the Sudbury Impact Layer (SIL) has been observed: Coleraine (Huber, et al., 2014; Nashwauk (Cannon, et al., 2017, this study); Eveleth (Addison, et al., 2005, this study); Erie (this study). The SIL on the Mesabi Iron Range consists of millimeter-scale ejecta particles expelled from the large crater near Sudbury, and coarser fragments, of sedimentary rocks, some greater than 3 cm diameter, which were derived locally. The ejecta can be subdivided into two categories: A- devitrified glass (Fig. 2), and B- millimeter-scale mineral grains and rock fragments displaying shock metamorphic features (Fig. 3). Figure 2. A-microtektite in matrix of coarse secondary dolomite. Original glass devitrified to K-mica. Vesicles are filled with dolomite. B-delicate bubble structures preserved in secondary dolomite. Bubble walls are mostly K-mica and chlorite. C-angular fragment of flattened vesicular glass, now mostly chlorite. D-rounded particle composed of fine K-mica. Spheres of vesicular glass and their fragments are common (Fig. 2A), including thinwalled hollow structures (Fig. 2B). They are now composed of micron-scale K-mica and chlorite. Irregularly shaped glass shards, mostly composed of chlorite, are also abundant (Fig. 2C). Many are larger than typical spherules and are probably far-flung bits of impact melt rather than broken spheres. Most are flattened into bedding. Also common are rounded grains composed of sub-micron K-mica with relict vesicles (Fig. 2D). These are distinct in having been 17 �sufficiently strong to have avoided flattening. Other glass particles are molded around them. They were likely droplets of melt with very uniform K-Al-Si composition. Quartz and feldspar grains with multiple sets of planar deformation features and zones of devitrified impact glass attest to the intense shock unique to meteor impacts. Small rock fragments with intense shock features are also common (Fig. 3). Figure 3. A-quartz with one well-developed set of planar deformation features and two weaker sets (red lines). B- intensely shocked quartz with “toasted” appearance and zones of devitrified glass. Ccathodoluminesence image of B showing complex shock-induced internal features. D-intensely shocked polycrystalline orthoquartzite fragment. Abundant glass spherules in the Nashwauk ejecta appear to be microtectites. Such particles are widely interpreted to form by condensation from impact vapor plumes above the atmosphere and can be distributed worldwide. At Nashwauk they are mixed with small rock and mineral particles from the outermost margins of the impact ejecta curtain. These were transported either (or both) on ballistic trajectories, or by intense impact-generated winds beyond the ejecta curtain. Many have strongly developed shock features attesting to their derivation by crater excavation near Sudbury. Notably missing from the SIL on the Mesabi Iron Range are accretionary lapilli, a hallmark of more proximal sites where ballistic ejecta and ground surges were the dominant transport mechanism for ejecta. The SIL at Nashwauk is very similar to that at the three other occurrences along the Mesabi Iron Range which, together, document a broad transition zone, between about 900 to 1000 kilometers from the impact point. Here the most distal ballistic ejecta persisted as millimeter-scale grains into a zone where ejecta plume material was becoming dominant. Along the Mesabi Iron Range ejecta was deposited in a shallow sea where fine-grained laminated silt and chert were being deposited, both before and after the impact. Strong impact-generated tsunamis reworked the ejecta and underlying sediments within hours or days of the impact to produce the intermixing of fine-grained ejecta particles with much coarser rip-up clasts from the pre-impact seabed. REFERENCES Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, 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. doi: https://doi.org/10.1130/G21048.1 Cannon, W.F., Woodruff, L. J., Jirsa, M., and Everett, W, 2017, New observations on distal ejecta from the Sudbury impact in the central Mesabi Iron Range, northern Minnesota, Institute on Lake Superior Geology, v. 63, Proceedings Part 1, Program with abstracts, p. 19-20. Huber, M.S., McDonald, I. and Koeberl, C., 2014, Petrography and geochemistry of ejecta from the Sudbury impact event, Meteoritic and Planetary Science: v. 49. p. 17491768. https://doi.org/10.1111/maps.12352 Jirsa, Mark, Chandler, V.W., and Lively, R. S., 2005, Bedrock geologic map of the Mesabi Iron Range Minnesota, Minnesota Geological Survey Miscellaneous Map Series map M-163. 18 �Updates on the Minnesota Department of Natural Resource’s Drill Core Library CARTER, Matt1 1 Minnesota Department of Natural Resource, Division of Lands and Minerals, 1525 3rd Ave E, Hibbing, MN, 55746 USA The Minnesota Department of Natural Resource’s (DNR) Drill Core Library (DCL) in Hibbing, MN is the only state-owned facility for archiving drill cores and other geological materials from Minnesota. The DCL was first established in 1972 when Building 1 (B1) was constructed. It was expanded in 1979 when Building 2 (B2) was constructed. Building 3 was first constructed in 1989 and expanded in 1995 and 2009. The facility currently stores around 3.5 million linear feet of drill core and contains material and/or data for over 20,000 drillholes. In 2023, it was identified that original shelving units installed in B1 and B2 needed to be replaced, and other safety issues needed to be resolved. The DNR diligently prepared to move all drill cores and other noncore geological materials from B1 to replace the shelving units. This was accomplished by assessing materials for deaccession, creating a box index of its holdings, applying barcodes to over 38,000 drill core boxes, as well as inventorying and barcoding noncore materials. Boxes were moved box by box by hand onto roller tables to an intake station where tracking information and digital images were captured. Boxes were then palletized and placed into temporary storage, which involved 712 pallets and 40 storage containers. The captured digital images have created a new and accessible digital record for B1 cores. Once the original shelving units were removed from B1, the DNR upgraded its lighting, and a new racking system was installed. Reverse flow of materials to B1 is anticipated to be completed by the end of May. Similar preparation activities are being applied on materials in B2. Instead of placing materials into temporary storage, rack space will be freed up through deaccession activities and materials will be rearranged within B2 to create new egress space. Over 210,000 iron ore boxes will be repackaged and moved onto new rack units. Lighting upgrades have been implemented in portions of B2, with the remaining lights to be replaced later this year. The DCL remains partially open to visitors, but users should be aware that materials from B1 and B2 may be unavailable until the project is completed on or before June 30, 2026. The DCL is nearing its facility-wide storage capacity and there is very limited space to accept additional materials. The DNR is only accepting deliveries on a case-by-case basis, and it is expected that any materials turned over to the state will become public upon delivery. In 2017, recognizing that the DCL was rapidly filling up, the DNR designed and actively sought funding for a fourth building to double the facility-wide storage capacity and quadruple view room space. This project is shovel-ready, but construction remains on hold until funding is legislatively secured. 19 �20 �Geology and Mineralization of the Plover Au Prospect, Marathon County, Wisconsin CASPER, Andrew A.1, LODGE, Robert W.D.1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA The Plover Au Prospect, located in Marathon County, WI, is hosted in Paleoproterozoic metaandesites, schist, and felsic/mafic intrusive units of the Wausau Volcanic Complex (LaBerge & Myers, 1983). Gold prospects in this region are bound to the southeast and northwest by large faults within the Eau Claire Deformation zone and the Wolf River Batholith (Lynott et al, 2022) (Figure 1). Rocks have undergone potassic and sericite alteration, greenschist to amphibolite grade metamorphism, and multiple stages of deformation. Research on the formational history of gold mineralization, in combination with its geochemical footprint, is essential for establishing a regional geologic setting of gold-forming events. Previous mineral exploration on this area has focused on exploring high concentrations of gold (Au) within volcanic units and sulfide vein networks at the Reef Deposit (Figure 1). With its proximity to the larger Reef Deposit, a more complete understanding of the Plover Prospect can add to a better regional context to the Au mineralizing system and potentially improve mineral exploration models. For this study, two holes (PL-76-1 & PL-76-4), totaling ~1,180 linear feet of core were chosen based on their relative locations and lithologic variation to fully characterize the range of units hosting mineralization. Representative volcanic strata and intrusive rocks were sampled and characterized through petrographic and geochemical analyses. The Plover deposit is primarily composed of andesitic/basaltic volcanics and gabbro/diorite intrusive units deposited sequentially showing sharp and, in some cases, brecciated contacts with one another. Brittleductile deformation is indicated by zones of brecciation present within the volcanic units. These structures include vein networks containing boudins and vugs containing sulfides and calcitechlorite alteration. It is probable that multiple deformational events occurred due to veins crosscutting foliation locally, and variation in the internal composition of veins. Hydrothermal alteration is suggested based on the presence of potassic alteration within the basaltic foliation and sericite-chlorite alteration in layers. Pyrite, chalcopyrite and pyrrhotite occur within vein networks. Since high Au concentrations are typically present within massive/semi-massive sulfide veins which contain brittle to brittle-ductile deformation, this mineralization likely occurred after Penokean deformation and metamorphism that formed the primary structural fabric in the rocks. The Reef gold-copper deposit has been researched extensively by various exploration companies since the 1990’s (Lynott et al, 2022). The deposit is located <1 mi east of the Plover deposit and has shown significantly higher Au concentrations. The deposit has been broadly classified as orogenic in origin and is claimed to have produced shear hosted vein-type gold and copper occurrences. Gold/copper mineralization occurs within stacked and relatively thin zones of quartz-sulfide veins and lenses; and sericite alteration within vein selvage typically accompanies gold mineralization within these areas. The primary lithology between the two deposits is similar, however, the Reef deposits proximity to the Wolf River batholith potentially influenced the degree of deformation and sericite, talc, tremolite, and pyrrhotite alteration. Future research should focus on more detailed comparisons between the Reef and Plover gold systems to better constrain potential genetic links between them. 21 �Figure 1. The relative geographic locations of the Plover and Reef deposits in the Penokean Volcanic Belt (PVB), central Wisconsin. Plover Au prospect is bounded to the east by the Eau Claire fault zone and the Wolf River Batholith. Figure has been adapted from Dematties (2022) and Lynott et al, (2022). REFERENCES LaBerge, G.L., and Myers, P.E., 1983a, Precambrian geology of Marathon County, Wisconsin: Information Circular, v. 45. Lynott, J.S., and Dematties, T.A., 2022, An Evaluation of the Reef Gold-Copper Deposit, Marathon County, Wisconsin, USA, NI 43-101 Technical Report, 402p. 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, p. 104489. 22 �Unusual early diagenetic structures in the Paleoproterozoic Gunflint Formation, Ontario, Canada CHURCHLEY, Sophie1, FRALICK, Philip2 1 Ontario Geological Survey, 435 James St S, Suite B002, ON P7E 6S7 Canada 2 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Newly identified microbial and diagenetic structures in the Gunflint Formation from the Thunder Bay area provide additional information to further our understanding of the environment in which these sediments were deposited and the diagenetic processes that affected them. Along the Current River, a horizon of carbonate lenses outcrop within a shale sequence. The structures are oblate in shape and range from approximately 0.5-1 m in diameter and 20-30 cm in height. Internally, the lenses are mostly calcite that displaces fine-grained siliciclastic laminae and preserves several interesting structures including cone-in-cone, inverted cuspate fenestrae and feather-like and braided fabrics. Some faces also display features that are more ‘fern-like’ in appearance and migrate up and across the surface of the oblate carbonate pods. We propose that these are oblate carbonate concretions formed via carbonate precipiation during diagenesis within a sequence of organic-rich laminated shales and siltstones. Several important features observed in the Gunflint Formation suggest that these structures formed via diagenetic processes including the similarities in stratigraphic placement within organic-rich shaley horizons and the light δ13Сcarb values recorded in the carbonate fraction. Cone-in-cone structures have been identified in both hand sample and thin section, displaying similarities to structures described from other locales (Figure 1A,B). Cone-in-cone structures occur in calcite-cemented sandstones or at the edges of disc-like to ellipsoidal concretions ranging in size from decimeters to meters long within shale beds in Erfoud, Morocco (Lugli et al. 2005). This stratigraphic positioning and size is consistent with what has been observed in the Gunflint Formation oblate concretions that host the cone-in-cone layers. Likewise, jagged ‘sawtoothed’ draping laminae that were identified in the Gunflint Formation are similar in appearance to those identified from cone-in-cone structure in the Devonian Middle Timan Formation in Russia (Figure 1C,D) (Shumilov 2020). The formation of cone-in-cone structure is still not well understood and numerous hypotheses have been proposed (see Lugli et al. 2005 and references therein). It is commonly associated with concretions and organic-rich sediments. The oblate carbonate horizons are located stratigraphically within a sequence of carbonaceous black shales near the base of the Upper Member of the Gunflint Formation with abundant evidence of microbially induced sedimentary structures (MISS) (Fischer and Fralick 2020). δ13Сcarb analyzed from the carbonate fraction in the Gunflint Formation samples displayed light values ranging from -12.29‰ to -0.17‰ with most values clustering near -10‰. These values are coincident with precipitation occurring in the zones of Fe, Mn, and/or sulfate reduction with relatively low rates of organic-carbonate oxidation and are consistent with ferruginous, reducing conditions (Mozley and Burns 1993). 23 �Figure 1. A. 1-3 cm scale cone-in-cone structure consisting of beige fibrous calicite draped by thin black siliciclastic films. B. Thin section close up of cone-in-cone structure from Permian carbonates in Thailand that is similar in appearance to those observed in the Gunflint Formation (see figure 3C; Chenrai et al. 2022). C. Hand sample of a carbonate-rich horizon displaying jagged ‘saw-toothed’ laminae near the top. D. Close-up image of similar jagged laminae from the Devonian Middle Timan Formation in Russia (see figure 9A; Shumilov 2020). At this location, the jagged laminae are associated with cone-in-cone structure. References Chenrai P., Assawincharoenkij T., Warren J., Sa-nguankaew S., Meepring S., Laitrakull K. and Cartwright I. (2022) The Occurrence of Bedding-Parallel Fibrous Calcite Veins in Permian Siliciclastic and Carbonate Rocks in Central Thailand. Front. Earth Sci. 9:781782. doi: 10.3389/feart.2021.781782. Fischer, S. and Fralick, P. (2020) Biological mats in siliciclastic sediments of the Paleoproterozoic Gunflint Formation, northwestern Ontario, Canada. Can. J. Earth Sci. 57: 947–953. Lugli, S., Reimold, W. and Koeberl, C. (2005). Silicified Cone-in-Cone Structures from Erfoud (Morocco): A Comparison with Impact-Generated Shatter Cones. doi: 10.1007/3-540-27548-7_3. Mozley, P.S. and Burns, S.J. (1993) Oxygen and carbon isotopic composition of marine carbonate concretions: an overview. Journal of Sedimentary Research 63, 73–83. Shumilov I.Kh. (2020) Сone-in-cone structure: New data. Litosfera, 20(1), 76-92. doi: 10.24930/16819004-2020-20-1-76-92. 24 �Alteration of magnetic mineralogy in the Giants Range Batholith by the Duluth Complex CORTOPASSI, Celia L., ALLERTON, Zsuzsanna P., FEINBERG, Joshua M. Department of Earth and Environmental Sciences, University of Minnesota, Suite 150, 116 Church St SE, Minneapolis MN 55455 During the Midcontinent Rift event (ca. 1.1 Ga) of the North American craton, the Duluth Complex (DC), a large mafic igneous intrusion, was emplaced into the Neoarchean Giants Range Batholith (GRB; ca. 2.7 Ga) in northeastern Minnesota, thermally altering the granitic country rock (Allison, 1925).The basal mineralized zone of the DC has been well-studied with regard to sulfide deposits, but the extent of alteration within the GRB footwall has not been as well constrained. Previous research has indicated the presence of sulfides at the DC-GRB contact, extending about hundred meters into the GRB (Steiner, 2014), and prior petrographic analysis has revealed textures consistent with contact metamorphism that diminish with distance from the contact (Pardi, 2024). This project seeks to define the magnitude of alteration within the GRB and to further characterize the orientation of the intrusion. This project utilizes a profile of 13 outcrop samples from the GRB that were collected systematically at distances between 100 and 4500 meters from the DC-GRB contact. We characterize changes in magnetic mineralogy as a function of distance from the DC-GRB contact using measured optical microscopy, electron microscopy, and magnetic properties (susceptibility and parameters calculated from hysteresis loops and backfield curves). These data reveal a distinguishable and consistent pattern in magnetic properties as a function of distance from the contact and distinct zones of textural alteration in oxide minerals (Figure 1). Patterns in smallscale magnetic properties broadly align with the large-scale trends seen in aeromagnetic data (Minnesota Geological Survey, n.d.), including changes in magnetic properties co-located with mapped faults (Jirsa et al., 2011). The orientation of the DC-GRB contact was examined using information from previously-drilled exploration drill holes (Minnesota Department of Health, n.d.) that penetrated through the DC and into the GRB. These observations, as well as outcrop measurements of modal layering and igneous foliation within the DC (Minnesota Geological Survey, 2023), constrain the orientation of the present-day DC-GRB contact to between 16-24° towards the east. The original depth of the modern day exposure of the DC-GRB remains unknown, as does any component of subsidence that occurred since the Midcontinent Rift event. Future work may include the collection of oriented samples for paleomagnetic studies, which would help constrain both the extent of thermal reheating of the GRB and postemplacement subsidence. Thermal modeling of the subsurface DC-GRB contact at various depths, alongside observed patterns in oxide mineral textures, could produce estimates of the extent of subsurface contact metamorphism. With these methods, we hope to better understand the conditions under which the DC was emplaced and accommodated, as well as estimate the thickness of Precambrian rock that has since been eroded away. 25 �Figure 1. A: Distribution of identified textures (A-E) as a function of distance from the DC-GRB contact. B: Measured magnetic properties as a function of distance from the DC-GRB contact. REFERENCES Allison, I. S.,1925. The Giants Range Batholith of Minnesota. The Journal of Geology, 33(5), 488–508. https://www.jstor.org/stable/30057863. Jirsa, M., Boerboom, T., Chandler, V. W., Mossler, J., Runkel, A., & Setterholm, D., 2011. S-21 Geologic Map of Minnesota-Bedrock Geology. https://conservancy.umn.edu/items/96de8d96-46ba441c-94ca-41080b4335be Minnesota Department of Health, n.d.. Minnesota Well Index (MWI). https://mnwellindex.web.health.state.mn.us/. Minnesota Geological Survey, n.d.. Collection of aeromagnetic data from Minnesota. https://doi.org/10.5066/P14LP38P. Minnesota Geological Survey, 2023. D-06, Structure Database. https://arcg.is/jfCLD. Pardi, L., 2024. Petrographic Analysis of the Giants Range Batholith in Northeastern Minnesota. Steiner, R. A., 2014. Genesis of sulfide mineralization within the granite footwall of the Maturi deposit of the South Kawishiwi intrusion, Duluth Complex, NE Minnesota. https://hdl.handle.net/11299/169376. 26 �Architecture of the Douglas Fault damage zone, northwest Wisconsin DANIELS, Nate, MCELLISTREM, Grace, VOGEL, Raeann, and BRAUNAGEL, Michael Department of Earth & Environmental Sciences, University of Minnesota Duluth, 1114 Kirby Drive Duluth, MN 55812 USA Major faults in the upper crust can be divided between the fault core, where most of the displacement is accommodated, and a surrounding damage zone (Faulkner et al., 2010). Fracturing in this damage zone occurs across a range of scales and intensity, varying from regularly spaced joint or deformation band sets to pervasive pulverization of the host rock. As such, fault damage zones can serve as fluid pathways, which control the migration of hydrothermal fluids and can alter the frictional strength of seismogenic fault systems. A number of processes are responsible for formation and evolution of a fault’s damage zone, including microfracturing within the process zone during fault propagation, localized wear along irregular fault surfaces, and volumetric changes associated with dynamic rupture propagation (Mitchell & Faulkner, 2009). As each process leaves a unique record in the fault system, the distribution and intensity of fault damage zones can provide insight into past fault activity and its relationship to fluid flow in the crust (Blenkinsop, 2008). This study presents preliminary observations of the fault-related damage surrounding the Douglas Fault from Amnicon and Pattison State Parks in northwestern Wisconsin. The Douglas Fault was activated during structural inversion of the Midcontinent Rift and previous work estimates its vertical displacement at ≳10 km (Grant, 1901; Cannon, 1994; Nicholson et al., 2006; Hodgin et al., 2024). At our study sites, the fault places basalts of the mid-continental rift Chengwatana volcanic group over post-rift siliciclastic sandstones of the Bayfield Group. Field and thin section observations along the fault system reveal pronounced damage zone asymmetry, with a hanging wall damage zone that is several times the width of the damage zone in the footwall. Chengwatana volcanics in the hanging wall are intensely fractured at the grain scale and cut by multiple generations of primarily calcite-filled opening mode veins. These veins and fractures broadly show two distinct orientations; one set striking generally NE to SW and the second characterized by NW to SE strikes. The damage-zone width is constrained by identifying changes in the slope of cumulative damage frequency plots, which shows high deformation frequency as a steep slope within an inner damage zone and less deformation decaying to background levels as a gentle slope in the outer damage zone of the Douglas Fault. Collectively, the full thickness of the hanging wall damage zone is >100 m (Grant, 1901). In contrast, sandstones of the Bayfield Group in the footwall exhibit lower frequency fracturing at the outcrop scale, no apparent grain-scale fracturing in thin section, and compressional deformation bands defined by porosity reduction. Bayfield sandstones in the footwall at these sites are also deformed by faultpropagation and drag folding that extend for tens of meters beyond the fault contact (Hodgin et al., 2024). Field and thin section scale observations of fault damage in both units are consistent with ultrasonic pulse velocities measured in samples collected from the fault zone with a Proceq Pundit Lab system. 27 �REFERENCES Blenkinsop, T.G., 2008. Relationships between faults, extension fracture and veins, and stress. Journal of Structural Geology, 30 (5), 622-632. Cannon, W.F., 1994. Closing of the Midcontinent Rift - A far-field effect of Grenvillian compression. Geology, 22 (2), 155-158. 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. Grant, U.S., 1901. Preliminary report on the copper-bearing rocks of Douglas County, Wisconsin (No. 3). Hodgin, E.B., Swanson-Hysell, N.L., Kylander-Clark, A.R.C., Turner, A.C., Stolper, D.A., Ibarra, D.E., Schmitz, M.D., Zhang, Y., Fairchild, L.M., and Fuentes, A.J., 2024. One billion years of stability in the North American midcontinent following two-stage Grenvillian structural inversion. Tectonics, 43 (9). Mitchell, T.M., and Faulkner, D.R., 2009. The nature and origin of off-fault damage surrounding strikeslip fault zones with a wide range of displacements: A field study from the Atacama fault system, northern Chile. Journal of Structural Geology, 31 (8), 802-816. Nicholson, S.W., Cannon, W.F., Woodruff, L.G., and Dicken, C., 2006. Bedrock geologic map of the Port Wing, Solon Springs, and parts of the Duluth and Sandstone 30’x60’ Quadrangles, US Geological Survey. 28 �The Archean Carney Lake gneiss complex in Michigan’s Upper Peninsula: Preliminary subdivisions with age constraints DeGRAFF, James1, DEERING, Chad1, and JONES III2, James 1 Department of Geological & Mining Engineering & Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 U.S.A. 2 U.S. Geological Survey, Alaska Science Center, 4210 University Drive, Anchorage, AK 99508 U.S.A. Much of the Precambrian bedrock in Michigan’s Upper Peninsula was last mapped at 1:24000 scale prior to modern tectonic concepts and advances in understanding related structural, magmatic, and metamorphic processes. A later decline in base and ferrous metal mining in the region reduced interest in commercial and scientific investigations, however recent concerns about the supply of critical minerals has renewed interest in developing an improved geologic framework for ore deposit exploration. The Archean Carney Lake gneiss complex (CLGC) and other granitegneiss complexes south of the Great Lakes tectonic zone are in the Minnesota River Valley subprovince of the southern Superior craton (Sims and Day, 1993). The CLGC, like the other complexes, is surrounded by Paleoproterozoic continental margin strata, partly older than and partly coeval with Penokean orogenesis (~1.85 Ga) that deformed the region (Schulz and Cannon, 2007). Bayley et al. (1966) describe the CLGC as predominantly felsic gneiss but with ~10% mafic inclusions and ~5% younger granodiorite and syenite intrusions by area. Mapping funded by the USGS Earth MRI program has revealed a wider variety of rocks than previously reported, differences in metamorphic grade, and new structural relationships (DeGraff et al., 2023). Felsic intrusions with little to no foliation are more abundant and varied than previously thought, ranging from granitic to tonalitic to locally syenitic. The original classification of gneiss based on mineralogy has been revised by also considering fabric characteristics. Consequently, we have identified an older EW-elongate core of thickly banded (≥2 cm) poly-deformed gneiss characterized by tightly folded banding, discordant banding across shear zones, and dismembered mafic pods (Fig. 1, area 1). Younger, less deformed, Archean rocks flank the older terrane, except on the north, and include the widespread felsic intrusions and thinly banded (≤1 cm), quartzo-feldspathic, gneissic rocks. The latter have quasi-planar, laterally continuous banding and local textures resembling cross-bedding and relict grains, suggesting derivation from a siliciclastic protolith. Boundaries between the older deformed gneiss terrane, the younger gneissic terrane with relict features, and areas with felsic intrusions are not yet well defined nor is their nature well understood. In addition to the above, at least four generations of mafic to ultramafic magmas have intruded the CLGC up to the late Mesoproterozoic. Our results, combined with those of others, indicate a long and complex tectonomagmatic history for the CLGC and adjacent rock units. The poly-deformed gneiss terrane includes rocks with inherited zircon cores dated at ca. 3750 Ma (Eoarchean) and recrystallized zircons and overgrowths dated at ca. 2750 Ma, the latter having formed during a Neoarchean thermal event (Ayuso et al., 2018). Neoarchean metamorphism of the Eoarchean gneiss, and perhaps much of its deformation, was accompanied by widespread felsic intrusions based on new zircon LA-ICPMS U-Pb dates ranging from ca. 2810 Ma to 2670 Ma (8 sites). Zircon trace-element analysis indicates that these magmas came from a hydrous oxidizing source and were contaminated while passing through a relatively thick crust, as is typical of magma generated during modern subduction. At the northern and eastern margins of the CLGC, relatively undeformed gneissic rocks were probably derived in part from siliciclastic protoliths of Neoarchean age. At the northern margin, 29 �however, NE-dipping beds of Paleoproterozoic Sturgeon Quartzite are parallel to well-defined layers of quartzo-feldspathic gneissic rocks along strike to the east. Field relationships and detrital zircon analysis suggest two scenarios: 1) a lateral facies change within Sturgeon Quartzite from meta-arkose on the east to meta-sandstone on the west, or 2) an onlapping relationship between younger quartzite and its parent Neoarchean meta-arkose. Figure 1: Preliminary subdivisions of the Archean Carney Lake gneiss complex (CLGC = 1, 2a, 2b, 3, Agu_clg). 1 = poly-deformed; 2 = metaigneous, 3 = meta-sedimentary; Agu_clg = undifferentiated; Xmrs = Paleoproterozoic Marquette Range Supergroup; Pz = Paleozoic clastic strata. Study area outlined in purple. 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, 64th Annual Meeting Proceedings, Part 1-Program and Abstracts, 64: 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: 1-96. DeGraff, J.M., Gannon, I.M., Deering, C.D., Smirnov, A.V., 2023. Bedrock geology of southeastern Dickinson County, Michigan: Vulcan 7.5’ quadrangle and adjacent parts of the Carney Lake, Cunard, Faithorn, Felch, Foster City, and Waucedah 7.5’ quadrangles. Michigan Geological Survey, Bedrock Geologic Map, 1:25,000 scale map sheet with explanatory text. Schulz, K.J. and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research, 157: 4-25. Sims, P.K. and Day, W.C., 1993. Great Lakes tectonic zone – revisited. U.S. Geological Survey, Bulletin 1904-S: S1-S11. 30 �Geochronology of lithium mineralization in the Florence pegmatite field, WI, USA DROUBI, Omar Khalil1, SCHOONOVER, Erik2, SIRBESCU, Mona-Liza3, GARBER, Joshua2, BONAMICI, Chloë1 Department of Geoscience, University of Wisconsin-Madison, 1215 W. Dayton Street, Madison, Wisconsin, 53706, USA 2 Department of Geosciences, The Pennsylvania State University, University Park, PA, USA 3 Geology Department, Central Michigan University, 314 Brooks Hall, Mount Pleasant, MI 48859, USA 1 Global and national progression toward decreasing reliance on fossil fuels will correlate with increasing the supply of mineral resources that contain high concentrations of elements like lithium, copper, or rare earth elements– “critical minerals” deemed essential for building low-CO2 technologies. Lithium-cesium-tantalum (LCT) pegmatites are a significant component of global lithium production; despite their importance, the tectonomagmatic mechanisms by which these pegmatites form are not completely understood. The two main models for LCT pegmatite formation are 1) as late-stage fractionation products of nearby peraluminous granites (“fractionation origin”) (e.g., Černý, 1991) or 2) directly from partial melts of Li-bearing highgrade metamorphic rocks (“anatectic origin”) (e.g., Knoll et al., 2023; Koopmans et al., 2023). These models have implications for LCT pegmatite exploration (i.e., mapping outward from plutons or anatectic zones in mountain belts) and testing them requires precise age estimates for the pegmatites and their neighboring magmatic and metamorphic rocks. This study provides new age constraints for models of LCT pegmatite mineralization in Florence County, WI, USA. (Falster et al., 2005; Falster et al., 1996; Sirbescu et al., 2008) by applying LA-ICP-MS U-Pb geochronology and trace-element analysis to apatite crystallized within the LCT pegmatites and titanite in the proximal wall rock and a separate, non-mineralized granitic pegmatite located >1.8 km away (Figure 1). The Florence LCT pegmatites were emplaced <2.5 km south of the Niagara Fault Zone and are hypothesized to be fractionated products from the nearby Bush Lake granite (undated, but hypothesized ~1835 Ma; Sims et al., 1985) or anatectic melts of the wall rock, the metavolcanic/metasedimentary Quinnesec Fm. (~1866 Ma; Sims et al., 1985). These models suggest pegmatite emplacement is broadly bracketed in space and time by the end of the Penokean orogeny (~1835 Ma) and Yavapai arc accretion (~1750–1700 Ma) (Figure 1). The apatite grains from the King’s X and Animikie Red Ace LCT pegmatites have U-Pb dates of 1446 ± 6 [29] Ma and 1432 ± 4 [29] Ma, respectively. The targeted apatite grains, which nucleated in the pegmatite chilled margin at the wall rock contact, are interpreted as magmatic based on oscillatory cathodoluminescence zoning (Sirbescu et al., 2009). Xenoblastic titanite from the Quinnesec Fm., sampled at distances <1 cm to ~150 m from the pegmatites, have U-Pb ages of 1473 ± 7 [29] Ma (<1 cm), 1466 ± 7 [29] Ma (<5 m), 1436 ± 13 [29] Ma (60 m), and 1471 ± 8 [29] Ma (150 m), but euhedral titanite grains from the non-mineralized granitic pegmatite have a U-Pb age of 1811 ± 10 [36] Ma. Our data indicate that the Florence LCT pegmatites did not result from fractionation of the Bush Lake granite nor anatexis during the Penokean or Yavapai orogenies and are instead coeval with emplacement of the ~1476 Ma Wolf River batholith further south. A revised age model for lithium mineralization in northern WI suggests involvement of the Wolf River batholith or far-field influence of the Mesoproterozoic Pinware-Baraboo-Picuris orogeny. 31 �Figure 1. Conceptual cross section (not to scale) showing age constraints for the Florence pegmatite field. Hypothesized ages based on the following references: Bush Lake granite and Quinnesec Fm. (Sims et al., 1985), metagabbro (Guice et al., 2023). U-Pb dates reported as: date ± internal 2s [external uncertainty=2% of date]. REFERENCES Bradley, D.C., McCauley, A.D., and Stillings, L.M., (2017), Mineral-deposit model for lithium-cesiumtantalum pegmatites: U.S. Geological Survey Scientific Investigations Report 2010–5070–O, 48 p., https://doi.org/10.3133/sir20105070O. Černý, P., 1991, Rare-element Granitic Pegmatites. Part II: Regional to Global Environments and Petrogenesis: Geoscience Canada, v. 18, p. 68–81, Falster, A. U., Simmons, W.B., and Webber, K.L. (2005), Origin of the pegmatites in the Hoskin Lake pegmatite field, Florence Co., Wisconsin, in Crystallization Processes in Granitic Pegmatites, International Meeting in Cavoli, Elba Island, Italy, May 23–28, 2005, edited by F. Pezzotta, Mineral. Soc. of Am., Chantilly, Va. Falster, A. U.; Simmons, Wm. B.; and Webber, K. L. (1996) The Mineralogy and Geochemistry of the Animikie Red Ace Pegmatite, Florence County, Wisconsin. In Pandalai, S. G., ed., Recent Research Developments in Mineralogy, 7-67. Guice, G. L., Viete, D. R., Holder, R. M., & Roy, S. (2023). A c. 1900 Ma Tethyan-type ophiolite in the Penokean Orogen, Pembine, Wisconsin (USA): Insights from the volcanic stratigraphy. Precambrian Research, 399, 107223. Knoll, T., Huet, B., Schuster, R., Mali, H., Ntaflos, T., & Hauzenberger, C. (2023). Lithium pegmatite of anatectic origin-A case study from the Austroalpine Unit Pegmatite Province (Eastern European Alps): geological data and geochemical model. Ore geology reviews, 105298 Koopmans, L., Martins, T., Linnen, R., Gardiner, N.J., Breasley, C.M., Palin, R.M., Groat, L.A., Silva, D., and Robb, L.J., (2023). The formation of lithium-rich pegmatites through multi-stage melting. Geology. Sims, P. K., Peterman, Z. E., & Schulz, K. J. (1985). The Dunbar Gneiss-granitoid dome: Implications for early Proterozoic tectonic evolution of northern Wisconsin. Geological Society of America Bulletin, 96(9), 1101-1112. Sirbescu, M. L. C., Hartwick, E. E., & Student, J. J. (2008). Rapid crystallization of the Animikie Red Ace Pegmatite, Florence county, northeastern Wisconsin: inclusion microthermometry and conductivecooling modeling. Contributions to Mineralogy and Petrology, 156, 289-305. Sirbescu, M. L. C., Leatherman, M. A., Student, J. J., & Beehr, A. R. (2009). Apatite textures and compositions as records of crystallization processes in the Animikie Red Ace pegmatite dike, Wisconsin, USA. The Canadian Mineralogist, 47(4), 725-743. 32 �Experimental Reproduction of Acidic Mafic-Ultramafic Hydrothermal Fluids with Implications for Linking Seafloor Lithology to Ore Mineral Solubility and Novel Geochemical Trapping Mechanisms EVANS, Guy N.1 and SEYFRIED JR., William E.1 1 Department of Earth and Environmental Sciences, University of Minnesota, 116 Church St SE, Minneapolis, MN, 55455, United States Ultramafic-hosted seafloor massive sulfide (UM-SMS) deposits constitute a distinct class of CuZn-Co-Ni-Au-rich seafloor hydrothermal deposits (Fouquet et al., 2010). However, ultramafichosted volcanogenic massive sulfide (UM-VMS) deposits have been historically overlooked, in part because the formation of UM-VMS deposits differs from traditional VMS genetic models based on basalt-hosted SMS deposits (Pattern et al., 2022). Adding to this complexity, ultramafic-hosted seafloor hydrothermal fluids span nearly the full range of pH and metal concentrations observed at active seafloor hydrothermal vents, from highly acidic (pH= 2.8), metal-rich (Fe > 20 mmol/kg) fluids observed at Rainbow Hydrothermal Field (Douville et al., 2002), to alkaline (pH = 10.5), metal-poor (Fe < .02 mmol/kg) fluids observed at Lost City Hydrothermal Field (Kelley et al., 2005; Evans et al., 2024). Here, we present results from recent experiments conducted at the University of Minnesota that for the first time reproduce acidic hydrothermal fluids from mixed mafic-ultramafic source minerals. The observed acidity of these fluids results from temperature-dependent fluid-rock reactions and superimposed geochemical and physical processes. We further highlight the implications of these findings for UM-VMS deposit models, including novel geochemical trapping mechanisms potentially relevant in areas exhibiting significant ultramafic volcanic/intrusive rocks. Regional examples include the Newton Belt (northeast Minnesota), Shebandowan Belt (northwestern Ontario), and Kidd-Monroe assemblage (eastern Ontario and Quebec). REFERENCES Douville, E., et al., (2002). The rainbow vent fluids (36 14′ N, MAR): the influence of ultramafic rocks and phase separation on trace metal content in Mid-Atlantic Ridge hydrothermal fluids. Chemical Geology, 184(1-2), 37-48. Evans, G. N. et al. (2024). Transition metals in alkaline Lost City vent fluids are sufficient for early-life metabolisms. Geochimica et Cosmochimica Acta, 385, 61-73. Fouquet, Y. et al. (2010). Geodiversity of hydrothermal processes along the Mid‐Atlantic Ridge and ultramafic‐hosted mineralization: A new type of oceanic Cu‐Zn‐Co‐Au volcanogenic massive sulfide deposit. Diversity of hydrothermal systems on slow spreading ocean ridges, 188, 321-367. Kelley, D. S. et al. (2005). A serpentinite-hosted ecosystem: the Lost City hydrothermal field. Science, 307(5714), 1428-1434. Patten, C. G. et al. (2022). Ultramafic-hosted volcanogenic massive sulfide deposits: an overlooked subclass of VMS deposit forming in complex tectonic environments. Earth-Science Reviews, 224, 103891. 33 �34 �Textural and chemical analysis of sphalerite ores from the Highland Subdistrict, Upper Mississippi Valley Zinc-Lead District, Wisconsin FITZPATRICK, William1 1 Wisconsin Geological and Natural History Survey, 3817 Mineral Point Rd. Madison, WI, USA Lead + zinc ± barite ± copper deposits are widespread in Ordovician carbonate rocks of southwestern Wisconsin and the bordering areas of Illinois and Iowa, commonly referred to as the Upper Mississippi Valley zinc-lead district (UMVD). Sphalerite, the primary zinc ore mineral in the UMVD, is known from other mining districts to contain valuable byproduct commodities as trace elements such as gallium, germanium, cadmium and silver. Several previous studies have examined sphalerite from the UMVD, but focused on samples from the southern part of the district (Hall and Heyl, 1968, McLimans and others, 1980). Zinc ores from other areas of the UMVD have received less attention, and little is known of their textural character and trace element content. This study presents new trace element data and textural observations of sphalerite ores from the Highland Subdistrict, the northernmost mining center in the UMVD. Two hand-picked sphalerite concentrates and thirty-five bulk ore samples were analyzed by whole rock geochemical methods, complemented by 282 in situ electron microprobe analyses on six thin sections from a mix of vein and disseminated ores. Textures in the sphalerite ores were also documented through scanning electron microscopy with the aim of understanding mechanisms that localized sulfide mineralization. Sphalerite from the Highland Subdistrict is characterized by alternating sequences of lighter, honey-colored bands and darker, reddish-brown bands in both the disseminated and vein hosted ores (Fig. 1). Microprobe analysis shows that darker bands tend to localize elevated iron and lower cadmium relative to lighter bands (Fig. 1). Silver content is variable, but tends to be higher in lighter bands, especially in the cores of disseminated grains. Comparing results from the whole rock and microprobe analyses from the Highland Subdistrict to sphalerite analyzed elsewhere in the UMVD, iron and cadmium are within known ranges, but silver is enriched to a significant degree (Hall and Heyl, 1968). Gallium and germanium abundances were too low to be detected in microprobe analyses, but whole rock analysis indicates they are towards the low end of the range observed in sphalerite from the UMVD (Hall and Heyl, 1968). Scanning electron microscope observation discovered abundant, texturally early framboidal pyrite intergrown with marcasite that is enveloped by later sphalerite. Framboidal pyrite has a well-documented association with sulfate reducing bacteria (e.g. Maclean and others, 2008), indicating bacterial processes were likely important in creating a reservoir of reduced sulfur within the carbonate host rocks. This in turn may have acted as a chemical trap for metals in migrating connate brines to form the zinc deposits. REFERENCES Hall, W.E., and Heyl, A.V., 1968, Distribution of Minor Elements in Ore and Host Rock, IllinoisKentucky Fluorite District and Upper Mississippi Valley Zinc-Lead District. Economic Geology, 63, 655-670. Maclean, L., Tyliszczak, T., Gilbert, P., Zhou, D., Pray, T., Onstott, T., and Southam, G., 2008, A high resolution chemical and structural study of framboidal pyrite formed within a low-temperature bacterial biofilm. Geobiology, 6, 471-480. McLimans, R.K., Barnes, H.L., and Ohmoto, H., 1980, Sphalerite Stratigraphy of the Upper Mississippi Valley Zinc-Lead District, Southwest Wisconsin. Economic Geology, 75, 351-361. 35 �Figure 1. Plots of iron, cadmium and silver along linear traverses through banded sphalerite crystals from the Highland Subdistrict. Top panel shows scans of the thin sections analyzed and locations of the analyses. Note the concentrically zoned disseminated grain (left) vs vein (right). 36 �Rare-element Geochemistry of the Eau Claire River Complex Pegmatites GRIES, Samara1, LODGE, Robert W.D1, HANEL, Sara1,2, HOOPER, Robert1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA 2 Current Affiliation: Department of Earth and Environmental Sciences, University of Minnesota Twin Cities, Suite 150, 116 Church St. SE, Minneapolis, MN 55455 Minerals, such as monazite and xenotime, are an important source of rare earth (La, Ce, Nd) and high field strength (Th, Nb, Zr) elements which are essential for modern energy, communication, and military technologies. These critical minerals are often sourced in pegmatites and are important exploration targets worldwide (Haque et al, 2014). The Paleoproterozoic Eau Claire Volcanic Complex (ECVC) is intruded by granitic pegmatite dikes that postdate peak metamorphism (Lodge et al, 2023), indicating they are unrelated to Penokeanaged orogenic events. The ECVC pegmatites are highly fractionated, garnet bearing, and contain a high concentration U, Th, La, Ce, and other rare earth elements. Based on major and trace element associations, the pegmatites in the ECVC are classified as NYF family pegmatites that contain Nb>Ta, REE, U, Th, Zr and are A- to I- types with peralkaline relationship (Cerny and Ercit, 2005). This study collected bedrock samples from several locations across the ECVC (Little Falls, North Fork, Muskeg Creek) (Figure 1). The pegmatite dikes can range in size from a few meters to 100 m in width near the North Fork of the Eau Claire River. They mainly intrude foliated and metamorphosed Paleoproterozoic to Archean tonalites, amphibolites, and gneisses. Samples from these pegmatites were analyzed for whole rock and mineral chemistries. Whole rock chemistry was analyzed on XRF and ICPMS whereas mineral chemistry was determined using SEM-EDS. All three locations have quartz, feldspar, plagioclase, biotite, and muscovite. The main mineralogy of Little Falls samples are albite and muscovite. Trace mineralogy of the Little Falls samples include Fe- and Mn-garnet, samarskite, columbite, zircon, and xenotime. Muskeg Creek samples contains both orthoclase and albite with biotite instead of muscovite. Trace mineralogy of the Muskeg Creek samples includes xenotime, monazite, and barite. The North Fork samples mainly contain albite with minor orthoclase and biotite. Trace mineralogy of the pegmatites in the North Fork area include in Fe- and Mn-garnets, monazite, xenotime, and thorite. The pegmatites from the ECVC are all low in Ca and have trace minerals with rare earth elements. They all contain with albite with low quantities of orthoclase and almost no anorthite. The North Fork and Muskeg Creek samples have more barium-rich minerals than Little Falls, which may be the result of fractionation of feldspars and plagioclase (Yu et al, 2007). Ba-rich minerals can also be a product of hydrothermal activity (Hanor, 2000), but there is no evidence of syn- to post-hydrothermal alteration of the pegmatites. Mn-rich garnets at Little Falls and North Fork indicate a higher degree of fractionation relative to Fe-garnets at Muskeg Creek (Hernández-Filiberto et al, 2021). North Fork and Muskeg Creek also had the largest crystals reaching over 20 cm in size. The North Fork is enriched in the heavy rare earth elements, U, and Th. In comparison, Little Falls has more light rare earth elements. Muskeg is also enriched in light rare earth elements in addition to an increased enrichment of heavy rare earth minerals like Gd and Dy. All three locations contain other metals such as Nb, Zr, Hf. 37 �Figure 1. Bedrock geologic map of the Eau Claire Volcanic Complex with site locations. North Fork depicts a pink pegmatite intruding into a grey tonalite. Muskeg Creek depicts a 6 m pegmatite dike with zoning. Little Falls shows a 13 m pegmatite dike. Map from Mudrey & Brown (1982). REFERENCES Cerny, P., and Ercit,T., 2005. The classification of granitic pegmatites revisited. The Canadian Mineralogist, 43: 2005-2026. Hanor, J.S., 2000, Barite-celestine geochemistry and environments of formation. In Alpers, C.N., Jambor, J.L., Nordstrom, D.K., eds. Reviews in Mineralogy and Geochemistry, 40: p. 193-275 Haque, N., Hughes, A.., Lim, S., Vernon, C., 2014, Rare Earth Elements: Overview of Mining, Mineralogy, Uses, Sustainability, and Environmental Impact: Resources, 3, p. 614-635 Hernández-Filiberto, L., Roda-Robles, E., Simmons, W.B., Webber, K.L., 2021, Garnet as Indicator of Pegmatites from the Oxford Pegmatite Field (Maine, USA): Minerals, 11(8), 802 Lodge, RWD, Weber, EM, Hooper, RL, 2023, Precambrian Geology of the Eau Claire River Valley: Rediscovering the Eau Claire Volcanic Complex. in Lodge, RWD (Ed.), Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 2 – Field Trip Guidebooks. v.69, part 2, p.47-70. Mudrey, M.G., Jr., Brown, B. A., Greenberg, J. K., 1982, "Bedrock Geologic Map of Wisconsin." Wisconsin Geological and Natural History Survey, scale 1:1,000,000 Yu, J.-H., O’Reilly, S. Y., Zhao, L., Griffin, W. L., Zhang, M., Zhou, X., Jiang, S.-Y., Wang, S.-Y., Wang, R.-C., 2007, Origin and evolution of topaz-bearing granites from the Nanling Range, South China: a geochemical and Sr-Nd-Hf isotopic study: Minerology and Petrology, 90, p. 271-300 38 �Revisiting Gravity and Magnetic Anomalies of the Baraboo Range HINZE, William J.1 and LONGACRE, Mark B.2 1 Purdue University, 30 Brook Hollow Ln., West Lafayette, IN 47906 2 MBL, Inc., 51 Captain Perry Dr., Phippsburg, ME 04562 The efficacy of gravity and magnetic methods of geological exploration have increased greatly since they were first used to investigate the Baraboo Synclinorium of Wisconsin nearly 75 years ago (Ostenso, 1953; Hinze, 1959). These methods and their associated technology are used for the first time since then to investigate the geology of the Mesoproterozoic Baraboo Synclinorium, its regional basement, and to illustrate the importance of modern data sets and analysis and interpretation methods. The latter are the result of computers for analysis, interpretation, and presentation of anomalies that were unavailable when the geophysical methods were first applied to mapping the Synclinorium. Analysis and interpretation of current gravity and magnetic anomaly data sets (Figure 1) indicate that the negative gravity anomaly associated with the Baraboo Synclinorium is not unique to the Synclinorium but is the southern termination of the Wisconsin Gravity Minimum (WGM) that covers a large portion of central Wisconsin including the Wolf River Batholith (WRB). The WGM is derived largely from felsic plutons in the upper crust extending outward from the ~1.5 Ga WRB. The lower density of the plutons compared to the metamorphosed orogenic rocks of the upper crust is the likely source of the negative gravity anomaly. The Synclinorium, located along an east-northeast trending gravity and magnetic lineament within the Yavapai orogenic province, occurs in a syncline of largely felsic volcanic rocks (Figures 2 and 3), the Sauk Syncline, that was likely deformed along with the Baraboo Synclinorium by south and southeast-verging thrusting during the Mazatzal and Picuris-BarabooPinware Orogenic events. Variations in thrusting has led to significant differences in the eastern and western portions of the Baraboo Synclinorium. Key results of the gravity and magnetic anomaly data analysis of the Synclinorium include: (1) The Baraboo Synclinorium’s negative gravity anomaly originates in upper crustal Yavapai and Wolf River Batholith felsic plutons that are the source of the Wisconsin Gravity Minimum. (2) The Synclinorium occurs within a synclinal structure resulting from deformation related to generally south-verging, thin-skinned thrust faulting that also produced the Baraboo Synclinorium. (3) The structure of the eastern and western portions of the Baraboo Synclinorium differ likely as a result of variations in the direction and intensity of thrusting during the Mazatzal and PicurisBaraboo-Pinware Orogenies (~1.63-1.41 Ga). REFERENCES Hinze, W.J., 1959. A gravity investigation of the Baraboo Syncline region. The Journal of Geology, 67(4), 417-446. Ostenso, Ned, 1953. Magnetic studies of the Baraboo Syncline. Unpublished M.A. thesis, University of Wisconsin-Madison. 39 �Figure 1. Gravity and magnetic anomaly maps of the Baraboo Synclinorium. Reduced to pole (RTP) total magnetic intensity anomaly map (right) eliminates the effect of the inclined earth’s magnetic field on the induced magnetization of the crustal rocks and the vertical gradient Bouguer gravity anomaly map of the Baraboo Synclinorium region (left) minimizes the regional gravity anomaly. The boundaries of the counties are indicated and the outline of the boundary of the Baraboo Synclinorium is the dashed white line. The white line interior to the Synclinorium is the boundary of the Freedom Formation. Color coding of both figures is non-linear. Figure 2. Tilt derivative of the Bouguer gravity anomaly map of the Baraboo Synclinorium region showing the outline of the Sauk Syncline in thick dashed white lines interpreted from the gravity and magnetic anomaly maps. The outline of the Baraboo Synclinorium is the thin dashed white line. The white line interior to the Synclinorium is the boundary of the Freedom Formation. Color coding is non-linear. Figure 3. High pass 10-km RTP magnetic anomaly map of the Baraboo Synclinorium region showing the outline of the Sauk Syncline in thick dashed white line interpreted from the gravity and magnetic anomaly maps. The outline of the Baraboo Synclinorium is the thin dashed white line. The white line interior to the Synclinorium is the boundary of the Freedom Formation. Color coding is non-linear. 40 �Emplacement of the Mesoproterozoic Wausau Syenite Complex, Wisconsin HULA, Linsey1 and CZECK, Dyanna1 1 Department of Geosciences, University of Wisconsin Milwaukee, Lapham Hall, Room 366, 3209 N. Maryland Ave. Milwaukee, WI 53211 The Wausau Syenite Complex (WSC) in Marathon County, Wisconsin is an intrusive complex of granitoids emplaced approximately 1.5 Ga (Dewane and Van Schmus, 2007). It is traditionally considered part of a major anorogenic ferroan granite magmatic event that affected the southern margin of Laurentia circa 1.4 Ga. Recent studies have recognized a Laurentian-scale accretionary margin between 1520-1340 Ma (Fig. 1), including the Pinware Orogeny in the northeast, the Picuris Orogeny in the southwest, and the most recently attributed section, the Baraboo Orogeny centered in Wisconsin (Daniel et al., 2023). This new hypothesis provides intriguing opportunities to reconsider the origin and tectonic setting of WSC emplacement as well as other Mesoproterozoic granitoids in Wisconsin, including the larger 1.4 Ga Wolf River Batholith (Dewane and Van Schmus, 2007). This research project will use the orientation of magnetic fabrics within the WSC to better understand how the batholith was emplaced. Figure 1: Simplified geologic map of Precambrian crustal provinces including the Mesoproterozoic accretionary margin of the Picuris, Baraboo, and Pinware Orogenies. The 1.48-1.35 Ga ferroan granites, including the Wolf River Batholith, are highlighted in white and the ~1.5 Ga Wausau Syenite Complex is added. Modified from Medaris et al., 2021, originally based on (Whitmeyer and Karlstrom, 2007). The project will consist of an anisotropy of magnetic susceptibility (AMS) survey and thin section analysis of each granitoid within the WSC. With these data, the magmatic flow directions and any subsequent tectonic overprint can be determined, which can be used to constrain the location of the magmatic feeder and the tectonic environment of emplacement. For the purpose of this abstract, three possible outcomes are proposed: 41 �1. Radial magmatic fabrics are preserved, indicating that the WSC was emplaced and cooled prior to the Baraboo Orogeny, with deformation accommodated by the surrounding weaker country rock (Fig. 2A). 2. Magmatic fabrics show a preferential flow pattern parallel to the tectonic margin caused by differential stress from the Baraboo Orogeny, suggesting syntectonic emplacement (Fig. 2B). 3. Only solid-state deformation fabrics are present, implying that the WSC was emplaced before or at the onset of the Baraboo Orogeny and had fully cooled before significant deformation occurred (Fig. 2C). By focusing on these oldest known Mesoproterozoic ferroan granites in the region, we can learn about the timing and geometry of the earliest Baraboo orogenesis. This study will address the question of how these enigmatic granites fit into the overall tectonic history of the Great Lakes Region. Figure 1: Schematic diagram of the WSC showing three possible outcomes of the AMS study. A) Radial magmatic fabric. B) Magmatic fabric with preferential flow parallel to the tectonic boundary. C) Solid state fabric. REFERENCES Daniel, C.G., Indares, A., Medaris Jr., L.G., Aronoff, R., Malone, D., and Schwartz, J., 2023. Linking the Pinware, Baraboo, and Picuris orogens: Recognition of a trans-Laurentian ca. 1520–1340 Ma orogenic belt, in Whitmeyer, S.J., Williams, M.L., Kellett, D.A., and Tikoff, B. eds., Laurentia: Turning Points in the Evolution of a Continent, Geological Society of America, 175–190. Dewane, T.J., and Van Schmus, W.R., 2007. U–Pb geochronology of the Wolf River batholith, northcentral Wisconsin: Evidence for successive magmatism between 1484Ma and 1468Ma: Precambrian Research, 157, 215–234. Medaris, L.G., Singer, B.S., Jicha, B.R., Malone, D.H., Schwartz, J.J., Stewart, E.K., Van Lankvelt, A., Williams, M.L., and Reiners, P.W., 2021. Early Mesoproterozoic evolution of midcontinental Laurentia: Defining the geon 14 Baraboo orogeny: Geoscience Frontiers, 12, 101174. Whitmeyer, S.J., and Karlstrom, K.E., 2007. Tectonic model for the Proterozoic growth of North America: Geosphere, 3, 220–259. 42 �Mapping oxidation reactions in iron-rich rocks from northeast Minnesota, USA. JAROZEWSKI, Sarah1, DUFFY, Paige1, BARRÉ, Cole1, BRENGMAN1, Latisha, EYSTER2, Athena 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Heller Hall, 1114 Kirby Drive, Duluth, MN 55812, USA 2 Department of Earth and Climate Sciences, Tufts University, Lane Hall, 2 North Hill Road, Medford, MA 02155, USA Aqueous alteration and post-depositional mineral assemblage modification in Precambrian terranes are ubiquitous, but clear accounting of the relative timing of oxidation reactions at the landscape scale is limited. Here we synthesize observations of oxidation and hydration reactions in three iron-rich lithologies in northeast Minnesota, the Soudan Iron Formation, the Cuyuna Iron Formation, and the Partridge River Intrusion of the Duluth complex to contribute to building a compiled relative mineral redox history for the landscape. The Soudan Iron Formation (~2.7 Ga) is a greenstone-hosted metamorphosed chemical sedimentary unit primarily composed of alternating bands of magnetite, hematite and microquartz affected by at least two Archean deformation events, and later fluid alteration (Thompson, 2015). The Soudan Iron Formation primarily consists of mm-scale bands of authigenic microquartz and iron oxides that preserve in outcrop samples distal to the ore zone, and within the ore horizon at Soudan underground mine. Reflected light petrography of oxides in outcrop samples reveals magnetite replacement by hematite (Figure 1A). Within ore zone samples, complete hematite replacement and large platy hematite is common, similar to previous observations (Thompson, 2015). Clear metamorphic minerals that could indicate high temperatures, pressures, and P-T-path histories are absent from the unit. The younger Cuyuna Iron Formation (~1.9 Ga) is part of an intensely folded metamorphosed sedimentary sequence deformed during the Penokean orogeny (Schmidt, R.G., 1963). Combining new observations and previous data from historic samples (Melcher et al., 1996), oxidation of magnetite is prevalent but limited in samples from the Gloria drill hole (Figure 1B). Metamorphic stilpnomelane is common in the Cuyuna iron formation in contrast to the Soudan Iron Formation. Documented differences in metamorphic silicate mineralogy between these two iron formations may indicate key differences in precursor phases, as both units were affected by significant metamorphic deformation events. Yet, for both, oxidation of magnetite and replacement by hematite indicate both iron formations are similarly affected by post-depositional fluid alteration and oxidation. The Partridge River Intrusion (PRI) is part of the layered series troctolitic intrusions that form the base of the Duluth complex (Tyson and Chang, 1984). In its present geometry, the magmatic layered series PRI now intersects the current land surface. Clear evidence of aqueous alteration in the first few hundred feet of drill core 17700 includes mineral transformation of biotite and olivine to hydrous ferric oxides and secondary iron silicates (Figure 1C). These replacement reactions are limited in scale, and primary igneous mineralogy is still preserved. Leveraging cross-temporal comparisons of current iron-rich bedrock outcrop exposures and drill cores in north-east Minnesota to identify formational vs. post-formational mineralogy will allow for landscape-scale mapping of oxidation reactions and their extent in the subsurface. 43 �Figure 1. Reflected light photomicrographs of the Soudan iron formation, Cuyuna Iron formation and back-scatter electron image (BSE) of the Patridge River Intrusion of the Duluth complex. (A) Magnetite is partially replaced by hematite in the Soudan iron formation. (B) Magnetite oxidation to hematite in the Cuyuna iron formation. (C) Back-scatter electron image of altered olivine in the Partridge River Intrusion from the UMTC EPMA lab, CHARFAC facility.. REFERENCES Melcher, F., Morey, G. B., McSwiggen, P. L., Cleland, J. M., & Brink, S. E. 1996. RI-46 Hydrothermal Systems in Manganese-Rich Iron-Formation Of the Cuyuna North Range, Minnesota: Geochemical and Mineralogical Study of the Gloria Drill Core. Report of Investigations 46, ISSN 0076-9177, 1 - 45. Schmidt, R. G. 1963. Geology and ore deposits of the Cuyuna North range, Minnesota. U.S. Geological Survey Professional Paper 407, p. 96. Taylor, Richard B., 1964. Geology of the Duluth Gabbro Complex near Duluth, Minnesota. Bulletin No. 44. Minnesota Geological Survey, University Digital Conservancy. Thompson, A. 2015. A hydrothermal model for metasomatism of neoarchean Algoma-Type banded iron formation to massive hematite ore at the Soudan Mine, NE Minnesota. University of Minnesota, Duluth. P. 1-59. Tyson, R. M., and Chang, L, L, Y. 1984. The Petrology and sulfide mineralization of the Partridge River Troctolite, Duluth Complex, Minnesota. Canadian Mineralogist, v. 22, p 23-38. 44 �Geology and Geochemistry of the Mesoproterozoic Round Lake Intrusion and associated Ti-Mineralization, Northern Wisconsin JEUTTER, Renee O.1, LODGE, Robert W.D.1 1 Department of Geology & Environmental Science, University of Wisconsin-Eau Claire, 105 Garfield Avenue, Eau Claire, WI 54701, USA Modern technology and renewable energy require large amounts of metals that are currently imported, and there is a tremendous effort to domesticate our mineral extraction and processing. Several of these critical minerals, such as Ti, are found in Wisconsin, but little data is available to guide future mineral exploration efforts. The Mesoproterozoic Mid-Continent Rift and its satellite intrusions are known to host Ti-Fe oxide mineralization and Ni-Cu-PGE magmatic sulfide deposits (Woodruff, 2020). During the development of the Mid-Continent Rift, there is a temporal evolution of occurrences of mineral deposits. The plateau stage typically created Ni-Cu-PGE sulfide deposits, layered Ti-Fe oxide deposits, and alkalic hosted U-Nb deposits. The Round Lake Intrusion, like other intrusions discovered through a strong aeromagnetic anomaly (Mudrey et al. 2003), is an example of a layered Ti-Fe oxide deposit. Anorthosite layers alternate with magnetite troctolite layers approximately every 100 ft. Fractional crystallization throughout the evolution of the magma created the alternating “layers” of plagioclase rich and plagioclase poor segments but has minimal additional differences in mineral composition and presence (Stuhr, 1976). This study describes the petrology and geochemistry of the Round Lake intrusion and Timineralization using historic drill cores stored at the Wisconsin Geological and Natural History Survey core repository. Two holes were relogged, totaling ~1787 feet, and representative samples were obtained of host intrusive phases and mineralization types. The intrusion was characterized via transmitted-light petrography and whole rock geochemistry was determined via WD-XRF. Mineral chemistry of the intrusion and mineralization was determined using SEDEDS. The intrusion segregated into layers: anorthosite, upper magnetite troctolite, middle magnetite troctolite, magnetite, and lower magnetite troctolite, crosscut by an intrusive gabbro dike (Stuhr, 1976). The main intrusion hosting mineralization magnetite-ilmenite rich troctolite, ranging from 35-60% intergrown magnetite-ilmenite and 5-20% coarse grained plagioclase laths (Figure 1). Movement and flow of magmas during emplacement are indicated trachytic flow textures of aligned plagioclase crystals. The anorthosite has 55-90% euhedral plagioclase, 10-15% magnetite, and 5-15% clinopyroxene. The magnetite-ilmenite rich troctolite and anorthosite are crosscut by fine-grained gabbroic dikes. Within the magnetite-ilmenite troctolite unit, magnetitetitanomagnetite and lesser ilmenite assumes interstitial growth between silicates (Figure 1). Apatite is variably present. Olivine is variably altered to iddingsite and serpentine strips of magnetite forming within fractures in the crystal. Both the Round Lake intrusion and Clam Lake intrusion are intrusions rich in magnetite associated with the Mid-Continent Rift, and both are likely to be hosts of Ti-Fe ± V deposits and are known to contain large amounts of titanomagnetite with approximately 1.5% V (Woodruff, 2020). Future work is recommended on the Round Lake Intrusion and Ti-mineralization to better constrain the layering and economic potential of Ti-mineralization. 45 �Figure 1: (A) Geologic map of Northwestern Wisconsin region surrounding the Round Lake Intrusion, Digitized from Stuhr (1976). (B) Image of Magnetite-Ilmenite rich troctolite core sample showing textures and magnetite matrix filling features. (C) Image from SEM showing major magnetite and olivine textures within a sample. Magnetite matrix filling texture and fracture filling within olivine fractures. REFERENCES Stuhr, S. W., 1976, Geology of the Round Lake Intrusion, Sawyer County, Wisconsin [Master’s Thesis]: Madison, University of Wisconsin, 148 p. Woodruff, L. G., Schulz, K. J., Nicholson, S. W., Dicken, C. L., 2020, Mineral Deposits of the Mesoproterozoic Midcontinent Rift system in the Lake Superior region – A space and time classification: Ore Geology Reviews, v. 126, p. 1-21. Mudrey Jr., M.G., Ervin, C.P., Olmsted, 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. 46 �Geology and Geochemistry of the Ritche Creek Cu-Zn deposit, North central Wisconsin JOHANNESEN, Haley P. 1, LODGE, Robert W.D.1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA The Ritchie Creek deposit is a Volcanogenic Massive Sulfide (VMS) deposit located within the Paleoproterozoic Penokean Volcanic Belt (PVB) of northcentral Wisconsin (Figure 1). Mineral exploration efforts have demonstrated that the VMS mineralization at this site is concentrated on the western edge of a felsic volcanic center that is interpreted to have formed in a back arc or intra-arc rift environment within bimodal volcanic sequences (DeMatties, 1990). These interpretations were based on physical descriptions of units intersected in drill core and comparisons to other VMS deposits regionally and globally. Like many VMS deposits in the PVB, little research has been done at the Ritchie Creek prospect to link petrogenesis with largerscale VMS environments. This research aims to characterize and refine the geological and geochemical characteristics of the Ritchie Creek Cu-Zn deposit by re-examining historic drill core and representative volcanic stratigraphic units and provide a more comprehensive understanding of the tectonic environment that influenced mineralization. The study involved logging 1,000 linear feet of historic drill core from two holes and collecting 22 core samples from representative stratigraphic units for petrographic and geochemical characterization. The four sampled units include: (1) a medium grey, fine grained quartz mica schist with alternating coarse-grained quartz and K-feldspar bands and disseminated sulfides including pyrite and chalcopyrite, (2) a light green quartz mica schist, strongly altered by sericite and chlorite, containing disseminated chalcopyrite and pyrite, (3) an intermediate metafelsite unit, characterized by sericite and biotite, that grades into a rhyolitic tuff with angular felsic fragments, localized sulfide blebs, and quartz veins, and (4) a semi-massive to massive sulfide unit consists mainly of pyrite with minor chalcopyrite, in a sheared and brecciated matrix. Major and trace element geochemical data was generated via WD-XRF at the Material Science Center at the University of Wisconsin-Eau Claire. Major element geochemistry is highly variable because of varying degrees of hydrothermal alteration. Therefore, immobile trace elements are used to classify protoliths and discriminate tectonic settings. Least-altered volcanic strata were chemically classified as mafic volcanics (based on low Zr/Ti, high Cr), intermediate volcanics (based on elevated Zr/Ti), and felsic volcanics (based on high Zr/Ti). Mafic volcanic strata have high Zr, consistent with calc-alkalic magmatic affinities. Felsic volcanic strata are FII type felsic magmas and have low Nb and Y consistent with volcanic-arc felsic magmas. The quartz-sericite altered rocks have trace element chemistry consistent with the intermediate volcanic strata and alteration indices indicate a potassic-dominated alteration. These characteristics suggest an oceanic arc-backarc bimodal-mafic petrochemical association (Piercey, 2011) and provide a more comprehensive understanding of VMS mineralization in the PVB. 47 �(A) (B) Figure 1. (A) A regional map of the Ritche Creek VMS Deposit located in North Central Wisconsin. (B) A Cross-section view of the Ritche Creek VMS deposit, showing drill hole locations and stratigraphic units, faulting and alteration zones. This cross section focuses on (RC5) a drill holes that intersects significant sericite alteration and massive sulfide mineralization zones. REFERENCES DeMatties, T.A., (1990), The Ritchie Creek Main Zone: A Lower Proterozoic CopperGold Volcanogenic Massive Sulfide Deposit in Northern Wisconsin. Economic Geology Vol. 85, 1990, pp. 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, p. 104489. Piercey SJ (2011) The setting, style, and role of magmatism in the formation of volcanogenic massive sulfide deposits. Mineralium Deposita 46:449-471. 48 �Geologic implications of detrital zircon U-Pb ages from Archean and Paleoproterozoic strata in central Minnesota and the Gogebic Range of Wisconsin and Michigan, USA JONES, James V.1, SALERNO, Ross2, CANNON, William F.2, and O’SULLIVAN, Paul4 1 U.S. Geological Survey, Anchorage, AK 99508, USA jvjones@usgs.gov U.S. Geological Survey, Reston, VA 20192, USA; 3 U.S. Geological Survey, Denver, CO 80225, USA 4 GeoSep Services LLC, Moscow, ID 83843, USA 2 Archean and Paleoproterozoic metasedimentary successions in the Lake Superior region of the northern United States record the assembly and breakup of southern Superia and the subsequent transition to long-lived accretionary orogenesis along the southern Laurentia margin. The successions are difficult to correlate for reasons that include contrasts in thickness, facies, and variable amounts of erosion, similarities in depositional environment through hundreds of millions of years of sedimentation, and variable overprinting by younger tectonic events. Detrital zircon U-Pb geochronology is useful for correlating siliciclastic strata and for identifying provenance patterns that reflect past tectonic and sedimentary interactions. We present new data for samples collected from ca. 2.6–1.8 Ga strata from across the Lake Superior region that provide key insights into regional correlations and local to global tectonic histories. In the eastern Gogebic Range of Michigan, Archean volcanic and volcaniclastic rocks are mapped in a fault-bounded panel between the Watersmeet gneiss dome to the southeast and Neoarchean Puritan batholith to the northwest. One new sample of Archean metagraywacke from within the supracrustal succession yielded only Neoarchean detrital zircon with age populations ranging from ca. 2740 to 2590 Ma, indicating derivation from nearby gneisses but not from older sources such as the early Paleoarchean Watersmeet gneiss. A younger succession of Paleoproterozoic metavolcanic and metasedimentary rocks near Lake Gogebic overlies the Archean gneisses and supracrustal rocks. One sample of fine- to medium-grained slate and metagraywacke from the Copps Formation yielded a mixture of Archean and Paleoproterozoic detrital zircon dates. Archean grains were minor and included age populations of ca. 2649 and 2553 Ma that match the nearby Neoarchean metagraywacke and gneiss domains. Paleoproterozoic grains defined a ca. 1846 Ma age peak and a maximum depositional age of ca. 1829 Ma. We also collected samples of the Paleoproterozoic Palms and Tyler Formations that overlie Archean domains in the western part of the Gogebic Range. Fine-grained gray quartzite of the Palms Formation yielded detrital zircon age populations ranging from ca. 2976 to 2458 Ma and a prominent peak at ca. 2675 Ma. The age spectrum indicates input and (or) recycling of Archean sources and an absence of coeval magmatic sources in the region. In contrast, fine-grained argillaceous sandstone of the overlying Tyler Formation contained mostly Paleoproterozoic detrital zircon with major age peaks at ca. 1863 and 1827 Ma together with minor older age populations ranging from ca. 2780 to 1953 Ma. In central Minnesota, new samples were collected from the Paleoproterozoic Denham and Little Falls Formations. The Denham Formation sample was collected on the northern side of the McGrath gneiss dome and consisted of fine-grained biotite argillite with 1-2 mm horizons of coarser sandstone. The sample yielded chiefly Archean detrital zircon with a dominant age population at ca. 2603 Ma and minor older populations ranging from ca. 3409 and 2789 Ma. Archean age populations match previously published data from nearby samples of basal arkose and dolomitic arkose from the same unit (Craddock et al., 2013). However, that basal arkose also contained a distinct ca. 2101 Ma age population that established a potential correlation between the Denham Formation and the East Branch Arkose of the Dickinson Group in Michigan. The 49 �Little Falls Formation sample of garnet-staurolite-biotite schist was collected from the southern side of the McGrath dome, and it predominately contained Paleoproterozoic detrital zircon that define a dominant unimodal age population at ca. 1846 Ma. The age spectrum for the Little Falls sample is nearly identical our data from the Copps Formation and is also like our Tyler Formation data and to previously published data for other parts of the upper Animikie Group. The marked difference in the proportion of Archean and Paleoproterozoic grains between the Little Falls and Denham Formations suggests a major change in provenance across their contact. The Denham Formation appears to have been derived from the underlying Archean gneiss dome with lesser contribution from older gneisses elsewhere in the region. Circa 2.1 Ga sources are rare in the region but are found locally to the east in Dickinson County, Michigan. The Paleoproterozoic age population that dominates the Little Falls Formation indicates derivation from the Wisconsin magmatic terrane that was approaching from the south (present coordinates) prior to collision that defines the Penokean orogenic cycle in the region. Additionally, our data indicate a maximum depositional age of ca. 1846 Ma for the Little Falls Formation that contrasts with the inferred ca. 2101 Ma age of the underlying Denham Formation reported by Craddock et al. (2013). Published observations suggest a gradational contact between schist of the Little Falls Formation and dolomitic marble of the underlying Denham Formation (Boerboom and Chandler in Bauer et al., 2022). Boerboom and Chandler (2022) noted a 1-meter graphitic/carbonaceous argillite at the base of the Little Falls Formation that could represent a hiatus and then a major change in depositional environment above the arkosic conglomerate and dolostone. We previously reported similar geologic and provenance patterns from the Dickinson Group approximately 600 km to the east in Michigan (Jones et al., 2024). In that area, the East Branch Arkose contains a similar distribution of DZ ages: a mixture of Archean detrital zircon and a distinctive ca. 2099 Ma age population interpreted to have been derived from local granitic sources. The overlying Solberg Schist is made up of biotite-staurolite schist that contains prominent ca. 1.86–1.84 Ga age populations together with minor ca 2.5 and 2.3 Ga age populations. More work is needed to better constrain the stratigraphic position of the depositional age and provenance shifts in both successions and to better understand the tectonic setting and significance of the subtle unconformities and pronounced shift in zircon sources. Preliminary observations and data suggest that the two successions are regionally similar but also distinct from surrounding strata. Thus, the Denham and Little Falls Formations may provide a distinctive and unique record of the transition from Superia rifting to Penokean orogenesis. REFERENCES Bauer, Emily J; Chandler, V.W.; Boerboom, Terrence J; Knaeble, Alan R; Nguyen, Maurice K; Lively, R. S.; Setterholm, Dale R; Steenberg, Julia R. (2022). C-52, Geologic Atlas of Aitkin County, Minnesota. Retrieved from the University Digital Conservancy, https://hdl.handle.net/11299/253808. 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, https://doi.org/10.1086/673265. Jones, J., Cannon, B., Drenth, B., and O’Sullivan, P., 2024, Geologic and tectonic implications of detrital zircon U-Pb age from the Dickinson Group in the western Upper Peninsula of Michigan, USA: Institute on Lake Superior Geology, “Institute on Lake Superior Geology: Proceedings, 2024,” Archives & Digital Collections at Lakehead University Library, accessed April 9, 2025, https://digitalcollections.lakeheadu.ca/items/show/10352. 50 �Zircon Petrochronology of Wisconsin’s Volcanogenic Massive Sulfide Deposits, Northcentral Wisconsin KWIATKOWSKI, Aidan O. 1, LODGE, Robert W.D.1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA Northern Wisconsin’s Paleoproterozoic Penokean Orogen, one of the classic Precambrian orogenic belts in North America, is known to host multiple volcanogenic massive sulfide (VMS) deposits which are important sources of Cu, Zn, Pb, Ag, and Au globally. Despite known large and potentially economic VMS deposits, limited outcrop exposure has hindered detailed reconstructions of the VMS-hosting environment to guide future exploration. Historic U-Pb geochronology indicates that volcanism occurred between 1889-1835 Ma (Sims et al. 1989) with the majority of VMS-hosting strata constrained between 1875-1873 Ma (Quigley 2016). Schultz & Cannon (2007) attribute the main VMS forming event ca. 1875 Ma to extension in a developing back arc basin, with a second later magmatic pulse around 1830 Ma being attributed to post-tectonic stitching plutons. However, a newer model by Zi et al. (2022) shows two VMS forming events around 1875 Ma and 1845 Ma suggesting a regime consisting of alternating compressional and extensional environments caused shifting subduction angles. Zircon petrochronology (U/Pb, Lu/Hf isotopic data and trace elements) can not only better constrain the timing of VMS formation but can also allow for a more complete understanding of the geological evolution and metallogeny of Wisconsin VMS deposits. This study sampled felsic igneous rocks from several VMS deposits to determine the timing and tectonic settings of VMS environments in the western Penokean Orogen. Samples were studied from the Flambeau, Eisenbrey, and Lynne deposits of the Ladysmith-Rhinelander Volcanic Belt (Figure 1a). Samples from the Flambeau and Eisenbrey deposits consist of felsic volcaniclastic units associated with sulfide mineralization and the sample from the Lynne deposits consists of a granodiorite which has intruded into the VMS deposit and volcanic strata. Samples were pulverized and heavy mineral separates were obtained by various magnetic and density separation techniques. The zircon mineral grains were imaged by cathodoluminescence prior to isotopic (U/Pb, Lu-Hf) and trace element analyses via LA-ICPMS at the Mineral Exploration Research Centre at Laurentian University, Sudbury, Ontario. U/Pb isotopic data constrains timing of magmatism. Trace elements and Lu-Hf data constrain the tectonic setting and crustal architecture. Preliminary results indicate two distinct VMS-forming magmatic events during the Penokean Orogeny that have similar tectonic and magmatic styles. All samples show a bimodal distribution of U/Pb ages centered on 1830-1835 Ma and 1870-1875 Ma (Figure 1). Trace element geochemistry of zircons reveals little petrogenetic difference between the magmatic events. Negative ƐHf(i) values, indicating interaction with Archean basement, is consistent amongst all samples and between magmatic events. The VMS-forming extensional event at ca. 1835 Ma contradicts the Schulz and Cannon (2007) model where collisional tectonics are dominant at this time. While the timing of VMS-formation more closely aligns with the Zi et al. (2022) model, the accordion-like tectonics cannot explain the lack of variation in magmatic setting or crustal architecture observed in our data. Therefore, additional data is needed to fully understand the tectonic and metallogenic significance of this younger extensional event. 51 �Figure 1. A) Generalized geologic map of the Penokean Orogen illustrating major tectonostratigraphic subdivisions and the location of sampled and major VMS occurrences. Figure modified from Schulz and Cannon (2007) and DeMatties (1994). Subdivisions of Pembine-Wausau terrane from DeMatties (1994, 2018). LRVC = Ladysmith-Rhinelander volcanic complex. B) U/Pb concordia diagram of older magmatic zircons. C) U/Pb concordia diagram of younger magmatic zircons interpreted to be crystallization age of sample. D) Weighted mean diagram distribution of ages and analyzed grains. Inset image shows frequency distribution of ages in samples. REFERENCES DeMatties TA (1994) Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Econ Geol 89: 1122-1151. DeMatties TA (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 Geol Rev 95: 216-237. Quigley A (2016) Setting of the volcanogenic massive sulfide deposits in the Penokean Volcanic belt, Great Lakes region, USA. Colorado School of Mines, Masters Thesis. 95 p. Schulz KJ, Cannon WF (2007) The Penokean orogeny in the Lake Superior region. Precambrian Res 157: 4-25. Sims PK, Van Schmus WR, Schulz KJ, Peterman ZE (1989) Tectonostratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean orogen. Can J of Earth Sci 26: 2145-2158. Zi J-W, Sheppard S, Muhling JR, Rasmussen B (2021) Refining the Paleoproterozoic tectonothermal history of the Penokean Orogen: New U/Pb age constraints from the Pembine-Wausau terrane, Wisconsin, USA. Geol Soc Am Bull 134: 776-790. 52 �Geologic Interpretation of Filtered Gravity and Magnetic Anomalies of the Baraboo Range LONGACRE, Mark B.1 and HINZE, William J.2 1 MBL, Inc., 51 Captain Perry Dr., Phippsburg, ME 04562 Purdue University, 30 Brook Hollow Ln., West Lafayette, IN 47907 2 Investigations over the past decade have made significant advances in our geologic knowledge of the Mesoproterozoic Baraboo Synclinorium and adjacent region of south-central Wisconsin (e.g., Medaris, Jr. et al., 2021; Stewart et al., 2021; Marshak et al., 2023). To further the geologic information of this feature and nearby region we have filtered their gravity and magnetic anomaly maps to identify geologic formations and structures in the crystalline basement. The filtered maps isolate specific attributes of the anomaly fields which are useful in interpretation especially when combined with constraining geological information. These maps have identified a buried geologic structure to the east of and immediately adjacent to and along strike of the Sauk Syncline (Figures 1 and 2) which encompass the Baraboo Synclinorium. The buried structure is a near mirror image of the Sauk Syncline and thus is referred to as the Twin Syncline. Unlike the Sauk Syncline and the Baraboo Synclinorium the eastern structure is south rather than north of a geological lineament within the Yavapai orogenic province that marks the southern boundary of the Wisconsin Gravity Minimum. The Twin Syncline is notable in the magnetic anomaly map because of the positive anomaly associated with a magnetite-rich formation that is likely an extension of the lower portion of the Freedom Formation of the Baraboo Synclinorium. The elliptical trace of this anomaly and the steep gradients of the outer margin support the synclinal nature of the structure. We interpret this structure to be a result of south-verging thrusting with steeply dipping thrusts along the northern and southern margins of the Twin Syncline similar to the situation of the Sauk Syncline. This structure is not as tightly folded as the Baraboo Synclinorium suggesting that the thrusting to the east of Baraboo Synclinorium was less intense. The Syncline can be identified on the filtered maps as can other quartzite synclinoriums of the region by the subdued geophysical anomalies of the underlying felsic volcanic rocks because of their burial beneath the non-magnetic quartzite. The magnetic anomalies of the magnetic lower half of the Freedom Formation are also a useful marker for detailing the structure within the Baraboo Synclinorium and defining the limits of the Freedom Formation within it. Additionally, two parallel intrusives on strike with the Denzer Diorite that crops out along the southwestern margin of the Synclinorium extend northnortheasterly within the sub-quartzite basement across the western portion of the Synclinorium. They are associated with anticlines within the Synclinorium that may have resulted from differential deformation caused by variations in the rheology of the basement rocks. These and other interpretations of the filtered gravity and magnetic anomalies suggest that revisiting the studies of basement rocks of Wisconsin and adjacent regions is in order using the available improved analysis, interpretation, and presentation methods and modern data gravity and magnetic data sets. REFERENCES Marshak, S., Wilkerson, M.S., and DeFrates, J., 2023. Kinematic and tectonic implications of crenulation cleavage, kink bands, and mesoscopic folds in the Baraboo Syncline, Wisconsin (∼1.45 Ga Picuris Orogen). Journal of Structural Geology, 178, 105007. Medaris, Jr, L.G., Singer, B.S., Jicha, B.R., Malone, D.H., Schwartz, J.J., Stewart, E.K., Van Lankvelt, A., Williams, M.L., and Reiners, P.W., 2021. Early Mesoproterozoic evolution of midcontinental Laurentia: Defining the geon 14 Baraboo orogeny. Geoscience Frontiers,12(5), 101174, 17 p. 53 �Stewart, E.K., Brengman, L.A., and Stewart, E.D., 2021. Revised provenance, depositional environment, and maximum depositional age for the Baraboo (< ca. 1714 Ma) and Dake (< ca. 1630 Ma) Quartzites, Baraboo Hills, Wisconsin. The Journal of Geology, 129, 1-31. Figure 1. Total horizontal derivative of the RTP total magnetic anomaly map of south-central Wisconsin. The outlines of the Sauk (left) and Twin (right) Synclines are shown by the thin dashed white lines and the Baraboo Synclinorium by a dashed white line. The Baraboo Lineament of the Yavapai province is shown by the wide broadly dashed line. Color coding is non-linear. Figure 2. Tilt derivative of the Bouguer gravity anomaly map of south-central Wisconsin that emphasizes the short wavelength components. The outlines of the Sauk (left) and Twin (right) Synclines are shown by the dashed white lines and the Baraboo Synclinorium by a thin dashed white line. The Baraboo Lineament of the Yavapai province is shown by the wide broadly dashed line. Color coding is non-linear. 54 �An Informal Review of the ILSG Field Excursion to Hawaii, January – February, 2025 MACTAVISH, Allan1, HINZ, Peter1, HUDAK, George1, LARSON, Phil1, AUBUT, Allan1, BOERBOOM, Terry1, CHILTON, Vern1, DeGRAFF, Jim1, ERICKSON, Tom1, FAULKNER, Barb1, SERRANO, Isabel1, and ZANKO, Larry1 1 Members of the 2025 ILSG Field Trip to Hawaii, 2025 Between January 24, 2025 and February 5, 2025, twelve members of the Institute on Lake Superior Geology participated in a geological field excursion to investigate the geology of the island of Hawaii, with a focus on observing field relationships, outcrop characteristics and geomorphology to better understand the characteristics of modern basaltic volcanism in a hotspot environment. The field excursion was led by Allan MacTavish, Peter Hinz, George Hudak and Phil Larson. A new field trip guidebook and glossary of geological terms (MacTavish and Hudak, 2024) was prepared and utilized during the thirteen-day long trip. This presentation will review key features and take-aways from the excursion, which included investigations of five of the seven volcanoes associated with the island of Hawaii. Investigations took place via examinations of various outcrops, hikes through the Hawaiian wilderness, and a helicopter tour. Various eruption types, volcano types, coherent (lava flow) and volcaniclastic deposit types and features, different types of volcanic products and hydrothermal alteration facies, and observations of historical and cultural artifacts and natural phenomena will be discussed. Challenges and surprises associated with field studies of Hawaii will also be presented. REFERENCES MacTavish, A., and Hudak, G., 2024, The Volcanoes of the Island of Hawaii – Field Trip Guide: Institute on Lake Superior Geology Special Publication 3, 200 p. 55 �56 �Refining the Age and Occurrence of Basement Rocks in Northwest Iowa: Implications for Precambrian Tectonics and Magmatic Evolution of the Laurentian Midcontinent MALONE, Jack1, MALONE, David2, ANDERSON, Raymond1, CLARK, Ryan1 1 Iowa Geological Survey, University of Iowa, Iowa City, IA 52242 USA 2 Geography-Geology, Illinois State University, Normal, Illinois 61790 Precambrian basement rocks in northwest Iowa reveal an Archean and two Paleoproterozoic tectonic sutures (Figure 1; 1.9-1.8 Ga Trans-Hudson/Penokean and 1.8-1.7 Ga Yavapai; Bickford et al., 1986; Holm et al., 2007). Here we present four new U-Pb (LA-ICPMS) ages for drill cores of basement rocks along the Transcontinental Arch in northwest Iowa, USA (Figure 2). The cores are on repository at the Iowa Geological Survey. The Camp Quest migmatite gneiss was sampled at a depth of 1,078 ft from the Camp Quest D-21 core (W25498; z=38). The weighted mean and Concordia ages were both 1845 Ma, which is the first TransHudson/Penokean age recognized in Iowa. This core is located south of the Spirit Lake tectonic zone (SLTZ) which is interpreted as the suture between Yavapai terrane rocks to the south and Archean Superior province rocks to the north. Nine inherited zircons are mostly Archean in age and interpreted as xenocrysts, indicating Archean crust occurs at depth south of the SLTZ. Granite was also sampled at a depth of 660 ft from the Hawarden D-7 core (W27270; z=35). The zircon age spectrum reveals three age clusters at ~2895, ~2683, and ~1800 Ma. The older, inherited age clusters are consistent with ages of the Minnesota River Valley terrane and the greater Superior Province, respectively. The ~1800 Ma age is similar to the nearby 1803-1810 Ma Matlock “keratophyre” and the distant 1805 Ma Humboldt granite (northern Michigan), representing the initiation of north-directed Yavapai subduction and granitic melt production into Archean and previously accreted Trans-Hudson/Penokean rocks north of the SLTZ (Kilburg, 2024). Granodiorite was sampled at a depth of 915 ft from the Harris D-13 core (W27270; z=41). The weighted mean and Concordia ages are ~1780 Ma, suggesting that Yavapai rocks intrude older Trans-Hudson/Penokean or Archean rocks north of the SLTZ. A late-stage granitic dike was sampled at a depth of 1,611 ft from the Spencer BX-2 core (W16223; z=7), which is from a tabular noritic body within the Spencer intrusive complex just south of the SLTZ. The sampled interval yielded sparse zircons; however, the weighted mean age of 1238 Ma is the first Grenville age recognized in Iowa. This age suggests an obscure early Grenvillian thermal resetting or reactivation in the upper Midcontinent which postdates anorthositic/noritic magnetism concentrated along the SLTZ at Spencer. New complementary whole rock WDXRF major oxide and ICP-OES trace element geochemical analyses (n=163) from Precambrian units in northwest Iowa reveal a complex tectonic and crustal growth configuration. Intermediate to felsic intrusions are generally LREEenriched and have I-type volcanic arc-like trace element patterns. The origin of anorthositic to mafic-ultramafic occurrences are less straightforward but are characterized by slight to significant Ce, Sm, Eu, and Lu anomalies, indicating basaltic to mantle fractionation, differential partial melting at depth, and/or derivation from Fe-rich residual melts. These new results provide significant insight into the tectonomagmatic evolution of the southernmost Superior Province during the final assembly of the Laurentian craton. 57 �Figure 1: Geological map of Precambrian basement rocks in the northern midcontinent and northwest Iowa. Top: Red dots indicate previously published U-Pb ages and white dots are new (this study). Bottom: New geochronologic ages indicated with stars are CQ = Camp Quest, HW = Hawarden, HA = Harris, SP = Spencer. Figure 2: Weighted mean, probability density, and Concordia plots of newly dated Precambrian units in northwest Iowa. REFERENCES Bickford, M.E., Van Schmus, W.R., and Zeitz, I., 1986. Proterozoic history of the midcontinent region of North America. Geology, 14(6), 492-496. Holm, D.K., Anderson, R., Boerboom, T.J., Cannon, W.F., Chandler, V., Jirsa, M., Miller, J., Schneider, D.A., Schulz, K.J., & Van Schmus, W.R., 2007. Reinterpretation of Paleoproterozoic accretionary boundaries of the north-central United States based on a new aeromagnetic-geologic compilation. Precambrian Research, 157(1-4), 71–79. Kilburg, N., 2024. Age and petrogenesis of the Matlock ‘Keratophyre’ in northwest Iowa [M.S. Thesis]: Iowa City, University of Iowa, 129 p. 58 �Post-Penokean and Pre-Yavapai Magmatism and Sedimentation in Central Wisconsin (Southern Lake Superior Region) MEDARIS, Gordon Jr.1 and MALONE, Dave2 1 Dept. of Geoscience, University of Wisconsin-Madison, Madison, WI 53706 2 Dept. of Geography, Geology, and the Environment, Illinois State University, Normal, IL 61790 The principal Precambrian domains in Wisconsin are the Penokean Province, consisting of the Marshfield Terrane, Wausau-Pembine Terrane, and Craton margin, which include 2450-1770 Ma craton margin and foreland basin sediments and 1890-1830 Ma volcanic arc associations, 1760 Ma rhyolite and granite of the Yavapai Province, <1643 Ma quartzite of the Baraboo Interval, 1484-1468 Ma granitic rocks of the Wolf River batholith, and 1109-960 Ma igneous and sedimentary rocks of the Midcontinent Rift (Fig. 1). In addition to these five major domains, small outcrops of post-Penokean and preYavapai igneous and sedimentary rocks are scattered across central Wisconsin, which have been investigated in detail at Hamilton Mounds (Medaris et al., 2007), Biron Dam (Holm et al., 2020), and Brokaw (this report) (Fig. 1). Two sedimentary successions occur at Hamilton Mounds: an older arkose and a younger quartzite correlative with the Baraboo quartzite. The arkose is a gray, fine- to medium-grained, feldspathic sandstone (CIA = 59.0; Fig. 2). Detrital monazite in arkose yields a total Pb median age of 1850 Ma, with the youngest detrital grain at 1757 Ma, signifying post-Penokean deposition of the arkose. An upper age for the arkose is provided by the intrusion of 1762 ± 7 Ma granite (Yavapai), whose age is within error of the youngest detrital monazite grain. Muscovite in the younger quartzite yields a 40Ar/39Ar plateau age of 1470 ± 11 Ma, reflecting the widespread thermal effect of the Wolf River batholith throughout central and southern Wisconsin. At Biron Dam, trachybasaltic diabase dikes (Fig. 2) intruded Archean gneiss and Penokean tonalite, granodiorite, and granite. Zircon grains in three samples of diabase yield 207 Pb/206Pb ages within error of each other, with a weighted mean age of 1817 ± 2 Ma, which demonstrates post-Penokean and pre-Yavapai emplacement of the dikes. The diabase dikes have been metamorphosed under amphibolite-facies conditions; hornblende in metadiabase yields a 40 Ar/39Ar plateau age of 1672 ± Ma, possibly representing a Mazatzal influence. At Brokaw, polymictic conglomerate, feldspathic sandstone (CIA = 59.0; Fig. 2) and siltstone were intruded by rhyolite, which contains inherited zircon with ages between 2125 Ma and 3565 Ma. The sandstone contains detrital zircon with a 207Pb/206Pb median age of 1850 Ma and an age of 1810 Ma for the youngest subset of grains with overlapping errors, demonstrating post-Penokean deposition of the Brokaw sedimentary rocks. Primary structures and textures of the Brokaw igneous and sedimentary rocks have been preserved on the macroscopic scale, but such rocks have been pervasively recrystallized to greenschist-facies mineral assemblages on the microscopic scale, as seen for example in rhyolite, in which plagioclase was replaced by albite and epidote, and hornblende, by epidote (Fig. 3). The age of such recrystallization has not yet been determined, but is presumed to be related to the nearby Wolf river batholith. It is now recognized that igneous rocks were emplaced and sedimentary rocks were deposited over much of central Wisconsin in the interval 1817-1757 Ma after the Penokean orogeny, perhaps as a precursor to the Yavapai orogeny. 59 �Figure 2. Chemical compositions of Biron Dam diabase, Brokaw rhyolite and sandstone, and Hamilton Mounds sandstone in terms of Al (Al2O3), Ca* (CaO), N (Na2O), and K (K2O); CIA: Chemical Index of Alteration. Figure 1. Map of the major Precambrian geological units in the southern Lake Superior region. Star symbols: Brokaw (BK), Biron Dam (BD), and Hamilton Mounds (HM) localities; B: Baraboo Interval sedimentary rocks. Figure 3. Photomicrograph (crossed polarizers) of recrystallized Brokaw rhyolite; ab, albite; ep, epidote REFERENCES Medaris, L.G. Jr., Van Schmus, W.R., Loofboro, J., Stonier, P.J., Zhang, X., Holm, D.K., Singer, B.S., and Dott, R.H. Jr., 2007. Two Paleoproterozoic (Statherian) siliciclastic metasedimentary sequences in central Wisconsin. Precambrian Research, 157, 188-202. Holm, D., Medaris, L.G. Jr., McDannell, K.T., Schneider, D.A., Schulz, K., Singer, B.S., and Jicha, B.R., 2020. Growth, overprinting, and stabilization of Proterozoic Provinces in the southern Lake Superior region. Precambrian Research, 339, Article 105587. 60 �US Steel Corporation / Ralph W. Marsden iron ore collection MOOERS, Howard1, SEVERSON, Mark2, JONGEWAARD, Peter3, LARSON, Phillip4 1 Department of Earth and Environmental Sciences, University of Minnesota Duluth, Duluth, MN 55812 2 2122 W 22nd St., Duluth, MN 55811, USA 3 7009 Three Lake Rd., Canyon, MN 55717, USA 4 1613 14th Ave. East, Hibbing, MN 55746, USA By the time Ralph W. Marsden joined Oliver Iron Mining Division of US Steel Corporation (USSC) in 1951 he was already one of the World’s experts on iron ore. From 1953-1964 he managed the Geologic Investigations Unit in Duluth, MN. During this time, Ralph was one of the co-founders of the Institute on Lake Superior Geology in 1954. In 1964 Ralph was transferred to the Pittsburgh corporate office as Manager of Geologic Investigations, Iron Ore, however, Ralph wanted to return to Minnesota, and in 1967 he left USSC and moved to the University of Minnesota Duluth (UMD) Department of Geology as Professor and Head. USSC had an active, worldwide exploration program for iron ore from the 1920s into the 1960s, and a large number of the samples collected were housed in Duluth, MN. When USSC closed its Duluth, MN, office, this iron ore sample collection was to be discarded. Ralph “rescued” the collection of iron ore samples and moved them to the University of Minnesota Duluth. In 1986, Ralph died suddenly while attending the Geological Society of America Annual Meeting in San Antonio, TX. The collection of iron ore samples sat in a service tunnel at UMD for 40 years. This globally significant collection of iron ore samples was recently inventoried, photographed, and placed in storage containers that are readily accessible. The inventory of the 483 samples, complete with photographs, is cataloged on the University of Minnesota Digital Conservancy (https://hdl.handle.net/11299/265081). Many of these samples are from localities that are no longer accessible, are from closed mines, or are from areas of the World that simply cannot be visited because of political and social issues. This collection of iron ore samples dates from 1926 to the 1960s and has samples from 25 countries and 30 US states and Canadian provinces. The individual sample boxes are labeled, and many have great detail on the origin of the samples. Most of the samples are also individually labeled, with sample numbers and descriptions. There are photographs of the contents of each box, and where possible supporting documents are shown in the photos. For further information or to request access to samples contact the Department of Earth and Environmental Sciences, University of Minnesota Duluth or Howard Mooers (hmooers@d.umn.edu). 61 �Countries represented: USA, Angola, Australia, Brazil, Canada, Chile, Colombia, Congo, Costa Rica, Cuba, Gabon, Germany, Guatemala, Honduras, India, Ivory Coast, Liberia, Mexico, Nicaragua, Namibia, Portugal, South Africa, Sudan, Sweden, Venezuela. US States and Canadian Provinces represented: Alabama, Alberta, Arizona, British Columbia, California, Idaho, Illinois, Massachusetts, Michigan, Minnesota, Missouri, Montana, Nevada, New Jersey, New Mexico, New York, Newfoundland, North Carolina, North Dakota, Ontario, Oregon, Puerto Rico, Quebec, South Dakota, Utah, Virginia, Washington, Wisconsin, Wyoming. Figure 1. Example of samples from Liberia, West Africa, with supporting documentation. REFERENCES University Digital Conservancy, University of Minnesota Duluth, (2024). List of Samples for US Steel Corporation / Ralph W. Marsden Iron Ore Collection. Retrieved from the University Digital Conservancy, https://hdl.handle.net/11299/266399. 62 �Lithogeochemical Characterization of Manganese Mineralization at the Cuyuna Range, Central Minnesota PALIEWICZ, Cory1, THAKURTA, Joyashish1 1 Natural Resources Research Institute (NRRI), University of Minnesota Duluth, 5013 Miller Trunk Hwy, Duluth, MN 55811 The Paleoproterozoic Cuyuna Range of central Minnesota contains elevated levels of manganese when compared to other Banded Iron Formations in the Lake Superior region. The total tonnage is estimated at 49 million metric tons at 7.84 percent Mn (Kilgore and Thomas, 1982). The Cuyuna Range consists of a Penokean fold-and-thrust belt divided into the Emily District, North Range, and South Range. These are separated by structural and stratigraphic discontinuities which make each area geologically distinct (Southwick et al., 1988; Morey, 1990). Although prior work has documented a variety of textural and sedimentary associations, this study will provide new lithogeochemical data to further characterize the manganese-bearing lithologies across the Cuyuna Range in support of ongoing research for manganese and other critical minerals in Minnesota as part of the USGS Earth MRI program. A total of 201 drill core samples were collected from 37 drill holes across the Emily District, North Range, South Range, and Glen Lake Sulfide Deposit (Figure 1). To date, all samples have been studied in hand-sample and sent for bulk geochemical analysis, 40 samples have been analyzed in thin section, and whole-rock geochemical results of 60 samples from 16 drill holes have been received from the USGS. Although lithologic features of both ironformations and non-iron-formations are variable across the range, the deposits also share many attributes. As such, we find it useful to texturally classify the collected samples into granular, banded, and irregular types while still recognizing the special characteristics of each individual mineral association. This study will present petrographic and whole-rock geochemical data, with particular emphasis on rocks from the Emily District, which from past studies is known to be mostenriched in Mn-content. In addition, Mn-bearing country rocks throughout the Cuyuna Range are also characterized and compared to historic drill logs and prior work (e.g., Morey et al., 1991, Dahl et al., 1992). In this way, new insights on lithological variation, manganese distribution, and other potential critical minerals at the Cuyuna Range may further be addressed and incorporated during the Earth MRI program. 63 �Drill Hole Sampled Figure 1: Regional geologic map of the Cuyuna Range showing approximate drill hole locations sampled for this study. Modified from Southwick et al., 1988 and Cleland et al., 1996. REFERENCES Dahl, L.J., Brink, S.E., Blake, R.L., Tuzinski, P.A., and Adamson, N.R., 1992, Site characterization of Minnesota manganese deposits to evaluate the potential for in-situ leach mining: Littleton, Colorado, Society for Mining, Metallurgy and Exploration, Inc. Preprint 92-243, 31 p. Cleland, J.M., Morey, G.B., and McSwiggen, P.L., 1996, Significance of tourmaline-rich rocks in the North Range Group of the Cuyuna Iron Range, east-central Minnesota: Economic Geology, v. 91, no. 7, p. 1282-1291 Kilgore, C.C., and Thomas, P.R., 1982, Manganese availability-Domestic: U.S. Bureau of Mines Information Circular 8889, 14 p. Morey, G.B., 1990, Geology and manganese resources of the Cuyuna iron range, east-central Minnesota: Minnesota Geological Survey Information Circular 32, 28 p. Morey, G.B., D.L. Southwick, and S.P. Schottler, 1991, “Manganiferous Zones in Early Proterozoic Iron Formation in the Emily District, Cuyuna Range, East Central Minnesota.” Minnesota Geological Survey Report of Investigations 39. 42 pp. Southwick, D.L., Morey, G.B., and McSwiggen, P.L., 1988, Geologic map (scale 1:250,000) of the Penokean orogen, central and eastern Minnesota, and accompanying text: Minnesota Geological Survey Report of Investigations 37, 25 p., 1 pl. 64 �Michigan Geological Survey’s Contributions to the USGS Earth MRI National Mine Waste Inventory Effort PEARSON, Sara1, GAMET, Nolan2, SHALIFOE, Molly 1, QUIGLEY, Ashley2, and MAHIN, Robert2 1 Michigan Geological Survey, Western Michigan University, 5272 W. Michigan Ave. Kalamazoo, MI 49009 2 Michigan Geological Survey, Western Michigan University, 416 Avenue C Gwinn, MI 49841 In the mid-19th century, the discovery of rich copper and iron deposits in Michigan’s Upper Peninsula (U.P.) led to intense mining, resulting in hundreds of abandoned mine waste sites. Both published and unpublished geological literature suggests that some of these legacy mine waste sites have the potential to host critical minerals, such as manganese and graphite, that were previously overlooked during production. The Michigan Geological Survey (MGS) is contributing to the United States Geological Survey’s (USGS) national effort to build a comprehensive national inventory of mine wastes, their compositions, and potential critical minerals. The MGS team has completed an inventory and submitted 120 mine waste sites from 6 counties across the western U.P. to the USGS for a final review and inclusion in the national mine waste database (Figure 1). These 120 sites are further subdivided into 216 individual mine waste features that met the minimum 2,000m2 size requirement. Finalized point and polygon layers for each mine site were accompanied by corresponding geology, resource, and reference attribute tables. The process consisted of creating an ArcGIS Pro project, adding all available mine-related state and federal datasets, LiDAR-derived DEMs (digital elevation models), published maps, and an ArcGIS geodatabase template containing feature classes and related attribute tables required by the USGS. Initial mine waste inventory work focused on searching for and digitizing mine waste features throughout the western U.P that exceeded the 2,000m2 size requirement. The MGS team originally located and digitized 441 mine waste features by utilizing LiDAR-derived, 1-meter DEMs, 2024 ESRI areal imagery, and published geologic maps. This process is depicted by a simplified workflow shown in Figure 2. The mine waste features were then filtered based on their size. Those smaller than 2,000m2 were omitted from the master dataset. Corresponding attribute tables were then populated with data from publicly available literature, websites, state and federal datasets, and information archived in the state’s drill core repositories. The final databases will ultimately comprise the most up-to-date record of the volume, tonnage, grade, and mineralogy of Michigan’s legacy mine waste sites. Future MGS work within the scope of the Earth MRI Mine Waste Cooperative Agreement is a mine waste characterization effort, which aims to sample and evaluate nonfuel mine waste sites that potentially contain critical minerals. This project will begin in 2025 and continue through 2026. 65 �Figure 1. Map displaying all mine waste features inventoried and submitted to the USGS for the fiscal year 2023 Priority 1 funding represented as purple points and polygons. Figure 2. Simplified process to locate and digitize the mine waste features using ArcGIS Pro coupled with online sources. A.) ESRI imagery (Esri, 2024); B.) Bedrock geology of central Dickinson County, MI (James and others, 1961); C.) 1-m QL2 LiDAR DEM model; D.) Digitization of mine waste features. REFERENCES Esri, 2024, World imagery: Esri, https://services.arcgisonline.com/ArcGIS/rest/services/ World_Imagery/MapServer. 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, 1:24,000. 66 �Critical Mineral Potential of the Northern Margin of the Watersmeet Gneiss Dome, MI USA QUIGLEY, Ashley K.1, MAHIN, Robert A. 1, and GAMET, Nolan G. 1 1 Michigan Geological Survey, Western Michigan University, 416 Avenue C Gwinn, MI 49841 Precambrian gneisses and schists on the northern margin of the Watersmeet Dome in Michigan have been shown to be unusually enriched in rare earth elements, fluorite and incompatible elements including U, Th, Hf, and Zr (Barovich et al., 1991; Sims, 1990). The area is within two Earth Mapping Resources Initiative (EMRI) critical mineral focus areas for IOCG/IOA and magmatic REE deposits (Dicken and others, 2022). To further assess the potential for critical minerals, the Michigan Geological Survey (MGS) is conducting detailed geologic mapping and sampling, as well as collecting geophysical and geochronological data. The project area is roughly 36 square kilometers on the border of Gogebic and Ontonagon Counties and 10 kilometers northwest of the town of Watersmeet, MI. Field work began in July of 2024 with a projected completion date in early 2026. During the 2024 field season, the MGS mapped, described and recorded 620 outcrops in the project area using ArcGIS Field Maps and submitted 124 samples for whole rock and trace element geochemistry. An RS-230 BGO gamma-ray spectrometer was used to take over 600 total gamma (K/U/Th) measurements from outcrop. Additionally, a drone-borne, high resolution magnetic survey was flown over areas where permission was granted by landowners. Preliminary field observations include the presence of fluorite in outcrop spatially associated with magnetic and gamma count anomalies. The results of the lithogeochemistry show a strong spatial correlation between fluorine, uranium, thorium, and total REEs. When rare earth element concentrations were converted to industry standard rare earth oxides (REOs), 14 samples had total rare earth oxide (TREO) values greater than 1,000 ppm (Hellman and Duncan, 2018). Geophysical, analytical, and field data also identified an anomalous magnetic high approximately 500m x 300m associated with previously undescribed REE-bearing, magnetic, fine-grained schists. In 1982, Rocky Mountain Energy (RME) conducted exploration drilling for uranium based on anomalous gamma radiation in outcrop. A reexamination of the core found chalcopyrite in close association with fluorite. The presence of anomalous F, Cu, U, REE and magnetite is suggestive of an IOA/IOCG footprint (Hitzman, 2000). This will be investigated using IOCG discrimination diagrams such as Montreuil and others (2013). Barovich et al. (1991) observed that the elevated REEs, fluorite and incompatible elements were tied to a gneiss and schist unit with an interpreted Paleoproterozoic age between 1.9 and 1.7 Ga, much younger than the Archean aged rock units that make up most of the Watersmeet gneiss dome. Because of the apparent link between rock age and critical minerals, confirming existing ages with modern U-Pb dating techniques, as well as adding ages from new locations, is an important piece of this study. Five samples were submitted for U-Pb geochronology of zircon grains. Results are pending. 67 �REFERENCES Barovich, K.M., Patchett, P.J., Peterman, Z.E., and Sims, P.K., 1991. Neodymium Isotopic Evidence for Early Proterozoic Units in the Watersmeet Gneiss Dome, Northern Michigan. U.S. Geological Survey Bulletin 1904-G: G1-G7. Dicken, C.L., Woodruff, L.G., Hammarstrom, J.M., and Crocker, K.E., 2022, GIS, supplemental data table, and references for focus areas of potential domestic resources of critical minerals and related commodities in the United States and Puerto Rico (ver. 2.0, April 2024): U.S. Geological Survey data release, https://doi.org/10.5066/P9DIZ9N8. Hellman, P.L. and Duncan, R.K., 2018, Evaluating Rare Earth Element Deposits. ASEG Extended Abstracts. 2018. 1. 10.1071/ASEG2018abT4_3E. Hitzman, M.W., 2000, Iron oxide-Cu-Au deposit: What, where, when, and why, in Porter, T.M., ed., Hydrothermal iron oxide copper-gold and related deposits a global perspective: Adelaide, Australian Mineral Foundation, p.9–26. Montreuil J-F., Corriveau L., Grunsky E., 2013. Compositional data analysis of IOCG systems, Great Bear magmatic zone, Canada: To each alteration types its own geochemical signature. Geochem. Explor. Environ. Anal. 13:219–247. Sims, P.K., 1990, Geologic map of Precambrian rocks, Marenisco, Thayer, and Watersmeet 15-minute quadrangles, Gogebic and Ontonagon counties, Michigan, and Vilas County, Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map I-2093, scale 1:62,500. 68 �Plume control on the initiation of Mid-Continent Rift breakup using Unconformities: Implications for the Tectono-magmatic evolution and mineral deposits ROHRMAN, Max1 1 DECAN Geosolutions, PO Box 131148, Houston, TX 77219 Regional unconformities from the stratigraphic record interpreted on existing Multi Channel Seismic (MCS) data obtained by Grant Norpac/Argonne (red numbered) and the GLIMPCE program (red lettered) (Figure 1A), are used for temporal and spatial control on MidContinent Rift (MCR) evolution. This allows identification of key events in the evolution of the rift, whereas potential field data, seismic refraction and Rayleigh waves, help constrain spatial and quantitative constraints. Based on magmatic stage definition, two regional unconformities were interpreted from MCS data: MU (Magmatic Unconformity), at the top of the Main stage (~ 1100 – 1089 Ma), signaling the end of major flood basalt magmatism, and BU (Breakup Unconformity) representing the Latent stage (~ 1104 – 1100 Ma). The latter is observed as a sequence at Mamainse Point (Figure 1), rather than an unconformity, stressing the importance of spatial control on events. Magmatic crustal thicknesses and lower crustal seismic velocities obtained from MCS and refraction data (Shay and Trehu, 1993) are used to constrain relative importance of important parameters in melt production, such as: potential temperature, active mantle upwelling and lithospheric thinning. Together, these data suggest that the MCR originated from an earlier NW-SE pre- or proto-rift (blue, Figure 2A) recognized from outcrop (Figure 1A) and MCS, further reconstructed by aligning Archean granitic blocks such as White Ridge (WR), Grand Marais (GM) and Wawa-Abitibi (purple, WA) from gravity lows (Figure 1B, 2A). The area was affected by a plume constrained by a Rayleigh Wave Low Velocity Anomaly (RWLVA) (Foster et al., 2020) (Figure 1A). This generated uplift in central Lake Superior focused on a region around the Coldwell Complex (Figure 1A). Subsequently, Earlystage (~ 1110 - 1104 Ma) magmatism in the proto-rift generated by NE-SW extension along strike slip faults such as the Thiel Fault (TF) (Figure 1A), in the central and eastern arm of the MCR. By the end of the Early-stage, the plume was deeply embedded in the lithosphere and initiated the start of a thick N-S crustal ridge or proto-hotspot track in central Lake Superior during the late Early- to Latent stage (Figure 2B,C). After a break in activity recorded by the Breakup Unconformity (BU), the plume moved relatively southward during the Main-stage and possibly influenced stress re-orientation to N-S (Figure 2D). This locked the eastern arm and locally, new thick oceanic crust formed along the syncline in central Lake Superior, generating the western rift arm. However, magmatism and breakup terminated shortly after as a result of Grenvillian compression, evidenced by the Magmatic Unconformity (MU). During the Main stage, active upwelling and anomalously thick oceanic crust formation was highest on the crustal ridge (black dash-dot line, Figure 2D), measured at line A, just north of the WA block (purple arrow, Figure 2D) and decreasing toward line C (purple arrow). Further west, at St Croix, upwelling rates approach unity and no oceanic crust formation took place. Pulsing and waning of the plume stem/conduit through time (Figure 2) is recorded in the unconformities, suggesting a drop in potential temperature and upwelling rate around BU time (Latent stage) (Figure 2C). 69 �Figure 1: A. Geological map with seismic lines (red). Numbering refers to onshore geological sections. B. Gravity map. Abbreviations: MB Marquette Basin, KP Keweenaw Peninsula, HVB High Velocity Body. Figure 2: Tectono-magmatic evolution. South shore (between yellow cubes) is mobile, North shore is kept fixed. EPC Early Plume center, LPC Latent Plume Center, MPC Main Plume Center. REFERENCES Foster, A., Darbyshire, F., and Schaeffer, A., 2020. Anisotropic structure of the central North American Craton surrounding the Mid-Continent Rift: evidence from Rayleigh waves. Precambrian Research, 342: 105662. Shay, J., and Trehu, A., 1993. Crustal structure of the central graben of the Midcontinent Rift beneath Lake Superior. Tectonophysics, 225: 301-335. 70 �Constraining the timing of crustal exhumation following the Penokean orogeny using U-Pb, Sm-Nd, and Lu-Hf geochronology and microstructural analysis SALERNO, R.,1 CANNON, W.F.,1 SOUDERS, A.,2 THOMPSON, J. M.,2 VERVOORT, J.,3 1 U.S. Geological Survey, Reston, VA 20192, 2U.S. Geological Survey, Denver, CO 80225, 3Washington State University, Pullman, WA 99164. Precambrian terranes in the Lake Superior region have complex igneous, metamorphic, and deformational histories spanning the Eoarchean to the Neoproterozoic. In this sequence, the Penokean orogeny (1880–1830 Ma) is the first collisional event in a long-lived subduction system on Laurentia’s southern margin, marking a transition in the style of Laurentian assembly from the amalgamation of disparate Archean cratons to growth by accretion of juvenile arcs. The metamorphic and structural history of the corridor of Archean gneiss domes south of Lake Superior is typically attributed to the Penokean orogeny. However, recent 40Ar/39Ar geochronology calls this relationship into question as ~1760 Ma cooling ages across the region indicate the deformation and metamorphism coincident with dome uplift is markedly younger (Schneider et al., 2004; Tinkham and Marshak, 2004; Holm et al., 2005; Schulz and Cannon, 2007). To correctly distinguish the effects of the Penokean orogeny and more accurately reconstruct the Paleoproterozoic tectonic history of the Upper Midwest, we present new U-Pb, Sm-Nd, and LuHf geochronology and microstructural analyses for a suite of metamorphosed and deformed rocks within and adjacent to several gneiss domes (Fig. 1). Titanite U-Pb ages and trace element compositions reflect Archean metamorphism at 2550 ± 46 Ma (2SE), and variable degrees of recrystallization in the Paleoproterozoic (Fig. 2). Apatite and monazite U-Pb ages, along with garnet Lu-Hf ages of metamorphosed supracrustal rocks directly outside of domes, record the onset of peak conditions by 1837 ± 7 Ma that continued beyond the end of the Penokean orogeny until 1782 ± 15 Ma. The garnet Sm-Nd ages of several samples are ~70 Ma younger than the Lu-Hf ages, reflecting a period of cooling and exhumation between 1752 ± 10 and 1738 ± 9 Ma. This exhumation interval overlaps with the U-Pb ages of synkinematic titanite at 1713 ± 32 Ma and the 1750 ± 6 Ma Lu-Hf age of re-equilibrated pre-kinematic garnets. U-Pb ages of apatite in one sample reflect much later reheating of the system at 1592 ± 26 Ma. These data show that deformation and metamorphism related to the uplift of gneiss domes in the Lake Superior region can only be partially linked to tectonic events between 1880–1830 Ma. Peak metamorphic conditions lasting until 1782 Ma indicate the persistence of thick orogenic crust well after the end of the Penokean orogeny—perhaps supported by continued convergence or an unrecognized collisional event along the margin. Exhumation beginning at 1752 Ma coincided with subduction farther south during the Yavapai orogeny (1760-1720 Ma), whereas uplift may be related to crustal extension above the downgoing slab, aided in part by gravitational forces acting on overthickened crust. Extension during this time would also have played a role in the generation and spatial accommodation of Yavapai-age granite intrusions across the region (e.g., East-Central Minnesota batholith). The youngest apatite U-Pb age at 1592 Ma likely represents distal thermal effects of the Mazatzal orogeny (1650–1600 Ma) farther south. These data reveal the gneiss dome structures in the Upper Midwest are the result of a protracted history including several Paleoproterozoic metamorphic, deformational, and uplift events spanning more than 70 m.y.. 71 �Figure 1: Left, geologic map showing gneiss domes in northern Michigan with geochronology sample sites. Cities shown – Marquette (M), Watersmeet (W), Republic (R), and Hardwood (H). Modified from Tinkham and Marshak (2004). Figure 2: Right, ages at 2SE precision. Vertical bars represent the timing of the Sacred Heart (S), Penokean (P), Yavapai (Y), and Mazatzal (M) orogenies. Hatched fields represent durations of metamorphic prograde and cooling intervals. Sm-Nd ages of UPMI 10 23 and UPMI 8 23 have high uncertainties from mineral inclusions that could not be removed prior to analyses and therefore are not used to define the duration of the cooling interval. Diagrams below show the Archean-Mesoproterozoic tectonic evolution of southern Laurentia. Yellow star shows study area location. REFERENCES Holm. D., Van Schmus, W., MacNeill, L., Boerboom, T., Schweitzer, D., Schneider, D., 2005, U-Pb zircon geochronology of Paleoproterozoic plutons from the northern mid-continent, USA: Evidence for subduction flip and continued convergence after geon 18 Penokean orogenesis: Geol. Soc. Am. Bull. 117, 259-275. Schneider. S., Holm. D., O’Boyle. C., Hamilton. M., Jercinovic. M., 2004, Paleoproterozoic development of a gneiss dome corridor in the southern Lake Superior region, USA: GSA Special Paper 380, 339357. Schulz. K., Cannon. W., 2007, The Penokean orogeny in the Lake Superior Region: Precambrian Research, 157, 4-5. Tinkham. D., Marshak. S., 2004, Precambrian dome and keel structure in the Penokean orogenic belt of northern Michigan, USA: GSA Special Paper 380, 321-338. 72 �Identifying Abandoned Mine Surficial Features Using Mask R-CNN, Upper Peninsula Michigan. SHALIFOE, MaryElizabeth1, VOICE, Peter1 1 Department of Geological and Environmental Sciences and Michigan Geological Survey, Western Michigan University, 1903 W Michigan Ave, Kalamazoo MI, 49008-5241, USA From the 1840s to the 1980s, iron, and copper mining in Michigan's Upper Peninsula thrived, leaving behind numerous surficial features from the early underground mining practices. Even today the Eagle Mine located in Marquette County is still active Mining both copper and nickel. Today's demand for rare earth minerals has sparked interest in exploring locations near these primary ores including the tailing piles (Demas A., 2023). Mapping old mine features using optical satellite imagery is challenging in Michigan's Upper Peninsula due to dense vegetation and snow cover – instead we need to use techniques that allow us to see through this cover. This study aims to assess the performance of object detection Deep Learning Models (DLMs) in mapping potential mine features using high-resolution terrain data (LIDAR-derived 1meter Digital Elevation Models) produced through the 3D Elevation Program. Dickinson County was chosen as the study area due to its rich history of 52 known abandoned mines within the East Menominee Iron Range (Figure 1). This study targeted various features, such as prospect pits, open pits, lateral ditches, and waste piles, resulting in a total of 946 identified features used for training the DLMs. The object detection methods available within ArcGIS software were evaluated including Feature Classifier, Faster R-CNN, and Mask R-CNN. Our initial evaluation has shown that Mask R-CNN performed better than the other methods, due to the Mask R-CNN method that enables pixel-level segmentation in addition to object detection (Maxwell A. E., et al., 2020. Our ongoing work is focused on the refinement of the model parameters to better locate surface features related to historic mining. Once the model is completed, it will be tested on various locations in northern Michigan within the mining ranges of the Marquette Iron Range, Menominee Iron Range, Gogebic Iron Range, and the Copper Ranges within Ontonagon and through the Keweenaw. This will then be ground-truthed, by going out into the field to verify the locations of the features or using historical topographic maps to verify the existence of features that may be inaccessible. REFERENCES Department of Environment, Great Lakes, and Energy, (2024) EGLE Geowebface; mining and minerals, State of Michigan, https://www.egle.state.mi.us/geowebface/#btnToolNavInfo Demas A. (2023). Bipartisan Infrastructure Law Funds Geologic Mapping in Michigan, by Bipartisan Infrastructure Law Investments, USGS, https://www.usgs.gov/special-topics/bipartisaninfrastructure-law-investments/news/bipartisan-infrastructure-law-funds-6 Maxwell, A. E., Pourmohammadi, P., & Poyner, J. D. (2020). Mapping the Topographic Features of Mining-Related Valley Fills Using Mask R-CNN Deep Learning and Digital Elevation Data. Remote Sensing, 12(3), 547. https://doi.org/10.3390/rs12030547 73 �Figure 1: Study Area in Dickinson County, showing the distribution of underground mines (Department of Environment, Great Lakes, and Energy, 2024). Figure 2: Mine Features located near East Central Vulcan Mine in Dickinson County, DEM sourced USGS TNM, 2016. (Department of Environment, Great Lakes, and Energy, 2024). 74 �Basaltic rocks of the Animikie Group in Ontario: Geochemical characteristics and tectonic significance SMYK, Mark1,3, HOLLINGS, Pete1, METSARANTA, Riku2, CUNDARI, Robert3, KISSIN, Stephen1 and KURCINKA, Colleen3 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada 2 Ontario Geological Survey, Ministry of Mines, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5 Canada 3 Ontario Geological Survey, Ministry of Mines, 435 James St. South, Thunder Bay, ON P7E 6S7 Canada The Paleoproterozoic Animikie Group in Ontario records a history of continental sedimentation and minor volcanism on the southern margin of the Superior Craton between ca. 1.88 Ga and 1.82 Ga. Both the chemical sedimentary rock-dominated Gunflint Formation and overlying, siliciclastic sedimentary rock-dominated Rove Formation contain significant intervals of tuffs and basalt flows. Copper-bearing amygdaloidal basalts were noted in Crooks and Blake townships (Coleman 1900); basalt flows and tuffs were identified by Gill (1925) and Goodwin (1960) in the Mink Mountain area, and by Tanton (1931) in Oliver Township. Tanton (1936) mapped “Rove basalt” in Devon Township. In 2022, a 774 m diamond drill hole (DDH ST-2201), completed by Metal Energy Corp. in Hartington Township, provided a complete section from Rove Formation into Archean basement. New geochemical, petrographic and stratigraphic data gleaned from this drill core and recent field work have provided insights into the nature of the basaltic rocks. The lowermost volcanic unit occurs in the middle of the Gunflint Formation, exposed near Mink Mountain; its base is ~53 m above Archean basement. Approximately 21 m thick, it consists of several distinctive, typically massive, locally pillowed, vesicular/spherulitic basalt flows. An isolated outcrop of amygdaloidal basalt in Oliver Township, approximately 40 km northeast of Mink Mountain, shares similar petrographic characteristics, stratigraphic position and geochemistry. Limited geochemical data gleaned from amygdaloidal basalts in Crooks Township are similar to those of the aforementioned Gunflint lavas. Further work is required to elucidate the nature and stratigraphic position of these flows. Basaltic flows, exposed on top of Rove shales and wackes in Devon Township (Cundari, 2010) had recently been considered part of the Mesoproterozoic Midcontinent Rift, based mainly on a Keweenawan reversed paleomagnetic mean direction and equivocal stratigraphic constraints (Cundari et al., 2012). However, a mafic interval, approximately 510 m above Archean basement and ~4 m thick, occurs within DDH ST-22-01 and displays a variolitic, chilled basal contact and spherulitic, vesicle-like features, similar to those displayed by the lowermost Devon flows. Similar trace element geochemistry further supports the contention that the mafic rocks intersected in drilling may be correlative with the Devon basalts and with other, similar rocks exposed in an isolated outcrop in Hardwick Township, ~30 km northwest of the Devon basalts. The Gunflint basalts are characterized by moderate La/SmCN ratios (~1.9 to 3.9), negative Nb-Ta and Ti anomalies and relatively flat Gd/YbCN ratios (~1.3 to 1.7). The Devon basalts are characterized by moderate La/SmCN ratios (~2.8 to 3.5), negative Nb and Ti anomalies and moderate Gd/YbCN ratios (~3.0-3.6). In a Penokean tectonic context, the Gunflint basalts may represent limited back-arc volcanism (cf. Kissin and Fralick, 1994), contemporaneous with the older phase of volcanism in the Pembine domain of the Pembine-Wausau terrane (PWT; ca. 1875 Ma, Zi et al. 2022). The Devon basalts may represent relatively deeply sourced, crustally contaminated, OIB-like magmas generated after ca. 1840 Ma, at the same time as renewed volcanism in the PWT. 75 �REFERENCES Coleman, A.P. 1900. Copper and iron regions of Ontario; in Ninth Report of the Bureau of Mines, 1900; Ontario Bureau of Mines, Annual Report, pp.143-191. Cundari, R. 2010. Geology and Geochemistry of the Devon volcanics, south of Thunder Bay, Ontario; unpublished HBSc. thesis, Lakehead University, Thunder Bay, 68p. Cundari R., Piispa, E., Smirnov, A.V., Pesonen, L.J., Hollings P. and Smyk, M. 2012. Geochemistry and paleomagnetism of the Devon township basalt, Ontario, Canada; in Mertanen, S., Pesonen, L. J. and Sangchan, P. (eds.). Supercontinent Symposium 2012 – Programme and Abstracts; Geological Survey of Finland, Espoo, Finland, p.30-31. Gill, J. E. 1925. Gunflint iron-bearing formation; Geological Survey of Canada, Summary Report 1924, pt.C, pp.28-88; https://doi.org/10.4095/103167. Goodwin, A.M. 1960. Gunflint iron formation of the Whitefish Lake area; Ontario Department of Mines, Annual Report, vol.69, pt.7, pp.41-63. Kissin, S.A. and Fralick, P.W. 1994. Early Proterozoic volcanics of the Animikie Group, Ontario and Michigan, and their tectonic significance; 40th annual Institute on Lake Superior Geology, Houghton, MI, Proceedings, vol.40, pp.18-19. Tanton, T.L. 1936. Pigeon River area, Thunder Bay District; Geological Survey of Canada, Map 354A, sheet 1, scale 1:63 360; https://doi.org/10.4095/107549. Tanton, T. L. 1931. Fort William and Port Arthur, and Thunder Cape map areas, Thunder Bay District, Ontario; Geological Survey of Canada, Memoir, 167, 222. https://doi.org/10.4095/100799. 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; GSA Bulletin; March/April 2022; v. 134; no. 3/4; p. 776–790; https://doi.org/10.1130/B36114.1; 8 figures; 1 supplemental file. published online 1 July 2021. 76 �Sedimentologic and geochemical evidence of marine incursion to the Oronto Group basin, southern Lake Superior region, at ca. 1.08 Ga STEWART, Esther K.1, 2, TAPPA, Michael 1, BAUER, Ann1, BRENGMAN, Latisha3, and PRAVE, Anthony 4 1 Department of Geoscience, University of Wisconsin-Madison, Madison, Wisconsin 53705 2 Wisconsin Geological and Natural History Survey, University of Wisconsin-Madison Division of Extension, Madison, Wisconsin 53705 3 Department of Earth and Environmental Sciences, University of Minnesota-Duluth, Duluth, Minnesota 55812 4 School of Earth and Environmental Sciences, University of St. Andrews KY16 9TS, Scotland/UK The late Mesoproterozoic Oronto Group (Copper Harbor Conglomerate, Nonesuch, and Freda Formations), Wisconsin and Michigan, preserves a continuous record of depositional environment and related microbial habitat. Over three kilometers of siliciclastic sediments with minor authigenic, molar tooth calcite record physical and biogeochemical processes acting within the Oronto Group basin at the time of deposition and early diagenesis. Combined sedimentologic and geochemical evidence motivates reevaluation and refinement of evolving depositional conditions (Stewart, 2025). Sedimentary facies indicate a shallow marine, tidal influence on deposition, requiring marine incursion to the Laurentian interior at ca. 1.08 Ga (Stewart et al., 2024). The degree of marine connectivity is investigated using C, O, and Rb-Sr isotope compositions of calcite microspar in molar tooth structures and carbonate laminae of the Nonesuch Formation. Molar tooth structures and laminae were milled from thick sections, and one split of sample powder was analyzed for C and O isotopes while the other underwent a multistep chemical separation process to isolate Rb-Sr isotopes from calcite. Carbon isotope (δ13C) values (-3.9 to -2.0‰) of earliest diagenetic calcite reflect organic matter remineralization driven by in situ microbial carbon cycling (e.g. Gilleaudeau and Kah, 2013). Values of δ18O (-6.7 to -3.6‰) measured in the calcite microspar of molar tooth structures overlap the isotopic signature of marine carbonates from other late Mesoproterozoic evaporative marine basins (e.g. Kah, 2000). The 87Sr/86Sr of least-altered calcite (~0.7068 to 0.7069) reflects marine mixing with continental runoff. Combined, these data reflect deposition within a restricted-marine epeiric setting. In addition to isotopic evidence for marine connectivity, conditions of salinity, redox, and productivity are evaluated using whole rock geochemistry of fine-grained siliciclastics and rare earth element + yttrium (REY) distributions of calcite microspar. Whole rock geochemistry was compiled from published sources and new data was collected from two cores in Wisconsin. Calcite REY distributions were analyzed from aliquots of the same sample material processed for Rb-Sr isotopes. Shale geochemistry, including Mo and U enrichment and stratigraphic trends in proxies for detrital input (Zr, Al), redox (S, TOC) and productivity (Ba, P) reveal deposition within an oxidized basin with a deep, fluctuating chemocline and expansion of anoxic and euxinic conditions during maximum flooding and base level lowstand. REY distributions of calcite microspar preserve an early diagenetic estuarine signal characterized by muted, positive La and Y/Ho anomalies and heavy REE enrichment. Shale geochemistry and carbonate REY distributions bring into focus the prevalence of particle shuttling between the water column and shallow sediments that likely enhanced and focused nutrient P bioavailability, analogous to modern estuarine nutrient cycling. Collectively, these data provide a richer understanding of late Mesoproterozoic environmental conditions that influenced early eukaryote ecology. 77 �Figure 1: Images from core (A, C-D, F) and thin section (B, E) of the Nonesuch Formation highlighting sedimentary structures indicative of tidal influence and molar tooth calcite microspar targeted for geochemistry. A & C: photos and line drawings showing close association of fine-grained sandstone (light color) and shale (dark color). Note mud drapes on bi-directional ripple laminae (red arrows, A), flame structures (red arrow, C), and structureless mud layers indicative of fluid mud deposits. B: Photomicrograph (cross-polarized light) showing bedding deflecting around molar tooth structure (arrow) and brittle deformation of molar tooth structures (1 displaced from 2). D: molar tooth structure (MT) cross-cutting carbonate-rich layers (CR) in drill core. E: Photomicrograph (plane-polarized light) highlighting characteristic molar tooth microspar texture. F: Core photo showing ~2 cm diameter mud ball with subangular rhyolite clast at its core. Scale bars are 1 cm unless otherwise noted. REFERENCES Gilleaudeau, G. J., and Kah, L. C., 2013. Carbon isotope records in a Mesoproterozoic epicratonic sea: carbon cycling in a low-oxygen world. Precambrian Research, 228, 85-101. Kah, L. C., 2000. Depositional δ18O signatures in Proterozoic dolostones: constraints on seawater chemistry and early diagenesis. SEPM Special Publication 67, 346 – 360. Stewart, E. K., Bauer, A. M., and Prave, A. R., 2024. End-Mesoproterozoic (ca. 1.08 Ga) epeiric seaway of the Nonesuch Formation, Wisconsin and Michigan, USA. Geological Society of America Bulletin, 136, 7-8, 2940-2960. https://doi.org/10.1130/B37060.1 Stewart, E.K., 2025. Sedimentologic and geochemical markers of marine incursion to the interior Laurentian Oronto Group basin at ca. 1.08 Ga. Ph.D. dissertation, University of WisconsinMadison. 78 �Midcontinent Rift extension ceased and the rift inverted due to the Grenvillian orogeny 1 2 3 SWANSON-HYSELL, Nicholas , HODGIN, Eben B. , ALEMU, Tadesse , FUENTES, 4 2 5 4 Anthony , ZHANG, Yiming , SLOTZNICK, Sarah and FAIRCHILD, Luke 1 Department of Earth and Environmental Sciences, University of Minnesota, Minneapolis, MN, USA Department of Earth, Environmental and Planetary Sciences, Brown University, Providence, RI, USA 3 Department of Geology and Environmental Science, University of Wisconsin, Eau Claire, WI, USA 4 Department of Earth and Planetary Science, University of California, Berkeley, CA, USA 5 Department of Earth Sciences, Dartmouth College, Hanover, NH, USA 2 The cessation of rifting within the Midcontinent Rift was a key event in the evolution of the Lake Superior region. If rifting had continued and led to the formation of an ocean basin, the subsequent geologic and paleogeographic history would have been profoundly different. In a 1994 paper, Bill Cannon used emerging geochronology from the Midcontinent Rift and the Grenville orogen to conclude that the closing of the Midcontinent Rift was a far-field effect of compression associated with the Grenvillian orogeny (Cannon, 1994). An alternative proposal was put forward by Stein et al. (2014) who proposed that the Midcontinent Rift is an abandoned rift segment associated with successful rifting along Laurentia’s margin. In this contribution, we leverage improved chronostratigraphy within the volcanics and sedimentary rocks of the Midcontinent Rift (e.g. Fairchild et al., 2017; Hodgin et al., 2024) combined with rich new records of metamorphic chronology associated with the Grenvillian orogeny (reviewed in Swanson-Hysell et al., 2023) to revisit this question and gain fresh insight. The transition from active rift extension to post-rift thermal subsidence is recorded by the Brownstone Falls angular unconformity in northern Wisconsin. The thinning of the Copper Harbor Conglomerate from >2,200 m thick on the Keweenaw Peninsula of Michigan to pinching out against the unconformity implies topographic relief at the onset of post-rift sedimentation that is comparable to that in the modern-day East African rift. The end of active extension (ca. 1090 to 1085 Ma) is coincident with early prograde metamorphism associated with the Grenvillian orogeny, whose metamorphic imprint extends from the Blue Ridge inliers of the eastern US up through the Grenville Province of eastern Canada. This timing is consistent with the onset of continent-continent collision resulting in the cessation of extension in the rift. Figure 1: The start and end of Midcontinent Rift extension compared with U-Pb dates from Grenville Province metamorphic chronometers (blue diamonds: zircon; red pentagons: monazite). The rift developed during an interval of tectonic quiescence on the margin. Extension ceased with the onset of the Grenvillian orogeny and the rift contractionally inverted during the peak of the Ottawan stage. 79 �Following the end of Midcontinent Rift extension, deposition of the Oronto Group continued until ca. 1045 Ma (Hodgin et al., 2024; Fuentes et al, in review). This deposition resulted from post-rift thermal subsidence prior to contractional deformation associated with the Grenvillian orogeny propagating into the continental interior. Paleomagnetic records from the Oronto Group, including recently published data from the Nonesuch Formation (Slotznick et al., 2024) and new unpublished data from the upper Freda Formation, reveal that Laurentia’s plate motion dramatically slowed coincident with the onset of Grenvillian orogenesis. Preceding rapid motion was associated with ocean basin closure leading up to continent-continent collision that changed the force balance and slowed the plate. Oronto Group deposition ended when contractional deformation associated with the Grenvillian orogeny propagated into the Midcontinent. This deformation occurred in two phases with major exhumation occurring during the peak of the Ottawan phase of the Grenvillian orogeny and a second more minor phase of ca. 1000 to 980 Ma contraction associated with the Rigolet phase (Hodgin et al., 2024). This final interval of contraction is associated with the ca. 990 Ma deposition of the Jacobsville-Bayfield Group (Hodgin et al., 2022; Alemu et al., 2023). Following 130 Myr of tectonic excitement from ca. 1110 to 980 Ma, stability returned to Laurentia’s Midcontinent region. While the comings and goings of inland seas and the occasional impact crater have left their mark on the geological record, there has been only very minor tectonism over the past billion years. REFERENCES Alemu, T.B., Hodgin, E.B., and Swanson-Hysell, N.L., 2023. Grooving in the midcontinent: A tectonic origin for the mysterious striations of L’Anse Bay, Michigan, USA. Geosphere, 19(5), 1291–1299. Cannon, W.F., 1994. Closing of the Midcontinent Rift—A far-field effect of Grenvillian compression. Geology, 22(2), 155–158. 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, 9(1), 117–133. Hodgin, E.B., Swanson-Hysell, N.L., DeGraff, J.M., Kylander-Clark, A.R.C., Schmitz, M.D., Turner, A.C., Zhang, Y., and Stolper, D.A., 2022. Final inversion of the Midcontinent Rift during the Rigolet Phase of the Grenvillian orogeny. Geology, 50(5), 547–551. Hodgin, E.B., Swanson-Hysell, N.L., Kylander-Clark, A.R.C., Turner, A.C., Stolper, D.A., Ibarra, D.E., Schmitz, M.D., Zhang, Y., Fairchild, L.M., and Fuentes, A.J., 2024. One billion years of stability in the North American Midcontinent following two-stage Grenvillian structural inversion. Tectonics, 43(9). Slotznick, S.P., Swanson-Hysell, N.L., Zhang, Y., Clayton, K.E., Wellman, C.H., Tosca, N.J., and Strother, P.K., 2024. Reconstructing the paleoenvironment of an oxygenated Mesoproterozoic shoreline and its record of life. Geological Society of America Bulletin, 136(3–4), 1628–1642. Stein, C.A., Stein, S., Merino, M., Keller, R.G., Flesch, L.M., and Jurdy, D.M., 2014. Was the Midcontinent Rift part of a successful seafloor-spreading episode? Geophysical Research Letters, 41(5), 1465–1470. Swanson-Hysell, N.L., Rivers, T., and van der Lee, S., 2023. The late Mesoproterozoic to early Neoproterozoic Grenvillian orogeny and the assembly of Rodinia: Turning point in the tectonic evolution of Laurentia. In: Whitmeyer, S.J., Kellett, D.A., Tikoff, B., and Williams, M.L. (Eds.), Laurentia: Turning Points in the Evolution of a Continent. Geological Society of America Memoir 220, 337–356. 80 �Ni-Cu-PGE Mineralization at the Mineral Lake Intrusive Complex, northern Wisconsin THOMPSON, Bekah R. 1, LODGE, Robert W.D.1 1 Department of Geology & Environmental Science, University of Wisconsin-Eau Claire, 105 Garfield Avenue, Eau Claire, WI 54701, USA The Mineral Lake Intrusive Complex (MLIC), near Mellen, Wisconsin, is a 1.1 Ga layered and differentiated mafic intrusive complex within the Mesoproterozoic Mid-Continent Rift in the Lake Superior region (Siefert et al., 1992). This intrusive complex hosts Ni-Cu-PGE mineralization discovered in the 1960’s via electromagnetic geophysical surveys and at least 16 drill holes were completed (Bakheit, 1981). With an increase in demand for domestic critical minerals to supply metals for energy, communication, and military infrastructure, underexplored prospects like the Mineral Lake Ni-Cu-PGE prospect are increasingly important. This project aims to describe the mineralogy of the sulfide inclusions and the host intrusion geochemistry to better understand the geological characteristics of PGE-mineralization within the MLIC. Two drill holes were re-logged (WIS-12 and WIS-11), totaling ~950 linear feet of core, and sixteen samples were collected from representative intrusive phases and mineralization types. Micron-scale PGE-bearing mineral phases are described using the SEM-EDS. Whole rock geochemistry of the MLIC was completed via X-ray Fluorescence (WD-XRF). Silicate and sulfide mineralogy was determined by transmitted and reflected light petrography. Mineralization is hosted in either medium-grained, equigranular olivine gabbro, olivine norite and troctolite phases of the intrusion and are found as mm-scale sulfide segregations composing 1-10% of the rock. Weak foliation and alteration along fractures are observed along brittle-ductile shears resulting in serpentinization of olivine. Contacts between intrusive phases are generally gradational over a few centimeters. Sulfide inclusions contain varying amounts of chalcopyrite, pyrrhotite, and pentlandite and are not obviously correlated with any specific intrusive phase. Graphite, both fracture-associated and matrix-associated, were observed in the Troctolite and Olivine norite phases. Sulfide inclusions are comprised of primarily pyrrhotite with variable amounts of chalcopyrite and pentlandite. Analysis on the SEM-EDS has shown PGE mineralization is commonly hosted as micron-scale inclusions within pyrrhotite and pentlandite. These PGEbearing mineral phases include rhenium-bearing molybdenite (Mo,Re)S2, padmaite (PdBiSe) (found within silicates), argentopentlandite Ag(Fe,Ni)8S8 , sperrylite (PtAs2), rhenite (ReS2), naldrettite (Pd2Sb). PGE minerals are typically ~5 microns. Notably large, 30-micron sperrylite (PtAs2) grains and 25-micron rhenite (ReS2) grains were observed (Figures 1C and 1D). PGE’s are most abundant hosted in sulfide minerals whereas the notable critical elements (Bi, Mo, Sb, Te, Ob, Se) tend to be hosted in the silicates. These results are comparable to other conduit-type and contact-type MCR intrusions, although the age of the MLIC is coeval with contact-type mineralization. Dunka road of the Duluth complex is a contact type Ni-Cu-PGE sulfide deposit. Phases include norite-hosted disseminated sulfides, troctolite-hosted disseminated sulfides, PGE-rich disseminated sulfides, and chalcopyrite rich disseminated sulfides (Theriault and Barnes, 1998). Since the MLIC is a large, differentiated intrusion that is coeval with other contact-type mineralization in the MCR, future exploration efforts and research should focus on the lower parts of the intrusion where dense sulfides may accumulate. 81 �Figure 1. (A) Regional map of Mineral Lake area. Map modified from Cannon and Ottke (1999). Inset map from Mudrey & Brown (1982). (B) Rhenium-bearing molybdenite (Mo,Re)S2 (white) under SEMEDS, (C) Rhenite (ReS2) (white) under SEM-EDS, (D) Sperrylite (PtAs2) under SEM-EDS (white). REFERENCES Bakheit, A.K., 1981. Petrography of Cu-Ni mineralization in the Mineral Lake area, Ashland County, Wisconsin. Unpublished M.S. thesis, University of Wisconsin-Madison. Cannon, W.F. and Ottke, D., 1999. Preliminary digital geologic map of the Penokean (Early Proterozoic) continental margin in northern Michigan and Wisconsin (No. 99-547). The Geological Survey of America. Middlemost EAK (1994) Naming materials in the magma/igneous rocks system. Earth Sci Rev 37:215– 224. doi:10.1016/0012-8252(94)90029-9 Siefert, K.E., Peterman, Z.E., Thieben, S.E. 1992. Possible crustal contamination of the Midcontinent Rift igneous rocks: examples from the Mineral Lake intrusions, Wisconsin. Canadian Journal of Earth Science, 29. 1140-1153. Thériault, R.D., Barnes, S.-J., 1998. Compositional variations in Cu-Ni-PGE sulfides of the Dunka Road deposit, Duluth complex, Minnesota: the importance of combined assimilation and magmatic processes. Can. Mineral. 36, 869–886 82 �A Porphyry in a Rift? Constraining the Petrogenesis of the Jogran Porphyry, Mamainse Point, Ontario, Canada: Insights from Zircon and Melt Inclusion Geochemistry. TOLLEY, James1, HANLEY, Jacob2, CROWLEY, James3, TSAY Sasha4, ZAJACZ Zoltan4, and HOLLINGS, Pete1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, Canada. Department of Geology, Saint Mary’s University, 923 Robie Street, Halifax, Nova Scotia, B3L 2Y5, Canada. 3 Isotope Geology Lab, Department of Geosciences, Boise State University, 1910 University Drive, Boise, Idaho, 83725-1535, USA. 4 Department of Earth Sciences, University of Geneva, Rue des Maraichers 13, Geneva, 1205, Switzerland. 2 The Jogran quartz-monzonite porphyry, located near Mamainse Point, Ontario, Canada, on the northeastern shoulder of the ~1.1 Ga Midcontinent Rift System (MRS), hosts unique porphyrystyle Cu-(Mo) mineralization in an intra-plate, rift-related large igneous province setting (Perelló et al., 2020). Combining high precision zircon geochronology with zircon and melt inclusion (MI) geochemistry refines the timing of emplacement and offers constraints on the crystallization temperature, oxygen fugacity (fO2), and melt composition (including ore metal tenor) during the magmatic evolution of the deposit. A new high precision 206Pb/238U zircon age of 1090.90 ± 1.27 Ma (CA-TIMS) constrains the formation of the Jogran porphyry to the waning of the main Rift Stage (1102-1090 Ma) and synchronous with the transition to the Late-Rift stage (1090-1083 Ma), as defined by Woodruff et al., (2020). Zircon geothermometry (Crisp et al., 2023) and oxybarometry (Loucks et al., 2020) suggest crystallisation conditions of 900-670 °C and a fO2 range of ∆FMQ = -1.3 to +0.6. As temperatures decrease, ΔFMQ values increase along a trend subparallel to the SO₂-H₂S buffer. The presence of sulfide inclusions in zircon, confirms sulfide saturation during crystallization. The analysed zircon crystals are zoned. They display an increase in [Yb/Gd]n ratios (1220) and concomitant depletion in Th/U (1.0-0.4) in the rims relative to the cores ([Yb/Gd]n = <12; [Th/U] = >1). This zonation infers that the parental magma underwent a single stage of fractionation and crystallisation upon emplacement. Melt inclusions (MIs) range in composition from 65-70 wt.% SiO₂ with 5.5-8.3 wt.% K₂O and K₂O/Na₂O ratios of ~1.5-3.5, suggesting the parental melt was alkalic to shoshonitic. Low Cs concentrations, coupled with high Rb, Ba, and Nb, in MIs indicate minimal crystal fractionation of a near-primitive, mantle-derived composition. In contrast, whole-rock data show lower alkali contents (4.0 wt.% K2O) and have a subalkalic affinity, suggesting crustal contamination or alteration obscured the primitive magmatic signature. A new, precise U-Pb zircon age constrains the felsic magmatism and porphyry-style mineralization at Jogran to the period of maximum lithospheric weakening/crustal thinning during the shift from extensional tectonics to thermal subsidence in the late stages of the MRS. This study suggests that early partitioning of metals and sulfur into magmatic fluids played a key role in ore formation. However, the conditions required remain ambiguous, as the tectonic environment at Jogran differs markedly from the subduction-related settings upon which most porphyry models are based. Porphyry deposits are increasingly recognised across a broader range of tectonic settings (e.g., southeast China [Richards, 2021]; and central Europe [Drew, 2006]). Jogran highlights the potential for porphyry-style mineralisation in non-subduction tectonic contexts and underscores the need to better understand metallogenic pathways beyond the traditional subduction models. 83 �Quartz-Feldspar Porphyry (K-Ar) 1 Tribag Breccia (K-Ar) 2 Mamainse Point Rhyolites (Rb-Sr) 3 Mamainse Point Volcanics (U-Pb) 4 Mamainse Point Tuff (U-Pb) 5 Jogran Porphyry Error bars represent the reported uncertainties in respective studies. Satellite Mineralisation (Re-Os) 6 Porphyry Stock Mineralisation (Re-Os) 6 References 1 Norman and Sawkins (1985) 2 Roscoe (1965) 3 Van Schmus (1971) 4 Davies et al. (1995) 5 Swanson-Hysell et al. (2014) 6 Perelló et al. (2020) Figure 1: New U-Pb zircon age (1090.90 ± 1.27 Ma) for the Jogran porphyry (diamond), published age data (circles) and MRS stages defined by Woodruff et al., (2020) – Early (green), 1109–1104 Ma; Latent (orange), 1104–1098 Ma; Main (blue), 1098–1090 Ma; and Late (purple), 1090–1083 Ma. REFERENCES Drew, L.J. (2005). A tectonic model for the spatial occurrence of porphyry copper and polymetallic vein deposits - Applications to central Europe: U.S. Geological Survey Scientific Investigations Report 2005-5272. Crisp, L. J., Berry, A. J., Burnham, A. D., Miller, L. A. & Newville, M. (2023). The Ti-in-zircon thermometer revised: The effect of pressure on the Ti site in zircon. Geochimica et Cosmochimica Acta 360, 241–258. Loucks, R. R., Fiorentini, M. L. & Henríquez, G. J. (2020). New magmatic oxybarometer using trace elements in zircon. Journal of Petrology, 61, egaa034. Perelló, J., Sillitoe, R. H. & Creaser, R. A. (2020). Mesoproterozoic porphyry copper mineralization at Mamainse Point, Ontario, Canada in the context of Midcontinent rift metallogeny. Ore Geology Reviews, 127, 103831. Richards, J.P, (2021). Porphyry copper deposit formation in arcs: What are the odds? Geosphere, 18, 130– 155. Woodruff, L. G., Schulz, K. J., Nicholson, S. W., and Dicken, C. L. (2020). Mineral deposits of the Mesoproterozoic Midcontinent Rift system in the Lake Superior region - A space and time classification. Ore Geology Reviews, 126, 103716. 84 �Evaluating Ni in Olivine as a Prospectivity Indicator for Magmatic Ni-Cu-(PGE) Deposits: A Preliminary Study from the Midcontinent Rift System. TOLLEY, James1, HOLLINGS, Pete1, MEXIA DURAN, Kevin1 and HARDING, Myles1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1, Canada. Nickel content of olivine [(Mg,Fe)2SiO4] can serve as an important petrogenetic marker in mafic igneous systems. Nickel’s concentration in olivine is controlled by several factors, including: (1) the Ni content of the parental magma; (2) the partition coefficient of Ni between olivine and the silicate melt; and (3) variable parameters such as temperature, pressure and fO2 of the melt (Li et al., 2007). More recently, Ni content in olivine has been studied as a potential fertility indicator for magmatic Ni-Cu sulfide deposits as well as providing information about the original composition of the magma. Olivine crystallizing from sulfide-saturated magmas will exhibit lower Ni contents relative to olivine crystallized from sulfide-undersaturated melts. This premise was assessed by Barnes et al. (2023) across Ni-Cu-(PGE) deposits globally, but there was a notable paucity of olivine data from Ni-Cu deposits within the Midcontinent Rift System (MRS). This study presents 700 new electron probe microanalyses (EPMA) of Ni and other major elements in olivine from five magmatic Ni-Cu-(PGE) deposits in the MRS: Sunday Lake, Steepledge, Escape, Current and Hele. These data have been integrated with published datasets from the mineralised Seagull, Eagle and East Eagle intrusions to produce the first regional-scale dataset of olivine chemistry from the MRS. Curation of this data aims to assess: (1) the deposit scale variability of olivine chemistry across the MRS; (2) the utility of Ni in olivine as a regional prospectivity indicator for Ni-Cu deposits within the MRS; and (3) the implications for primary melt evolution across the MRS. Preliminary results show that olivine forsterite (Fo) contents (i.e., 100*Mg/[Mg+Fetotal], mol %) range from Fo72.5-85 across most intrusions, except for the Hele intrusion, which has a much wider range (Fo44.0-82.5; Fig. 1). Across the entire dataset, Ni concentrations in olivine vary significantly (600-2500 ppm) and generally increase with higher Fo values. The range of Ni in olivine values can be vary up to 1000 ppm from a single deposit, over a narrow Fo range (e.g., Current Intrusion – Fo79.7-81.7). Furthermore, concentric zoning between Mg-rich cores relative to the Mg-depleted rims is frequently observed – most notably at Eagle East, where an average core analysis displays Fo80 vs. average rim value of Fo77. This preliminary compilation of olivine compositions across the MRS both reveals the variability of olivine compositions within a single intrusive complex and highlights fractionation trends regionally. The integration of the MRS data with the global compilation of Barnes et al. (2023) highlights the similarities between the signatures of unmineralized and mineralized intrusions and that there is no universal evidence for consistent Ni depletion in olivine from mineralised deposits. Placing the MRS olivine data within the context of other Ni-Cu-(PGE) systems may elucidate previously unrecognized potential within the MRS, and similarly these data can contribute to the global understanding of magmatic processes that culminate in economically viable deposits. 85 �Figure 1: Ni concentrations (ppm) in olivine as a function of forsterite content (Fo#) from a suite of maficultramafic Ni-Cu intrusions located in the Midcontinent Rift System. Grey field denotes the global array of ‘barren’ intrusions as defined by Barnes et al. (2023). Published datasets comprise: (1) Eagle and East Eagle Intrusion – Ding et al. (2010); (2) Seagull Intrusion – Heggie (2005); (3) Coldwell, Two Duck Gabbro – Good (1992); (4) Coldwell, Eastern Gabbro – Shaw (1997). REFERENCES Barnes, S. J., Yao, Z. S., Mao, Y. J., Jesus, A. P., Yang, S., Taranovic, V., & Maier, W. D. (2023). Nickel in olivine as an exploration indicator for magmatic Ni-Cu sulfide deposits: A data review and reevaluation. American Mineralogist, 108, 1-17. Ding, X., Li, C., Ripley, E. M., Rossell, D., & Kamo, S. (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). Good, D.J. (1992). Genesis of copper-precious metal sulphide deposits in the Port Coldwell Alkalic Complex, Ontario; unpublished Ph.D. thesis, McMaster University, Hamilton, Ontario, 203p. Heggie, G.J. (2005). Whole rock geochemistry, mineral chemistry, petrology and Pt, Pd mineralization of the Seagull Intrusion, northwestern Ontario. Unpublished M.Sc. thesis, Lakehead University, Thunder Bay, Ontario, 156. Li, C., Naldrett, A. J. & Ripley, E. M. (2007). Controls on the Fo and Ni Contents of Olivine in Sulfidebearing Mafic/Ultramafic Intrusions: Principles, Modeling, and Examples from Voisey’s Bay. Earth Science Frontiers 14, 177–183. Shaw, C. S. (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, 40(2-4), 243-259. 86 �Origin of magnetic black sand found on the south Shore of Lake Superior Verhoeven, J.D1., and Zowada, Tim2 1 Iowa State University, Emeritus Prof., Iowa State University, Levering MI 49755, jver@iastate.edu, 2 Custom Knifemaker, Boyne Falls, MI, timzowada@gmail.com Many of the beaches on the shores of Lake Superior contain black sand which is magnetic. This sand can be smelted into iron using the ancient bloomery process which produces small chunks of iron called blooms. They consist of iron containing a low level of carbon. The chunk of iron is filled with cavities containing remnant slag produced in the smelting process. Recent experiments [1] have shown that often but not always the resultant iron of the blooms contain significantly levels of Ti and that one of the microconstituents in the slag is the mineral ulvöspinel. The authors of [1] had assumed that the magnetic black sand came from erosion of banded hematite-magnetite iron formations (BIF) which are the source of the iron mined in the Lake Superior region. Finding Ti in some of the blooms shows that there is likely an alternate source of the iron in the black sands, namely the Fe–Ti oxide-bearing ultramafic intrusions (OUIs) deposited in the lake bottom from the 1.1Ga Midcontinent Rift (MCR) that runs through the lake region. This talk presents a comparison of the composition of the ulvöspinel constituent found in bloom slags of black sand smelts with the composition of the ulvöspinel constituents found in a recent study [2] of drillings from the Coldwell Complex region located at the north central region of Lake Superior which contain Fe-Ti magnetite-ilmenite intergrow deposits from the MCR. The results present strong evidence that the some of the magnetic black sand on Lake Superior’s shores comes from source rocks of MCR deposits in the Coldwell Complex and some from BIF deposits in the lake bottom. Additional evidence that the Fe-Ti source rock is the Coldwell Complex is that the location of the black sand used in the study is in the same region of the south shore of Lake Superior near White Fish Point where yooperlite rocks have been found. Literature data [3] shows that the source rock of the yooperlite is the Coldwell Complex. REFERENCES 1 Zowada T., Straszheim W., Chumbley S. and Verhoeven. J.D., 2025. A study of the carbon distribution and alloy composition of iron blooms made from two different batches of black sand collected from Lake Superior, accepted for publication in JMMA. 2 Brzozowski M.J., Samson I.M., Gagnon J.E., Linnen R.L. and Good D.J., 2021. Effects of fluid-induced oxidation on the composition of Fe–Ti oxides in the Eastern Gabbro, Coldwell Complex, Canada: implications for the application of Fe–Ti oxides to petrogenesis and mineral exploration, Mineralium Deposita 56, 601–618. 3 Laughlin, R. and Carlson A., 1987. A new find of fluorescent sodalite, Mineral News 34, no 5. 87 �88 �Zircon Petrochronology of the Eau Claire Volcanic Complex in the Marshfield Terrane of the Penokean Orogen, Northcentral Wisconsin VICKERS, Lyndsie A.1, LODGE, Robert W.D.1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA The Eau Claire Volcanic Complex (ECVC) serves as a type locality for Penokean-age magmatism and volcanism associated with the Marshfield terrane in the Penokean Orogeny (Figure 1A). This volcanic event is central to tectonic models that describe the collision of the Pembine-Wausau oceanic arc terrane and Archean crustal fragments of the Marshfield terrane with the southern margin of the Superior Craton (Shultz and Cannon 2007). A defining feature of these models is the proposed "double" subduction zone system, which is thought to have overprinted the Archean Marshfield terrane with younger Penokean volcanism and magmatism during ocean closure. Newer tectonic models suggesting accordion-like tectonics (Zi et al, 2022) still rely on historic interpretations of the ECVC where only physical outcrop descriptions in the literature (Myers et al, 1980). Despite its significance, the ECVC has remained understudied due to extensive Paleozoic and Quaternary cover which obscures outcrops and little mineral exploration and drilling. To address these challenges, this study focused on remote outcrops of the ECVC along the Eau Claire River in Wisconsin, aiming to better constrain tectonic models and clarify terrane boundaries in the southern Penokean Orogen. Field mapping yielded samples that were processed to isolate zircon grains for U/Pb radiometric dating and petrochronological analyses. These zircons were analyzed using a Laser Ablation Inductively Coupled Plasma Mass Spectrometer (LA-ICP-MS) at Laurentian University, providing the only modern geochronological and petrochronological data (U/Pb, Lu/Hf, zircon trace elements) from this region in the orogen. The results challenge long-standing interpretations of the area’s stratigraphy. Rocks previously classified as Paleoproterozoic volcanic units have Archean U/Pb ages and are now redefined as part of an Archean greenstone belt, significantly altering the geological narrative of the region. This study confirmed the presence of Paleoproterozoic intrusions (Figure 1C), but Lu-Hf isotopic analyses revealed that magmas did not have isotopic inheritance from the Archean basement (Figure 1D). This suggests the Paleoproterozoic magmas are in structural contact with Archean rocks. Additionally, Paleoproterozoic metasedimentary samples exhibited a diverse array of sedimentary sources (Figure 1-B), including Penokean, Marshfield, and a 2.2 Ga provenance, hinting at potential links to the Chocolay and Huronian groups which are continental rift assemblages formed during the breakup of an Archean supercontinent (Shultz & Cannon, 2007). As the first comprehensive petrochronological dataset from the Penokean Orogen, this study not only redefines the age and origin of key outcrops but also shows the complexity of the region’s tectonic and magmatic evolution. The discovery of previously unrecognized Archean basement rocks necessitates a reassessment of regional stratigraphy, particularly for classic outcrops historically attributed to Paleoproterozoic activity. Furthermore, the potential connection between the Marshfield terrane’s sedimentary sources and those of the Superior Craton’s rift assemblages raises questions about the terrane’s origins, suggesting it may represent a southernmost fragment of the Superior Craton. 89 �Figure 1: (A) Geologic map of the North Fork of the Eau Claire River adapted from Brown (1988). (B) Histogram displaying the distribution of zircon ages from a metasedimentary sample (C) Weighted mean diagram for intrusive sample showing a uniform range of zircon 207Pb/206Pb ages. Grey bars represent outliers and were excluded from age calculation. (D) ƐHf(i) versus 207Pb/206Pb age comparing ECVC intrusion to other Penokean intrusions in the Marshfield Terrane (Weber et al., 2023). 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. Schulz K.J., Cannon W.F., 2007. The Penokean orogeny in the Lake Superior region. Precambrian Research 157:4-25. Weber, E.M., Lodge, R.W.D., Marsh, J.H., 2023. U/Pb geochronology and zircon petrochronology of Paleoproterozoic magmas from the Marshfield terrane, Penokean Orogen, Wisconsin. Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 1-Program and Abstracts, p. 97-98. 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: GSA Bulletin, v. 134, p. 776–790. 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 90 �Geospatial Learning Resources to Explore Relationships with Keweenaw Geology VYE, Erika1, and LIZZADRO-MCPHERSON, Daniel2 1 Great Lakes Research Center, Michigan Technological University, 1400 Townsend Drive, Houghton, MI, 49931, United States 2 Geospatial Research Facility, Michigan Technological University, 1400 Townsend Drive, Houghton, MI, 49931, United States The globally significant geologic processes and features of the Keweenaw have fostered relationships with land and water for millennia. We have created three geospatial, digital resources that express the deep relationships between the underpinning geology and the scientific, educational, cultural, economic, and aesthetic significance of publicly accessible geosites in the Keweenaw region. These geospatial resources serve as living databases that will evolve over time in order to support formal and informal learners in understanding the fundamental role geology plays in our varied relationships with land and water. All resources are hosted and shared publicly on the Geospatial Research Facilities’ Enterprise Geospatial Research Portal at Michigan Technological University. 1) The Keweenaw Coastal Geoheritage StoryMap was created as a teaching and learning resource for local K-12 educators to explore the rock types of the Keweenaw at geosites along the shores of Lake Superior (Fig. 1). This resource: a) provides an overview of the main lithologies in the Keweenaw region, b) shares where federal, state, local government, and nonprofit organizations are working to preserve the rich geologic landscape and fragile wetlands of the Keweenaw, and c) provides a virtual learning experience to explore over 30 geologically significant sites along Lake Superior (Lizzadro-McPherson & Vye, 2023). 2) The Keweenaw Geoheritage Geoatlas is a knowledge directed exploration geospatial data hub that integrates physiographic landscape-wide feature coverages with a variety of downloadable GIS datasets. The repository of maps and data articulate the geoheritage of the region; the data hub supports educators, students, the scientific community, local tourist entities, land use planners, and the broader public in learning more about specific geosites in the Keweenaw region (Cowling, et al., 2024). 3) The Keweenaw Geoheritage geodatabase and web-viewer provide an innovative way of exploring the relationships between the bedrock geology and how this influences current and future education, conservation, and sustainable economic development initiatives in the Keweenaw region (Fig. 2). Each site expresses: a) a brief description of how the site contributes to the rich geoheritage of the Keweenaw, b) a 360-photo, and c) a description of the scientific (specific to the geologic phenomena), educational, cultural, economic, and aesthetic significance of the site (Lizzadro-McPherson & Vye, 2024). These resources are intended to support the co-stewardship of cultural heritage, restoration of legacy mining sites, conservation issues, and the development of sustainable economic opportunities based on the region’s globally significant geologic underpinnings. Further, they serve as the foundation for an evolving community participatory geoheritage mapping project in the Keweenaw. Through innovative, interactive geospatial resources we aspire to engage the broader public in sharing and exploring their relationships with the Keweenaw landscape (e.g. stories, valued geosites, photos, and curiosities). 91 �REFERENCES Cowling, R., Lizzadro-McPherson, D.J., Verissimo, L. & Vye, E.C. (2023). Keweenaw Geoheritage Geoatlas. DOI: 10.13140/RG.2.2.30945.28005 Lizzadro-McPherson, D.J., and Vye, E.C. (2024). Keweenaw Geoheritage Geodatabase. Michigan State Geological Survey; U.S. Geological Survey, National Cooperative Geologic Mapping Program (Award #G23AC00285 FY23). Lizzadro-McPherson, D. J. & Vye, E.C. (2023). Keweenaw Coastal Geoheritage StoryMap. DOI: 10.13140/RG.2.2.12680.74242 Fig. 1: Keweenaw Coastal Geoheritage StoryMap Fig. 2: Keweenaw Geoheritage Viewer 92 �Battle between the bands: competitive precipitations lead to bands in banded iron formations Xu, Huifang, and Zhou, Tianyu Department of Geoscience, University of Wisconsin–Madison, Madison, WI 53706, USA Banded iron formations (BIFs) are massive chemical deposits composed of alternating layers of chert and iron-rich minerals (such as hematite, magnetite and siderite), with three scales of bandings: microbands, mesobands (1 mm - 10 cm) and macrobands. Their abundance in the Archaean/early Proterozoic era and their absence thereafter suggest that chemical conditions and iron transport pathways on the early Earth surface were different from those after 1.7 billion years ago. Thermodynamic calculations show that Fe-silicate metal complex can be generated by hydrothermal leaching of low-Al oceanic crustal rocks such as komatiites, which suggest that the presence of low-Al ultramafic rocks (for example, komatiitic rocks) in the early oceanic crust were the reason for both the formation of BIFs and their abundance in the Archaean/early Proterozoic era (Wang et al., 2009). This is consistent with the findings that the ages of komatiites are correlated strongly, at the 99% confidence level, with the ages of BIFs (Isley and Abbott, 1999). We used the PHREEQC geochemical modeling package was used to test the chemical reactions that may have led to the banding pattern in the BIFs based on competitive precipitation of ferrihydrite (precursor of hematite and magnetite) and silica gel (precursor of chert) (Zhou et al., 2024). After aqueous ferrous silicate decomposition in O2-sufficient condition (pO2 ≥ 10-4), the faster Fe2+ oxidation and precipitation rate led to Fe-rich layer preceding Si-rich layer with ferrihydrite and amorphous silica as the precursor to the hematite and quartz, respectively. Episodic Fe(H3SiO4)2 input resulted in successive cycles of layering (Fig. 1). O2-deficient environments (pO2 < 10-4) results in jaspilite (no bands). The kinetic model also works well for the formation of siderite bands under O2-deficient environments when pCO2 is high. In summary, the precipitation process model proposed in this study offers an alternative abiotic explanation for the formation of distinct bands within the BIFs. 93 �Figure 1: Schematic depositional model of felsic volcanism associated BIF-like Iron Formations under different surface oxygen levels in a shallow hot spring lake (O2-deficient: pO2 < 10-4; O2-sufficient: pO2 ≥ 10-4). When the O2 level is high, the mix of aerobic lake water and Fe(H3SiO4)2-bearing spring fluid leads to the ferrihydrite-rich layer and amorphous silica-rich layer precipitating successively. But ferrihydrite and silica coprecipitate when O2 is deficient and there is no layering. The ferrihydrite-rich layer would convert to hematite-rich layer and amorphous silica-rich layer transforms into Si-rich layer. The surface water level is regulated by precipitation, evaporation and seepage from surrounding rock without visible inflow or outflow. DOI:10.1016/j.chemgeo.2024.122091) REFERENCES Isley, A. E. & Abbott, D. H., 1999. Plume-related mafic volcanism and the deposition of banded iron formation. J. Geophys. Res. 44, 15461-15477. Wang Y., Xu, H., Merino, E., and Konishi, H., 2009. Generation of banded iron formations by internal dynamics and leaching of oceanic crust. Nature Geoscience, 2, 781-784. Zhou, T., Hill, T., Roden, E. E., and Xu, H., 2024. The Felsic Volcanism Associated BIF-like Iron Formations: Their Origin and Implication for BIFs. Chemical Geology, 656, 122091. 94 �Broadly coeval but migrating deformation, plutonism and deposition in the northeastern Superior Province, Québec: evidence of hot accretionary orogeny and oroclinal folding in the late Archean? ŽÁK, Jiří1, TOMEK, Filip 1, 2, KACHLÍK, Václav 1, VACEK, František 1, 3 SVOJTKA, Martin 2, and ACKERMAN, Lukáš 2 1 Institute of Geology and Paleontology, Faculty of Science, Charles University, Albertov 6, Prague, 12843, Czech Republic 2 Institute of Geology of the Czech Academy of Sciences, Rozvojová 269, Prague, 16500, Czech Republic 3 Czech Geological Survey, Klárov 3, Prague, 11821, Czech Republic The James Bay Road in Québec provides a unique crustal-scale transect across several principal lithotectonic belts of the northeastern Superior Province. From north to south, these belts are Bienville (plutonic), La Grande (ʽgrayʼ gneisses, metaplutonic), Opinaca–Némiscau (metasedimentary), and Opatica (mostly volcano-plutonic). This assemblage has been controversially interpreted to record non-plate vertical tectonics driven by mantle plume activity or as resulting from the step-wise accretion of these belts to the northerly proto-cratonic core. We present here new structural and anisotropy of magnetic susceptibility (AMS) data from all the units along the James Bay Road transect. The data indicate a multistage fabric evolution: (1) an early fabric F1 is preserved only in isolated domains across the La Grande and Opinaca belts and is at a high angle to boundaries between the individual belts; (2) the F2 fabric seems to record a progressive reorientation (folding) towards an E–W direction; (3) the regionally dominant F3 fabric indicates regional NNE–SSW shortening across all units and is coeval with pluton emplacement and anatexis; (4) the last major ductile event is represented by localized dextral shear zones. The AMS indicates that magnetic foliations in general match well the mesoscopic foliations, whereas magnetic lineations vary from steeply plunging to subhorizontal, interpreted as recording a transition from vertical stretching during folding to horizontal stretching during shearing. The latter interpretation is further supported by a more detailed analysis of the ca. 2712–2697 Ma Radisson pluton, which is a syntectonic intrusion at the Bienville–La Grande boundary. Its magmatic to solid-state fabrics analyzed through the AMS also suggest a strain evolution from vertical magma stretching during regional shortening overprinted by later dextral shearing. In conjunction with the previously published U–Pb geochronology, the structural data suggest a short time span and north-to-south migration of plutonism, deposition, and contractional/transpressional deformation, altogether favoring a modern-style plate tectonics operating in the NE Superior Province in the late Archean. Furthermore, the relict F1 and F2 fabrics overprinted by F3 are interpreted as being compatible with changing block/microplate convergence vectors and crustal-scale folding of the outboard La Grande and Opinaca–Némiscau belts. In conclusion, the northeastern Superior Province may have been assembled as large, hot accretionary supra-subduction orogen, oroclinally folded, and finally dextrally sheared. Were this interpretation correct, a key question arises what was the geodynamic cause and mode of the oroclinal folding, whether with or without hard collision, taking the Alaskan terrane wreck or Mongolian orocline as prime examples, respectively. 95 � https://digitalcollections.lakeheadu.ca/files/original/772524cda5601210f7360a628832f9e3.pdf ed3bb63e47c2b1557a98f43e22c97dfb PDF Text Text Volume 71, Part 2 71st ANNUAL MEETING Mountain Iron, Minnesota, May 14-17, 2025 PART 2—Field Trip Guidebook �Meeting Co-Chairs Amy Radakovich, Allison Severson, Eric Nowariak, Stacy Saari, Aaron Hirsch Special thanks to field trip leaders: Zsuzsanna Allerton Terry Boerboom Kevin Boerst Latisha Brengman Annia Fayon George Hudak Mark Jirsa Phil Larson Dean Peterson Cullen Phillips Laurie Severson Mark Severson Alex Steiner i �71st Institute on Lake Superior Geology Volume 71 consists of: Field Trip 1 – Transect of the Quetico Subprovince ................................................................................... 1 Field Trip 2 – Drill Core from three Cu-Ni Deposits of the Duluth Complex .......................................... 15 Field Trip 3 – How Do You Make Iron and/or Manganese Ores in Proterozoic Iron Formation?............ 46 Field Trip 4 – New Geological Insights into the Genesis of Iron Ores at Lake Vermilion – Soudan Underground Mine State Park..................................................................................................................... 74 Field Trip 5 – Neoarchean Alkalic Intrusions in the Wawa and Quetico Subprovinces ......................... 108 Field Trip 6 – Unique Keweenawan Inclusion (Colvin Creek) in the Duluth Complex ......................... 136 Field Trip 7 – Classic Outcrops of Northeastern Minnesota ................................................................... 151 Field Trip 8 – Glacial Lake Norwood and the Koochiching Lobe…. ..................................................... 188 ii �Trip 1 – Quetico FIELD TRIP 1 Transect of the Quetico Subprovince Eric Nowariak1 and Mark Jirsa (retired)1 1 Minnesota Geological Survey, College of Science and Engineering, University of Minnesota, 2609 Territorial Road, St. Paul, MN 55114 Introduction This trip will examine exposures of the metasedimentary, migmatitic, and intrusive rocks of the Neoarchean Quetico subprovince from north of Mountain Iron to near Crane Lake and along part of the Echo Trail. It will attempt to “unpack” the primary components of deposition, magmatism, deformation, and metamorphism that likely spanned 40 million years (~2700-2660 Ma). The latter is based in part on newly acquired geochronologic analyses (Jirsa and others, 2020; Salerno, 2017). The trip will also address the challenge of creating meaningful geologic maps of this and similarly complex terranes, and the apparent lithologic and temporal link between Quetico metasediments and those associated with successor basins in the region. Figure 1-1. Complex migmatite exposed at field trip stop # 3. 1 �Trip 1 – Quetico Figure 1-2. Geologic Map of Central St. Louis County. This draft version of the St Louis County Precambrian bedrock map (superceded by Jirsa, 2020) portrays parts of the Neoarchean Wawa and Quetico subprovinces of Superior Province, and the approximate location of field trip stops. Wawa subprovince colors: greens=volcanic and volcaniclastic rocks; blues=metasedimentary rocks (primarily metagraywacke); reds=iron-formation; pinks=granitoid rocks; yellows=epiclastic and volcaniclastic sedimentary rocks. Quetico subprovince is labeled: BS=biotite schist (metagraywacke); SM=schist-rich migmatite; GM=granite-rich migmatite; TM=tonalite-rich migmatite. Pale pink Lac La Croix granite is more magnetic, darker pink is less so. Bold line marks the approximate boundary between subprovinces—a fault in some places, an inferred unconformity in others. 2 �Trip 1 – Quetico GEOLOGIC SETTING The Quetico is one of a number of east-trending, largely metasedimentary subprovinces in the Superior Province. It is bounded on the south by the Wawa volcanoplutonic subprovince, and on the north by the Wabigoon subprovince. It consists of schist derived from turbiditic sedimentary rocks and a complex suite of granitic intrusions and associated migmatite. In northeastern Minnesota, the subprovince displays a roughly symmetrical distribution of metasedimentary rocks on the north and south, that grade irregularly through zones of schist-rich migmatite, to a central axial zone composed largely of polyphase granitoid migmatite and younger granite. To some extent, metamorphic grade mimics this symmetry, with generally higher grade rocks in the central axis and lower grade near the bounding volcanoplutonic subprovinces. The accretionary prism model of Williams (1990) implies deposition of sediment shed from the craton to the north, and the subducting island arc to the south by submarine fans and abyssal turbidites. The Rainy Lake-Seine River Fault zone at the southern margin of the Wabigoon subprovince is thought to mimic the subduction front. The southern boundary against the Wawa subprovince is interpreted as an unconformity in some locales, and a fault in others. DEPOSITIONAL HISTORY Timing of deposition of the clastic sedimentary rocks of the Quetico subprovince in Minnesota has been constrained by populations of the youngest detrital zircons at 2690 +/- 12 Ma (Salerno, 2017). Similarly, geochronologic studies of the Quetico metasedimentary sequences in Canada has been constrained to 2698 Ma near Atikokan, ON (Davis and others, 1990) and <2690 according to Zaleski and others (1999) near the Manitowage Greenstone Belt of the Wawa subprovince. Zaleski and others (1999) also proved deposition of graywacke units within the Manitowage Greenstone Belt were contemporaneous, if not genetically related. A similar temporal and possible genetic link between the metasedimentary rocks of the Lake Vermilion formation and turbiditic metasediments of the Quetico in Northern Minnesota, wherein immature, volcaniclastic rocks of the Lake Vermilion formation gave way to silicic, clastic sedimentation observed in the Quetico subprovince as the basin evolved from alluvial fan deposits to a deep-water, active margin depocenter as the Quetico basin developed (Davis and others, 1990). In addition to Neoarchean zircons, small populations of older zircons including Mesoarchean zircons have also been recognized (Salerno, 2017; Davis and others, 1990). Probable sources of these older zircons have identified from multiple terranes in the southern Superior province and a proximal source for the sediments is inferred. The composition of the Quetico metasedimentary rocks suggests the source region was shedding sediment from a mixture of sialic plutonic terranes and lesser juvenile volcanic terranes. In addition to the clastic metasedimentary rocks that dominate the bulk of the Quetico subprovince, thin, discontinuous amphibolitic layers are found interbedded in many areas; likely representing volcanoclastic deposits and rare flows from active volcanism occurring near the margins of the subprovince. VERMILION GRANITIC COMPLEX The migmatitic and plutonic rocks in the axial zone are known collectively in Minnesota as the Vermilion Granitic Complex (Southwick and Sims, 1980). Southwick and Ojakangas (1979) subdivided migmatite for mapping purposes as schist-rich and granite-rich components, depending on the ratio of paleosome to neosome. Subsequent mapping by Jirsa (2011) and Jirsa and others (2014), applied the same terms, based instead on the extent to which the predominant fabric in the rock is controlled by neosome vs. paleosome and further distinguished units based on neosome composition. In this nomenclatural system, schist-rich migmatite is schist containing intrusions of granitoid neosome as both delaminating and crosscutting bodies; granite-rich migmatite is neosome with inclusions of paleosome. These can be equated generally with the terms metatexite (low degree of partial melting) and diatexite (nearly complete fusion), respectively, of Sawyer, (2008). Field relationships within the complex (Southwick, 1991) indicate that earliest granitoid phases are leucogranite, tonalite, granodiorite, and trondhjemite, which make up the leucosome of a broad area of migmatite across the western portion of the Vermilion Granitic Complex. For this field trip guide, these granitoids, broadly of TTG affinity, are referred to as “neosome 1”. The 3 �Trip 1 – Quetico migmatite is interlayered at all scales with paleosomes of biotite schist, paragneiss, orthogneiss, and amphibolite and often carries an internal fabric similar to the paleosome. Salerno (2017) obtained a discordant U-Pb zircon age of a granodiorite phase of neosome 1 at 2684 +/- 23 Ma. The migmatite is cut by poorly to non-foliated dikes, sills, and irregular masses of two mica leucogranite with accessory garnet, and a slightly younger biotite granite and pegmatite that contain minor magnetite. The presence of magnetite within the granitic rocks related pegmatites has proved to be an important mapping tool as these intrusive bodies create conspicuous aeromagnetic anomalies (Fig. 3b). The latter forms a large granitic mass, known as the Lac La Croix granite nearest the US/Canadian border, and apophosial intrusions that flare and pinch westward, producing aeromagnetic anomalies that highlight broad fold structures. These intrusive “fingers” generally decrease in thickness and continuity westward, suggesting that the western portion of the complex may represent the roof- or floor-zone of the batholith cored by massive granite. For simplicity within this field guide, the leucogranitic and granitic rocks of the Lac La Croix granite are referred to as “neosome 2” and represents the youngest granitic intrusive units of the Vermilion Granitic Complex. Geochronologic analyses of leucogranite and granite of the Lac La Croix granite has dated the crystallization of this unit with U-Pb zircon ages of 2658.71 +/- 0.47 Ma and 2668 +/10 Ma (Jirsa and others, 2014; Salerno, 2017). a b Figure 1-3. Maps of the Crane Lake and Brule Narrows 30’X60’ quadrangles (US and Canada) illustrating the connections between attributes of structure, lithology, topography, and magnetite content in this area of abundant near-surface bedrock. (a) 30m lidar land surface topographic grid; low areas darker. Topography defines major fold structures, and massive granitic vs. foliated orthogneissic and schistose bedrock. Prominent NNW-trending linear low areas are fault and fracture systems, many of which are occupied by rivers and lakes (named). (b) First vertical derivative map of aeromagnetic data. Magnetic highs (lighter colored) typically are more granitic; lows, more schist-rich. Like the topographic map, the magnetic data identify folds and faults. Linear, NW-trending highs are normally polarized diabase dikes of the Paleoproterozic Kenora-Kabetogama dike swarm. Linear lows are coincident with topographic lows, implying oxidation by meteoric, or more likely hydrothermal fluids along fractures. Some field evidence indicates that rock adjacent to fractures is chemically weathered, and hence more easily eroded. The subparallelism of dikes with fractures may indicate that oxidizing hydrothermal fluids were temporally related to dike emplacement. Neosome 1 and neosome 2 are readily distinguished in the field based on mineralogy and textural characteristics described above. Day and Weiblen (1986) used simple geochemical plots and CIPW normative mineralogy to visualize these differences (Fig. 1-4). Both plutonic suites are characterized as calc-alkaline, metaluminous to weakly peraluminous, magnesian granitoids. Neosome 1 tonalitetrondhjemite-granodiorite intrusions are inferred to have been sourced from partial melting of mafic crust. 4 �Trip 1 – Quetico Geochemical evidence indicates that the early neosome 2 migmatite was derived from partial melting of a metasedimentary protolith (Day and Weiblen, 1986). Southwick (1991) and Day and Weiblen (1986) suggested that the younger Lac La Croix-type granite of neosome 2 may represent further distillation of granitic liquid from partial melting of the combined older migmatite and metasedimentary rocks. Figure 1-4. From Day and Weiblen (1986). (A) CIPW normative mineralogy for Vermilion Granitic Complex. Q – quartz; Pl – albite+anorthite; Or – orthoclase. “Early Plutonic Suite” is equivalent to neosome 1 of this guidebook. (B) AFM diagram of same data (Irvine and Baragar, 1971). DEFORMATION AND METAMORPHIC HISTORY The Quetico subprovince has undergone a complex deformation history over a relatively contracted tectonic history between deposition of sediments ca. 2690 Ma and intrusion of the Lac La Croix Granite related pegmatite dikes ca. 2658 Ma. As summarized by Bauer and others (2011), three main phases of deformation have been recognized. D1 produced tight to isoclinal, recumbant folds plunging to the southwest and locally overturned to the southeast and produced a weak, bedding parallel axial planar foliation. Hinges of these recumbent folds are rare, and recognition of this event are often limited to overturned bedding and bedding-parallel foliation that is crenulated by subsequent deformational events. This event may locally have produced recumbent folds over a broad region (Bauer, 1985; Poulsen and others, 1980). It occurred shortly after deposition and involved burial to produce metamorphic conditions of moderate pressure and temperature (Valli and others, 2004). Fralick and others (2006) suggested D1 deformation was contemporaneous with development of the Quetico basin as an accretionary wedge. D2 deformation was synchronous with peak regional metamorphism to upper greenschist facies in the Wabigoon subprovince and amphibolite facies in the adjacent Quetico subprovince, and produced the dominant structural grain observed in the Minnesota segment of the Quetico subprovince. Folding associated with D2 deformation produced tight to isoclinal upright folds that plunge to the E-NE 10-30°. The intrusion of neosome 1 occurred slightly prior to or contemporaneous with D2 as veins of neosome 1 are commonly folded and occupy gently to moderately plunging D2 related fold hinges. Peak metamorphism presumed to be contemporaneous with D2 deformation within the Vermilion Granitic Complex has been dated by U-Pb monazite geochronology by Salerno (2017) and has constrained to ca. 2675 Ma. Continued contractional and transpressional deformation during D3 has been noted as ductile, eastnortheast trending transpressional shear zones and coaxial refolds of D2 related structures. Folding associated with D3 deformation is better developed near plutons of the Lac La Croix granite and related granitoids, suggesting early stages of neosome 2 intrusions were contemporaneous with deformation or used these structures as conduits (Bauer and others, 1992). D2-D3 is interpreted to be a result of the accretion of the Quetico subprovince to the Wabigoon subprovince to the north. Metamorphic indicator minerals within the Vermilion Granitic Complex including garnet, sillimanite, and locally cordierite have been well documented (Day, 1990; Tabor, 1988; Salerno, 2017). Limited thermobarometric studies have determined peak metamorphism reached amphibolite facies in the axial core of the Minnesota segment of the Quetico subprovince and upper greenschist facies along the northern margin of the subprovince, with garnet-biotite thermometry revealing metamorphic temperatures between 500-600°C and 430-475°C, 5 �Trip 1 – Quetico respectively (Bauer and others, 1992; Salerno, 2017). It is unknown whether the intrusion of the Lac La Croix granite and other neosome 2 intrusions produced a significant metamorphic overprint, however Tabor (1988) recognized kyanite in metamorphic assemblages of the Quetico subprovince along its northern boundary along the Rainy Lake-Seine Fault; which may represent an earlier, relatively higher pressure metamorphic regime prior to intrusion of the Lac La Croix granite. Subsequent deformation including brittle-ductile faulting with associated planar fabrics developed locally near fault zones and minor open folds reorienting existing structures has been ascribed to continued contraction post-dating metamorphism and major plutonism has been ascribed to D4 deformation. One of the most prominent and through-going features of the Quetico subprovince in Minnesota is the Vermilion Fault—a northwest-trending structure that truncates metamorphic zones and folds that are apparent on aeromagnetic maps is likely a product of late D3 and/or D4 deformation. Based largely on geophysical maps, dextral offset along this fault is on the order of 40 km. The fault can be traced from the extreme NW corner of the state for some 250 miles southeastward to near Ely, Minnesota. There it appears to veer to the northeast, manifest as a complexly splayed, post-metamorphic thrust-system known collectively as the Burntside Lake Fault. A summary of the deformational and metamorphic features observed in the Minnesota segment of the Quetico subprovince is shown in Table 1-1 below. Event General Features Associated Fabrics Metamorphic Features Timing Source(s) D1 Recumbant folds Bedding parallel foliation N/A ca. 2690 Ma Fralick and others (2006), Bauer (1985) D2 Upright to inclined tight to isoclinal folds, axes plunge to the northeast and southwest Axial planar cleavage, strong hinge-parallel lineation Upper greenschist to amphibolite facies ca. 2675 Ma Bauer and others (1992) and references therein, Salerno (2017) D3 Upright tight to isoclinal folds – coaxial to D2 folding, east-northeast trending shear zones Axial planar cleavage, strong hinge-parallel lineation, shear fabrics proximal to fault zones Amphibolite Facies 2675-2668 Ma Bauer and others (1992) and references therein, Jirsa (2014) D4 Brittle-ductile faulting, broad folding Planar fabrics proximal to shear zones Hydrothermal alteration along fault zones </=2668 Ma Bauer and others (1992), Recent unpublished mapping Table 1-1. Summary of deformational and metamorphic features observed in the Quetico subprovince within Minnesota. CONSIDERATIONS FOR GEOLOGIC MAPPING Complex geologic terranes recording multiple, interdependent geologic processes including sedimentation, multiple phases of deformation, and diverse polyphase intrusive histories like that of the Quetico subprovince represent a unique challenge in creating meaningful, consistent geologic maps and map units. Multiple attempts to properly portray the complicated geology of the Minnesota segment of the Quetico subprovince have used varied approaches, which have proved to require the incorporation field observation, petrography, aeromagnetic and gravity anomalies, LiDAR and aerial photography, and magnetic susceptibility measurements. Early iterations of geologic maps in the area focused on the proportional differences between the paleosomatic and neosomatic components of the migmatitic rocks within the subprovince to distinguish geologic units (Southwick and Ojakangas, 1979), while other authors have decided to incorporate the textural, compositional, and petrophysical characteristics of paleosomes and neosomes to further distinguish coherent map units (Jirsa, 2011; Jirsa and others, 2014). In addition to traditional field observations, thousands of magnetic susceptibility measurements recorded for units across the subprovince have been used to varied effect (Chandler and Lively, 2014). While petrophysical characteristics of the host 6 �Trip 1 – Quetico rocks are not sufficient to determine many units, the extent and morphology of some distinct units, namely late magnetite bearing granites and pegmatites, have been found to correlate with higher magnetic susceptibilities and resultant aeromagnetic anomalies (Fig. 1-3b). The structural complexities observed in this trip are preserved from the outcrop to map-scale. Many map-scale structures are discernable in aeromagnetic derivative maps and have been used in conjunction with the magnetic characteristics described above to outline geometric and temporal relationships between deformation and intrusive intervals where outcrop exposure is insufficient. Careful observations at individual outcrops has been found to be beneficial in comparison to regional lithologic mapping. Many geologic structures may be more readily identified by field checking and correlating the roughness and patterns of exposed outcrops using lidar and aerial photos. Recognition of post-intrusive faults visible in LiDAR derived maps and as linear magnetic lows in aeromagnetic maps have also helped reconcile locales where map patterns would be otherwise difficult to align with the known structural character of the area. FIELD TRIP STOP DESCRIPTIONS It should be noted that this trip derives from several years of field work to produce two geologic maps of the western-most exposed portions of the Quetico subprovince in Minnesota (Jirsa, 2011; Jirsa and others, 2014); and refinement by more recent field work to create maps of St. Louis and Koochiching Counties (Jirsa and others, 2020; Nowariak and others, in preparation). Mapping focused largely on structural and magnetic attributes that could yield a “meaningful” geologic map of this very complex terrane, and little analytical work was conducted; though ongoing work in Koochiching county has begun to tackle this. As a result, this field trip lacks details of metamorphism, petrology, and geochemisty. Instead, the focus was largely structural, in an attempt to reconcile prominent geophysical anomalies and topographic trends with field observations. The associated maps incorporate structural data, field relationships, and thousands of magnetic susceptibility measurements to ascertain the connections between lithology and magnetite content. Because glacial sediments are thin to absent in much of the area, mapping was also influenced by 10-meter (and subsequent 1-meter, for more recent mapping) LiDAR imagery (Fig. 1-3a). Mapping in the Quetico subprovince on which this field trip is based was supported by grants from the U.S. Geological Survey STATEMAP element of the National Geologic Mapping program, and by the Minnesota Environmental and Natural Resources Trust Fund. NOTE: All locations are denoted in UTM coordinates, NAD 83, Zone 15N STOP 1 – Feldspathic graywacke of the Lake Vermilion Formation Location: 526342E/5288565N, (47.74991°, 92.64856°), Highway 53 Northbound, 0.4 miles north of Heino Road (County 467) Description: This stop examines the feldspathic metasedimentary and meta-volcanogenic sediments of the Lake Vermilion formation, formally part of the Wawa Subprovince. Here, the Lake Vermilion formation is composed of feldspathic graywackes and tuffaceous slates and wackes. The stratigraphy Figure 1-5. Pavement exposure of laminated feldspathic metagraywacke of the Lake Vermilion generally tops to the north and is folded, with locally formation. well-developed axial planar cleavage and thin shear bands. This stop serves as a reference in comparing the composition and character of the metasedimentary 7 �Trip 1 – Quetico rocks of the uppermost units of the Wawa subprovince and the metasedimentary rocks of the Quetico subprovince. Directions: From the Mountain Iron Community Center, head east on highway 169 and turn north onto highway 53, continue north 20 miles and pull-off on the right side of the highway. STOP 2 – Alkalic and Lamprophyric Intrusive Rocks, Gheen Pluton Area Location: 515162E/5306034N (47.74992°, -92.64856°) (2a); 514380E/5306809N (-92.80754°, 47.91444°) (2c); Highway 53, ~6.5 miles northwest of Cook, MN Description: This series of outcrops examines exposures of alkalic granitoids and lamprophyric rocks intruded into metasediments of the Quetico Subprovince. Stop 2a (515162E/5306034N): This outcrop preserves outstanding porphyritic textures within pyroxene syenite and syenodiorite of the Gheen Pluton (Fig. 1-6). Evidence for multiple phases of intrusion and magma mingling are observed throughout. Very coarse phenocrysts exhibit compositional zoning and local magmaticflow features. The main phase of syenite and syenodiorite contains inclusions of, and is cross-cut by medium grained, amphibole-phyric gabbro and pyroxenite. Late aplitic and pegmatitic dikes represent the youngest intrusive components of the Figure 1-6. Porphyritic syenodiorite with abundant outcrop. Chloritic slickensides are apparent on feldspar phenocrysts at stop 2a. fracture faces, locally. Stop 2b (514530E/5306605N): This outcrop of the west side of the highway, USE CAUTION WHEN CROSSING THE ROAD. Here, pyroxene-biotite phyric lamprophyric rocks are exposed (Fig. 1-7). Beyond the dominant biotite and pyroxene, the mineralogy includes prismatic hornblende, feldspar, apatite, and trace chalcopyrite. Limited work to characterize these rocks has determined they are best described as augite bearing kersantites and spessartites (Le Bas, 2007). The mineralogy and texture vary within the a b Figure 1-7. Representative examples of lamprophyric rocks exposed at stops 2b-2d. (a) Biotite-pyroxene bearing kersantite with inclusions of wallrock. (b) Pyroxene-hornblende phyric spessartite with plagioclase dominated groundmass. Blocky, prismatic pyroxene dominates the modal mineralogy here. 8 �Trip 1 – Quetico outcrop at multiple scales, where complex structural and intrusive relationships juxtapose and include multiple mineralogic and lithologic phases. Stop 2c (514380E/5306809N): This outcrop, on the east side of the highway, is composed of similar lamprophyric rocks as stop 2b, but include blocks of lamprophyric rocks of varied composition and the schist wall-rock. The schist here is commonly altered and is cross-cut by small dikes and veinlets of lamprophyric mineralogy. Schist inclusions become more abundant to the north. Stop 2d (514262/5306920): Continuing to the north from stop 2c, the dominant lithology transitions to well foliated biotite-muscovite schist cross-cut by sulfide bearing quartz veins and discontinuous dikes and veinlets of lamproid parallel to and cross-cutting foliation. Directions: From the Stop 1, continue north along highway 53, continue north ~14 miles and pull-off on the side of the highway. STOP 3 – Polyphase, granitoid rich migmatite Location: 512642E/5316320N, (48.00005°, -92.83053°), Highway 53, ~2.5 miles north of Gheen Corner Figure 1-8. Multiphase migmatite at stop 3 showing representative intrusive relationships between the host biotite schist (dark-grey to black), tonalitic neosome 1 (grey),and granitic neosome 2 (tan-pink). Schist preserves crude structural grain. 9 �Trip 1 – Quetico Description: This extensive roadcut exhibits the complex features common throughout much of the migmatitic core of the Quetico subprovince. Here, we will observe and discuss the structural and intrusive relationships and geophysical properties between the metasedimentary quartz-biotite schist paleosome, early granodioritic and tonalitic neosome 1 intrusions, and granitic neosome 2 intrusions. The complexity observed here begs the question of how to create coherent geologic maps in similarly complex regions across the central Quetico Subprovince. Stop 3a (512642E/5316320N) Here, paleosomes of biotite schist have been strongly recrystallized and exhibit a granoblastic texture with faint foliation defined by biotite orientation. Multiple phases of neosome intrusions, both mafic and felsic, include lenses and irregular, blobby bodies of biotite-hornblende granodiorite ascribed to neosome 1 affinity. All units are cross-cut by pink, coarse-grained biotite granite and syenogranite with abundant pegmatitic veins and segregations. The intrusive relationships seen here generally apply to the regional evolution of magmatic rocks within the Vermilion Granitic Complex and Quetico Subprovince, at large. Stop 3b (512645E/5316400N) The agmatic migmatite here includes mafic and silicic paleosome blocks which are disaggregated by the intrusion of both neosome 1 and neosome 2 (Fig. 1-8). Although intrusive phases of the migmatite dominate the outcrop, the structural grain of the paleosomes is preserved as relict bedding and faint foliations. One may note that the paleosomes of differing compositions are difficult to distinguish on the outcrop. Silicic, quartz-biotite schist paleosomes are strongly recrystallized and exhibit an almost massive granular texture. Mafic paleosomes are locally present and are characterized by poorly foliated hornblende (+/- pyroxene) bearing assemblages along with coarsened biotite. Mafic paleosomes seen here may represent thin layers of primary, mafic protoliths or may be restitic components of in-situ melting of the migmatitic host rock. The exposure here is representative of many of the outcrops within the migmatitic core of the Quetico subprovince and highlights the difficulty of creating meaningful geologic maps in the region. How would you map this outcrop? Stop 3c (512644E/4135316N) Small, biotite-pyroxene lamproid intrusion, similar to those inspected at stop 2. Here, acicular, prismatic pyroxene is supported in potassium feldspar-rich segregations (Fig. 1-9). Chalcopyrite is present in trace amounts. Stop 3d (512645E/5316450N) Throughout the core of the Quetico Subprovince, migmatites are locally associated with pyroxenite and pyroxene-hornblende rich gabbroic dikes. Here, a set of pyroxenite dikes with sheared, biotite rich margins crosscut the biotite schist and neosome 1 wallrock and have mutually cross-cutting relationships with granitic neosome 2 intrusions. Stop 3e (512612E/5316833N) On the northern end of the roadcut, the complex multi-stage migmatite gives way to bedded biotite schist with graded beds and crosscutting dikes of late neosome 2 granite and Figure 1-9. Acicular, prismatic pyroxene within pegmatite. Beds here are stratigraphically facing up, potassium feldspar-rich matrix at stop 3c. This small based on fining upward sequences in graded beds, intrusion is similar to alkalic rocks observed at stop 2. and dip 45 to the south-southeast. Directions: From stop 2, continue north along highway 53, continue north ~6.5 miles and pull-off on the right side of the highway. 10 �Trip 1 – Quetico STOP 4 – Schist and schist-rich rich migmatite near Myrtle Lake Location: 523750E/5324590N, (-92.68116°, 48.07414°), Highway 23, ~7.5 miles east of Orr Description: Here, quartz-feldsparbiotite schists and schist-rich migmatite of the Quetico subprovince are exposed. This outcrop preserves moderately dipping beds (30° to the E-SE) of turbiditic metasedimentary rocks with graded beds. Though obscured by metamorphism, bedding here is interpreted to be upright with graded beds observed as decimeter scale, subtle, rhythmic changes in the amount of micaceous minerals. The base of individual beds is marked by coarse grained sandy layers, which transition to biotite rich schist marking the top of the beds. Fine grains of garnet are present locally in beds Figure 1-10. Biotite schist with faint, relict bedding intruded by with appropriate composition. The schist is boudinaged and lit-par-lit dikelets of tonalitic neosome 1. intruded by boudinaged dikes and veins up to 1 meter thick and lit-par-lit injections of tonalitic and granitic neosome 1 (Fig. 1-10). Rare dikes of coarse-grained to pegmatitic, pink, granitic neosome 2 crosscut bedding and dominant fabric of the outcrop and mark the latest intrusive event. Directions: From stop 3, continue north on Highway 53 to the town of Orr and make a right turn on OrrBuyck Road (Highway 23) and continue 7.5 miles east to the roadcut. LUNCH AND STOP 5 – Vermilion Falls ***No Hammers*** Location: 531860E/5345460N, (48.26155°, -92.57072°), Picnic area off Vermilion Falls Rd (USFS 491) Description: This picturesque waterfall cuts through quartz-plagioclase-biotite schist and schist rich migmatite. Both upstream and downstream of the falls, the Vermilion River runs parallel to the dominant regional fabric defined by the orientation of the underlying bedded and foliated metasedimentary rocks and generally foliation parallel intrusions of neosome 1 before draining into Crane Lake. Vermilion Falls occupies a N-NW trending, post-metamorphic and post-intrusive fracture and fault zone orthogonal to the dominant internal fabric of the Precambrian bedrock (Fig. 1-3, Fig. 1-11). These fracture and fault networks are ubiquitous throughout the Vermilion Granitic Complex and the Quetico Subprovince and strongly influence the surface topography and outcrop exposure in the area. Little is known about the timing and relative offsets along these fault zones, though correlation of map units on either side of these features suggests only minor relative motion. Near the upper portion of the falls, tonalitic neosome 1 dikes and sills delaminate the schist along bedding and sub-parallel foliation planes and occupy mesoscopic fold hinges. Downstream, tonalitic neosome 1 dikes are discordant and cross-cut the dominant fabric in the rock. 11 �Trip 1 – Quetico Figure 1-11. Geologic map of the Vermilion Falls area, after Jirsa and others (2011) draped over LiDAR hillshade. NW trending, post-metamorphic and post-intrusive fractures and faults have been highlighted with dashed lines. "GM" - granite rich migmatite, "SM” – schist rich migmatite, “LLC” – Lac La Croix granite, “TTG” – tonalitetrondhjemite-granodiorite gneiss. Directions: From Stop 4, continue east on Orr-Buyck Road (Highway 23) for 8.5 miles, continue straight along Crane Lake Road (Highway 24) at the village of Buyck for 9.5 miles, turn left onto Vermilion Falls Road (USFS 491) for 5.6 miles, turn left onto single-lane access road to turnaround at the picnic area. STOP 6 – Echo Lake Quarry Location: 549810E/5324630N, (48.07300°, -92.33132°), Quarry off USFS 200 Description: NOTE: This is an active quarry, permission to access this site needs to be granted from the quarry operator prior to visiting. The photogenic exposures at this dimension stone quarry include washed, glacially scoured outcrops and fresh blast faces of taxitic, red-pink to pinkish grey granitic gneiss and granite with abundant mafic inclusions (Fig. 1-12). This unit is mapped as a gneissic phase of the Lac La Croix of the Vermilion Granitic Complex, temporally related to neosome 2 seen at other stops (Jirsa, 2011). Gneissic layering here is chaotic and boundaries between gneissic phases are diffuse. Abundant mafic inclusions and schlieren ranging from a few centimeters to multiple meters in size and are randomly oriented. Mafic inclusions are delaminated along planar features and have diffuse, fringed boundaries. Increases in the abundance of biotite and 12 �Trip 1 – Quetico hornblende on the margins of mafic inclusions represent restitic rinds developed during assimilation and interaction with the host granitic melt. Directions: From Stop 5, return to Crane Lake Road (Highway 24) along Vermilion Falls Road (USFS 491) and turn right. Continue southward on Crane Lake Road (Highway 24) for 5.5 miles and turn left onto Echo Trail. Continue along Echo Trail for 8.3 miles and turn right onto USFS 200 for 5 miles. Turn Left onto unnamed forest road near Gustafson Lake. RETURN TO MOUNTAIN IRON COMMUNITY CENTER Figure 1-12. Gneissic granitoid with partially digested amphibolite inclusion. Directions: From Stop 6, return to Echo Trail via USFS 200 and turn left on Crane Lake Road. Continue south on Crane Lake Road/Orr-Buyck Road to the town of Orr. Turn left onto Highway 53 and continue 44.1 miles to the Highway 169 exit ramp. Turn left onto Enterprise Drive. 13 �Trip 1 – Quetico REFERENCES Bauer, R.L., 1985, Correlation of early recumbent and younger upright folding across the boundary between an Archean gneiss belt and greenstone terrane, northeastern Minnesota: Geology, v. 13, p. 657-660. Bauer, R.L., Czeck, D.M., Hudleston, P.J., and Tickoff, B., 2011, Structural geology of the subprovince boundaries in the Archean Superior Province of northern Minnesota and adjacent Ontario: Geological Society of America Field Guide 24, p. 203-241. Bauer, R.L., Hudleston, P.J., and Southwick, D.L., 1992, Deformation across the western Quetico subprovince and adjacent boundary regions in Minnesota: Canadian Journal of Earth Sciences, v. 29, p. 2087-2103. Chandler, V.W., and Lively, 2014, Rock Properties Database; Minnesota Geological Survey web-accessible file data (http://www.mngs.umn.edu/). 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, 99(3), pp.195-205. Day, W.C., 1990, Bedrock geologic map of the Rainy Lake area, northern Minnesota: U.S. Geological Survey Miscellaneous Investigations Series I-1927, scale 1:50,000. Day, W.C. and Weiblen, P.W., 1986. Origin of late Archean granite: geochemical evidence from the Vermilion Granitic Complex of northern Minnesota. Contributions to Mineralogy and Petrology, 93(3), pp.283-296. Fralick, P., Purdon, R.H., and Davis, D.W., 2006, Neoarchean trans-subprovince sediment transport in southwestern Superior Province: sedimentalogical, geochemical, and geochronological evidence: Canadian Journal of Earth Sciences, v.43, p. 1055-1070. Jirsa, M.A., 2011, Bedrock geology of the Crane Lake and Brule Narrows 30’X60’ quadrangles, northern Minnesota: Minnesota Geological Survey, Miscellaneous Map M-192, scale 1:100,000. Jirsa, M.S., Block, A.R., Boerboom, Chandler, V.W., and Peterson, D.M., 2020, Bedrock geology of St. Louis County, Minnesota: Minnesota Geological Survey County Geologic Atlas C-51, Part A, Plate 2—Bedrock Geology; scale 1:200,000. [contains ancillary digital files including geophysics and geochronology] Jirsa, M.A., Boerboon, T.J., and Chandler, V.W., 2014, Bedrock geology of the International Falls-Little Fork 30’X60’ quadrangles, northern Minnesota: Minnesota Geological Survey Miscellaneous Map M-197, scale 1:100,000. Le Bas, M., 2007. Igneous rock classification revisited 4: Lamprophyres. Geology Today, 23(5), pp.167-168. Poulsen, K.H., Borradaile, G.J., and Kehlenbeck, M.M. 1980. An inverted Archean succession at Rainy Lake, Ontario: Canadian Journal of Earth Sciences, v. 17, p. 1358-1369 Salerno, R.A., 2017. Neoarchean Deposition, Metamorphism, And Intrusion In Rapid Succession, Vermilion Granitic Complex, Superior Province Of Northern Minnesota. Master's thesis, University of Minnesota – Duluth. Sawyer, E.W., 2008. Atlas of migmatites (Vol. 9). NRC Research press. Southwick, D.L., 1991, On the genesis of Archean granite through two-stage melting of the Quetico accretionary prism at a transpressional plate boundary: Geological Society of America Bulletin v. 103, p. 1385-1394. Southwick, D.L., and Ojakangas, R.W., 1979, Geologic map of Minnesota, International Falls sheet: Minnesota Geological Survey, scale 1:250,000. Southwick, D.L., and Sims, P.K., 1980, The Vermilion Granitic Complex—A new name for old rocks in northern Minnesota: U.S. Geological Survey Professional Paper 1124A, p. A1-A11. Tabor, J.R., 1988, Deformational and metamorphic history of Archean rocks in the Rainy Lake District, Northern Minnesota, [Ph.D. thesis]: Minneapolis, University of Minnesota, 224 p. Valli, F., Guillot, S., and Hattori, K.H., 2004, Source and tectono-metamorphic evolution of mafic and pelitic metasedimentary rocks from the central Quetico metasedimentary belt, Archean Superior Province of Canada: Precambrian Research, v. 132, p. 155-177. Williams, H.R., 1990, Subprovince accretion tectonics in the south-central Superior Province: Canadian Journal of Earth Sciences, v. 27, p. 571-581. 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(6), pp.945-966. 14 �Trip 2 – Cu-Ni Duluth Complex FIELD TRIP 2 Drill Core from three Cu-Ni Deposits of the Duluth Complex Mark Severson1,2 (retired), Cullen Phillips3, and Kevin Boerst4 1 (1988–2012) Natural Resources Research Institute, University of Minnesota, Duluth, 5013 Miller Trunk Hwy, Duluth, MN 55811 2 (2013–2018) Previously Teck American, then Teck Resources Unlimited, now NewRange (joint venture between Teck and PolyMet Mining Inc.) 3 NewRange Copper Nickel, 6500 Kensington Dr., Hoyt Lakes, MN 55750 4 Twin Metals Minnesota, 400 Miners Drive East, P.O. Box 329, Ely, MN 55731 Diagram from Peterson (2010) modified from plots of Eckstrand and Hulbert (2007). This guidebook is modified and updated from a guidebook published in 2016 for the 62nd Institute on Lake Superior Geology (pdf). Severson, M., Ware, A., Boerst, K., and Geerts, S., 2016, Cu-Ni-PGE Deposits of the Duluth Complex. Proceedings of the Institute on Lake Superior Geology, Volume 62, Part 2-Field Trip Guidebook, Trip 3, P. 27-78 15 �Trip 2 – Cu-Ni Duluth Complex INTRODUCTION The Duluth Complex, located in northeastern Minnesota, is a series of tholeiitic intrusions of Keweenawan age (1.1 billion years ago) that formed with coeval flood basalts along a portion of the Midcontinent Rift. The Midcontinent Rift system developed during crustal extension during the Mesoproterozoic era and is traceable in a broad arc that begins in northeastern Kansas extending northward through the axis of Lake Superior and then southeastward into Michigan. The Duluth Complex and associated Keweenawan intrusions constitute one of the largest mafic complexes in the world. These rocks cover an arcuate area over 3,000 square miles (5,000 square kilometers) extending from the city of Duluth northward 170 miles (275 km) to the Canadian border. The northwest, convex edge of the complex defines its basal contact, which dips to the southeast towards the rift. Along this contact the complex is successively underlain by Neoarchean granites (Giants Range granitic rocks) and greenstones (Vermilion District) to the north, and Paleoproterozoic sediments (Virginia Formation and Biwabik Iron Formation) to the south. Roof rocks to the Duluth Complex consist of Mesoproterozoic intrusive and volcanic rocks of the Beaver Bay Complex and North Shore Volcanic Group, respectively. Once recognized as a single large lopolithic intrusion, the complex has since been established to be collectively comprised of numerous smaller subintrusions (Figure 2-1) that were episodically emplaced into the base of a comagmatic volcanic edifice between 1108 and 1098 million years ago. Figure 2-1. Generalized geologic map of northeastern Minnesota (modified from Miller et al., 2002). 16 �Trip 2 – Cu-Ni Duluth Complex The Duluth Complex hosts several known large low-grade disseminated Cu-Ni occurrences (Figure 2-2), all of which are located within the basal portions of the Partridge River (PRI), Bathtub (BTI) and South Kawishiwi (SKI) sub-intrusions. A cursory study by Listerud and Meineke (1977) estimated 4.4 billion tons of material averaging 0.66% Cu and 0.20% Ni, using a 0.5% Cu cutoff, in at least nine deposits. Five of these Cu-Ni deposits have recent NI 43-101 reports that estimate a combined mineral inventory well over that amount using lower cutoff values. Known resources (Measured and Indicated, and Inferred) for several of the deposits in Duluth Complex are shown in Table 2-1. Copper-to-nickel ratios generally range from 3:1 to 4:1. Primary mineralization is magmatic. Sulfur source is probably both local (from the footwall sediments) and magmatic. Sulfur isotope studies indicate that most of the sulfur was derived from the Virginia Formation. Most of the mineralization is in the basal portions of these intrusions but there are also local disseminated zones higher in the intrusions. The latter tend to be much more discontinuous except for continuous mineralized horizons, termed “Magenta” style mineralization, that are present at NorthMet and Mesaba, and to a lesser degree, at South Filson Creek. The mineralization styles at each of the Cu-Ni deposits are varied. The general geology and mineralization of the Partridge River, Bathtub, and South Kawishiwi intrusions, as well as the deposits that they host, are presented below. Figure 2-2. Distribution of Cu-NiPGE deposits (in red) and potential titaniumenriched ultramafic pipes (OUIs in blue) in the Partridge River, Bathtub, and South Kawishiwi intrusions. Note that the Mesaba deposit is mostly contained in the Bathtub intrusion. 17 �Trip 2 – Cu-Ni Duluth Complex Table 2-1. Known Resources for the Various Duluth Complex Cu-Ni-PGE Deposits at various Cut-Offs. Average values for Co and Ag are available for some of the deposits but are not shown in the table. Deposit Tons (st) millions Cu % Ni % Pd ppb Pt ppb Au ppb Maturi – Measured and Indicated 1,233 0.58 0.19 334 147 80 0.30% Cu Twin Metals Minnesota 43-101 AMEC Oct-14 Maturi – Inferred 563 0.49 0.16 305 134 68 0.30% Cu Twin Metals Minnesota 43-101 AMEC Oct-14 Birch Lake – Indicated 100 0.52 0.16 515 235 115 0.30% Cu Twin Metals Minnesota 43-101 AMEC Oct-15 Birch Lake – Inferred 239 0.46 0.15 370 180 87 0.30% Cu Twin Metals Minnesota 43-101 AMEC Oct-14 480 0.43 0.16 0.30% Cu Twin Metals Minnesota 43-101 AMEC Oct-14 425 0.41 0.14 0.20% Cu Encampment Resources Non 43101 Amax 1979 NorthMet – Measured and indicated 702 Open Pit 0.25 0.07 234 67 34 0.20% Cu New Range Cu-Ni 43-101F1 M3/HRC Dec-22 NorthMet – 441 Open Pit 0.25 0.07 243 67 34 0.20% Cu /New Range Cu-Ni 43-101F1 M3/HRC Dec-22 Mesaba – Measured and Indicated 2,207 Open Pit 0.43 0.10 97 34 25 NSR $12/ton New Range Cu-Ni 43-101F1 IMC/JDS Nov-22 Mesaba – Inferred 1,423 Open Pit 0.37 0.09 143 43 26 NSR $12/ton New Range Cu-Ni 43-101F1 IMC/JDS Nov -22 Spruce Road – Inferred Serpentine Inferred Cut-Off Company Source Partridge River intrusion The Partridge River intrusion (PRI) is exposed in an arc-shaped area ~10x20 miles (16x32 km) that extends from the southern edge of the Mesaba deposit on the northeast to the Water Hen deposit on the southwest as shown in Figure 2-2. Footwall rocks include the Virginia Formation and very locally the Biwabik Iron Formation. The basal stratigraphic section (Figure 2-3) was first described by Severson and Hauck (1990) and is briefly summarized below. Unit I (PR1) The lowest troctolitic unit of the PRI consists of intermixed troctolite and augite troctolite that locally grade to olivine gabbro. Most of the unit is sulfide-bearing with a PGE-bearing horizon at the top (Red Horizon of Geerts, 1991, 1994). Unique to PR1 are extreme variations in modal mineral percentage and average grain size. Due to this heterogeneous texture, numerous internal contacts divide PR1 into several subunits that are probably related to continuous magma replenishment. Hornfels inclusions of the Virginia Formation are most commonly present within PR1. Near the basal contact the intrusive rocks of PR1 have undergone silica contamination and norite and gabbronorite are often the dominant rock type in the bottom zone. Unit II (PR2) This unit is characterized by sulfide-poor, texturally-homogenous, troctolite that locally grades to augite troctolite and leucotroctolite. PR2 grades downward into a persistent ultramafic horizon(s) defined by melatroctolite, with local peridotite zones, that generally exhibits a sharp contact with PR1. 18 �Trip 2 – Cu-Ni Duluth Complex Figure 2-3. Stratigraphy of the Partridge River intrusion at the Mesaba, NorthMet, Wetlegs, and Wyman Creek deposits (note that the Bathtub intrusion is denoted by the BT-series units in the lower right corner). From Severson and Hauck (2008). Unit III (PR3) Unit III is the most distinctive “marker bed” of the PRI at the NorthMet, Wetlegs, and southern Mesaba deposits. This unit is fine-grained and is characterized by leucotroctolite that locally grades to troctolite and augite troctolite. In all cases, the rock presents a mottled appearance due to the presence of very coarse-grained (>2 cm) olivine oikocrysts that are irregularly distributed throughout the rock. This mottled-texture and fine-grained nature make PR3 unique relative to all the other units of the PRI. PR3 exhibits variable thicknesses at each of the Cu-Ni deposits and pinches out to the west of Wetlegs and to the southeast of Mesaba. The extreme thickness range for PR3 and its physical attributes (poikilitic) have suggested to several geologists that it may be associated with an earlier Anorthositic Series intrusive phase. In this scenario, PR3 may have been intruded earliest along the Virginia Formation-North Shore Volcanic contact and was later underplated by PR1 and PR2 in an early-formed magma chamber subject to continuous magma replenishment. Unit IV (PR4) Unit IV of the PRI is characterized by thick intervals of texturally-homogeneous troctolite and/or augite troctolite. In many areas, PR4 grades upward into a persistent zone of augite-rich augite troctolite and olivine gabbro, which in turn, grades upward into leucotroctolite that is characteristic of PR5. At its base, PR4 has a semi-persistent ultramafic horizon that contains one or more melatroctolite and/or peridotite layers. In some areas a thin semi-massive oxide layer containing very fine-grained chromium titanomagnetite is present immediately above the upper contact of PR3. Unit V (PR5) Unit V is generally an easily recognizable unit in that it is characterized by thick intervals of texturally-homogeneous, medium- to coarse-grained leucotroctolite (dominantly anorthositic troctolite). Another feature that aids in distinguishing PR5 is a highly gradational bottom contact into augite troctolite at the top of PR4. The upper contact of PR5 is sharp against one or more ultramafic horizons that mark the base of the overlying PR6. 19 �Trip 2 – Cu-Ni Duluth Complex Unit VI (PR6) Leucotroctolite (anorthositic troctolite to troctolitic anorthosite) is the most common rock type in PR6. However, near equal amounts of troctolite and augite troctolite are more common in some drill holes at Mesaba. Overall, PR6 becomes more heterogeneous, consisting of multiple rock types, toward the southern and eastern margins of the Mesaba deposit. The base of PR6 is usually marked by a fairly persistent ultramafic horizon. Unit VII (PR7) This unit consists almost wholly of homogeneous leucotroctolite at the NorthMet deposit, but it is characterized by a potpourri of rock types at Mesaba with leucotroctolite being slightly more common. Overall, PR7 becomes more heterogeneous, consisting of multiple rock types, toward the southern and eastern margin of the Mesaba deposit. PR7 contains a basal ultramafic horizon(s) in most drill holes. Unit VIII (PR8) The uppermost PRI unit that has been drilled at Mesaba is referred to as PR8 that consists of a multitude of rock types with no consistent pattern except that leucotroctolite is slightly more dominant. PR3-like inclusions are excessively common to this unit. Oxide-bearing Ultramafic Intrusions (OUIs) Several plug-like, late-stage, oxide-bearing ultramafic intrusions have been delineated in the PRI, the Bathtub intrusion (BTI), and elsewhere within the Duluth Complex (Figure 2-4). The OUIs are intrusive into all units of the PRI and BTI and range in size from large bodies (>200 feet thick, >60 meters thick) to small bodies/lenses (<30 feet thick, <9 meters thick). Rock types are characterized by coarse- to very coarse-grained peridotite and dunite to pegmatitic clinopyroxenite and locally minor orthopyroxenite. These rock types contain varying amounts of ilmenite and titanomagnetite ranging from 5% to massive oxide zones (>80% oxides). The OUIs are in sharp contact with the surrounding troctolitic rocks and are clearly younger. In almost all instances the OUIs are spatially arranged along linear trends suggesting that structural control was important to their genesis. Two of the OUIs are currently being evaluated for their titanium potential and include: 1. Longnose with a NI 43-101 inferred resource of 65.3 million tonnes of 16.4 TiO2; and 2. Titac with a NI 43-101 inferred resource of 45.1 million tonnes of 15% TiO2 (Farrow, 2012). A third OUI, Skibo, is being evaluated for its high-grade Cu-Ni-PGE potential where two vein stockwork zones have been identified by historic drill holes (Inco – up to 6.42% Ni in a 1 foot-thick massive sulfide and other intervals in the hole) and recent drilling by Encampment Minerals (Green Bridge Metals press release, Feb. 6, 2024 at www.greenbridgemetals. com). NorthMet Deposit (NewRange Copper Nickel) The NorthMet deposit is located in the PRI as shown in Figures 2-2 and 2-5. This deposit was initially drilled by United States Steel Corporation (USSC) at what they called the Dunka Road deposit. More recent drilling was conducted by PolyMet Mining Incorporated at the now renamed NorthMet deposit that is being developed by NewRange Copper Nickel (Glencore and Teck joint venture). The geology of the deposit consists of seven igneous units, originally defined by Severson and Hauck (1990) as shown in Figure 2-6. 20 �Trip 2 – Cu-Ni Duluth Complex Mineralization Trends at NorthMet Two open pits are currently planned at the NorthMet deposit - an East Pit and a West Pit (shown in Figure 2-7). The majority of economic mineralization at NorthMet occurs in three scenarios: 1. All of Unit I is mineralized at the East Pit (see cross-section in Figure 2-8); 2. mostly the upper portion of Unit I is the best mineralized in the West Pit (see cross-section in Figure 2-9) and the bottom portions of Unit I will not be mined; and 3. the cross-cutting Magenta Zone (Figs. 2-9 and 2-10) is located well above the basal contact. Grades are generally highest at the top of Unit I and decrease going down hole. However, there are exceptions, and the middle of Unit I contains the highest grades in the center of the deposit. PGE-enriched zones at NorthMet Geerts (1991, 1994) found that the top of Unit I often hosts a PGE-bearing zone that he referred to as the Red Horizon (also referred to as Red Zone) which was determined to be approximately 10 meters thick with an average of 0.57% Cu and 986 ppb Pt+Pd. Geerts also found two more PGE-bearing zones within Unit I referred to as Orange and Yellow horizons. All three of these horizons are positioned beneath ultramafic layers suggesting that they are the result of recharge events and magma mixing whereby a new influx of primitive, PGE-bearing magma was injected into the chamber creating Figure 2-4. Distribution of the Oxide-bearing Ultramafic the ultramafic layers (crystal settling) before Intrusions (OUIs) within the Partridge River, Western mixing with the resident magma (sulfideMargin and Boulder Lake intrusions. Note the linear bearing) and forming the PGE-enriched zones arrangement of OUI along various trends. beneath them. The continuity of these three PGE-bearing zones in more recent PolyMet/NewRange drilled holes is unknown and the three zones are not specifically mentioned in any NI 43-101 reports or field trip guidebooks. It is important to note that the Red Horizon/Zone has been documented to be present at the top of PR1 at Mesaba (Severson and Hauck, 2003). Mineralized Magenta Zone at NorthMet In addition to the Red Horizon, at the top of PR1, there is another PGE-bearing horizon that has been referred to as the Magenta Horizon (or Magenta Zone). This zone is unique in that it crosses several lithologic contacts and progressively downcuts through Units 6, 5, 4, and 3 in a northerly direction (Figure 2-10). The total resource volume of the Magenta zone relative to the rest of the deposit has not been documented in any NI 43-101 reports. Initially, Geerts (1991, 1994) found the Magenta Zone in six holes wherein it averaged about 0.72% Cu and 1,488 ppb Pd+Pt in an over 8 meters thick zone. Cu:Ni ratios in the Magenta Zone are reported to be 3.9-4.1:1. More recent drilling by PolyMet and NewRange has documented the presence of the PGE-enriched Magenta Zone in additional drill holes that are positioned in 21 �Trip 2 – Cu-Ni Duluth Complex the western half of the deposit as shown in Figure 2-8. The Magenta Zone is also present in the PRI along the southern edge of Mesaba deposit to the east where it is referred to as the PRU zone by NewRange CuNi. Figure 2-5. Location and geology of the NorthMet deposit relative to the nearby Mesaba deposit. Modified from combined maps of Miller and Severson (2005) and Severson and Miller (2005). Mesaba Deposit and the Bathtub intrusion (NewRange Copper Nickel) In 1990, the Natural Resources Research Institute (NRRI) was the first to define and describe the igneous stratigraphy of the PRI (Severson and Hauck, 1990). This same stratigraphy was documented to be present in portions of the Mesaba deposit in 1995 (then referred to as the Babbitt deposit). However, this stratigraphy applied to only the deep drill holes along the extreme southern portion of Mesaba and all 22 �Trip 2 – Cu-Ni Duluth Complex Figure 2-6. Igneous stratigraphic section recognized by NewRange Copper Nickel at their NorthMet deposit (not that these same units are also present along the southern margin of the Mesaba deposit where they are referred to as PR1, PR2, PR3 etc.) 23 �Trip 2 – Cu-Ni Duluth Complex Figure 2-7. Geologic map of the NorthMet deposit showing outlines of the two planned open pits. The location of the mineralized Magenta zone is present in the southern half of the West Pit. Figure 2-8. Cross-section illustrating mineralization trends in NorthMet’s East Pit. Note that all of Unit I is mineralized down to the footwall Virginia Formation. 24 �Trip 2 – Cu-Ni Duluth Complex Figure 2-9. Cross-section illustrating mineralization trends in NorthMet’s West Pit. Note that the top portion of Unit I will be mined as it exhibits the best mineralization. Note also that the Magenta mineralized zone is present in a downcutting relationship in Units 5, 4, and 3. Figure 2-10. Typical cross-section at NorthMet (facing east) showing mineralized zones and modeled units. The “Upper Zone Mineralization” in this diagram is also referred to as the Magenta zone. Note how this zone progressively transects downward into the lower geologic units in a northerly direction. 25 �Trip 2 – Cu-Ni Duluth Complex Figure 2-11. Geologic map (circa 2015) showing distribution of major igneous units in the Bathtub intrusion (BTI) and adjacent Partridge River intrusion (PRI) of the Mesaba deposit. attempts to carry this stratigraphy to the north into the majority of Mesaba were not conclusive. Through several iterative follow-up logging campaigns by the NRRI, the Bathtub intrusion (BTI) was finally recognized as a separate intrusion (Severson and Hauck, 2008). A geologic map of the Mesaba deposit showing the geologic units (per the igneous stratigraphy) is shown in Figure 2-11. At least five criterions, listed below and discussed in Severson and Hauck (2008), were initially used to help separate the PRI from the newly named BTI: 1. Abrupt terminus of the PR3 Unit (major marker bed in the PRI) northward into the BTI at the Mesaba deposit 2. Thicker sections of heterogeneous-textured rock in the BTI relative to the adjacent PRI 3. Lack of PGE-enrichment at the top of a specific unit in the BTI (BT1 Unit) relative to PGEenrichment at the top of a similar unit (PR1) in the adjacent PRI 4. The best mineralization at Mesaba is near the base of the BTI (base of the BT1 unit) where it is characterized by high Cu grades associated with pyrrhotite- and cubanite-rich zones. In contrast, the best mineralization at the majority of the NorthMet deposit is present at the top of the PR1 unit where it is associated with chalcopyrite-rich zones 5. Use of a hornfels-rich zone, termed the Hidden Rise, was used to help separate the BTI from the PRI. The current igneous stratigraphy of the Bathtub intrusion, as defined by the NRRI, is summarized in Figure 2-12, and is described in the sections below. 26 �Trip 2 – Cu-Ni Duluth Complex Figure 2-12. Stratigraphic section at the Mesaba deposit showing the relationships between major units (and their corresponding subunits) in the BTI and PRI. Modified from Severson and Hauk (2008). BT1 Unit The lowest unit of the BTI consists of intermixed troctolite and augite troctolite with localized leucotroctolite zones. Unique to BT1 are extreme variations in modal mineral percentage and average grain size; both change rapidly over zones that vary from a few feet to tens of feet thick. Due to this heterogeneous texture, numerous internal contacts subdivide the BT1 into several subunits that often cannot be correlated from drill hole to drill hole. Thus, BT1 is a mixture of various troctolitic subunits that are probably related to continuous magma replenishment. Most of this unit is sulfide bearing. Hornfels inclusions of Virginia Formation are common within the BT1 Unit, especially closer to the basal contact. The BT1 Unit has been further subdivided into several internal subunits based on the dominant presence of one rock type over other rock types. Contacts between these rock types vary from highly gradational to abrupt with locally measurable sharp contacts. The various subunits of the BT1 Unit, and hornfels-rich zones in the BT1, are presented in Figure 2-12 and some are briefly discussed below. • BT1-c – At the base of the BT1 unit there is significant silica contamination of the magma, due to assimilation of the footwall rocks, and orthopyroxene rather than olivine crystallized to produce noritic rocks. Thus, rock types that dominate in the BT1-c subunit range from norite to gabbro norite; especially near either the basal contact or surrounding common hornfels inclusions. Overall, the BT1-c subunit spatially occurs as a rind or coating along the basal contact of the BTI where it ranges anywhere from a foot-thick to over 650 feet-thick. • “The Rise” – along the extreme northern edge of the entire Mesaba deposit, the basal contact of the BTI rises steeply toward the surface. However, in one area, called “The Rise,” the basal contact actually subcrops at the surface and then drops off again in a northerly direction beneath the South Kawishiwi intrusion (see Figure 2-11 for location); A pyrrhotite-rich and graphite-rich unit within the Virginia Formation in “the Rise” has been informally termed the Bdd Po or BDPO unit. • The “Hidden Rise” – the “Hidden Rise” unit is a loosely-defined zone situated along the crest of the Local Boy anticline (Figures 2-11, 2-12 and 2-13) wherein scattered hornfels inclusions, and associated noritic rocks, are fairly common. Like the BT1-c unit, the Hidden Rise shows evidence 27 �Trip 2 – Cu-Ni Duluth Complex of mixing and contamination with the Virginia Formation. This unit is indicative of strong magma contamination and assimilation of what once may have been a magma chamber wall initially separating the BTI and PRI. Thus, the Hidden Rise is used to both define this hornfelsbearing zone and to artistically, and conveniently, divide the BTI from the PRI. Figure 2-23. Projected distribution of the Hidden Rise at Mesaba relative to structural features. The projected location of the shaft and drifts of the Local Boy ore zone are shown in red. The southern edge of the Hidden Rise is approximated due to a paucity of drill holes. BT4 Unit The uppermost unit of the Bathtub intrusion is referred to as the BT4 Unit. It was originally correlated with PR4 of the PRI. However, BT4 is distinctly different from PR4 in that the BT4 Unit is heterogeneous-textured at all scales (though less heterogeneous than BT1 overall), composed of many alternating rock types, and is locally sulfide-bearing. The BT4 Unit appears to grade into thicker, more homogenous troctolitic packages toward the extreme east of the deposit. The BT4 Unit has been further subdivided into several more internal subunits based on the dominant presence of one rock type over other rock types. All these various subdivisions of the BT4 Unit are shown Figure 2-12 and are discussed below. • “± Picrite – the base of the BT4 is defined by a semi-persistent ultramafic layer and/or package, consisting of melatroctolite to peridotite ± troctolitic beds that is referred to as the "± Picrite.” The "± Picrite is present in about 60% of the drill holes in the Bathtub Intrusion and acts as a local horizon that defines the BT1-BT4 contact; however, in many instances the "± Picrite is absent and the BT1-BT4 contact is arbitrarily chosen based on its presence in nearby drill holes. 28 �Trip 2 – Cu-Ni Duluth Complex • Bathtub Layered Interval (BTLI) – this subunit designates zones (see Figures 2-12 and 2-14) where ultramafic layers are extremely common within the BT4 Unit. The ultramafic layers may represent repetitious cyclic layers and can be correlated in drill holes as an overall rock package. The inclination of internal contacts and modal bedding associated with the ultramafic layers are highly variable, ranging from 5° to 80° (even within a single drill hole). Individual ultramafic beds cannot be traced with certainty between drill holes; however, correlations of packages of the BTLI can be traced. This dichotomy for individual ultramafic beds indicates that the bedding relationships are extremely complex in the third dimension and may be related to rapid pinch-out of individual beds. In addition, the BTLI package fades out to the north with increased distance away from the Hidden Rise. If the Hidden Rise represents a magma chamber wall, the BTLI may have crystallized against it via either a static crystallization method or by current-driven crystal settling against the wall. Figure 2-34. Spatial distribution of the BTLI (in solid green hatch) relative to Bathtub syncline and the Hidden Rise (cross-hatched zone). This map is circa 2015 and changes have been made by NewRange Copper Nickel based on newer information. 29 �Trip 2 – Cu-Ni Duluth Complex Footwall Rocks The footwall rock types at both the NorthMet and Mesaba deposits consist mainly of the Virginia Formation, Biwabik Iron Formation (BIF), and very locally the Pokegama Quartzite. All are Paleoproterozoic in age (approximately 1.9-1.8 billion years ago) and collectively comprise the Animikie Group. The rock types of the Virginia Formation and BIF have undergone metamorphism and partial melting that was produced during emplacement of the Duluth Complex. The metamorphic variants of the footwall rocks are schematically portrayed in Figure 2-15 but are not discussed individually herein. Figure 2-45. General Relationships of the Metamorphosed Footwall Rocks beneath the Duluth Complex at the Mesaba, NorthMet, Wetlegs, and Serpentine Deposits. See Severson and Hauck (2008) for more information. Structural Features There are several prominent structural features at Mesaba that were important to formation of the BTI and possibly to mineralization trends. These major features are shown in Figure 2-16 and discussed below (features such as the Rise and the Hidden Rise have been discussed previously). Local Boy anticline and Bathtub syncline The most prominent structural features at the Mesaba deposit are a pair of east-west trending parallel folds, defined by contouring the top of the footwall Biwabik Iron Formation (Figure 2-17), that are informally referred to as the Local Boy anticline and Bathtub syncline. Both of these folds probably exerted strong controls on the style of emplacement of the BTI and its basal contact mimics the form of the anticline and syncline. The trend of the Hidden Rise, the possible wall once separating the BTI and PRI, also correlates with these two fold axes. The structural history regarding the Local Boy anticline and Bathtub syncline appears to be extremely complicated and long lived. Grano Fault Along the far eastern edge of the Mesaba deposit is the north-trending Grano Fault (Fig. 2-16), so named for the abundant and sometimes voluminous amounts of associated late granitoid lenses and OUIs that are associated with the fault zone (Severson, 1994). Both types of late intrusive lenses are interpreted to be steeply oriented and to have been injected along subsidiary fault zones parallel to, and immediately west of the Grano Fault. These late intrusives cross-cut the troctolitic rocks and thus, demonstrate that the 30 �Trip 2 – Cu-Ni Duluth Complex Figure 2-56. Major structural features at the Mesaba deposit. Note the orange-outlined zone to the immediate west of the Grano Fault is a zone wherein late stage subvertical lenses of granitoid and OUI (cyan outlines) commonly cross-cut the troctolitic rocks of the PRI and BTI. OUI (outlined in cyan) are also common along the inferred trace of the South Minnamax Fault. Figure 2-67. Contoured top of the footwall beneath the Mesaba deposit relative to sea level. The contour interval is 100 feet. Note that the contoured lines in this map are derived from Severson and others (1994) and do not take into account any of the more recent drill holes; however, the overall trends would remain basically the same. 31 �Trip 2 – Cu-Ni Duluth Complex fault was active during and after emplacement of the PRI, BTI and SKI. The Grano Fault is thought to be a primary feeder structure for the BTI and possibly the massive sulfides at the Local Boy ore zone (Severson and Hauck, 2008). South Minnamax Fault The South Minnamax Fault is an east-west trending fault along the extreme southern edge of the Mesaba deposit. Several OUIs occur at the surface along the trend of the fault (Figure 2-16). Displacement of the fault, based on correlations and projections of units between only six drill holes, is generally 100200 feet (30-61 meters), but in one area a displacement of over 400 feet (122 meters) is indicated. Mineralization The Mesaba deposit is characterized by disseminated sulfide mineralization that occurs most commonly as fine- to coarse-grained, intercumulus disseminations of chalcopyrite, cubanite, pentlandite, and pyrrhotite. The most important continually mineralized zone at Mesaba is a basal zone with disseminated sulfides that is present within all or portions of the BT1 unit, and locally in the bottom of the BT4 unit (Figure 2-18). This zone commonly ranges between 200 and 600 ft (61 to 183 m) in thickness. Higher in the intrusive package, often overlapping the BT1-BT4 unit boundary, are thinner, secondary zones of erratic and discontinuous, disseminated sulfide mineralization referred to as “cloud zones.” Increased sulfide contents with depth are obvious in drill core and are manifested mainly by increasing amounts of pyrrhotite and cubanite. This dramatic increase in pyrrhotite and cubanite with depth appears to be related to contamination from the footwall rocks and has been classified as occurring mainly in the basal contaminated BT1-c unit but there are exceptions. Talnakhite [Cu9(Fe,Ni)8S16] is present in numerous holes coincident with the axis of the Bathtub syncline (and north of the Hidden Rise), as well as, in the massive sulfides at Local Boy, as shown in Figure 2-19. Talnakhite occurs as exsolution lamellae with chalcopyrite and cubanite. Talnakhite is difficult to distinguish from chalcopyrite in freshly drilled core. However, talnakhite tarnishes rapidly, sometimes within 10-15 minutes depending on the relative humidity, to a purplish brown or peacock blue similar to bornite (orange-brown color in polished sections). Figure 2-78. Typical cross-section at the Mesaba deposit showing mineralization throughout most of BT1 and in portions of BT4. Note the discontinuous “cloud zone” occurrences in the upper portions of the BT4 unit. 32 �Trip 2 – Cu-Ni Duluth Complex Figure 2-89. Distribution of holes that contain significant amounts of Talnakhite based on tarnished relationships observed on drill core. This map is circa 2015. Note that the Local Boy ore zone, shown in lower right red ovoid, also contains significant talnakhite. 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 1,700foot-deep exploratory shaft (Minnamax shaft), and in 1977, completed four drifts (A, B, C, and D; Figures 2-20 and 2-21). Underground Fan drilling (217 holes) was completed in 1978 to further define the massive sulfide distribution. Potential ore resources for Local Boy are presented in Table 2-2; high PGE values (up to 11 ppm Pd and up to 8 ppm Pt) are locally present in the ore. Sulfide minerals include pyrrhotite, pentlandite, chalcopyrite, talnakhite, cubanite, maucherite (nickel arsenide), sphalerite, bornite, late mackinawite, chalcocite, covellite, godlevskite, and native silver (Severson and Barnes, 1991). Table 2-2. Grade/tonnage data for Cu and Ni in the Local Boy ore zone. These values are for geologic resources, not mineable ore. From Severson and Barnes, 1991. The Local Boy ore zone is also situated over the Local Boy anticline. The majority of massive sulfide ore zones, hosted mainly by the Virginia Formation (Severson and Barnes, 1991), are broadly 33 �Trip 2 – Cu-Ni Duluth Complex coincident with the axis of the anticline. The contoured top of the BIF in the Local Boy area is shown in Figure 2-20 (left). Similar anticline geometries are also present for the basal contact as shown in Figure 220 (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. Figure 2-20. Contoured top of the Biwabik Iron Formation at Local Boy (left) and the contoured top of the basal contact between the Virginia Formation and the intrusive rocks at Local Boy (right). 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 (mostly norite) 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. Figure 2-21 is an attempt to show, in 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 2-21 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 and progressively became more Cu and PGE enriched as it moved through the footwall rocks in an east-to-west direction. 34 �Trip 2 – Cu-Ni Duluth Complex Figure 2-29. 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 cumulative thickness of the massive sulfides (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. 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 2-21 right), close to the Grano Fault (feeder) and along the axis of the Local Boy anticline (structurally-prepared site); 2) the VirgSill, at the base of the Virginia Formation, 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 Curich (5-25% Cu) 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. 35 �Trip 2 – Cu-Ni Duluth Complex Wetlegs Deposit The Wetlegs deposit (Figures 2-2 and 2-3) was drilled by Bear Creek (13 holes) and Exxon (12 holes). Exxon determined that there were 38 million tons of material at a 0.57% Cu equivalent (files at DNR) but details regarding their cursory calculations are unknown. Most of the igneous units that are present at NorthMet are also present at Wetlegs except: Unit II thins down to a single ultramafic horizon positioned immediately below Unit III (Figure 2-3), and Unit I contains abundant ultramafic layers that are referred to as the Wetlegs Layered interval (Miller and others, 2002). The top of Unit I (aka Red Horizon of the NorthMet deposit) contains scattered anomalous concentrations of PGEs up to 3,132 ppb Pd+Pt (Severson and Hauck, 2003). The Magenta Zone is also present at Wetlegs, but is only known in one drill hole (A4-11) with up to 6,072 ppb Pd+Pt. No work has been conducted on this property since 1998. Wyman Creek Deposit (Encampment Minerals) The Wyman Creek deposit (Figures 2-2 and 2-3) is located at a turning point in the basal contact of the PRI – the contact trends northeast to the east of Wyman Creek and then exhibits a drastic change to a north-south orientation to the south of the deposit. This area was initially drilled by Bear Creek, followed by more extensive drilling (21 holes) by United States Steel Corp. (USSC), and very limited drilling by Exxon. USSC determined (literally a back-of-the-envelope calculation) an open pit potential of 14 million tons of material containing 0.30% Cu and 0.18% Ni (Severson and Heine, 2007). South Kawishiwi intrusion The South Kawishiwi intrusion (SKI) is exposed in an arc-shaped area ~5x20 square miles (8x32 square km) that extends from the Serpentine deposit on the southwest to the Spruce Road deposit on the northeast as shown in Figure 2-2. Footwall rocks include the Virginia Formation, Biwabik Iron Formation, and granitic rocks of the Neoarchean Giants Range granitic complex; the latter is the dominant footwall rock type. The basal stratigraphic section (shown in Figure 2-22) is known in detail from studies of historic drill core (Severson, 1994; Zanko and others, 1994) and is divided into 17 different units that are present over a strike-length of 19 miles (31 kilometers). Figure 2-102. Generalized igneous stratigraphy of the basal zone of the SKI (Severson, 1994). The Lowermost 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 1; AT-T = Anorthositic Troctolite to Troctolite; UW = Up dip Wedge; Main AGT = Main Augite Troctolite; AN-G Group = Anorthositic Series inclusion with internal gabbroic lenses. 36 �Trip 2 – Cu-Ni Duluth Complex The lowermost units are unevenly distributed along the strike-length of the intrusion in a compartmentalized fashion, suggesting a complicated intrusive history. The stratigraphy, as defined by Severson (1994), has been documented to be present in all the holes drilled recently by Twin Metals Minnesota (TMM) at the Birch Lake and Maturi deposits but it has since been simplified by TMM. According to Severson and Hauck (2008), a few salient features of the SKI include: • • • • • The vast majority of sulfide mineralization is confined to the BH, BAN, and U3 units - all of these are collectively referred to as BMZ by TMM at the Maturi deposit). The PEG unit, though not particularly mineralized except locally, is also included in the BMZ by TMM Major marker beds include three horizons that contain abundant ultramafic layers (U1, U2 and U3) and a pegmatite-bearing unit (PEG unit – originally recognized by Foose, 1984). The U3 unit is unique in that it contains several massive oxide pods (titanomagnetite-rich and locally Cr-bearing) as well as recognizable inclusions of bedded Biwabik Iron Formation. 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 a high degree of partial melting of the iron-formation. This relationship is the most prevalent at Birch Lake The U3 unit contains the vast majority of high PGE values; however, high PGE values are locally present in the overlying PEG unit. High PGE values are also present well above the base of the SKI at the South Filson Creek deposit A large inclusion of anorthosite of formidable size (3,500 feet thick) is present at Maturi and was referred to as the AN-G Group by Severson (1994). Peterson (2001) suggested that the high PGE contents within the BMZ unit (beneath the inclusion) formed as a result of confined turbulent magma flow, and thus an increased R-factor, beneath a “pillar” of anorthosite. Peterson (2001) further hypothesized that a portion of a Nickel Lake Macrodike, which served as a feeder to the nearby Bald Eagle intrusion, may have projected beneath the Maturi deposit and also served as a feeder to the SKI Four of the Cu-Ni deposits within the SKI historically held by TMM are shown in Figure 2-23. These four deposits and others within the SKI are discussed in the following sub-sections. The information given for each of the deposits is based on various NI 43-101 reports, field trip guidebooks (Patelke and others, 2009; and Severson and others, 2016), various NRRI reports, oral presentations at professional meetings, and personal knowledge. Maturi Deposit (Twin Metals Minerals) The very first exploration drill hole in search of Cu-Ni deposits in the Duluth Complex was cored in the Maturi deposit by Fred S. Childers (prospector) and Roger V. Whiteside (investor) in 1951. Eventually, the International Nickel Company (Inco) picked up the property and outlined a sizeable, but low grade, Cu-Ni deposit. A shaft was sunk on the property during 1966 to 1967 to collect material for metallurgical tests. Inco took their bulk sample from pyrrhotite-rich material, which is more prevalent near the basal contact, and decided the grade was too low to support an underground mine. Three other companies (Bear Creek, Duval and Newmont) put down several scattered drill holes on the periphery of the deposit and intersected good mineralization at great depth but also determined it was too low grade to support an underground mine. 37 �Trip 2 – Cu-Ni Duluth Complex Figure 2-113. Twin Metals Minnesota (TMM) deposits and Resource Classification. Mineralization at Maturi, and an extension referred to as Maturi SW, is present in the lower portion of the South Kawishiwi intrusion (SKI) in what TMM refers to as the Basal Mineralized Zone (BMZ). Note that the BMZ collectively consists of the PEG, U3, BH and BAN units of Severson (1994). While the Severson units are recognized by TMM at Maturi they are not consistently present throughout the deposit and for simplicity-sake TMM combined them into the BMZ unit. In 2008, Dean Peterson from Duluth Metals (the original company from which TMM was created) theorized that the initial SKI magmas at Maturi were intruded as sulfide-bearing, crystal-laden (olivine- and plagioclase-rich), crystal slurries. Based on this new interpretation, TMM combined Severson’s (1994) basal units into the BMZ unit that they believed originated by physical sorting of the crystal slurries (possibly top down) and by melting of the footwall granitic rocks (bottom up) to create the heterogeneous lithologies and textures of the BMZ as is shown in Figure 2-24. Mineralization at the Maturi deposit consists of a tabular sheet of disseminated Cu-Ni-Fe sulfides that averages 215 feet thick (65 meters) with a range of 5 to 865 feet thick (1.5 to 260 meters) with the thickest range towards the north end of the deposit. Dips of the BMZ vary from 35 to 55 degrees with a N60E plunge along the contact. Higher grades are concentrated in the upper 100 feet (30 meters) of a zone that has been traced laterally by drilling for approximately 2.2 miles (3.5 km) and open at depth. While mineralization is mostly restricted to the BMZ, exceptions are locally present in the overlying PEG unit and in the footwall granitic rock. TMM reports that mineralization within the footwall granite occurs in appr38 �Trip 2 – Cu-Ni Duluth Complex Figure 2-124. Simplified crystal-liquid slurry model for the SKI in the Maturi area. oximately one-quarter of the holes drilled to date with about 80% of these holes showing mineralization in the overlying BMZ that continues directly downward into the footwall mineralization with little or no breaks. Mineralization typically consists of 1-5% disseminated chalcopyrite, talnakhite, cubanite, pyrrhotite, and pentlandite. Bornite, covellite, and millerite occur in subordinate amounts. Better grades of Cu, Ni and PGE are associated with more mafic units located near the top of the BMZ. Modeling of the ore deposits by Duluth Metals, TMM, and AMEC indicated that the mineralization at Maturi can be characterized by several distinct patterns as shown in Figure 2-25. Figure 2-135. Mineralization trends of the BMZ and adjacent rocks within the Maturi area. 39 �Trip 2 – Cu-Ni Duluth Complex The three stages of mineralization within Maturi’s BMZ zone include: • Stage 1 Mineralization (S1, with top and bottom zones): barren to very low-grade mineralization showing low variability. • Stage 2 Mineralization (S2, with top and bottom zones): moderate grade mineralized intervals showing low variability. Cu:Ni ratios are 3.0 to 3.2:1. Cu-sulfides are the dominant sulfide but pyrrhotite becomes increasingly present with depth. Cubanite and pentlandite decrease in abundance with depth. Normalized chalcopyrite/chalcopyrite+cubanite ratios are approximately 0.61 to 0.65 (Hoffmann and others, 2015). • Stage 3 Mineralization (S3): higher grade mineralized intervals that are commonly bounded by low grade selvages and, interestingly, contains ultramafic units (aka U3 unit of Severson, 1994). Cu:Ni ratios are 3.0 to 3.2:1. Cu-sulfides are the dominant sulfide. Normalized chalcopyrite/chalcopyrite+cubanite ratios are approximately 0.58 to 0.67 (Hoffmann and others, 2015). Birch Lake Deposit (Twin Metals Minerals) The Birch Lake deposit lies to the south of Maturi (Figures 2-2 and 2-23), with mineralization also hosted at the bottom of the SKI. The area was first drilled by Duval in the 1970s but remained dormant until high PGE values were found in drill hole Du-15 by state agencies in the mid-1980s (Sabelin and Iwasaki, 1985). This discovery marked the start of serious PGE exploration in the Duluth Complex. Ernest Lehmann formed several joint ventures, the last known as Franconia Minerals LLC or Beaver Bay Joint Venture, and several holes were drilled on the property intermittently during 1988 through 2010. TMM acquired the property in 2011 and drilled 30 holes from 2011 to 2012. A total of 114 holes have been drilled at Birch Lake (excluding 154 wedge holes that were drilled mainly to obtain material for metallurgical testing). The geology is very similar to Maturi except for the common occurrence of more (and often thicker) ultramafic layers, assimilated BIF inclusions, and discontinuous oxide-rich horizons/pods that are inferred to represent BIF “restites”, all of which are present in the U3 unit. The continuity of these U3 rock types is extremely heterogeneous in 3D as revealed by wedge drilling. Mineralization is associated with what TMM also refers to as the BMZ which consists of the U3, BH, and BAN units of Severson (1994). The BMZ averages about 100 feet thick (30 meters) but is as thick as 515 feet (157 meters). The main footwall unit at Birch Lake is the Neoarchean Giants Range granitic complex, but Paleoproterozoic rocks are exposed at the surface in the Dunka Pit mine located <1 km to the southwest. The four inferred mineralization types at Birch Lake in the BMZ and GRB are shown in Figure 2-26 (non-mineralized material below the GRB_M is identified as GRB_B for barren footwall rocks). Figure 2-26. Igneous stratigraphy according to mineralization trends at Birch Lake. 40 �Trip 2 – Cu-Ni Duluth Complex Mineralization trends at Birch Lake are very similar to Maturi, with four inferred types: 1. Melatroctolite / BL_MT Unit (similar to S3 at Maturi and U3 unit of Severson): an upper melatroctolite to mafic troctolite unit that hosts the highest grade mineralization and is correlative across the deposit. The top and bottom of this unit are typically based on high Mg contents with values generally greater than 6% Mg. The Cu:Ni ratio is about 3.3. The base of the BL_MT unit is gradational downward into the BL_T unit. Almost all of the significant Cu-Ni and precious metal mineralization is hosted by this unit but the total volume, or percentage of the mineral resource, has not been published 2. Troctolite / BL_T Unit (similar to S1 at Maturi and BH and BAN units of Severson): a lower troctolitic unit with lower grades that is also correlative across the deposit. Mg contents are in the 3.5-4.5% range. Locally the top of BL_T is more mineralized and there are small, mineralized zones near the base 3. Basal Hybrid Zone / BL_HX: a basal hybrid rock sequence, with localized oxide-rich layers, that shows similarities to both BL_T and underlying metasomatized Giants Range granitic rocks. This hybrid sequence is marked by an abrupt increase in P and erratic Sr, Ba, Mg, and V concentrations. Iron ranges from 2% Fe to upwards of 45% Fe (largely because of assimilated BIF inclusions) 4. Mineralized GRB / GRB_M: consists locally of mineralized Giants Range granitic rocks as well as locally mineralized Virginia Formation and BIF. Average grade is about 0.28% Cu and 0.16% Ni with a Cu:Ni ratio of about 2.3. Local massive sulfide bodies are present and contribute significantly to the average grade The thickness of the four units is quite variable, but the stratigraphic succession does not vary across the deposit. Any one or more of the units, however, can be missing locally from a specific drill hole. Geologic modeling indicated that there is a sinuous, channel-like body of persistent and higher Cu grades that traverse the length of the deposit and follows the thickest portion of the BL_MT unit as shown in Figure 2-27. The origin of the channel is not well understood but it may be related to a magma conduit. Spruce Road Deposit (Twin Metals Minerals) The Spruce Road deposit lies to the northeast of Maturi (Figures 2-2 and 2-23). Mineralization is also present at the base of the SKI. It was at this deposit that the first good indications of Cu-Ni mineralization were uncovered while constructing a forest access road in 1948. From 1954 to 1971, Inco drilled the deposit on 200-foot centers (61 meters) for a total of 232 holes (the vast majority of which are no longer preserved after they were destroyed in a fire at Sudbury, Ontario). In 1997, Inco’s subsidiary, American Copper and Figure 2-147. Birch Lake magma channel superimposed on average copper grade base map. 41 �Trip 2 – Cu-Ni Duluth Complex Nickel Company (ACNC), joint ventured the property with Wallbridge Mineral Company Limited (from which Duluth Metals was eventually created). Wallbridge eventually drilled two holes on the property during 1999 to 2000 in search of high-grade footwall veins but failed to find significant mineralization in the footwall rocks. In 2002, Franconia Minerals Corp. entered into an agreement with Beaver Bay Joint Venture to acquire the Spruce Road and Maturi properties from ACNC but conducted no work. TMM acquired the property in 2011 and drilled 57 drill holes totaling 65,635.5 ft between September 2012 and January 2014 and a prefeasibility study technical report was issued on the TMM project in August 2014. The geology at Spruce Road is vastly different than at either Maturi or Birch Lake. Publiclypreserved core from historic drill holes are extremely limited for this deposit in that only six Inco holes are preserved along with two Wallbridge holes. From this limited data, Severson (1994) determined that most of the igneous units that typify the SKI elsewhere are not present at Spruce Road. Rather, the mineralization appears to be present in a much thicker BH unit (also referred to as the BMZ unit by TMM) consisting of a heterogeneous mix of troctolitic rocks with common hornfelsed inclusions of basalt (North Shore Volcanic Group). Also present are extremely localized noritic rocks associated with hornfelsed sedimentary rocks (Virginia Formation and Biwabik Iron Formation) and at the basal contact with the Giants Range granitic complex. The U3 unit and massive oxide zones are also locally present. Mineralization does not appear to correlate with any specific igneous lithology and there are no known marker horizons. Historic drill logs indicate that there may be an igneous mega-breccia unit that is referred to as “Spruce Road breccia.” South Filson Creek Deposit (Encampment Minerals) The South Filson Creek deposit, located to the east of Spruce Road, as previously presented in Figure 2-2, was initially drilled by the Hanna Mining Company (23 holes) in the late 1960s. There, the CuNi mineralization is hosted by troctolitic rocks (AT&T unit of Severson, 1994), both in outcrop and in the tops of several drill holes situated well above the basal contact. In 1987, encouraging high PGE values (>1.0 ppm) were reported in these “cloud” zone sulfides by Steve Hauck of the NRRI. A subsequent study of the PGE mineralization (Kuhns and others, 1990) indicated that the PGE were concentrated by a latestage hydrothermal event that concentrated the PGE in extremely fine, discontinuous, microscopic veinlets that were inferred to be associated with a NE-trending fault zone. Encampment Minerals drilled an additional 27 holes on the property, but results are largely unknown. Serpentine Deposit (Encampment Minerals) The Serpentine deposit, shown in Figure 2-26, is located to the north of the Mesaba deposit. The deposit was initially discovered by Bear Creek Mining Company in 1967 as part of a follow-up drilling campaign of an airborne electromagnetic conductor (Kulas, 1979). The name “Serpentine” was chosen for this deposit due to the presence of a sinuous-trending massive sulfide located at the base of the SKI. When the next owner, Amax, began working on the deposit, they calculated that the deposit contained 250 million tons of resources (not NI 43-101 compliant) grading 0.41% Cu, 0.14% Ni and 1.96% S at a 0.20% copper cut-off, with a higher-grade portion of over 7 million tons, at a 0.60% copper cut-off, with a grade of 0.88% Cu, 0.30% Ni and 5.67% S (Kulas, 1979, Zanko and others, 1994). The presence of such voluminous pyrrhotite-rich massive to semi-massive sulfide at the basal contact at Serpentine makes this an unusual deposit (Figure 2-29). There, the massive sulfide is closely related to the BDPO which provided a local sulfur source. Empirical evidence in drill core is evidenced by partially melted BDPO (present in the footwall and in inclusions) that transitions upwards and downwards into massive sulfide near the basal contact. The massive sulfide is also located close to the projected location of the Grano Fault that may have played a role in its origin. Another feature that may be related to the Grano Fault at Serpentine is a northerly-trending zone wherein subvertical olivine-rich ultramafic dikes were emp42 �Trip 2 – Cu-Ni Duluth Complex -laced in the troctolitic host rocks while the host rocks were still solidifying. Encampment Minerals drilled eight holes at the Serpentine deposit, but no data are known regarding their results. Figure 2-168. Location of the Serpentine deposit in relation to the Mesaba deposit. Note drill holes posted in red are holes that intersected massive sulfides at or slightly above the basal contact (larger red dots intersected significantly more and thicker massive sulfide zones). Figure 2-159. Trend of BDPO unit relative to basal massive sulfide mineralization at the Serpentine deposit (from Zanko and others, 1994). 43 �Trip 2 – Cu-Ni Duluth Complex Field Trip Stops Drill core will be displayed at NewRange’s Babbitt core facility, for both the NorthMet and Mesaba deposits, and at Twin Metals Ely office for the Maturi deposit. Cross-sections displaying the geology will be posted, as well as Cu-Ni-PGE grades, for the appropriate holes. At this point in time, the core displayed will be determined by the companies. NewRange core facility: 578810 / 5284970, (47.71330°, -91.94930°) Twin Metals Ely office: 585230 / 5306550, (47.90661°, -91.85949°) References Barber, J., Parker, H., Frost, D., Hartley, J., White, T., Martin, C., Sterrett, R., Poeck, J., Eggleston, T., Gormely, L., Allard, S., Annavarapu, S., Radue, T.,Malgensini, M and Pierce, M., 2014:, Twin Metals Minnesota Project, Ely, Minnesota, USA: NI 43-101 Technical Report on Pre-feasibility Study prepared by AMEC E&C Services, Inc. for Duluth Metals Limited, October 2014, Project 176916. Bennett, A., Dempers, N., Neff, D., Radue, P.E., Roth, D., Schwering, R., Tahija, L., Uble, J.S. and Welhener, H.E., 2022, NorthMet Copper-Nickel Project, Feasibility Update National Instrument 43-101F1 prepared for PolyMet Minerals Corp. by M3 Engineering and Technology Corp., Project M3-PN220283, Dec. 30th, 2022, 248 p. Eckstrand, 0.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 Types, District Metallogeny, The Evolution of Geological Provinces, and Exploration Methods: Geological Association of Canada, Mineral Deposits Division, Special Publication No.5, p. 205-222. Farrow, D, and Johnson, M, 2011, January 2012 National Instrument 43-101 Technical Report on the Titac Ilmenite Exploration Project, Minnesota, USA. SRK Consulting (Canada) Inc. SRK Project Number 2CC031.004. Cardero Resources Corp. Farrow, D, and Johnson, M, 2012, January 2012 National Instrument 43-101 Technical Report on the Longnose Ilmenite Exploration Project, Minnesota, USA. SRK Consulting (Canada) Inc. SRK Project Number 2CC031.004. Cardero Resources Corp. Foose, M., 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 Geerts, S.D., 1991, Geology, stratigraphy, and mineralization of the Dunka Road Cu-Ni prospect, northeastern Minnesota: Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN, Technical Report NRRI/TR-91/14, 63 p. Geerts, S.D., 1994, Petrography and geochemistry of a platinum group element-bearing horizon in the Dunka Road prospect, (Keweenawan) Duluth Complex, northeastern Minnesota: University of Minnesota Duluth, Unpublished M.S. thesis, 100 p. Kuhns, M.P, Hauck, S.A, and Barnes, R.J, 1990, Origin and occurrence of platinum group elements, gold and silver, in the South Filson Creek copper-nickel deposit, Lake County, Minnesota: Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN, Technical Report NRRI/GMIN-TR-89-15, 60 p. Kulas, J.E., 1979, Serpentine Reserve – Minnamax project: unpublished AMAX Company report on file at the Minnesota Department of Natural Resources, Lands and Minerals Division, Hibbing, MN, 5 p. 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, Div. of Minerals, Hibbing, MN, Report 93, 49 p. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., and Peterson, D.M., 2001, Geologic Map of the Duluth Complex and Related Rocks: Minnesota Geological Survey, Miscellaneous Map Series, Map M-119. 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 northern Minnesota: Minnesota Geological Survey, Report of Investigations RI-58, 207 p. Miller, J.D., and Severson, M.J., 2005, Bedrock Geology of the Babbitt Southwest Quadrangle, St. Louis County. Minnesota: Minnesota Geological Survey, University of Minnesota, Miscellaneous Map Series, M-161. 44 �Trip 2 – Cu-Ni Duluth Complex Patelke, R, Peterson, D, Severson, M, Jefferson, T, and Lehmann, E., 2009, Cu-Ni-PGE Deposits of the Duluth Complex, Geology and Development; 55th Annual Institute on Lake Superior Geology, Ely, MN, Part 2 Field Trip Guidebook, p. 1-80. Peterson, D.M., 2001, 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., 2010, The Nokomis Cu-Ni-PGE Deposit, Minnesota. Prospectors and Developers Association of Canada. Annual Meeting, Powerpoint presentation. 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, Duluth, MN, Technical Report NRRI/TR-93/34, 210 p. 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, Duluth, MN, Technical Report NRRI/GMIN-TR-99-11, 236 p. Severson, M.J. and Hauck, S.A., 1997, Igneous stratigraphy and mineralization in the basal portion of the Partridge River intrusion, Duluth Complex, Allen Quadrangle, Minnesota: Natural Resources Research Institute, University of Minnesota Duluth, Technical Report NRRI/TR-97/19, 102 p. Severson, M.J. and Hauck, S.A., 2003, Platinum group elements (PGEs) and platinum group minerals (PGMs in the Duluth Complex, Natural Resources Research Institute, University of Minnesota Duluth, 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): Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN Technical Report NRRI/TR-2008/17, 68 p. + 94 plates. Severson, M.J. and Heine, J.J., 2007, Data compilation of United States Steel Corporation (USSC) exploration records in Minnesota; Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN, Technical Report NRRI/TR-2007/25, 98 p. Severson, M.J. and Miller, J.D., 2005, Bedrock Geology of the Babbitt Quadrangle, St. Louis County. Minnesota: Minnesota Geological Survey, University of Minnesota, Miscellaneous Map Series, M=159. Severson, M.J., Patelke, R.L., Hauck, S.A., and Zanko, L.M., 1994, The Babbitt copper-nickel deposit, Part B: Structural datums: Natural Resources Research Institute, University of Minnesota Duluth, Duluth, MN, Technical Report NRRI/TR-94/21b, 48 p. Severson, M, Ware, A, Boerst, K, and Peterson, D, 2016, Cu-Ni-PGE Deposits of the Duluth Complex, Geology and Development; 62nd Annual Institute on Lake Superior Geology, Duluth, MN, Part 2 Field Trip Guidebook, p, 27-78. Welhener, H. and Crowie, S.T., 2022, NI 43-101F1 Technical Report on Mesaba Project, Mineral Resource Statement prepared for PolyMet Mining Corp. by Independent Mining Consultants, Inc. and JDS Energy and Mining, Inc., November 2022 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, Duluth, MN, Technical Report NRRI/GMIN-TR93-52, 90 p. 45 �Trip 3 – Proterozoic Fe & Mn Formations FIELD TRIP 3 How Do You Make Iron and/or Manganese Ores in Proterozoic Iron Formation? Dean Peterson1, Alex Steiner1, and Latisha Brengman2 1 Big Rock Exploration, 2505 W. Superior St., Duluth, MN 55806 Earth and Environmental Sciences, Swenson College of Science and Engineering, University of Minnesota, Duluth, 1114 Kirby Dr., Heller Hall 229, Duluth, MN 55812 2 Introduction Iron formations are among the most important rocks for our modern industrial world. Their extraordinary iron content facilitates the manufacture of steel, while their manganese content is of crucial importance as a steel-alloy product and a critical component of battery technologies. Fueling modern technology requires efficient production of iron resources, exploration of manganese resources, and determination of enrichment processes that lead to ore formation. This field trip will explore ore-formation processes that turn otherwise uneconomic iron formations into valuable resources of iron and manganese. The trip may include an optional stop at the Hibbing Core Library where participants will examine drill core of the Biwabik taconite ores. Participants will explore sedimentary features, diagenetic reactions, and weathering reactions that contribute to iron grade and iron distribution within ore-horizons of the Biwabik. We will then travel to the North Star Manganese/Electric Metals core logging facility in Emily, MN to look at four recent (2023 drilling) drillholes where we will discuss the formation and subsequent redistribution of manganese within the Emily Iron Formation. Participants will have the opportunity to observe highgrade manganese oxide drill core, primary iron-manganese carbonate facies iron formation, and breccia horizons possibly associated with the 1.85 Ga. Sudbury impact. The trip will then proceed to the Mary Ellen mine, a former natural ore pit, where the oxidation and weathering of the Biwabik was central to ore formation and early mining efforts on the range. Participants can observe primary features such as stromatolites and sedimentary structures as well as oxidation-weathering features. If time allows, we will wrap up the trip at the Biwabik outcrops in Virginia near the new Highway 53 bridge over the historic Rouchleau natural ore (hematite) mine before heading back to Mountain Iron. Regional Geologic Setting To gain a true understanding of the geology and origin of the high-grade Paleoproterozoic iron and manganese resources of the Mesabi and Cuyuna Ranges of northern Minnesota (Fig. 3-1), it is best to start with an understanding of the regional-scale geologic setting and its contained ferrous mineral resources. These Paleoproterozoic iron ranges include several categories of marine chemocline mineral systems outlined in recent USGS publications (Schulz et al., 2017 and Hofstra and Kreiner, 2020). These categories include: 1) Superior-iron deposits (Mesabi Iron Range and the Emily District of the Cuyuna Iron Range) and 2) Algoma-type iron-manganese deposits (Cuyuna North and South Iron Ranges). Superior Type Iron Resources of the Mesabi Iron Range Superior type iron formation resources of Minnesota are exemplified by the long-standing mining of iron resources of the Biwabik Iron Formation along the length of the Mesabi Iron Range. The Mesabi Iron Range is largely located in St. Louis and Itasca counties and has been the most important iron ore district in the United States since ~1900. The Mesabi Iron Range is 120 miles long, averages one to two miles wide, and is comprised of rocks of the Paleoproterozoic Animikie Group. The Animikie Group on 46 �Trip 3 – Proterozoic Fe & Mn Formations the Mesabi Iron Range consists of three major conformable formations: Pokegama Formation at the base; Biwabik Iron Formation in the middle; and the overlying Virginia Formation. On the Mesabi Iron Range, these three formations generally dip gently to the southeast at angles of 3-15 degrees. Figure 3-1. Location map of identified ferrous mineral resources in Minnesota. Since the early 20th century, the Biwabik Iron Formation has been subdivided into four informal members referred to as (from bottom to top): Lower Cherty member, Lower Slaty member, Upper Cherty member, and Upper Slaty member (Wolff, 1917). The cherty members are typically characterized by a granular (sand-sized) texture and thick-bedding (beds ≥ several inches thick); whereas the slaty members are typically fine-grained (mud-sized) and thin-bedded (≤1 cm thick beds). The cherty members are largely composed of chert and iron oxides (with zones rich in iron silicate minerals), while the slaty members are composed of iron silicates and iron carbonates with local chert beds. Both cherty and slaty iron-formation 47 �Trip 3 – Proterozoic Fe & Mn Formations types are interlayered at all scales, but one rock type or the other predominates in each of the four informal members, and they are so-named for this dominance Severson et. al. (2009). Leached and iron enriched direct ores (or natural ores) were the first materials mined, with the first shipments beginning in 1892, from strongly oxidized pockets along fault and fracture zones and the blanket oxidation of the iron formation at the surface. Taconite, which is the material that is mined today using magnetic separation methods, constitutes most of the iron formation and pertains to the hard, non-oxidized portions of the iron-formation. Production has been dominantly controlled by vertically integrated steelmakers since 1901, and therefore the mining and utilization of these ores have been dictated largely by US ironmaking capacity and demand. Taconite typically contains 30-35% iron and 40-50% SiO2, plus other components (Morey, 1992). The Biwabik Iron Formation is around 175-300 feet thick in the extreme eastern end of the Mesabi Iron Range at Dunka Pit, 730-780 feet thick in the central Mesabi Iron Range/Virginia Horn area near Eveleth, around 500 feet thick in the western Mesabi Iron Range near Coleraine, and eventually exhibits a “nebulous ending about 15 miles southwest of Grand Rapids” (Marsden et al., 1968) on the extreme western end of the Mesabi Iron Range. Maps of currently active taconite mining operations on the Mesabi Iron Range are presented in Figure 3-2 and compiled grade/tonnage ore reserve calculations for these operations are given in Table 3-1. Table 3-1. Reported grade/tonnage of active taconite mines. operations. Geology of the Cuyuna Iron Range The Cuyuna iron range is about 160 km west-southwest of Duluth in Aitkin, Cass, Crow Wing, and Morrison Counties (Fig. 3-1). It is part of an Early Proterozoic geologic terrane which occupies much of east-central Minnesota. The Cuyuna iron range is traditionally divided into three districts, the Emily district, the North range, and the South range (Fig. 3-3). The Emily district extends from the Mississippi River northward through Crow Wing County and into southern Cass County and comprises an area of about 1,165 square kilometers. Although exploration drilling has been extensive in the Emily district, mining never commenced. The North range, a much smaller area about 19 km long and 8 km wide, is near the cities of Crosby and Ironton in Crow Wing County. 48 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-2. Bedrock geology and iron mining features of the Mesabi Iron Range. 49 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-3. Bedrock geologic map of the Cuyuna Iron Range of Minnesota, illustrating the locations of the Emily District, North Range, and South Range. Since their discovery in 1904, it has been recognized that the iron-formations and associated ore deposits of the Cuyuna iron range in east-central Minnesota contained appreciable quantities of manganese, and large quantities of manganese were extracted as ferromanganese ores from several mines on the North range from 1911 to 1984. The presence of this manganese resource sets the Cuyuna range apart from other iron-mining districts of the Lake Superior region. Although relatively small, the North range was the principal site of mining activity (Fig. 3-4), which had largely ceased by 1970. The South range, where one small open pit and only a few underground mines were operated, in the 1910s and 20s, comprises an area of northeast-trending, generally parallel belts of iron-formation extending from near Randall in Morrison County northeast for about 100 km. In addition to the three named districts, numerous linear magnetic anomalies occur east of the range proper, and may indicate other, but currently poorly defined, beds of iron-formation. Three major insights regarding the geology of the Cuyuna range have emerged from the geologic mapping (Schmidt, 1963) and associated studies which utilized geophysical and drilling data (Southwick et al., 1988). First, there is clear evidence that iron sedimentation occurred at several different times and under varying geological conditions. This observation invalidates the stratigraphic premises of Morey (1978). Major iron-formations are associated stratigraphically with volcanic rocks in the South range, with black shale, argillite and rare volcanic rocks in the North range, and with shallow-water deposits of sandstone and siltstone in the Emily district. 50 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-4. Bedrock geology and open pit Fe-Mn mine map of the North Range of the Cuyuna Iron Range. Second, the iron-rich strata of the Emily district are correlative with the Biwabik Iron Formation of the Mesabi Range, as inferred by Marsden (1972) and Morey (1978). However, they and the other sedimentary rocks of the well-known Animikie Group occur above a major deformed unconformity that cuts across previously deformed, somewhat older sedimentary and volcanic rocks of the North range. There, a prominent iron-rich unit named the Trommald Formation, as well as several other units beneath the unconformity, forms part of a locally twice-deformed sequence. Therefore, the rocks of the North range and the Emily district cannot be correlative but are separate stratigraphic entities. Because the stratigraphic succession of folded sedimentary rocks on the North range comprises a distinct stratigraphic entity, Southwick et al., (1988) referred to it informally as the North Range group with the understanding that a formal name may be justified later. As defined by Schmidt (1963), the stratigraphic sequence in the North range consists of a quartz-rich lower sedimentary unit named the Mahnomen Formation, a middle iron- and locally manganese-rich sequence assigned to the Trommald Formation, and an upper greywacke shale interval called the Rabbit Lake Formation. Third, Southwick et al., (1988) recognized several geophysically defined structural discontinuities in the southern part of the Cuyuna iron range, within and southeast of the South range. These discontinuities are marked by demonstrable contrasts in metamorphic grade, by differing structural styles, and by different lithic components. One of the most pronounced of these, the Serpent Lake structural discontinuity, passes along the south edge of the North range. This discontinuity is interpreted as a tectonic boundary, probably involving major thrust faults between slices of folded rocks. Thus, it seems certain that the iron-rich strata of the South range are not correlative with either the Trommald Formation of the North range or the ironrich strata of the Emily district. The fact that iron-formation occurs within three different stratigraphic and structural contexts in the Cuyuna iron range is of considerable importance to the ultimate development of manganese resources. Since we now recognize that the Emily district, the North range, and the South range are separate entities, we can no longer develop regional syntheses that extrapolate mineralogical and structural attributes from one entity to another. 51 �Trip 3 – Proterozoic Fe & Mn Formations Cuyuna Iron Range Manganese Resources Several attempts have been made over the last 70 years to estimate the size of the manganese resources of the Cuyuna iron range. For example, Lewis (1951) estimated that 455 million metric tons of manganiferous iron-formation containing from 2 to 10 percent manganese were available to open-pit mining to a depth of 45 meters. Dorr et al., (1973) used that estimate to establish that the Cuyuna range contains approximately 46 percent of known manganese resources in the United States. US Steel geologist Richard Strong (1959) estimated iron and manganese resources from several well-drilled deposits in the Emily District and Beltrame et al., (1981) estimated a minimum of 170 million metric tons of manganiferous rock with an average grade of 10.46 weight percent manganese. All historic grade/tonnage estimates (Lewis, 1951, Strong, 1959, and Beltrame et al., 1981) should be considered with a certain amount of skepticism for at least two reasons. First, the manganese data used to make these estimates were, for the most part, by-products of data that were acquired originally by various mining companies as they explored for iron. Second, the various estimates were prepared for different reasons at different times, using different databases and different methodologies. Therefore, the results of these estimates are neither comparable, nor do they necessarily reflect the actual resource. A table listing the grade and tonnage from properties that Strong (1959) and Beltrami et al. (1981) estimated manganese resources is given in Table 3-2 and a location map of these properties is presented in Figure 3-5. Table 3-2. Manganese grade and tonnage estimates from reports by Strong (1959) and Beltrame et al. (1981). 52 �Trip 3 – Proterozoic Fe & Mn Formations Despite their problematic nature, the estimates of Lewis (1951) and Beltrame et al., (1981) do show that the Cuyuna range contains a large, but low- to moderate-grade manganese resources remaining. This large size, combined with the fact that the manganese deposits are in an established mining district, makes the Cuyuna range an ideal place to study geological and technological factors needed to evaluate this and other sedimentary manganese deposits in the United States. Especially important are studies of the geologic habit of the manganese and the controls on its distribution and subsequent concentration into deposits of minable size. Figure 3-5. Bedrock geology and location map of properties outlined in reports by Strong (1959) and Beltrame et al. (1981) that includes manganese grade-tonnage estimates. Labeled parcels correlate with the MAP ID column in Table 3-2 and are those with an estimated resource greater than 100,000,000 pounds of manganese metal. 53 �Trip 3 – Proterozoic Fe & Mn Formations Penokean Orogeny The Penokean orogeny began at about 1880 Ma when an oceanic arc, the Paleoproterozoic Pembine–Wausau terrane, collided with the southern margin of the Archean Superior (Laurentia) craton marking the end of a period of south-directed subduction. The docking of the buoyant craton to the arc resulted in a subduction jump to the south and development of back-arc extension both in the initial arc and adjacent craton margin to the north. Synchronous extension and subsidence of the Laurentia craton resulted in the development of broad shallow seas overlapping the Archean craton. The classic Superior-type banded iron-formations of the Lake Superior District, including those in the Marquette, Gogebic, Mesabi, and Gunflint Iron Ranges, formed in that sea. The newly established subduction zone caused continued arc volcanism until about 1850 Ma when a fragment of Archean crust, now the basement of the Marshfield terrane, arrived at the subduction zone. The convergence of Archean blocks of the Superior and Marshfield cratons resulted in the major contractional phase of the Penokean orogeny. Rocks of the Pembine–Wausau arc were thrust northward onto the Superior craton causing subsidence of a foreland basin in which sedimentation began at about 1850 Ma in the south (Baraga Group rocks) and 1835 Ma in the north (Rove Formation). A thick succession of arc-derived turbidites constitutes most of the foreland basin-fill along with lesser volcanic rocks. In the southern fold and thrust belt, tectonic thickening resulted in high-grade metamorphism of the sediments by 1830 Ma. At this same time, a suite of post-tectonic plutons intruded the deformed sedimentary sequence and accreted arc terranes marking the end of the Penokean orogeny. A regional geologic map of the Penokean orogen, modified from Schulz and Cannon (2007), is given in Figure 3-5. Figure 3-5. Generalized geologic map of the Penokean orogen. Abbreviations: ECMB - East-central Minnesota batholith; EPSZ - Eau Pleine shear zone; MD - Malmo discontinuity; NFZ - Niagara fault zone. Modified from Schulz and Cannon, 2007. The Penokean deformation in Minnesota includes a southern intensely and complexly deformed series of thrust panels (Cuyuna North, Cuyuna South, Moose Lake, McGrath-Little Falls panels) that gives way northward to progressively more weakly and simply deformed rocks (Emily District) across a belt 54 �Trip 3 – Proterozoic Fe & Mn Formations about 100 km wide. Farther north strata in the Mesabi and Gunflint Iron Ranges are essentially undeformed (Holst, 1991). It should be noted that the “more weakly and simply deformed rocks” of the Emily District have been shortened ~250% into a series of shallowly east-plunging anticlines and synclines. Substantial progress has been made in deciphering the structure of the poorly exposed rocks of the Minnesota foreland through the use of aeromagnetic and gravity data and drillhole information. Southwick and Morey (1991) and Southwick et al. (1988) have presented syntheses of this information. The complex thrust panels on the south, like comparable structures in Michigan, appear to be thinskinned slices without Archean basement. However, as in Michigan, this area of thin-skinned thrusting is also the area where Archean-cored gneiss domes developed during post orogenic collapse of the Penokean orogen (Holm and Lux, 1996; Schneider et al., 2004). Farther north, basement-cover relations are not well known except for the Mesabi Range where Paleoproterozoic strata are mostly nearly flat lying above an undisturbed unconformity with Archean basement rocks. A schematic north-south geologic cross section of the Penokean orogeny in Minnesota, modified from Southwick and Morey (1991) is presented in Figure 3-6. Figure 3-6. Schematic diagram illustrating the interpreted tectonic setting of the Penokean orogen in Minnesota. A) continental margin sedimentation, and B) thin-skinned thrusting and deformation related to the Penokean orogeny. Modified from Southwick & Morey, 1991. Post Penokean Weathering and Erosion Perhaps the most important component in the formation of the high-grade iron and manganese ores on the Mesabi and Cuyuna ranges is the vast amount of time (measured in hundreds of millions of years) upon which the newly-formed and uplifted Penokean mountains of the southern Laurentia craton weathered and eroded. As plate tectonic forces moved Laurentia across the globe to its current position on planet Earth, there were long periods of time when it resided within the tropical weathering zone (+30° to -30° latitude) near the Earth’s equator. It is believed that the supergene enrichment of iron (to >60 wt.% elemental Fe) and manganese (to >50 wt.% elemental Mn) on the Mesabi and Cuyuna largely formed during the protracted periods of time that the area resided within the tropical weathering zone. A paleogeographic reconstruction of the location of Laurentia on planet Earth is given in Figure 3-7. 55 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-7. Paleogeographic reconstruction of the Laurentia craton from the Paleoproterozoic to present times. FIELD TRIP STOPS Four field trip stops have been selected to showcase selected geological features associated with supergene weathering of primary Paleoproterozoic Superior-type iron formation into high-grade of Fe and Mn ores. The locations of the field trip stops are shown in Figure 3-8 and briefly described below. 1) DNR drillcore library in Hibbing: Drillcore review of unoxidized Biwabik Iron Formation, 2) Emily deposit core shed: Drillcore review of high-grade supergene Mn ores, primary Mn-Fe carbonate-facies iron formation, Overlying Sudbury Impact breccias & accretionary lapilli? 3) Mary Ellen Mine: Walk into a historic natural-ore (hematite) open pit iron mine, sampling of the classic Mary Ellen stromatolites, and 4) Large roadcut of the partially oxidized Biwabik Iron Formation adjacent to the historic naturalore Rouchleau Mine Complex. Stop 1: DNR Drillcore Library, Hibbing Minnesota Longitude/Latitude: 47.432412°N, -92.941811E UTM NAD 83 Zone 15N: 504388E, 5253220N At our first stop on this field trip, we will examine sections of two different drill cores of the Biwabik iron formation to directly compare depositional features to post-depositional features. The Drill Core Library is maintained by the Minnesota Department of Natural Resources, Lands and Minerals Division, and provides direct access for visitors to examine publicly owned geologic materials and exploration data. This incredible 56 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-8. Location map of field trip stops and the collar location of the two drill holes looked at Stop 1. repository contains over 7,000 mineral exploration cores, 1,500 roadway and bridge foundation cores, and 500 cores collected during scientific, governmental, and academic research, and their curation activities support researchers, exploration geologists, and engineers from around the world. Depositional features and mineralogy in Precambrian chemical sedimentary rocks like iron formations have long been of interest to the scientific community as they may record information about Earth’s early surface conditions. However, recovering data that links to depositional conditions requires the reconstruction of post-depositional mineral reactions and quantification of geochemical exchange. For this reason, the original mineralogy and geochemistry of iron formations has been the subject of numerous investigations. Critical early investigations recognized the mineral greenalite - postulating its authigenic origin as a “chemical oceanic precipitate”, and hypothesizing its role in forming the iron ore deposits in Minnesota, USA (Irving, 1886; Irving and Van Hise, 1892; Leith, 1903; Van Hise and Leith, 1911; Aldrich 1929; Gruner, 1946; Tyler, 1949; James, 1954; White, 1954; Goodwin, 1956; Gundarson and Schwartz, 1962; LaBerge; 1964). Key initial evidence for a primary or early origin for greenalite in the Superior craton included: (1) its abundance in low temperature, well-preserved assemblages, and comparative absence in metamorphosed iron formations in the same region (LaBerge, 1964); (2) the existence of submicroscopic 57 �Trip 3 – Proterozoic Fe & Mn Formations greenalite in the cores of circular to elliptical sand-sized grains as either the main mineral phase, <0.05 mm spherules, or as dusty nano-scale (submicroscopic) particles (Goodwin, 1956); intergranular relationships between phases where greenalite is crosscut by other phases in the same sample; (LaBerge, 1964; French, 1973; Klein and Fink, 1976), and even distribution throughout primary bedding (LaBerge, 1964). Recent mineralogical investigations (e.g, Duncanson et al., 2024, Muhling et al., 2025 and references therein) highlight that the earliest forming minerals in iron formation are commonly found within silica-cemented horizons, where abundant chert cement silicified the sediments at or near the sediment water interface. Such silica-cemented horizons preserve incomplete reactions and allow for identification of direct mineral relationships and local element exchange. Common mineral reactions observed in the Biwabik iron formation include the transformation of greenalite to minnesotaite, minnesotatite to stilpnomelane, greenalite to magnetite, siderite to magnetite, and magnetite to hematite. Of the mineral reactions that commonly occur in iron formation, transformations of Fe2+-containing silicates like greenalite to mixed valence state minerals like magnetite and further oxidation of magnetite to hematite, contributed to formation of iron ores in the Biwabik. To illustrate some of the many post-depositional reactions that occur in iron formations worldwide, we will examine a small section of two drill cores, MGS 8 from the western end of the Mesabi iron range, and LWD-99-01 from near the Virigina horn area (see Figure 3-8 for locations). LWD-99-01 clearly preserves depositional features, while MGS-8 documents abundant post-depositional oxidation throughout. Stop 2: North Star Manganese Inc Drillcore Shed, Emily Minnesota Longitude/Latitude: 46.753571°N, -93.973496E UTM NAD 83 Zone 15N: 425650E, 5178240N Historic exploration and drilling in the 1940’s and 1950’s by Pickands Mather and US Steel identified iron and manganese-bearing mineralization within the Emily Iron Formation. US Steel developed but did not implement a preliminary mine plan for mining of the Emily Deposit. Following approximately 50 years of inactivity, Cooperative Mineral Resources (subsidiary of Crow Wing Power) pursued a pilot mining operation using pressurized water that was ultimately unsuccessful. As a follow up investigation into the outcomes of pilot mining, a small-scale drill program was accomplished in 2010-2012. A drilling program was designed and executed by Big Rock Exploration, LLC, in 2022-2023. A total of 29 drill holes were completed to extend mineralization and refine the previous resource estimates. A total of 13,107 feet of drilling was completed for this program. Data collected for this project includes lithological, structural, geotechnical, geochemical and geophysical data from the drill core. Through interpretation of legacy, recent and new drilling data, Big Rock Exploration identified coherent zones of high-grade manganese mineralization (30 to ≥40 wt.% Mn) over a 1.25-kilometer strike length. Mineralization is comprised of horizons of secondary manganese oxide minerals, as well as locally present primary iron-manganese carbonate mineralization. An ore deposit model has been developed that incorporates the oxidation of primary thin-bedded manganese-iron carbonates into massive manganese oxide through early folding and prolonged periods of weathering, oxidation, and erosion. This ore deposit model and associated geological model have been used to support an updated and expanded mineral resource estimate (Table 3-3) for the Emily Deposit that was completed and published by Forte Dynamics (Hulse et al., 2024) on May 24, 2024. We’ll first begin the review of important geological features revealed in the four Emily Mn deposit drillholes on display in the core shed by elucidating our current understanding of the geology and structure for the whole Cuyuna Iron Range and then focus specifically on the stratigraphy and ore-forming processes at Emily through a series of figures and descriptive text written into a technical report (Steiner et al., 2024) and orally presented at the 2024 ILSG conference (Peterson and Steiner, 2024). 58 �Trip 3 – Proterozoic Fe & Mn Formations Table 3-3. Mineral resource estimate for the Emily Manganese deposit. Figure 3-9. Continent-scale initial condition framework of Paleoproterozoic iron formations of Minnesota. Initial Conditions Deposition of Paleoproterozoic iron formations of the Lake Superior district all owe their origins to the ~2.4 – 2.1 Ga. rifting of the Wyoming Province craton off of the southern Superior Province craton (Figure 3-9). This rifting set the stage for the development of environments of deposition conducive to the formation of thick sequences of both Algoma- and Superior-type iron formations (Figure 3-10). During the Penokean orogeny (see Figure 3-6) these variable environments of iron formation deposition were transposed northwestward via thin-skinned tectonics into a fold & thrust belt (Cuyuna North and South range thrust panels) over a series of thrust-front folds (Emily District) that was bounded by a basal decollement. Outcomes of the Penokean orogeny in the Cuyuna Range of central Minnesota are shown in an idealized cross sectional view Figure 3-11 and how ~1.8 billion years of erosion has left it today in Figure 3-12. 59 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-10. Schematic model for the variable environments of deposition of Paleoproterozoic iron formations of the Cuyuna and Mesabi ranges of Minnesota. Figure 3-11. Idealized cross section of Minnesota’s Penokean Mountains of central Minnesota approximately 1.83 billion years ago. Figure 3-12. Schematic representation of the results of deep weathering and erosion of the Penokean Mountains in central Minnesota. 60 �Trip 3 – Proterozoic Fe & Mn Formations Geologic Maps Geologic maps are the foundation upon which geologists interpret Earth processes and depict on a piece of paper the final outcomes of such processes in plan form. As such, the authors have added annotated bedrock geology map of the Cuyuna Range in Figure 3-13 and more specifically for the Emily District of the Cuyuna Range in Figure 3-14. Figure 3-13. Annotated bedrock geologic map of a portion of the Cuyuna Range, central Minnesota. Clipped from the map of Peterson (2022). Figure 3-14. Annotated bedrock geologic map of the Emily District, Cuyuna Range, central Minnesota. Modified after Peterson, 2022. 61 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-15. Schematic stratigraphic section through the Emily Manganese deposit, after Peterson & Steiner, 2024. Stratigraphy Relogging of historic drill core and logging of new core drilled during the 2023 exploration drilling program has led to the identification of a predictable stratigraphic sequence at the Emily Deposit (Fig. 315). The four formations recognized at the Emily Deposit include three units of the Paleoproterozoic Animikie Basin unconformably overlain by Quaternary glacial drift. Descriptions of these stratigraphic units from the oldest to youngest are given below: 1. Pokegama Formation – White-grey, tan, or cream-colored indurated argillite and quartzite. The Pokegama formation at the Emily deposit is typically a clayey siltstone, though mudstones and quartz-arenites are common. Bedding is quite variable ranging from massive siltstones and quartzarenites to thick, medium, and thin bedded or even finely laminated clayey-siltstones and mudstones. The Pokegama formation does not host significant manganese mineralization but when observed manganese minerals occur in trace amounts in veinlets or as small patches with iron oxides. 2. Emily Iron Formation – See below for subdivision descriptions. 3. Virginia Formation – Grey-brown in color, red when oxidized, fine-grained, well bedded clastic sediments commonly forming turbidite sequences. Fine-scale bedding, graded beds, and sandy lenses are common. Thin horizons of lean iron formation (subunit Pvif) composed of ferruginous chert occur locally. The basal 20-40 feet is characterized by highly disrupted and fragmented turbidite clasts sed in a poorly sorted massive matrix. This horizon has been hypothesized to be landslides associated with the Sudbury Impact. 4. Glacial Overburden – Unconsolidated glacial material including well sorted sands, lacustrine clays, and unsorted glacial till. The preservation of earthy hematite and saprolitic materials immediately below the basal angular unconformity indicates that overlying Laurentide ice sheet was not eroding its base in the immediate deposit area. Composition of the overburden is inferred from drill returns during tri-cone drilling. Emily Iron Formation is further divided into five sub-units. Criteria for subunit designation requires that a given interval be sufficiently distinctive in petrologic character to be easily identified, and laterally extensive enough to be intercepted in multiple boreholes. Distinctive petrologic characteristics may be texture (e.g., banded or granular iron formation, composition (e.g., chert, carbonate), or unique 62 �Trip 3 – Proterozoic Fe & Mn Formations characteristics such as stromatolites. The subdivisions of the Emily Iron Formation are as follows from bottom to top: 1. Peif1 – This unit is located as the base of the Emily Iron Formation lies conformably atop the Pokegama quartzite. The base is commonly cherty or stromatolitic before giving way to grain stones. The majority of the unit is a red-brown to black, granular iron formation (GIF) or ferruginous quartzose sandstone. The upper part (above the Peif1r marker, see below) is granular iron formation composed of silicious, hematitic granules that range from <1 to 2mm in size. Granular iron formation is weakly bedded. The lower part of Peif1 (below the Peif1r marker horizon) contains abundant quartzose sands cemented by iron oxides giving the rock a purple appearance. Sands are composed of well-rounded fine-grained quartz and are interbedded with granular iron formation. Manganese oxide mineralization is most intense within the Peif1 unit. Manganese oxides occur in multiple styles including massive Mn-oxide that replaces all original textures (may be within a bed or cross bedding), interstitial to grains (replacing the original iron cement?), and as veinlets. The most intense manganese oxide mineralization within Peif1 occurs adjacent to (above and below) the Peif1r marker horizon, though mineralization may occur throughout the unit. The lower Peif1 commonly exhibits a pock-mark texture when strongly mineralized. A thin stromatolite horizon (Peifbs) typically occurs at the base of Peif1. 2. Peif1r – This unit is located within the Peif1 unit. Usually <3m thick, the Peif1r is composed of finegrained hematite-chert banded iron formation. The base of the Peif1r hosts distinctive digitate stromatolites. 3. Peif2 – This unit lies conformably atop Peif1, usually gradually transitioning from granular iron formation (Peif1) to banded iron formation (Peif2) over a meter. The Peif2 is characterized by finegrained well-bedded banded iron formation though the composition of the iron formation is variable. The most common composition for Peif2 is a hematite-chert banded iron formation, though this appears to be a secondary, altered composition. The primary composition is iron-manganese carbonate facies type iron formation. The carbonate facies are cream to greenish and gradually becomes red with increased oxidation. Fresh carbonate facies contain much more manganese than oxidized material, with the manganese found in rhodochrosite. This represents a very different manganese host than the oxide mineralization found in the granular iron formation units (Peif1 and Peif3). 4. Peif3 – Peif3 is characterized by interbedded medium to fine grained GIF and fine grained, narrow BIF lenses. This unit lies conformably atop Peif2 where the contact is a graduation from BIF to GIF dominant facies. Both GIF and BIF are weakly to moderately strongly bedded and silicious in composition. Peif3 is commonly mineralized manganese oxides, only subordinate in manganese endowment to Peif1. 5. Peif4 – White to grey, massive chert with mottled patches of iron-oxides. The sharp basal contact with the underlying Peif3 is often “sheared” possibly from bedding parallel slip. Some parts of the massive chert contain faint outlines of granules while most is simply massive white chert. The mottled iron oxides include large irregular pods and stringers, sometimes reaching a meter in width. Manganese is rarely found within the oxide pods. 6. Peif5 – Well-bedded banded iron formation consisting of 2-5cm beds of chert and iron oxide in gradational contact with the underlying Peif4 massive chert. This unit is the least spatially consistent, due to a lack of drillhole intercepts and difficulty identifying it as a result of intense oxidation. Supergene Enrichment of Manganese The unique manganese endowment of the Emily Iron Formation is attributed to the primary deposition of Mn-carbonates in a shallow water environment (Figures 3-10, 3-16 and 3-17). However, 63 �Trip 3 – Proterozoic Fe & Mn Formations subsequent weathering, erosion, and oxidation has redistributed much of the manganese from the carbonate unit to other areas. The majority of the manganese mineralized material at the Emily deposit is composed of manganese oxides including manganite, jacobsite, and cryptolomene/hollandite. However, these phases are not the stable manganese phase predicted by geochemical modelling of early oceans (Mitra et al., 2022). Instead, manganese carbonates are the predicted stable phase. Therefore, the manganese oxide minerals that constitute the majority of the orebody must have formed at a later stage. The tectonics during and immediately after deposition of the Animikie basin sediments provide a plausible explanation for the extremely high-grade manganese oxide formation. Figure 3-16. Carbonate facies iron formation where the gradual oxidation of primary carbonates can be observed from left to right. Note the increasingly hematite rich BIF from left to right. Primary silicate and carbonate minerals in iron formations are well documented to be unstable under oxidizing, near surface conditions. For example, the formation of direct ship ores of the Biwabik Iron Formation on the Mesabi Iron Range has been ascribed to deep weathering of primary Fe-minerals (e.g., greenalite and siderite) over hundreds of millions or a billion years. The direct ship ores were composed of hematite and goethite. The Emily Iron Formation, having been deposited contemporaneously with the Biwabik Iron Formation, would have endured at least as much weathering over that period. The weathering and subsequent supergene enrichment of manganese is related to the hydrogeologic and geochemical interaction between interbedded banded (BIF) and granular (GIF) iron formation. Manganese in the Emily Figure 3-17. Emily deposit drill core that seemingly documents that the oxidation of carbonate-facies (rhodochrosite-siderite-chert) BIF generates classic thin-bedded hematite-jasper BIF as well as being the primary source of Mn3+ that forms the massive manganese oxide zones in permeable GIF horizons. 64 �Trip 3 – Proterozoic Fe & Mn Formations Iron Formation was originally co-precipitated with iron carbonate minerals (rhodochrosite MnCO3 and siderite FeCO3) within the banded iron formation lithotype (Peif2 subunit). Primary carbonates are observed at various stages of oxidation in several boreholes (e.g., NSC-23005). Like the Biwabik Iron Formation, the Emily Iron Formation underwent a protracted period of weathering and oxidation. Exposure of carbonate facies iron formation to oxidizing waters over that period is hypothesized to be the causative mechanism for supergene manganese enrichment at the Emily deposit. Oxidized meteoric water percolating through the carbonate iron formation reacts with and dissolves the carbonates, liberating manganese from rhodochrosite and converting siderite to hematite. The restite lithology appears very similar to hematiterich banded iron formation (Fig. 3-17). The now manganese enriched waters redistribute manganese downslope to other subunits of the Emily Iron formation. The second important litho-type, granular iron formation, is recognized as the primary manganeseoxide ore hosts at the Emily Deposit. Granular iron formation is composed of granules of varying compositions (e.g., Fe-silicate, chert, Fe-carbonates) with pore space found between the granules. That pore space creates permeability that drives fluid flow through the granular iron formation thereby moving and redepositing manganese from the enriched waters leaving the carbonate facies banded iron formation. The observations from drillcore logging at Emily indicate that manganese oxides are not found in significant concentrations within the banded iron formations, but manganese oxides are abundant in the granular iron formation. The migration of manganese-rich waters from the banded iron formation into the granular units is the primary redistribution mechanism for manganese. Once manganese enriched waters enter the granular units, it is unclear by what mechanism the precipitation of manganese occurs. However, manganese oxide minerals are observed in the interstices between granules indicate direct precipitation from the pore fluids. It is unclear whether this interstitial manganese is the result of filling otherwise empty pore space if it is the result of replacement of prior GIF matrix. Additionally, pock marked textures in manganese-rich units suggest that granules may be replaced by the manganese-rich fluids, though the mechanism by which this may occur is unclear due to a lack of mineralogical constraints. Stratigraphic and Structural Controls on the Distribution of Secondary Manganese The compression associated with the Penokean Orogen uplifted the rocks Emily deposit. Folding and subsequent uplift of these originally shallow water subaqueous rocks into the Penokean mountains has important hydrogeological implications by greatly lowering the water table and exposing the Emily Iron Formation to oxidizing meteoric waters. The Emily deposit is on the northernmost anticline of the Penokean fold and thrust belt, specifically within a parasitic syncline along the norther limb of the larger anticline. The structural geometry established during the Penokean provides a hydrogeologic “slope” that meteoric waters can migrate down under the influence of gravity. In particular, the parasitic syncline that hosts the Emily Deposit (see Figure 3-14), likely acted like a funnel or gutter that focused fluid flow through the rocks that now constitute the deposit. A schematic stepwise ore genesis model for the Emily deposit is presented in Figure 3-18. On a deposit scale and within this structural setting, the two major litho-types play an important role in the movement of fluids due to their contrasting hydrogeological characteristics. In particular, the Peif1r stromatolite horizon represents an aquitard that seemingly focused fluid along its margins. The focusing of fluids along these margin manifests as exceptionally high Mn-grades (often massive manganese oxides) at the upper and lower contacts with flanks GIF of Peif1. Similarly, but to a lesser extent, the basal contact with the Pokegama formation and the contact between Peif2 and Peif3 represent areas of contrasting hydrogeologic characteristics that may concentrate mineralizing fluids. Both areas are observed to host massive manganese oxide mineralization, supporting such a relationship. 65 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-18. Emily deposit Mn-Oxide ore deposit model that incorporates stratigraphy, permeability, processes, and time, after Peterson & Steiner, 2024. Drillholes on Display Four drillholes from the 2023 exploration program at Emily will be on display for the 2025 ILSG field trip. These holes include: 1) NSC-23002A - high-grade Mn-oxide ores at the bedrock interface in Peif1, 2) NSC-23004 - an almost complete stratigraphic section through the Emily IF, 3) NSC-23005 - Peif2 with carbonate facies IF and Sudbury Impact breccias in unit Pvf, and 4) NSC-23050 – The Western-most drillhole, nearly complete section of the Emily IF. A detailed bedrock geology and drillhole location map of North Star Manganese Inc’s Emily Project is presented in Figure 3-19, and striplogs of the four holes on display are given in Figures 3-20, 3-21, 3-22, and 3-23. 66 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-19. Detailed bedrock geology and drillhole location map of North Star Manganese Inc’s Emily project. Note that the collar location of the four drillholes on display are highlighted by the small yellow circles. Modified after Steiner et al., 2024. 67 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-20. Striplog for drillhole NSC-23002A. Figure 3-21. Striplog for drillhole NSC-23004. 68 �Trip 3 – Proterozoic Fe & Mn Formations Figure 3-22. Striplog for drillhole NSC-23005. Figure 3-23. Striplog for drillhole NSC-23050. 69 �Trip 3 – Proterozoic Fe & Mn Formations Stop 3: Mary Ellen Mine, Biwabik Minnesota Longitude/Latitude: 47.527677°N, -92.366645E UTM NAD 83 Zone 15N: 547675E, 5264000N The historic natural ore Mary Ellen mine (Figure 3-24) near Biwabik is probably most well-known today as the source of Mary Ellen Jasper, a world-class type-locality of Precambrian stromatolites. Rocks and polished slabs of stromatolites from the Mary Ellen mine can be found in natural history museums throughout the world, and you’ll get to find and take-home pieces yourself during this field trip. According to several annual mining directories, the Stanley Iron Mining Company operated the Mary Ellen Mine between 1924 and 1928, with stockpile shipments occurring in 1929 and 1930. It actively mined the property again between 1948 and 1951. Beginning in 1952, the Pioneer Mining Company worked the mine, continuing to do so through 1961. The Pittsburgh Pacific Company operated it for one final year, in 1962. Its cumulative output of natural ore was 4,574,973 long tons, again according to an annual mining directory. An interesting quote from the 2015 book titled: Stromatolites Ancient, Beautiful and EarthAltering, by Bruce Stinchcomb & Bob Leis copied below strongly hints that some oxidation-related processes that formed the high-grade earthy hematitic iron ores were similar to those outlined for the highgrade manganese oxide ores at Emily. The quote is as follows, "In the early days of iron mining in Minnesota, the location of stromatolite material would indicate that an iron rich vein was close. The Mary Ellen Stromatolite material could be as much as 15 feet thick and would have to be removed before mining could commence. To the miners this material was considered a nuisance and a waste product." Figure 3-24. Simplified bedrock geology and iron mine map of the Mary Ellen mine area. 70 �Trip 3 – Proterozoic Fe & Mn Formations Stop 4: Rouchleau Mine Complex Bridge, Virginia Minnesota Longitude/Latitude: 47.516178 °N, -92.518663 E UTM NAD 83 Zone 15N: 536240 E, 5262640 N The Thomas Rukavina Memorial Bridge carries U.S. 53 over the Rouchleau Mine pit connecting Virginia, Minnesota with other cities to the south. U.S. 53 continues through Virginia to International Falls and Canada; International Falls is about 100 miles (160 km) north of Virginia. The bridge, opened in 2017, was named after Tom Rukavina in 2021 following his death in 2019. Rukavina was a state legislator from the Iron Range. At 204 feet tall, it is the tallest bridge in Minnesota. This bridge carries a traffic volume of about 22,200 cars per day, making it one of the most-traveled highway segments on the Iron Range. The bridge also features a bike lane and pedestrian walkway (the Mesabi Trail) that leads to trails connecting Gilbert and Virginia. In 1960, the state of Minnesota and the mining companies in the area came to an agreement that allowed the construction of U.S. 53 across lands held by the mining company without the state paying anything for the land. The agreement stipulated that after 1987, the state would be responsible for the costs involved with moving the roadway to allow for mining after given advance notice by the mining companies. The two owners of the land notified MnDOT of their intent to mine the site in 2010 which gave the state until 2017 to move the roadway. Cliffs Natural Resources, which had a nearby active mine, hoped to begin mining the site by 2017. After evaluating several more expensive options that involved longer bridges or routing US 53 across an active mine pit, an alignment was selected that resulted in the highest bridge in Minnesota. A route on level ground away from the mining formation was identified as too disruptive to development patterns in the area. The entire project estimated to cost $220 million with $159 million for construction of the bridge and diverted roadway. The bridge crosses the Rouchleau Mine pit.[9] The water filled pit also serves as Virginia's water supply. The final cost was $230 million with $30 million coming from the federal government and the remaining from the state. To prevent the need to move the bridge in the future, the state purchased the mineral rights for the land beneath roadway for $15 million. References Aftabi, A., Atapour, H., Mohseni, S., and Babaki, A., 2021, Geochemical discrimination among different types of banded iron formations (BIFs): A comparative review, Ore Geology Reviews, Volume 136. Aldrich, H. R., 1929, The geology of the Gogebic Iron Range of Wisconsin: Wisconsin Geol. Survey Bull. 21, 279 p. Beltrame, R.J., Holtzman, R.C., and Wahl, T.E., 1981, Manganese resources of the Cuyuna range, east-central Minnesota: Minnesota Geological Survey Report of Investigations 24, 22 p. Berg, T., Peterson, D.M., and Sweet, G., 2022, The Emily Manganese Deposit, Crow Wing County, Minnesota: A mineral resource evaluation for North Star Manganese Inc, Big Rock Exploration technical report BRE-TR2022-02, 33 pages, 3 appendices. Dorr, J.VN., II, Crittenden, M.D., Jr., and Worl, RG., 1973, Manganese, in Probst, D.A., and Pratt, W.P., eds., United States Mineral Resources: U.S. Geological Survey Professional Paper 820, p. 385-399. Duncanson S and 5 coauthors (2024) Reconstructing diagenetic mineral reactions from silicified horizons of the Paleoproterozoic Biwabik Iron Formation, Minnesota. American Mineralogist 109: 339-358, doi: 10.2138/am2022-8776. Goodwin, A.M., 1956, Facies relations in the Gunflint iron-formation: Ecox. GEOL., v. 51, p. 565-595. Gruner, 1946, Mineralogy and geology of the Mesabi range: Iron Range Resources and Rehabilitation, St. Paul, Minn., 127 p. Gunderson, J. N., and Schwartz, G. M., 1962, The geology of the metamorphosed Biwabik iron-formation, Eastern Mesabi District, Minnesota: Minnesota Geol. Survey Bull. 43, 139 p. Hofstra, A.H., and Kreiner, D.C., Systems-Deposits-Commodities-Critical Minerals Table for the Earth Mapping Resource Initiative: U.S. Geological Survey Open-File Report 2020-1042, 24 p. Holm, D.K., Lux, D.R., 1996, Core complex model proposed for gneiss dome development during collapse of the Paleoproterozoic Penokean orogen, Minnesota, Geology 24, 343–346. 71 �Trip 3 – Proterozoic Fe & Mn Formations Holst, T.B., 1991, The Penokean orogeny in Minnesota and Upper Michigan, U.S. Geological Survey Bulletin 1904D, 10 pages. Hulse, D.E., Irons, A., and Malhotra, D., 2024, Electric Metals (USA) Limited Emily Manganese Project, NI 43-101 Technical Report, Project No. 219001, Forte Dynamics, 89 pages. Irving, R. D., 1886, Origin of the ferruginous schists and iron ores of the Lake Superior region: Am. Jour. Sci., v. 32 p, 255-272. Irving and Van Hise, C. R., 1892, The Penokee iron-bearing series of Michigan and Wisconsin: U.S. Geol. Survey Mon. 19, 534 p. James, H. L., 1954. Sedimentary facies of iron-formation. Economic Geology, 49(3), 235–293. Klein, C., and Fink, R.P., 1976, Petrology of the Sokoman Iron Formation in the Howells River area, at the western edge of the Labrador trough: Economic Geology, v. 71, p. 453–487. LaBerge, G.L., 1964, Development of magnetite in iron-formations of the Lake Superior Region: Econ. Geol., V. 59, p. 1313-1342. Leith, C. K., 1903, The Mesabi iron-bearing district of Minnesota: U.S. Geol. Sur. Mono. 43, 316 p. Lewis, W.E., 1951, Relationship of the Cuyuna manganiferous resources to others in the United States, in Geology of the Cuyuna Range Mining Geology Symposium, 3rd, Hibbing, Minnesota, Proceedings: Minneapolis, University of Minnesota, Center for Continuation Study, p. 30-43. Marsden, R.W., 1972, Cuyuna district, in Sims, P.K., and Morey, G.B., eds., Geology of Minnesota: A centennial volume: Minnesota Geological Survey, p. 227-239. Marsden, R.W., Emanuelson, J.W., Owens, J.S., Walker, N.E., and Werner, R.F., 1968, The Mesabi Iron Range, Minnesota, in Ridge, J.D. (ed.), Ore Deposits of the United States, 1933-1967: New York, American Institute of Mining, Metallurgical, and Petroleum Engineers, Inc., The Grafton-Sales Volume, v. 1, p. 518-537. Mitra, Kaushik, Eleanor L. Moreland, Greg J. Ledingham, and Jeffrey G. Catalano, 2023, Formation of manganese oxides on early Mars due to active halogen cycling, Nature Geoscience 16, no. 2, p. 133-139. Morey, G.B., 1978, Lower and Middle Precambrian stratigraphic nomenclature for east-central Minnesota: Minnesota Geological Survey Report of Investigations 21, 52 p., 1 pIate. Morey, G.B., 1992, Chemical composition of the eastern Biwabik Iron Formation (Early Proterozoic), Mesabi Iron Range, Minnesota: Economic Geology, v. 87, p. 1649-1658. Muhling, J., Brengman, L., & Johnson, J. Greenalite (2025, June issue): Cryptic mineral of ancient ferruginous oceans. Elements: Greenalite - Tiny crystal with a big story, Elements (in press). Peterson, D.M., 2022, Bedrock geology and manganese mineral resource assessment map of the Emily Manganese Deposit, Big Rock Exploration map BRE-MAP-2022-02, 1:50,000 scale. Peterson, D.M., and Steiner, A., 2024, The geology, history, and ore deposit model of the high-grade Emily Manganese Deposit, Cuyuna Range, Minnesota: Oral presentation, Institute on Lake Superior Geology conference, Houghton, Michigan. Schmidt, R.G., 1963, Geology and ore deposits of Cuyuna North range, Minnesota: U.S. Geological Survey Professional Paper 407, 96 p. Schneider, D.A., Holm, D.K., O’Boyle, C., Hamilton, M., Jercinovic, M., 2004, Paleoproterozoic development of a gneiss dome corridor in the southern Lake Superior region, U.S.A. In: Whitney, D.L., Teyssier, C., Siddoway, C.S. (Eds.), Gneiss Domes in Orogeny. Geol. Soc. Am. Spec. Pap. 380, pp. 339–357. Schulz, K.J. and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region, Precambrian Research, 157, p. 4–25. Schulz, K.J., DeYoung, J.H., Jr., Seal, R.R., II, and Bradley, D.C., eds., 2017, Critical mineral resources of the United States - Economic and environmental geology and prospects for future supply: U.S. Geological Survey Professional Paper 1802, 797 p., https://doi.org/10.3133/pp1802. 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, Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR- 2009/09, 173 p. + 37 plates. Southwick, D.L. and Morey, G.B., 1991, Tectonic imbrication and foredeep development in the Penokean orogen, east-central Minnesota; an interpretation based on regional geophysics and results of test drilling, U.S. Geological Survey Bulletin 1904-C, pp. C1–C17. Southwick, D.L., Morey, G.B., and McSwiggen, P.L., 1988, Geologic map (scale 1:250,000) of the Penokean orogen, central and eastern Minnesota, and accompanying text: Minnesota Geological Survey Report of Investigations 37, 25 p., 1 pIate. 72 �Trip 3 – Proterozoic Fe & Mn Formations Steiner, A., Peterson, D.M., Berg, E., Solie, J., Larson, M., Schaefbauer, E., and Sweet, G., 2024, The North Star Emily Manganese Deposit: Observations, Interpretations, and Recommendations Following the Initial 2023 Drilling Campaign, Big Rock Exploration Technical Report BRE-TR-2023-01, 47 pages, 4 appendices, 1 plate. Strong, R., 1959, Report on Geological Investigation of the Cuyuna District, Minnesota, 1949-1959, US Steel Internal Report, 318 pages. Tyler, S.A., 1949, Development of Lake Superior soft iron ores from metamorphosed information: Geol. Soc. Am. Bull., v. 60, p. 1101-1024. Van Hise, C. R., and Leith, C. K., 1911, Geology of the Lake Superior region. U.S. Geol. Survey Mon. 52, 641 p. White, D. A., 1954, The stratigraphy and structure of the Mesabi Range, Minnesota: Minnesota Geol. Survey Bull. 38, 92 p. Wolff, J.E., 1917, Recent geologic developments on the Mesabi Iron Range, Minnesota: American Institute of Mining and Metallurgical Engineers, Transactions, v. 56, p. 229-257. 73 �Trip 4 – Soudan FIELD TRIP 4 New Geological Insights into the Genesis of Iron Ores at Lake Vermilion – Soudan Underground Mine State Park George J. Hudak1,2,3, Zsuzsanna P. Allerton1, and Annia Fayon1 1 Department of Earth and Environmental Sciences, University of Minnesota Twin Cities, 116 Church Street SE, Minneapolis, MN 55455 2 Department of Earth and Environmental Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth, MN 55812 3 George Hudak Geosciences P.L.L.C., Duluth, MN 55804 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 metaintrusive rocks that extend northeastward through northwestern Ontario and Quebec (Figure 4-1). In Canada, this terrane hosts numerous volcanogenic massive sulfide deposits (e.g. Winston Lake, Geco, Noranda), gold-rich volcanogenic massive sulfide deposits (Horne (Noranda camp), Bousquet 2 – LaRonde 1, LaRonde-Penna; 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 cut by Neoarchean shear zones. 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; Lodge et at., 2015; Thompson, 2015) have indicated that evidence for volcanic, hydrothermal, and structural processes associated with these types of mineral deposits is present throughout the Vermilion District. The Vermilion District’s iron ore mining heritage is currently preserved at Lake Vermilion / Soudan Underground Mine State Park located near Soudan, Minnesota as well as within several historic mines west of and within Ely, Minnesota. The Soudan mine operated from 1882 until December, 1962 and produced approximately 15.5 tons of hematic iron ore. 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. The mine previously hosted several underground physics laboratories, including: 1) Soudan 1 (23rd level) which studied neutrino decay; 2) Soudan 2 (27th level), also to study neutrino decay; and 3) the MINOS (Main Injector Neutrino Oscillation Search) lab, which was built on the 27th level adjacent to Soudan 1 and studied the decay of neutrinos within the earth as they passed from Fermilab to Soudan. This popular tourist site continues to be the focus of a wide variety of research related to geology, geochemistry, hydrogeology, biology, biochemistry and physics. Lake Vermilion/Soudan Underground Mine State Park is Minnesota’s newest state park. 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 the former Soudan Underground Mine State Park. Lake Vermilion/Soudan Underground Mine State Park was established in June 2010 after over 3,000 acres land was purchased from U. S. Steel Corporation (Bakst, 2013). At the present time, considerable development 74 �Trip 4 – Soudan Figure 4-1. Regional geology of the Lake Superior region illustrating the wide variety of mineral deposit types (modified from Hudak and Peterson, 2014; D.M. Peterson, personal communication, 2013). has taken place in the eastern part of the park, including the 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. 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. This research is summarized in several recent Institute on Lake Superior Geology (ILSG) field trips (Hudak et al., 2004; Jirsa et al., 2004; Peterson and Patelke, 2003; Larson and Mooers, 2009; Peterson et al., 2009a; Jirsa and Hillman, 2009; Peterson et al., 2009b), as well as in a few recent journal publications (Lodge et al., 2013; Lodge et al., 2015). In 2010 and 2011, students and faculty from the University of Minnesota Duluth Precambrian Research Center conducted new, 1:5000 scale mapping of this park and several maps and reports were produced (Radakovich et al., 2010; Vallowe et al., 2010; Heim et al., 2011; Baumgardner et al., 2013; Hudak et al., 2016; Peterson et al., 2016). These findings are summarized in Hudak et al., 2014. In addition, geologists from the Natural Resources Research Institute (NRRI), the Minnesota Geological Survey (MGS), and the University of Wisconsin Eau Claire produced a 1:10000-scale map of the park as well as a project report and accompanying spatial databases (Peterson et al., 2016; Hudak et al., 2016). Over the past several years, students and faculty from the University of Minnesota Twin Cities Advanced Field Camp have refined the geological map in an area approximately one-half mile east of the Soudan Mine headframe. 75 �Trip 4 – Soudan The results of these studies have provided a solid foundation for geological research that is currently taking place in Lake Vermilion/Soudan Underground Mine State Park (e.g.). Recent masters and doctoral studies from the University of Minnesota Duluth (Thompson, 2015) and the University of Minnesota Twin Cities (Allerton, in prep.; Allerton et al., 2024a; Allerton et al., 2024b; Allerton et al., in review) have focused their research on understanding the absolute age of massive hematite mineralization at the Soudan Mine. This is a problem that has baffled geoscientists for over a century (e.g. Gruner, 1926; Klinger, 1960). In addition, students and faculty from the University of Minnesota Twin Cities Advanced Field Camp have conducted more recent geological mapping (1:5000 scale) in an area approximately one-half mile east of the Soudan Mine headframe for the past several years. This mapping has led to minor reinterpretations of the geology in the central part of Lake Vermilion/Soudan Underground Mine State Park that was depicted by Peterson et al. (2016). Recently, a grant from the Leaonardt Foundation was awarded to one of the co-authors (Fayon) to develop a new trail in the park that will focus on public education related to the ancient geology and geological processes that have taken place in the park. K-12 teachers are playing a major role in developing the curriculum and lessons that will be part of this trail project. The goals of this field guide are to illustrate to field trip participants the wide variety of geological processes that have taken place within Lake Vermilion/Soudan Underground Mine State Park. Morning field trip stops will focus on understanding the stratigraphy, structure, hydrothermal alteration and mineralization closely associated with the Soudan iron orebodies. The afternoon will focus on observing both geological features of the massive hematite orebodies, as well as recent advances in our geochemical and geochronological understanding of these iron ore deposits in the Montana stope, located on the 27th level of the Soudan Mine. Figure 4-2. Simplified correlation map of Neoarchean assemblages in Minnesota and northwestern Ontario (after Peterson et al., 2001; Hudak and Peterson, 2014). 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. 76 �Trip 4 – Soudan Regional Geologic Setting A simplified regional geological map of the Neoarchean terranes of northeastern Minnesota and adjacent Ontario is presented in Figure 4-2. Supracrustal rocks in the Vermilion district consist of volcanicdominated 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 based on 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 Lake Vermilion/Soudan Underground Mine 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). The Soudan belt (Figure 4-3) contains large, broad generally east-west trending folds involving calc-alkalic and tholeiitic volcanic strata overlain by, and locally interlayered 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 komatiite/basaltic komatiite flows and peridotite sills. The two belts are faultbounded, 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 4-1) 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 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). Regional chronostratigraphic correlations between the Vermilion district, the Wawa Greenstone (northwestern Ontario) and the Abitibi greenstone belt (eastern Ontario and Quebec) are indicated in Figure 4-4. Geochronological information for supracrustal and intrusive lithologies in the Vermilion District is relatively sparse (Figure 4-4). 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 Table 4-1. Lithostratigraphic units within the western Vermilion District (modified after Peterson and Jirsa, 1999; Peterson et al., 2009; Hudak et al., 2012). 77 �Trip 4 – Soudan Intrusive Rocks Late Intrusions 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). Vermilion Granitic Complex Granite, schist, amphibolite, and schist-rich migmatite Giants Range Batholith Granite, granodiorite, monzodiorite, and schist-rich migmatite. U-Pb zircon dates indicate a crystallization age ranging from 2640-2777Ma (Allerton et al., 2024a). Supracrustal Rocks Newton Belt Newton Lake Formation Tholeiitic and komatiitic basalt lava flows, intrusions, and clastic strata (deep subaqueous?) Bass Lake Sequence Tholeiitic basalt lava flows, iron-formation, and felsic porphyries (deep subaqueuous) Soudan Belt Knife Lake Group Graywacke, slate, conglomerate, and sheared equivalents Lake Vermilion Formation Graywacke, slate, dacitic tuff, minor conglomerate. Detrital zircons from planar bedded, normal-graded resedimented volcaniclastic rocks have UPb age dates of 2680-2690 Ma (Lodge et al., 2013; subaerial to subaquous) Gafvert Lake Sequence Dacitic to rhyodacitic tuff, lapilli-tuff, tuff-breccia, and iron-formation. Basal dacite tuff-breccia deposits in Lake Vermilion State Park have UPb age date of 2689.7 ± 0.8 Ma (Lodge et al., 2013; subaerial to subaqeous) Britt Sequence Tholeiitic basalt lava flows (deep subaqueous?) Upper Member – Ely Greenstone Tholeiitic basalt lava flows and iron-formation (deep subaqueous?) Soudan Member – Ely Greenstone Oxide-facies iron formation with intercollated basalt lava flows and felsic volcaniclastic rocks (deep subaqueous) Lower Member – Ely Greenstone Calc-alkaline and tholeiitic basalt-rhyolite lava flows, tuffs, epiclastic rocks, and minor iron-formation (shallow- to deep subaqueous) Central Basalt Sequence 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 (transition from shallow- to deep water environment) Fivemile Lake Sequence 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; interpreted as shallow subaqueous environment). Eagles Nest Sequence Algoma-type iron formation, basalt-andesite lava flows, hydrothermal exhalites, felsic tuffs. 78 �Trip 4 – Soudan Figure 4-3. Generalized geology of the Soudan belt in the vicinity of the Tower-Soudan anticline (modified after Peterson, 2001; Hudak et al., 2014; Hudak and Peterson, 2014). 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. The outline of the Lake Vermilion section of Lake Vermilion/Soudan Underground Mine State Park is shown in green. 79 �Trip 4 – Soudan Figure 4-4. Regional chronostratigraphic correlations between the Vermilion district (Minnesota), the Wawa greenstone belt (northwestern Ontario), and the Abitibi greenstone belt (eastern Ontario and Quebec; after Ayer et al., 2010). 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. Structural Geology The structural geology of the Vermilion District has been well described by Peterson et al. (2009). 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 Timiskamingtype 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 exhibiting dominantly 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 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 (Peterson, 2001; Hudak et al., 2004; Peterson et al., 2009). 80 �Trip 4 – Soudan The third deformation event (D3) is believed to be associated with the juxtaposition of the Wawa Abitibi and Quetico terranes (Peterson and Patelke, 2003). Structures associated with D3 include abundant NE- and NW-trending faults that dissect the stratigraphic assemblages and 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. Geology of Lake Vermilion/Soudan Underground Mine State Park Lake Vermilion State Park contains a variety of supracrustal and intrusive lithological units (Figure 4-5). Supracrustal rocks that can be observed in the park 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. Additionally, a wide variety of syn- and post-volcanic mafic and felsic intrusive rocks and several varieties of sheared rocks crop out in the park (Peterson and Patelke, 2003; Radakovich et al., 2010; Heim et al., 2011; Hudak et al., 2016; Peterson et al., 2016). These various lithologies are described below. Lithology Supracrustal rocks in Lake Vermilion/Soudan Underground Mine State Park were described by Hudak et al. (2016) based on lithological types rather than lithostratigraphic members and/or formations. Their lithological descriptions are included below. A summary of mafic supracrustal rocks that occur within the park include: • undivided mafic volcanic rocks, including gray-green to green massive basalt, pillow basalt, basalt tuff, bedded scoria tuff and lapilli-tuff, and foliated basalt rocks • massive basalt comprising green to dark green, aphyric to sparsely plagioclase-phyric basalt • • • • pillow basalt, including gray-green to green bun, mattress, and lobe morphologies using the pillow lava classification of Dimroth et al., (1978) basalt tuff, including green, massive to bedded, aphyric to sparsely plagioclase-phyric tuff. bedded scoria tuff and lapilli-tuff, composed of green, thin- to very thick-bedded, poorly-sorted, typically poorly-graded tuff and lapilli-tuff containing up to 65% <1-20cm scoria lapilli foliated basaltic rocks, made up of green, fine-grained, moderately to strongly foliated basalt comprising anastomosing bands of chlorite-rich phyllite separating domains of less deformed basalt Felsic volcanic rocks within the park include: • • • • epiclastic, intermediate-felsic volcanic-derived sedimentary rocks, composed of light gray to brownish gray polymict volcaniclastic matrix-supported conglomerates and sandstones containing clasts of felsic volcanic and volcanic strata, oxide facies iron formation, and chertrich iron formation. laminated felsic tuff, made up of white to dark gray, laminated- to very thinly bedded, aphyric to sparsely quartz- ± plagioclase-phyric dacite to rhyolite tuff. felsic tuff breccia, comprising light gray, very thickly bedded to massive, matrix-supported quartz- and plagioclase-phyric polymict dacite to rhyodacite tuff breccia containing 10-20% 110 cm quartz and plagioclase-phyric coherent dacite lapilli and blocks, 5-7% lens-shaped quartz- and plagioclase-phyric pumice lapilli up to 3 cm in diameter, 1% light- to dark-gray chert lapilli up to 3 cm in diameter, and 1-3% 0.5-5.0cm diameter black to dark gray to red magnetite-rich, hematite-rich, or jasper-rich banded iron formation lapilli. massive felsic lava flows composed of light gray to greenish gray, fine-grained, massive, aphyric to quartz-phyric rhyodacite to rhyolite lava flows. 81 �Trip 4 – Soudan Figure 4-5. Geologic map of Lake Vermilion/Soudan Underground Mine State Park (after Peterson et al., 2016). Detailed maps showing locations of field trip stops are provided in Figure 4-8, and optional stops are shown in Figure 4-12. 82 �Trip 4 – Soudan Figure 4-6. Regional stratigraphic correlations across the Vermilion District (after Hudak et al., 2012; Hudak et al., 2014; Hudak and Peterson, 2014). The sections are hung on the base of the Soudan Member of the Ely Greenstone Formation 83 �Trip 4 – Soudan • felsic tuff, made up of gray to tan, fine-grained, aphyric to quartz ± plagioclase-phyric rhyodacite to rhyolite tuff. Clastic sedimentary rocks within the park include: • graywacke-slate, made up of light gray, fine- to medium-grained, thin- to medium-bedded graywacke containing up to 3% <1-2mm quartz and plagioclase grains that are interbedded with dark gray, laminated to thin-bedded mudstone/slate. • graphitic argillite, composed of dark gray to black, laminated to thin-bedded graphite-bearing argillite Chemical sedimentary rocks that occur in the park include: • oxide-facies iron formation, made up of black (magnetite-rich), dark gray (magnetite- and/or hematite-rich, red (jasper-rich or hematite-rich), or gray (chert-rich) laminated to medium-bedded, planar bedded to chaotically soft-sediment folded, banded iron formation. Hydrothermal alteration of the oxide-facies iron formations has resulted in the genesis of the massive hematite ores that make up the numerous iron ore lenses of the Soudan Mine (Gruner, 1926; Klinger, 1960; Thompson, 2015; Allerton, 2024a, 2024b). • chert-rich iron formation, composed of light gray to black laminated to very thin bedded chert that is locally interbedded with subordinate laminated to very thin bedded oxide facies iron formation Both mafic and felsic intrusive rocks have been identified in the park. Mafic intrusive rocks include: • lamprophyre intrusions, including 1) massive gray-green intrusions containing scoria, chert and granite clasts within a fine- to medium-grained groundmass composed of up to 85% acicular amphibole; and 2) black, fine-grained massive hornblende-plagioclase-bearing intrusions containing up to 15% fine-grained hornblende needles and local rounded granite blocks greater than 25cm in diameter in a fine-grained gray-black to red groundmass (Peterson and Patelke, 2003). Felsic intrusive rocks in the park include: • diorite, comprising gray to gray-green, fine- to medium-grained, plagioclase- and hornblendephyric equigranular diorite (actinolite pseudomorphs of hornblende are common) • granodiorite, made up of whitish-pink to green-gray, medium-grained granodiorite and hornblende granodiorite that locally contains xenoliths of oxide-facied banded iron formation, chert, felsic epiclastic rocks, and mafic volcanic and volcaniclastic rocks • feldspar porphyry, composed of white to whitish-pink, medium-grained, holocrystalline dacite with 5-12% 1-4mm subhedral to euhedral tabular plagioclase feldspar phenocrysts, and locally, 2-5% 13mm dark green actinolite pseudomorphs of hornblende (Radakovich et al., 2010) • quartz feldspar porphyry, characterized by white to whitish-pink, light gray to pale green-gray porphyritic dacite and rhyodacite that contains 20-25% 1-5mm diameter subhedral to euhedral plagioclase feldspar phenocrysts and 5-15% 1-3mm diameter subhedral to euhedral pale gray to gray-blue quartz phenocrysts Sheared rocks that crop out in Lake Vermilion/Soudan Underground Mine State Park include: • chlorite-dominant schist, composed of dark green very fine- to fine-grained chlorite phyllite and schist (Peterson and Patelke, 2003) • sericite-dominant schist, made up of pale yellow to yellow-gray to yellow-green very fine- to finegrained sericite-bearing phyllite and schist (Peterson and Patelke, 2003) • green mica (fuchsite)-dominant schist, comprising pale yellow to yellow gray, very fine- to finegrained sericite-bearing phyllite that contains up to 20% emerald green disseminated porphyroblasts of green mica that are up to 5mm in length • Schist ‘n’ BIF, an enigmatic unit made up of interlayered chlorite-dominant phyllite and schists and sericite-dominant phyllites and schists that contain lens-shaped clasts of oxide facies iron formation ranging from 1mm – 1 meter in length 84 �Trip 4 – Soudan Stratigraphic correlations across the central part of the Vermilion district are illustrated in Figure 4-6 (Hudak et al., 2012). Structure Three distinctive types of fault zones have been identified during geological mapping within Lake Vermilion/Soudan Underground Mine State Park. These structures include: • • • Synvolcanic fault zones (D0), which formed at the time of volcanism associated with the genesis of the volcanic rocks in the State Park, and which possess higher concentrations of synvolcanic hydrothermal alteration mineral assemblages proximal to the synvolcanic structures (see Gibson et al. (1999) and Hudak et al. (2014) for a detailed explanation of synvolcanic fault zones). Two potential synvolcanic fault zones have been described in the north-central part of the former Lake Vermilion State Park by Hudak et al. (2014); Shear zones that are associated with the regional D2 deformation, and are characterized by linear zones of sheared rocks including chlorite-dominant schist, sericite-dominant schist, fuchsite (green mica)-dominant schist, and schist ‘n’ BIF. The Mine Trend and Murray shear zones (Peterson and Patelke, 2003; Peterson et al., 2016; Table 4-2) are examples of D2associated shear zones within the bounds of Lake Vermilion/Soudan Underground Mine State Park. Late faults are characterized by brittle deformation and associated offset of adjacent lithological units. Within Lake Vermilion/Soudan Underground Mine State Park, these D3associated structures are commonly expressed as northwest- to northeast-trending, minor displacement (generally less than one meter) brittle faults that offset sedimentary bedding and / or contacts between adjacent lithological units (D3-associated faults are clearly evident at field trip stop 1). Table 4-2. 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. Measurements of other planar (e.g. bedding orientation, orientations of geological contacts, foliation measurements) and linear (e.g. mineral lineations, glacial striations) geological structures were recorded during field mapping, and are included on the new geologic map of Lake Vermilion/Soudan Underground Mine State Park (Peterson et al., 2016; see Figure 4-5). Geochronology Geochronological information for supracrustal and intrusive lithologies in the Vermilion District is relatively sparse (refer back to Figure 4-4). 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. Allerton et al. (2024a) obtained a crystallization age of 2708 ± 25 Ma for the Purvis Pluton, which intrudes the Eagles Nest Succession of the Lower Ely Member and has been interpreted as a synvolcanic intrusion (Peterson, 2001). The age of the Upper Member of the Ely Greenstone formation is currently unknown. Jirsa (2016) obtained an age of 2715.74 ± 0.50 Ma for a felsic volcanic 85 �Trip 4 – Soudan unit within the Newton Lake Formation (Boerboom, T. J., 2020). 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 2 meters 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. The age of the orebodies at the Soudan Mine has eluded geologists for nearly a century. The genesis of the massive hematite orebodies was previously interpreted to be syn- or post-depositional to the formation of the Soudan Member of the Lower Ely Greenstone Formation (Gruner, 1926; Klinger, 1960; Thompson, 2015). Gruner (1926) believed the ores could be as young as the Mesoproterozoic Duluth Complex. Klinger (1960) found abundant evidence for post-iron formation depositional genesis of the massive hematite ores, but he could not determine a specific date for mineralization and concluded the orebodies were formed along with or after shear zones that are spatially associated with the ores. Thompson (2015) speculated based on geological, structural, and lithogeochemical data, that the ores were formed during the D2 deformation, but could not determine a specific date for the massive hematite mineralization. Recent U/Pb and (U-Th)/He radiometric dating of hematite by Allerton (2024b) suggests the massive hematite orebodies at Soudan formed during Paleoproterozoic time (1640.8 ± 47.2 Ma – 1740.4 ± 72.5 Ma) and have been overprinted by a Mesoproterozoic hydrothermal event at approximately 1100 Ma (1093.1 ± 16.4 Ma). Terminology Used for This Field Trip The terminology used on this field trip will be consistent with the terminology used by Hudak et al. (2014) for their “Walk in the Park” ILSG field trip and is described below. 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) as well as latitude/longitude. Section subdivisions read from smallest to largest quarter (e.g., “NW, SE” should be read “NW quarter of the SE quarter”). A geologic map with stop locations is given in Figure 4-8. A map of optional field trip stops is given in Figure 4-12. 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 to facilitate consistent and accurate describe these geological features. Volcaniclastic rocks contain abundant volcanic material irrespective of their origin or depositional environment (Fisher, 1966). Such rocks can form directly from volcanic eruptions (whether subaerial or subaqueous), resedimentation of non-lithified volcanic deposits (for example, resedimentation of pyroclasts prior to lithification), or 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): 86 �Trip 4 – Soudan • • • • 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 due to contact with water. Deposition of such deposits is is influenced by the continued emplacement of the lava in the presence of water, and the thicknesses of the hyaloclastite deposits can be dictated by the temperature of the magma, the effusion rate, and the distance from the volcanic vent (Cas and Wright, 1987; Gibson et al., 1999; Newkirk et al., 2001); and Peperite deposits, which are generated when magma intrudes into unconsolidated clastic material and mingles with (generally wet) debris to form a volcaniclastic deposit (McPhie et al., 1993). 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 lithified rocks. Epiclasts are volcaniclasts when the pre-existing rocks are volcanic. The terminology for volcaniclastic rocks has historically been somewhat confusing because many different classification schemes have been developed (for example Fisher, 1961; Fisher 1966; Schmid, 1981; Cas and Wright, 1987; McPhie et al., 1993; White and Houghton, 2006), and different classification schemes are preferentially used in different parts of the world. As a result, the terminology relating to volcaniclastic rocks is commonly misused or misinterpreted. Four classification schemes that have been used most in the recent geological literature include: • • • • 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”. 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!” Classification is also especially difficult in ancient volcaniclastic rocks because key aspects of classification can be obscured by subsequent hydrothermal alteration, 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). For this field trip guidebook, we will utilize Fisher’s (1966) classification (Figure 4-7) for volcaniclastic rocks. This classification scheme is based on the relative proportions of ash-sized material (< 2mm), lapilli-sized material (2-64mm), 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: 87 �Trip 4 – Soudan • • • 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 (2006) have developed a modified version of Fisher’s (1966) volcaniclastic classification scheme (Figure 4-7). 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). Figure 4-7. 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. Specific terms for bedding thicknesses are also used in this guidebook. The terminology for bedding thickness has been adopted from McPhie et al. (1993) and includes: • • • • • • Laminated Very thinly bedded Thinly bedded Medium bedded Thickly bedded Very thickly bedded <1 centimeters thick 1-3 centimeters thick 3-10 centimeters thick 10-30 centimeters thick 30-100 centimeters thick >100 centimeters thick 88 �Trip 4 – Soudan FIELD TRIP OVERVIEW NOTE: This field trip will require hiking along trails and through the bush in Lake Vermilion/ Soudan Underground State Park and includes observations on the 27th level of the Soudan Mine. Hiking boots and safety eyewear are strongly encouraged as traverses in the park may encounter slippery conditions and vegetation which can cause eye injuries. Field trip participants should plan to wear a jacket and gloves while underground as the temperatures in this location are typically around 50°F (10°C). Upon arriving at Lake Vermilion/Soudan Underground Mine State Park, we will park in the main parking lot located near the Park Manager’s office. After a coffee break, we will strap on our hiking boots and spend the remainder of the morning making field trip stops in the central part of Lake Vermilion/Soudan Underground Mine State Park along a more-or-less north-south traverse. These field trip stops (Figure 4-8) will illustrate the stratigraphy, structural, and hydrothermal alteration features associated with the massive hematite ores at the Soudan Mine. We will then head back to the visitor center and have lunch overlooking one of the original Soudan Mine ore pits. After lunch, we plan to continue the field trip by going underground. We will proceed to the mine shaft and travel 2341 feet underground to the 27th Level of the Soudan Mine. We will board a train (converted ore cars) and travel west for approximately three-quarters of a mile to the Montana Stope, the last active part of the mine. At this point we will climb vertically approximately 30 feet using a very tight spiral staircase. One in the Montana Stope, we will observe the massive hematite ore, the transitional region of non-ore iron formation, and will observe massive chlorite-rich schists associated with approximately east-west-trending D2-associated shear zones. Here we plan to discuss recent geological research that has been conducted to determine the absolute age of the massive hematite ores that reside there (Allerton et al., 2024a, 2024b; Allerton et al., in review). At the end of the tour we will proceed back down to the 27th level drift, board the train, and head back east through the drift to the shaft station where we will proceed back to the surface. Field trip participants may not be able to access the 27th level of the Soudan Mine due to flooding that occurred during summer, 2024. Should this happen, afternoon field trip stops will investigate outcrops that illustrate the rarely exposed geological contact between the Soudan Member iron formation and Gafvert Lake Sequence tuffs, lapilli-tuff and tuff-breccias, Gafvert Lake Sequence tuffs and lapilli-tuffs, and subvolcanic intrusive rocks related to the Gafvert Lake Sequence that occur in the northeastern part of Lake Vermilion/Soudan Underground Mine State Park. Descriptions of these outcrops are included in a section below called “Optional Outcrop Stops” which have been taken from a recent ILSG field trip titled “Field Trip 2 - A Walk in the Park: Neoarchean Geology of Lake Vermilion State Park” (Hudak et al., 2014). Following the completion of the field trip, we will board the vehicles and proceed back to the Mountain Iron Community Center, where the field trip will end. FIELD TRIP From the Mountain Iron Community Center, proceed 0.2west on Enterprise Drive S to Emerald Avenue. Turn north on Emerald Avenue and go 0.05 miles to Highway 169. Turn east on Hwy 160 and proceed 1.5 miles to the turn off for Hwy 169/Hwy 53N. Take Hwy169/Hwy53 approximately 6.1 miles, bear right, and continue north on Hwy1/169 toward Ely. Continue north/northeast on Hwy 1/169 for approximately 23.75 miles to the first turn-off to Soudan (this will be Main Street and you will see an ore car and a sign for the Soudan Mine at the intersection). Proceed north for 0.4 miles on Main Street, turn right, and continue on Main Street for approximately 0.9 miles until it intersects 1st Ave./Stuntz Bay Road. Turn north on 1st Ave/Stuntz Bay Road and proceed for approximately 0.4 miles until you see the dirt road (McKinley Park Road) that is the east entrance to the Soudan Mine. Turn west on the dirt road, go about 0.05 miles, and park near the Lake Vermilion/Soudan Underground Mine State Park Manager’s office. Here 89 �Trip 4 – Soudan we will check in with the Manager. Turn left (south) on the dirt road and follow it around the old mine infrastructure, parking near the intersection of McKinley Park Road and Stuntz Bay Road (approximately 0.1 miles). From our vehicles parked at the intersection of Mckinley Park Road and Stuntz Bay Road, we will walk north approximately 175 meters up 1st Ave/Township Highway 4598 to field trip stop 1. Please walk against traffic as we proceed to and from this location. Stop 1: Soudan Member Banded Iron Formation Longitude/Latitude: 47.820074°N, -92.2365908°E UTM NAD 83 Zone 15N: 557144E, 5296585N (NOTE: From Peterson et al., 2009A; Hudak and Peterson, 2014) 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 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. This classic exposure of the Soudan member of the Ely Greenstone Formation lies on the north limb of the Tower-Soudan anticline approximately 75 meters north of the stratigraphic contact with the Lower member of the Ely Greenstone. The outcrop displays two generations of tight folding in delicate laminae of chert (creamy white), chert-hematite jasper (red), and magnetite-chert (black to silver-colored). The second generation of folds (F2) is tectonic in origin, having subvertical axial surfaces that trend east, and steeply plunging axes. Most display Z-asymmetry. The earlier folds (F0-1) appear to have been sharply refolded to produce complex interference patterns. Lundy (1985) studied folding at this locality and concluded that some of the apparent interference structures are the product of early-formed sheath folds that did not involve refolding by D2. The F1 structures are predominantly intrafolial, and exhibit a great variety of styles and orientations; implying they formed by layer-parallel, soft-sediment slumping (Fig. 49). Lundy’s mapping of this outcrop is an interesting demonstration of how unraveling details at a single outcrop that led to recognition that D1 deformation was not systematic here, but likely the result of soft sediment folding. It is interesting to observe the rhythmic microlaminae (1 mm or so thick) in various cherty beds exposed here and speculate about the paleoenvironment - that is, whether these represent daily heating/cooling, tidal, climatic, annual, or some other repetitive influence (e.g. waxing/waning of a hydrothermal system) in the depositional environment. What is known about units of iron-formation in the Ely Greenstone, of which there are many, is that deposition occurred in deep water (below wave base) during periods of relative volcanic and tectonic quiescence by the slow subaqueous precipitation of chemical sediments. The deep excavations in this area are the early workings of the Soudan iron mine, the first in Minnesota. The mine produced about 16 mt of high-grade hematite ore (60-63 percent ironconverted to a park. Although some early production came from open pits, most of the ore was extracted from underground workings that began here in 1900, and which now can be visited on guided tours. The mine previously housed several underground physics research facilities. These include Soudan 1 (23rd level) which studied neutrino decay; 2) Soudan 2 (27th level), also to study neutrino decay; and 3) the MINOS (Main Injector 90 �Trip 4 – Soudan Figure 4-8. Geologic map of the central part of Lake Vermilion/Soudan Underground Mine State Park (after Peterson et al., 2016) illustrating locations of field trip stops. See Figure 4-5 for the description of map units. 91 �Trip 4 – Soudan Figure 4-9. Outcrop map showing bedding trajectories and multiple generations of folds and faults (from Lundy, 1985). F1 folds are non-systematic and include both nappe- and sheath fold geometries. Neutrino Oscillation Search) lab, which was built on the 27th level adjacent to Soudan 1 and studied the decay of neutrinos within the earth as they passed from Fermilab to Soudan. Follow the field trip leaders south along First Avenue/Township Highway 4598 to the blacktop-paved Mesabi Trail. Turn to the east and walk along the paved Mesabi trail for approximately 800m. There, turn north and proceed up the unpaved trail approximately 350 meters, where the trail intersects another trail that goes northeast. Turn right and proceed northeast along the trail for about 90 meters. We will then head into the bush and hike approximately 200 meters northeast to field trip stop 2. Stop 2: Soudan Member Basalt Pillow Lavas Longitude/Latitude: 47.82544775°N, -92.22434651°E UTM NAD 83 Zone 15N: 558055E, 5297191N Detailed mapping in the park by Peterson and Jirsa (1999), Peterson and Patelke (2003), Hoffman (2007), Radakovich et al. (2010), Vallowe et al. (2010), Heim et al. (2011), and Baumgardner et al. (2013) has shown that the Soudan member is dominantly composed of oxide facies iron formation horizons that are locally interlayered with massive and pillowed mafic lava flows and associated volcaniclastic rocks (e.g. pillow breccias). Basalt lava flows associated with the Soudan Member of the Lower Ely Greenstone Formation are characterized by a medium green to dark green color. They are typically aphyric- to sparsely plagioclase ± pyroxene (now actinolite)-phyric. Plagioclase phenocrysts vary from subhedral to euhedral tabular in morphology, are typically less than or equal to 1mm in length and are locally present in abundances up to 3%. 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. At this outcrop we will observe well-preserved 0.5-2m long, aphyric- to sparsely plagioclasephyric, massive- to sparsely amygdaloidal bun- and mattress-shaped pillow lavas. These pillows dip steeply to the north and strike approximately east-west. Interpillow hyaloclastite is locally well-preserved and is 92 �Trip 4 – Soudan composed of <1-5mm chlorite-rich cuspate shards that are pseudomorphs of original volcanic glass formed by quenching of the mafic magma by water. These pillow lavas share many characteristics with the underlying Central Basalt Sequence mafic lava flows that comprise the uppermost part of the Lower Ely Member of the Ely Greenstone Formation. Such characteristics include exceptional preservation of primary volcanic textures, medium- to dark green color, sparsely plagioclase ± pyroxene-phyric, and low vesicularity. Based on these features, the Soudan Member pillow basalts at this location and are interpreted to have formed in a “deep water” (e.g. below wave based) volcanic environment. Proceed approximately 200 meters southwest to the northeast-southwest trending trail. Walk approximately 90 meters southwest to intersect the main north-south trail that intersects the Mesabi Trail. Walk approximately 350 meters south back to the paved Mesabi Trail. Turn to the east and follow the field trip leaders through the bush for about 140 meters to field trip stop 3. Stop 3: Soudan Member Oxide Facies Banded Iron Formation Longitude/Latitude: 47.82139974°N, -92.225809591°E UTM NAD 83 Zone 15N: 557990E, 5296775N Within Lake Vermilion/Soudan Underground Mine State Park, the Soudan Member oxide-facies banded iron-formation is generally 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 syndepositional soft sediment deformation and subsequent tectonic deformation, are present. As indicated above, these iron formation deposits are locally intimately interbedded with basalt lava flows such that mapping individual iron-formation and basalt horizons is often impossible at 1:5000 scale (Peterson and Patelke, 2003; Hudak and Peterson, 2014; Hudak et al., 2016). This outcrop is composed of slightly- to moderately hematite-altered Soudan member oxide facies banded iron formation. The rock varies from locally non-magnetic to slightly magnetic due to alteration of magnetite to hematite/martite. Such alteration is common in areas within a few hundred meters of massive hematite ore and is commonly found in close proximity to D2-associated shear zones. The closest previously mined massive hematite orebody was located approximately 250 meters west-southwest of this location in an existing mine pit. D2-associated shear zones have been identified approximately 25 meters north and south of this outcrop. Here, the oxide-facies banded iron formation comprises interlayered planar horizons of gray oxiderich (hematite ± magnetite), red jasper-rich, and pale white (silica (chert)-rich that are laminated, thinly bedded, and locally medium bedded. Bedding orientations generally strike more or less east-west, although locally contorted layers may vary significantly in strike direction. Dips are generally steep (>75°) to the north, although locally dips may be steep to the south. Follow the field trip leaders southwest for about 85 meters to field trip stop 4. Stop 4: Mine Trend Shear Zone “Schist ‘n’ BIF” Longitude/Latitude: 47.82126659°N, -92.22607876°E UTM NAD 83 Zone 15N: 557725E, 5296740N The “Schist ‘n’ BIF” units at this location (Figure 4-10) are composed of sheared rocks comprising chlorite schist that are interlayered with, and locally contain fragments of red, jasper-rich banded iron formation and light gray to white chert. The chlorite schist is fine-grained (<1 mm) with a tan (chloriteankerite) to green (chlorite-rich) weathered surface. Common minerals include chlorite, ankerite, sericite, 93 �Trip 4 – Soudan Figure 4-10. Photographs of outcrop exposures at Stop 4. A) Image is showing the southern face of the “Schist ‘n’ BIF” unit. The greenish tan rock is chlorite schist, and the red and white layers are banded iron formation. B) This image is north of image A and is the near horizontal exposure of the “Shist ‘n’ BIF” outcrop. The dark red inset in this figure is the location of the photographic in Figure 4-10C. C) This image shows a sigma clast of banded iron formation enclosed by chorite schist. Although the banded iron formation pieces are locally broken off, the clast can be traced, and a sketch of the clast is shown in Figure 4-10D). D) Sketch of a somewhat intact banded iron formation clast (dark pink) surrounded by silicates and other chert fragments (light pink) and enclosed by green and tan chlorite ± ankerite schist. and siderite. Banded iron formation fragments occur as clasts or thin bedded layers between foliation planes of the schist. The foliation here strikes east-west and dips near-vertically. The previously mentioned D2 associated shearing has a dextral or right lateral sense of shear that is approximately east-west trending on a regional scale, although locally sinistral shear sense indicators are present locally. This outcrop is located southwest of the previously visited oxide facies banded iron formation. The construction of a new paved road in 2020 exposed the now southern face of the unit (Figure 4-10A), providing access to three planes for structural measurements. The shear plane (Figure 4-10B) contains banded iron formation/chert clasts that serve as kinematic indicators and are located conveniently under our feet due to the dip of the schist layers. Outcrop-scale kinematic indicators of sigma and delta clasts (Figures 4-10C and 4-10D) trend mostly east-west with dextral sense of shear, mimicking regional deformation trends. Follow the field trip leaders and walk west-southwest for approximately 800 meters along the paved Mesabi back to First Avenue/Township Highway 4598. We will then proceed back to the vehicles and head to the 94 �Trip 4 – Soudan main parking lot for the Soudan Mine. We will eat lunch at the Soudan Mine visitors’ center (bathrooms are available in the visitors’ center). Stop 5: Soudan Underground Mine Shaft Longitude/Latitude: 47.82126659°N, -92.22607876°E UTM NAD 83 Zone 15N: 556765E, 5296536N Following lunch, we will watch a short video, pick up hard hats from state park staff, and proceed to the Soudan Mine shaft. Make sure to bring a jacket and gloves underground as the temperature there is approximately 50°F (10°C). Once in the cages, we will travel 2341 feet underground to the 27th Level of the Soudan Mine. After exiting the cage we will board a train (converted ore cars) and travel west for approximately three-quarters of a mile to the Montana Stope, the last active part of the mine. It is important for everyone’s safety to stay in the train car for the duration of the trip, and to not raise your hands while traveling in the train car. At this point we will exit the train, and after a short walk, climb vertically approximately 30 feet using a very tight spiral staircase to the Montana Stope. The Montana Stope is the last active part of the Soudan Mine. Thompson (2015) conducted detailed mapping of the Montana ore zone (Figure 4-11) and noted the presence of several rock types, including: • • • • • • • Chlorite-dominant schist, which locally replaces sericite-dominant schist proximal to the ore (unit 5c) Chlorite + sericite schist, which locally replaces sericite-dominant schist (unit 5cs) Sericite-dominant schist composed of sericite + paragonite ± pyrophyllite that has a mylonitic texture and occurs intermediate to ore breccia zones (unit 5s) Sericite-dominant schist that is locally silicified and occurs in well foliated zones at the margins of ore bodies that locally contain disseminated iron-rich chlorite domains (unit 5sc) Hematite ore, predominantly composed of specular hematite with microplaty hematite occurring locally within fractures and vugs (unit 4o) Hematite ore breccia, composed of hematite-rich banded iron formation and brecciated massive hematite ore with abundant milky “bull” quartz and disseminated sulfides (pyrite ± chalcopyrite; unit Fbx) Hematite-jasper banded iron formation, which retains many of its primary sedimentary textures and represents an intermediate rock between fresh Soudan Member oxide facies banded iron formation and the altered hematite-rich iron ore (unit 4a) The absolute age and geological processes associated with the genesis of the Soudan (and other Vermilion district) massive hematite ores have baffled geoscientists for over a century (e.g. Gruner, 1926; Klinger, 1960). Gruner (1926) proposed that massive hematite mineralization occurred after deposition and lithification of the Soudan Member oxide facies banded iron formation and proposed that the mineralization occurred resulted from alteration of the original banded iron formation by ascending upwelling hydrothermal solutions that oxidized most of the iron, dissolved quartz, and precipitated secondary carbonates and sulfides. He did not specify an exact age for this mineralization process. Klinger (1960) noted the close association of massive hematite ore zones at the Soudan Mine to faults (shear zones) within the mine. He states that “the orebodies occur in the iron formation and their dimensions are controlled by its structure”. He also noted that the massive hematite ores showed little evidence of deformation and concluded that “a second generation of hematite appears to post-date structural movements which occurred after the main ore-forming period. These movements, and later hematite, are 95 �Trip 4 – Soudan Figure 4-11. Geological map of the Montana stope (modified from Thompson, 2015). both later in time than a sericite rock that has been dated at 1.67 billion years by the A40/K40 method”. He also indicated that “at least some of the hematite is younger than 1.67 billion years” and concluded that the ores were “post-Huronian to pre-Keewenawan” in age. Recent masters and doctoral studies from the University of Minnesota Duluth (Thompson, 2015) and the University of Minnesota Twin Cities (Allerton, in prep.; Allerton et al., 2024a; Allerton et al., 2024b; Allerton et al., in review) have focused their research on understanding the absolute age of massive hematite mineralization at the Soudan Mine. Based on detailed mapping, petrographic studies, and lithogeochemical studies, Thompson (2015) suggested that the massive hematite ores at Soudan were formed from a multi-stage process involving alteration of the original oxide-facies banded iron formation by a fluid-dominated synvolcanic sea-floor hydrothermal system followed by interaction with hydrothermal metamorphic fluids associated with the subduction of strata within the Vermilion district. Therefore, his model for the genesis of the Soudan massive hematite ores suggests a Neoarchean age ranging from the time of the original deposition of the oxide-facies banded iron formations (~2720 Ma) to the time spanning the D2 deformation (2674-2685 Ma (Boerboom and Zartman, 1993) which is likely associated transpression and the development of the D2 shear zones in which the ores occur. New research (Allerton et al., in review) utilizing petrographic observations and electron microprobe analyses shows that the massive hematite ore can be divided further into two distinct ore 96 �Trip 4 – Soudan textures comprising: 1) homogenous microcrystalline hematite-martite; and 2) heterogenous microcrystalline hematite. The fine-grained microcrystalline hematite-martite (martite comprises hematite pseudomorph replacing magnetite) locally contains minute vug spaces and larger fractures that are filled by microplaty hematite and silicates. The heterogeneous ore contains a minor amount of metallic and/or earthy microcrystalline hematite, but is predominantly composed of coarser-grained microplaty hematite and silicates. Microcrystalline hematite-martite predates microplaty hematite and silicates based on crosscutting relationships. U-Pb radiometric dating of hematite was used to establish the timing of ore mineralization at ca. 1.8-1.6 Ga. Our new model for the genesis of Soudan massive ore suggests that hydrothermal alteration related to mineralization is coeval with orogenic events generated by Proterozoic terrain accretion and associated magmatism. Additional geochemical analyses involving mass-balance calculations and Fe stable iron isotopes indicate that the upgrade of BIF to hematite ore was a two-stage process. Dense microcrystalline hematitemartite matrix yielding a homogeneous texture was produced during the first stage. The second stage resulted in the formation of a heterogeneous texture containing microplaty hematite and silicates in larger vugs and fractures in the microcrystalline hematite-martite ore. At the end of this stop, we will proceed back to the 27th level drift using another tight spiral staircase. We will board the train and proceed east back to the shaft station where we will board the cages and return to the surface. At the surface, we will reboard the vehicles and proceed back to the Mountain Iron Community Center via the directions below. From our parking spot at Soudan Mine, proceed down the hill on McKinley Park Road for approximately 0.4 miles to the intersection with Main Street. Turn south and drive for approximately 0.4 miles to the intersection with Hwy 1/169. Turn west and Hwy 1/169 and drive for about 23.7 miles and merge onto Hwy 53/169 South. Follow Hwy 1/169 the intersection with Hwy 53/Hwy 169. Merge on to Hwy 169 south and proceed for 1.5 miles to Emerald Avenue. Turn south and proceed on Emerald Avenue for approximately 0.1 mile. Turn east and proceed for approximately 0.2 miles back to the Mountain Iron Community Center. OPTIONAL OUTCROPS The following field trip stop descriptions have been taken from the 2014 ILSG Field Trip 2 “A Walk in the Park – Neoarchean Geology of Lake Vermilion State Park” (Hudak et al., 2014). A map showing the locations of the optional field trip stops is shown in Figure 4-12. From the original parking spot near the Manager’s office at Soudan Mine, proceed approximately 0.2 miles south on Stuntz Bay Road/1st Avenue to the intersection with Jasper Street. Go southeast on Jasper Street for about -.5 miles to the intersection of Hwy 1/169. Proceed east on Hwy 1/169 for 0.75miles to Vermilion Park Drive (this is the eastern entrance to Lake Vermilion/Soudan Underground Mine State Park and allows access to campsite near Cable Bay). Turn north on to Vermilion Park drive and proceed for 2.9 miles to Old Hwy 169. Turn west (left) on to Old Hwy 169 and follow it for 0.8 miles to Vermilion Ridge Road. Turn west on Vermilion Ridge Road and proceed for approximately 0.5 miles. Turn right and park near the restroom east of Cable Bay. We will depart the vehicles here and walk across the street to the Crosscut Trail. walk southeast along the Crosscut Trail for about 2200 meters. We will then take a short hike (approximately 30 meters) up the hill to Stop 6o. 97 �Trip 4 – Soudan Figure 4-12. Geologic map of the northeastern part of Lake Vermilion/Soudan Underground Mine State Park (after Peterson et al., 2016) illustrating locations of optional field trip stops. See Figure 4-5 for the description of map units. 98 �Trip 4 – Soudan Stop 6o (Optional)” Contact” Between Soudan Member Banded Iron Formation and Gafvert Lake Sequence Rhyodacite Polymict Lapilli-tuff/Tuff-breccia Longitude/Latitude: 47.834710°N, -92.211647°E UTM NAD 83 Zone 15N: 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 within the informally named Gafvert Lake Sequence can be observed (Figure 4-13). 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. This sequence is part of the Lake Vermilion Formation. 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 Algomatype oxide facies banded iron formations and associated massive- to bedded chert deposits ranging from 25-250 meters and up to 175 meters in stratigraphic thickness, respectively. Northwest of the Soudan Mine, 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, nongraded, 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 0.53cm in diameter; and 4) 1-3% 0.5-5cm dark gray to black to red magnetite-rich, hematite-rich, or jasperrich banded iron formation lapilli. These deposits are overlain by, and interbedded with, light gray, matrixsupported, non-sorted and non-graded quartz- and plagioclase-phyric dacitic to rhyodacitic tuff deposits (Figure 4-14) 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 quartzand plagioclase-phyric coherent dacite to rhyodacite lapilli and up to 5% locally quartz- and plagioclasephyric 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). Based on regional mapping, Sims and Southwick (1980), Southwick (1993), and Southwick et al. (1998) have suggested 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 IronFormation Member and the Lake Vermilion Formation occurs, bears out this 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 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 at this field trip stop produced a high precision U-Pb date of 2689.7 ±0.8 Ma using 99 �Trip 4 – Soudan Figure 4-13. Detailed (1:5000 scale) map (after Hudak et al., 2014) 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 start our investigation where Stop 6o is indicated, and traverse along the bedding and are locally folded. 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 7o. Figure 4-14. 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 daciterhyodacite lapilli-tuff. B) Close-up 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. 100 �Trip 4 – Soudan hermal 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 (refer back to Figure 4-13). The largest part 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 this outcrop, we observe a small break in the outcrop exposure. This break occurs directly above the contact between the Soudan Member iron formation (to the south) and the Gafvert Lake volcaniclastic rocks (to the north). 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. 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 4-14), 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%. As we move northwest then north down the hill, we will traverse several outcrops composed of Gafvert Lake Sequence tuff-breccia and lapilli-tuff deposits. We will traverse northwest then north down the hill (as shown in Figure 4-13) for about 80 meters back to the Crosscut Trail. We will then head northeast along the Crosscut Trail for approximately 900 meters. We will then traverse southeast through the bush for about 45 meters to Stop 7o. Stop 7o (Optional): Gafvert Lake Sequence Tuffs and Lapilli -tuffs Longitude/Latitude: 47.838875°N, -92.202496°E UTM NAD 83 Zone 15N: 559,675E / 5,298,700N We will stop here to observe several small outcrops of the Gafvert Lake Sequence tuffs and lapillituffs. These deposits comprise very thickly bedded, light gray, quartz- and 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 quartz- and plagioclase-phyric pumice lapilli; 3) <1mm dark gray to light gray angular chert lapilli ranging from 0.53cm in diameter; and 4) 1-3% 0.5-5cm dark gray to black magnetite-rich banded iron formation lapilli. We will traverse northwest for about 45 meters back to the Crosscut Trail. We will then proceed northeast along the Crosscut Trail for approximately 750 meters, then turn north for about 35 meters to Stop 8o. 101 �Trip 4 – Soudan Stop 8o (Optional): Quartz- ± Plagioclase-phyric Rhyodacite Sill (informally named the Gafvert Lake Intrusive Complex) Longitude/Latitude: 47.843117°N, -92.197887°E UTM NAD 83 Zone 15N: 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 (Schwierske et al., 2014; Figure 4-15) 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 performed to determine unambiguously if the GLIC and Gafvert Lake volcaniclastic rocks are indeed 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 2-10% 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 quartz-feldspar-porphyry intrusions in the Vermilion District. 102 �Trip 4 – Soudan Figure 4-15. Chemical classification of various lithologies within Lake Vermilion State Park (Schwierske et al., 2014) using the immobile element 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. We will return to the Crosscut Trail and proceed northeast on the trail back to the vehicles. Upon loading the vehicles, we will return to the Mountain Iron Community Center. Travel east on Vermilion Ridge Road for approximately 0.5 miles to the intersection with Old Highway 169. Turn right (south) and continue east on Old Highway 169 for 0.8 miles. Turn south at the intersection with Vermilion Park Drive and proceed south for 2.9 miles to the intersection with Hwy 1/169. Turn west and Hwy 1/169 and drive for about 25.4 miles and merge onto Hwy 53/169 South. Follow Hwy 1/169 the intersection with Hwy 53/Hwy 169. Merge on to Hwy 169 south and proceed for 1.5 miles to Emerald Avenue. Turn south and proceed on Emerald Avenue for approximately 0.1 mile. Turn east and proceed for approximately 0.2 miles back to the Mountain Iron Community Center. END OF FIELD TRIP 103 �Trip 4 – Soudan Acknowledgements The authors would like to thank Jim Essig (Manager, Lake Vermilion/Soudan Underground Mine State Park), James Pointer (former Interpretive Supervisor, Lake Vermilion/Soudan Underground Mine State Park), and Jim DeVries (Assistant Manager, Lake Vermilion/Soudan Underground Mine State Park) for their assistance over the past two decades while the authors have conducted research, teaching, and numerous field trips within the park. References Allerton, Z., in prep., Thermal and Hydrothermal Effects of Proterozoic Events in the Archean Terrain of Northeastern Minnesota: unpublished Ph. D. dissertation, University of Minnesota Twin Cities. Allerton, Z., Hudak, G., Teyssier, C., Fayon, A., Daniŝik, M., Courtney-Davies, L, and Larson, P., 2024b, Geochronology campaign in northeastern Minnesota: Institute on Lake Superior Geology, Proceedings Volume 70, Part 1 – Program and Abstracts, p. 2-3. Allerton, Z., Hudak, G., Teyssier, C., Fayon, A., Daniŝik, M., Courtney-Davies, L, and Larson, P., 2024b, Geochronology and geochemistry of hematite ore in northeastern Minnesota: Institute on Lake Superior Geology, Proceedings Volume 70, Part 1 – Program and Abstracts, p. 4-5. Allerton, Z. P., Courtney-Davies, L., Daniŝik, M., Hudak, G. J., Teyssier, C., Mitchell, J. T., and Larson, P., in review, Hematite double-dating defines Proterozoic mineralization and thermal history of Archean banded iron formations, NE Minnesota, USA: submitted to Geology. Ayer, J. A., Goutier, J., Thurston, P. C., Dube, B., and Kamber, B. S., 2010, Tectonic and Metallogenic Evolution of the Abitibi and Wawa subprovinces: Summary of Field Work and Other Activities, 2010, Ontario Geological Survey Open File Report 6260, p. 3-1 – 3.6. 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-park-minnesotas-newest-evolving-at. Bauer, R. L., 1985, Correlation of early recumbent and younger upright folding across the boundary between an Archean gneiss belt and greenstone terrane, northeastern Minnesota: Geology, v. 13, p. 657-660. Baumgardner, M., Brown, N., 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 Map Series, PRC/Map-2013-04, 1:10,000 scale. Boerboom, T. J., 2020, D-07, Geochronology Database: Minnesota Geological Survey – Open Data Site, https://mngsumn.opendata.arcgis.com/datasets/d7903b83244a4878bcfb31f362bf5787_0/explore?location=46.147285%2C92.317888%2C7.04. 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. Cas, R. A. F., and Wright, J. V., 1987, Volcanic Successions – Modern and Ancient: George Allen and Unwin, London, 528 p. 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. 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. 902-918. 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. 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. 104 �Trip 4 – Soudan Gruner, J. W., 1926, Hydrothermal alteration of iron ores of the Lake Superior type – a modified theory: Economic Geology, v. 32, p. 121-130. 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. Hooper, P., and Ojakangas, R., 1971, Multiple deformation in the Vermilion District, Minnesota: Canadian Journal of Earth Sciences, v. 8, p. 423-434. 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/TR-2002/03, 390 pages. 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. 42-43. Hudak, G. J., and Peterson, D. M., 2014, Non-Ferrous Mineralization Associated with the Wawa-Abitibi Terrane and Duluth Complex Cu-Ni-PGM Deposits, Northeastern Minnesota: Society of Economic Geologists, Guidebook Series, v. 47, 150 p. Hudak, G. J., Peterson, D. M., Radakovich, A, Pignotta, G., Schwierske, K., and the students from the 20120-2013 Precambrian Research Center Field Camp, 2016, Bedrock Geology of Lake Vermilion/Soudan Underground Mine State Park: Natural Resources Research Institute Technical Report NRRI/TR-2016-20, 23 p. Hudak, G. J., Radakovich, A., Pignotta, G., and Schwierske, K., 2014, Field Trip 2 – A Walk in the Park – Neoarchean Geology of Lake Vermilion State Park: Institute on Lake Superior Geology, Proceedings Volume 60, Part 2 – Field Trip Guidebook, p. 37-75. Hudleston, P. J., Schultz-Ela, D., and Southwick, D. L., 1988, Transpression in an Archean greenstone belt, northern Minnesota: Canadian Journal of Earth Sciences, v. 25, p. 1060-1068. Hudleston, P.J., 1976, Early deformational history of Archean rocks in the Vermilion district, north-eastern Minnesota: Canadian Journal of Earth Sciences, v. 13, p. 579-592. 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., 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. Jirsa, M.A., 2016, Preliminary geologic maps of Lake and St. Louis counties, northeastern Minnesota: Minnesota Geological Survey Open File Report OFR 16-4, https://conservancy.umn.edu/items/a68ec090-cf02-463a-8fcbe3c05926d80e. 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, Institute on Lake Superior Geology, Proceedings Volume 50, Part 2 – Field Trip Guidebook, p. 129-169. 105 �Trip 4 – Soudan 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. 110115. 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. Jirsa, M.A., Southwick, D.L., and Boerboom, T.J., 1992, Structural evolution of Archean rocks in the western Wawa subprovince, Minnesota: refolding of precleavage nappes during D2 transpression: Canadian Journal of Earth Sciences 29:2146-2155. Klinger, F. L., 1960, Geology and ore deposits of the Soudan Mine, St. Louis County, Minnesota: unpublished Ph. D. dissertation, University of Wisconsin, Madison, 96 p. 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., 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., 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. Lundy, J. R., 1985, Clues to structural history in the minor folds of the Soudan Iron Formation, northeastern Minnesota: Unpublished M. S. thesis, University of Minnesota, Minneapolis, 144 p. 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. 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. Newkirk, T., Hudak, G. J., and Hauck, S. A., 2001, Preliminary lava flow morphology studies at the Five Mile Lake VMS prospect, Vermilion District, NE Minnesota: Implications for volcanic processes, volcanic paleoenvironments, and VMS exploration [abstract/poster]: Geological Society of America Abstracts and Programs V. 33, No. 6. 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., 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., Hudak, G.J., Radakovich, A., Pignotta, G., and Schwierske, K., 2016, Geologic Map of Lake Vermilion/Soudan Underground Mine State Park: Precambrian Research Center Map PRC/Map-2016-01, 1:10,000 scale. 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. 106 �Trip 4 – Soudan 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. 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. 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., 2014, 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, p. 113-114. 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. Thompson, A., 2015, A Hydrothermal Model for Metasomatism, of Neoarchean Algoma-Type Banded Iron Formation to Massive Hematite Ore at the Soudan Mine, NE Minnesota: unpublished M. S. Thesis, University of Minnesota Duluth, 59 p. 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. 677-680. 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. 107 �Trip 5 – Alkalic plutons FIELD TRIP 5 Neoarchean Alkalic Intrusions in the Wawa and Quetico Subprovinces Terry Boerboom (retired)1 and Amy Radakovich1 1 Minnesota Geological Survey, College of Science and Engineering, University of Minnesota, 2609 Territorial Road, St. Paul, MN 55114 This field trip will visit several alkalic intrusions that have been mapped from a combination of outcrop, drilling, and geophysical data. Refer to Figure 5-1 for the pluton names and generalized regional geology. The stop descriptions are very brief, but a more thorough description of each pluton (as well as others not visited on this trip) are contained in the introductory text. These plutons were emplaced mainly into the Lake Vermilion Formation which is composed of volcanogenic sedimentary rocks sourced from Gafvert Lk rhyodacite tuff (2689.7±0.8 Ma) and also likely from felsic tuffs at the south limb of the Britt structure (2689.6±0.5 Ma). The Linden pluton has an age of 2681.00±0.29 Ma, and the Lost Lake pluton an age of 2675.1±0.5 (Boerboom et al., 2022). These ages are slightly older than the large Shannon Lake granite (2674±5 and 2674±27; Boerboom and Zartman 1993), and straddle two ages obtained on the Britt granodiorite (2681±4 and 2685±4 Ma; Boerboom and Zartman 1993). The Idington pluton is intruded by the Shannon Lake granite, consistent with the aforementioned age dates. The suite of alkalic plutons we will visit are located mainly in the Wawa subprovince but one (Gheen) is in the Quetico subprovince and will include the Side Lake, Morcom, Linden, Gheen, Idington, and Lost Lake plutons (Figure 5-1). All of these intrusions are similar in mineralogy with varied ratios of perthitic to antiperthitic feldspar and Na-plagioclase, hence are divided into those that are more syenitic vs. monzodioritic. All contain Na-rich aegirine/aegirine-augite with variable proportions of primary and secondary-deuteric hornblende, titanite, biotite, and minor oxides and apatite. Textures vary from uniformly medium-coarse grained to strongly and coarsely porphyritic, and they typically exhibit a flow-foliation defined by feldspar and subprismatic pyroxene and/or hornblende. All except the Linden are multi-phase with variations from ultramafic pyroxenite/hornblendite to intermediate syenite/monzodiorite, with relatively minor late-phase felsic phases; where multi-phase they generally show complex and conflicting intrusive relationships between the various phases. The following discussion, modified from Minnesota Geological Survey Report of Investigations 43 (Boerboom, 1994), covers the plutons we will visit as well as others that we will not visit. Note that some of the ideas presented in this report may have been modified or discredited based on newer geochronological data. 108 �Trip 5 – Alkalic plutons ALKALIC PLUTONS OF NORTHEASTERN MINNESOTA Minnesota Geological Survey Report of Investigations 43 By T. Boerboom ABSTRACT A series of alkalic plutons in northeastern Minnesota intrude metamorphosed sedimentary and volcanic rocks in the Wawa and Quetico subprovinces of the Archean Superior Province. The plutons generally fall into one of three categories-a syenitic clan, a monzodioritic clan, and a granitic clan. The main rock phases of the syenitic and monzodioritic clans are strongly porphyritic, coarse-grained, green and pink, quartz-poor syenite and diorite. Na-rich pyroxene is the predominant mafic mineral in these intrusions, and titanite is prominent in hand sample. Some of the syenitic intrusions contain melanite garnet, and at least one contains the feldspathoids nepheline and cancrinite. The granitic intrusions consist of variably porphyritic, coarse-grained, pink granite and monzonite, with hornblende as the dominant mafic mineral. Whereas these granitic plutons tend to be uniform in texture and composition, the syenite and monzodiorite plutons are characterized by abrupt internal variations in rock type ranging from dark-colored pyroxenite to light-pink leucocratic granite, syenite, and trondhjemite. The alkalic plutons range in size from 1.5 to 60 mi2 (3.9 – 155 km2) are oval to amoeboid in shape and elongate to the northeast, and are eroded to middle and upper levels. All of the alkalic plutons produce positive aeromagnetic anomalies; outcrops, although limited, confirm that these anomalies reflect the shapes of the plutons. Several unexposed plutons, whose shapes are inferred from aeromagnetic data, have been verified by test drilling. Field relationships show that these plutons are post-tectonic. Although chemical data are not available for all of the plutons, those with analyses plot as alkalic in terms of Na2O + K2O vs SiO2, but as mainly calc-alkalic on an AFM diagram. The syenitic and monzodioritic clans are generally neither nepheline-normative to neither quartz- nor nepheline-normative, with the exception of minor leucocratic phases. All are characterized by steep REE patterns and exceptionally high concentrations of Ba and Sr. INTRODUCTION Recent mapping in northern and northeastern Minnesota, including the Koochiching-ItascaBeltrami County area (Jirsa and Boerboom, 1990) and western St. Louis County (Jirsa and others, 1991), has delineated several previously unrecognized subalkalic to alkalic plutons. This report summarizes the lithological and intrusive relationships of several of these alkalic intrusions and briefly summarizes their geochemical characteristics. Some plutons in this group, such as the Snowbank and Kekekabic stocks and the Daisy Bay and Dead River plutons have been previously described (Geldon, 1972; Sims and Mudrey, 1972) and are not included in this report. Others (Coon Lake, Linden, Lost Lake plutons) have been briefly described in the literature (Sims and others, 1970, 1972; Sims and Mudrey, 1972), but are detailed here, as are others which have no published information or were unknown (Fig. 5-1A). Alkalic rock complexes similar to these are well known in Ontario (for example Sage, 1988a, 1988b, 1988c), but few have been described from Minnesota. Several other small alkalic plutons are inferred from aeromagnetic data, but are not exposed or have not been drilled (Jirsa and others, 1991). 109 �Trip 5 – Alkalic plutons Characteristics of the Alkalic Rock Suite The alkalic intrusions fall into three general categories – a syenitic group comprising the Coon Lake, Linden, Gheen, and Baudette plutons; a monzodioritic group including the Side Lake, Morcom, Idington, and Cook plutons; and a granitoid group containing the Bello Lake, Stingy Lake, and Rice River plutons (Fig. 5-1). Although most classify into one of the three clans, the many phases in each pluton (Table 5-1) produce considerable overlap. The syenitic and monzodioritic intrusions consist mainly of medium- to coarse-grained, porphyritic, pink and green syenite and monzodiorite, whereas the granitoid intrusions are typically medium-grained, variably porphyritic, pink quartz monzonite or granodiorite. The syenite and especially the monzodiorite plutons contain multiple erratic melanocratic to felsic phases with aegirineaugite as the predominant mafic mineral, whereas the granitoid plutons generally lack multiple phases, are more uniform in texture, and contain mainly hornblende as the mafic phase. Most of the alkalic plutons intrude metamorphosed volcanic and sedimentary rocks in the western Wawa subprovince, but some are within the Quetico subprovince (Card and Ciesielski, 1986; Fig. 5-1A). All were emplaced in the latest stages of the last major regional deformational event (Jirsa and others, 1992) or after it. Several of the alkalic plutons are cut by northwest-trending Late Proterozoic diabase dikes of the KenoraKabetogama swarm, which have been dated al 2,125 Ma (Rb-Sr; Beck, 1988) [NOTE: more recent U-Pb geochron ages of ca. 2067 -2070 Ma (Chamberlain and others, 2015; Schmitz and others 2006; Wirth and others , 1995). The plutons are exposed at various levels, and many, such as the Gheen, Side Lake, and Lost Lake, are exposed close to their roof zones. All of the alkalic plutons produce positive aeromagnetic anomalies which generally conform to the pluton shape (Fig. 5-1B). The major-element geochemistry of the alkalic plutonic rocks varies greatly as a result of their diverse mineralogy. However, except for minor proportions of felsic differentiates, they are low in SiO2 (49-62 wt. % for syenites, 47 to 58 wt. % for monzodiorites, 61-70 wt. % for granites; Table 5-2), and are mostly metaluminous to weakly peralkalic in composition (Fig. 5-2). The syenitic and monzodioritic rocks are generally quartz-free to nepheline-normative, whereas the granite from the Bello Lake pluton is mostly quartz-normative (Fig. 5-3). Except for one of the granites and a leucocratic differentiate of the Idington pluton, all plot as alkalic in terms of Na20 + K20 vs Si02, but as calc-alkalic to weakly alkalic on an AFM diagram. In all the plutons, Ba and Sr are in general highly enriched, but vary between the different phases. However, the Coon Lake pluton although slightly enriched, is surprisingly low in Ba and Sr, considering its extremely alkalic composition. The Linden pluton is extremely enriched in Ba and Sr, with Ba values of up to 13,000 ppm and Sr values up to 8,100 ppm reported from company drill cores. Chondrite-normalized REE patterns for the syenites and monzodiorites are fairly consistent, with moderately steep slopes and negligible Eu anomalies (Fig. 5 6). No REE data are available for any of the granitoid plutons. Table 5-1 (next page). Modal analyses of alkalic plutonic rocks; results in volume percent. Linden analyses from Sims and others (1972, p. 161); samples with KIB and CD prefixes from drill cores, all others from outcrops. Pyroxene includes aegirine to augite; n, points counted; est, estimate 110 �Trip 5 – Alkalic plutons Sample Quartz K-feldspar Plagioclase Pyroxene Hornblende Biotite Muscovite Chlorite Epidote Apatite Sphene Opaques Calcite, Fl* Nepheline Cancrinite Melanite n KIB-7 20 43 32 2 1 tr tr tr 2 est Bello Lake Coon Lake KIB-39 KIB-40 DL-61 KIB-6 3 32 46 50 79 53 40 9 tr 8 5 2 1 8 1 tr tr tr 1 tr tr tr tr tr tr 1 tr tr tr 3 3 tr tr tr 33 15 2 2 2 est est 1143 est Gheen Idington Linden C027 C029 C551X C552B C650 C561A Gnw7A Msw2A Gnw7-2 tr 41 2 tr 8 61 26 3 tr 3 77.6 56.2 37.1 21 7 29 55 32 88 1.9 5.6 0.4 12 31 49 16.8 25.9 57.1 46 19 8 6 7 5 7 4.1 0.4 4 4 1 1 tr tr 2 3 2 5 2 2 1368 1114 989 2 Fl* 1157 tr tr 2 2 tr 999 1 1 1 0.8 2.9 1.1 3.7 3.4 2.7 2.3 946 Side Lake satellites Morcom Stingy Rice R. Cook Side Lake Linden L-Sat Sample Gnw-7B CD-4 1242 C706B C533A C534A C543A C564A C603A CD-7** CD-7 CD-9 CD-17 CD-19 Quartz 26 13 K-feldspar 56.3 28 31 tr 13 2 28 1 29 7 20 22 30 Plagioclase 0.9 52 30 65 47 33 55 44 20 54 55 47 43 67 Pyroxene 27.1 16 17 22 26 30 47 21 5 Hypersthene 14 8 Hornblende 2 34 16 9 3 8 8 3 6 Biotite 9 7 4 1 3 20 1 7 10 1 Muscovite 1 14 3 1 1 4 Chlorite Epidote 5 1 1 1 2 183 Apatite 5.2 1 tr tr 2 2 1 1 0.5 tr Sphene 1.5 2 2 I 1 1 0.5 1 0.5 tr Opaques 1 1 tr 0.5 tr Calcite 1 n 1166 1151 est 976 1188 947 1115 1137 988 1152 836 960 est * Fluorite in sample C552B; ** Poikilitic and non-poikilitic phases, sample CD-7; L-Sat is intrusion beween Linden and Gheen 111 �Trip 5 – Alkalic plutons Figure 1-1. 112 �Trip 5 – Alkalic plutons SYENITIC PLUTONS Most of the syenitic plutons are northwest of the other plutons (Fig. 5-1A). The Gheen and Baudette plutons are within the Quetico subprovince, the Coon Lake pluton is in the Wawa subprovince, and the Linden pluton straddles the subprovince border. These plutons are distinguished by a preponderance of K-rich perthite, typically as trachytic, blocky phenocrysts in an aegirine-rich groundmass, or an amphibole-rich groundmass in the case of the Gheen pluton. . The Gheen and Linden plutons contain quartz, chlorite, apatite, epidote, titanite, opaque oxides, and pyrite as ubiquitous but generally minor constituents. The Coon Lake pluton differs from all others in that it contains substantial nepheline and cancrinite; the Baudette pluton lacks both feldspathoids and quartz. Melanite garnet is present in the Coon Lake and Baudette plutons, and in some phases of the Linden. A distinctive phase of spotted monzodiorite with centimeter-size poikilitic feldspar enclosing pyroxene, hornblende, plagioclase, biotite, and sphene is present in both the Linden and Gheen plutons and in plutons of the monzodiorite clan. Although the Gheen and parts of the Linden plutons are texturally similar to rocks of the monzodiorite clan, they differ by having phenocrysts of pink perthite instead of gray antiperthite. Trachytic fabric in the syenitic intrusions conforms to the pluton edges and dips steeply toward the pluton centers. However, outcrops are generally limited to the pluton borders, and the Baudette and Linden satellite intrusions are seen only in drill core. Aeromagnetic signatures correspond with intrusion shapes, whether it be a consistent oval like the Baudette, Coon Lake, and Linden plutons, or irregular and amoeboid like the Gheen and Linden satellite plutons (Fig. 5-l B). Coon Lake Pluton The Coon Lake pluton (Fig. 5-l; Jirsa, 1990; Jirsa and Boerboom, 1990) is a 48-mi2 subcircular pluton which intrudes mafic to felsic volcanic rocks metamorphosed to greenschist grade. A narrow aureole of amphibolite-grade metamorphism accompanied pluton emplacement. The pluton has a strongly magnetic border and internal lithological zonation is indicated by a circular, weakly positive magnetic anomaly within the pluton. Its north and northeast edges are exposed in scattered outcrops, and a single 10-foot-long vertical drill core was obtained from the pluton center (Boerboom and others, 1989). The main rock type in the exposed and cored portions of the Coon Lake pluton is pink to gray, medium- to very coarse grained, slightly to strongly porphyritic nepheline syenite, 50-79% microperthite, 15-33% nepheline {Ne76-80) 2-5% aegirine (Ac23Wo18En7Fs52), and as much as 9% plagioclase, 2% cancrinite, and 3% melanite, together with accessory sphene, apatite , biotite, magnetite, muscovite , and zircon (Tables 5-1 and 5-3, Fig. 5-7). Minor proportions of pyroxenite occur in ill-defined dikelets. String- and braid-textured microperthite forms rectangular crystals with minor inclusions of aegirine, sphene, cancrinite, and nepheline. Nepheline is typically anhedral but locally euhedral, up to 2 mm in size, and ranges from fresh to moderately altered to an unknown fibrous mineral of low birefringence. Prismatic, grass- green aegirine formed early in the crystallization sequence and is trachytic. Plagioclase and cancrinite are interstitial, the latter as colorless, highly birefringent fibrous grains. Melanite garnet forms subhedral, dark-brown grains up to 1 cm across with inclusions of aegirine and altered feldspar. In places the syenite consists of trachytic, purplishbrown, rectangular perthite crystals up to 7 cm in length, with minor nepheline, melanite, biotite, and aegirine. A syenite dike that cuts mafic volcanic rocks outboard of the main pluton contains an estimated 1% scolecite and a trace of blue corundum. Netlike anastomosing veinlets of white nepheline parallel to the vertical trachytic fabric of the feldspar in drill core from the center of the pluton imply that a late influx of volatiles affected the magnetic signature of the pluton's interior. 113 �Trip 5 – Alkalic plutons Table 5-2. Major- and select minor-element geochemical analyses of alkalic rocks [Major elements in wt% oxides, minor elements in ppm, blank – nod determined. See Boerboom 1994 for more information. Bello Lake Coon Lake Gheen Morcom Sample KIB-7 KIB-40 SiO2 Al2O3 CaO MgO Na2O K2O Fe2O3t FeO Fe2O3c MnO TiO2 P2O5 LOI Total Rb Sr Y Zr Nb Ba Ni Cu Zn Cs La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf V Cr Li B 69.7 15.7 1.87 0.51 5.67 3.51 1.84 0.4 1.4 0.04 0.2 0.09 0.7 100.2 165 1210 <10 90 <10 1340 63.9 17.4 2.16 1.28 6.13 4.47 3.32 1 2.21 0.07 0.35 0.18 0.77 100.4 135 1340 41 303 14 1470 F KIB39 60.2 17.2 3.55 1.85 6.3 4.4 3.94 1 2.83 0.08 0.42 0.32 1 99.8 129 2180 17 253 14 1970 KIB-6 CLP-1 I-561A DL-61 C029 C027 12, IC-2 CD-7 57.7 17.8 5.45 0.42 5.75 5.8 2.71 0.3 2.38 0.07 0.26 0.11 2.77 99.6 100 3700 81 64 15 2530 62.3 18.9 0.33 0.35 5.32 9.3 2.18 1.6 0.76 0.05 0.18 0.02 0.85 100.1 246 1300 <10 118 24 676 8 8.7 70.3 9 17.5 33 60.3 22.4 1.69 0.28 7.1 4.13 2.54 1.7 1.78 175 ppm 0.19 <10 ppm 1.31 100.1 132 806 10 404 27 500 <l 9.8 46.3 3 79.3 130 12.8 39.7 5.5 1.45 3.9 0.5 2.6 0.49 1.4 0.5 1.4 0.19 9 45 29 74 41 55.65 21.88 1.65 1.01 8.12 7.28 3.67 6.08 5 0.08 0.44 0.08 0.37 100.38 100 926 2 119 49.2 9.52 12.3 10.5 1.51 2.4 10.2 49.65 9.27 13.08 8.89 2.38 1.71 11.76 58.5 l5.4 5.01 3.58 5.9 3.8 5.22 0.16 0.88 0.33 2.39 99.5 60 290 14 84 13 697 94 40.8 80.1 2 20.8 46 0.56 0.89 0.06 1.99 99.93 79 1470 <10 0.1 0.48 0.29 1.08 99.7 215 7 16 30 4.5 27 58 56 14.1 6.48 2.55 3.57 6.56 6.19 3.2 1.66 0.12 0.71 1.03 1.85 99.9 120 2520 47 331 24 3710 78 5.5 84.5 1 130 328 33 5.1 1.21 186 34.7 9.3 24 5.7 1.9 0.4 2.3 0.7 0.65 0.1 4 2.3 0.2 8 1.7 0.2 3 11 39 <10 67 <10 <10 630 <10 <10 160 850 11.2 8.18 11 1.7 0.4 <0.5 22 26 36 <10 0.2 <0.1 4 20 10 18 30 114 Cs �Trip 5 – Alkalic plutons Idington Sample C552B SiO2 Al2O3 CaO MgO Na2O K2O Fe2O3t FeO Fe2O3c MnO TiO2 P2O5 LOI Total Rb Sr y Zr Nb Ba Ni Cu Zn Cs La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Hf V Cr Li B F 74.5 14.8 0.08 0.09 8.8 1.08 0.51 0.03 0.03 0.02 0.16 100.1 97 43 <10 36 30 81 <1 1.5 28.6 7 5.9 10 <5 0.2 0.2 <0.5 <0.2 <0.1 2 <10 4 <10 20 10, IC2- 11, IC-2 47.27 6.99 20.49 9.18 1.91 0.89 9.15 5.44 3.1 0.3 0.67 1.62 1.33 99.28 50.27 7.61 13.87 7.92 2.41 3.63 10.68 7.04 2.86 0.2 1.51 1.46 0.84 99.91 Linden L’ndn Sat Side Lk-Sat Side Lake Cook C650A C551X 8, IC-2 MN-10 CD-4-92 C564A 1242 C706B CD-19 48.3 7.24 16.6 9.58 2.12 1.23 IO.IO 6 3.43 0.19 1 1.25 0.77 98.6 59 772 37 262 15 817 76 185 125 9 164 348 41.4 171 27.9 6.99 18.2 2 8.8 1.46 3.2 0.3 2.6 0.32 7.6 256 140 231 14 2900 54 13.1 8.16 4.56 4.24 3.86 7.05 3.3 3.38 0.13 0.96 0.75 1.23 98.5 57 1770 26 181 9 1830 64 54.9 120 1 134 278 30.8 124 19.2 5.14 11.8 1.3 6.4 1.01 1.9 0.2 1.6 0.22 5 148 120 76 14 1400 60.21 16.28 4.76 2.21 3.78 6.32 4.29 1.62 2.49 0.08 0.56 0.23 0.73 99.76 57.1 10.2 10.1 4.43 2.64 6.52 6.51 2.83 3.36 0.15 0.64 1.07 62.2 15 4.09 1.62 6.57 4.73 3.98 1.4 2.42 0.09 0.48 0.24 0.54 99.9 94 510 10 246 14 2290 52.2 11.7 9.89 7.08 4.04 2.22 10.3 55.2 14.6 7.38 6.89 3.42 2.95 7.93 0.19 0.81 0.45 1.08 100.2 67 1080 <10 102 21 985 57 108 128 2 50.6 101 0.14 0.77 0.39 0.23 100.2 60 1030 20 30 10 1220 53.5 15.4 7.68 6.25 3.79 2.21 8.37 5.7 2.04 0.15 0.76 0.39 0.47 99.3 31 1190 20 82 2 1530 61 20.4 108 1 43.2 88 11.3 48.8 9.3 2.81 6.4 0.8 4.1 0.73 1.8 0.2 1.6 0.23 2.2 208 180 26 <10 680 53.6 17.7 7.59 3.26 5.4 1.79 7.07 2.6 4.18 0.13 0.71 0.36 1.54 99.5 39 1950 <10 110 15 882 99.1 164 2924 3574 406 183 28.1 6.62 16.6 46 8.9 2.4 0.6 1.53 0.262 115 1.4 0.2 4 210 80 84 <10 �Trip 5 – Alkalic plutons Figure 5-3. Modal (A) and normative (B) compositions of the alkalic plutonic rocks. Compositional fields from Streckeisen (1973), except “P” corner, which consists of albite and anorthite, used here to emphasize variations in K content. Q, quartz; F, feldspathoids; A, alkali feldspars; P, plagioclase. Circled symbols are feldspathoidal, not quartz. Figure 5-4. Geochemical discrimination diagrams. (A) Alkalic versus subalkalic discrimination diagram; modified from Irvine and Baragar (1971). (B) AFM diagram for the alkalic plutonic rocks; modified from Barker and Arth (1976). Figure 5-5. Harker diagrams for the alkalic plutonic 116 rocks, in wt % oxides recalculated to no loss on ignition. �Trip 5 – Alkalic plutons Figure 5-6. Chondrite-normalized rare-earth-element patterns for the alkalic plutonic rocks for which analyses are available. Figure 5-7. Pyroxene compositions from the Coon Like and Linden plutons. Pyroxene compositions from Poohbah Lake (Sage, 1988a) and compositions of pure aegirine (Deer and others, 1966, p. 107) shown for comparison. 117 �Trip 5 – Alkalic plutons Table 5-3. Microprobe analyses of minerals from the Coon Lake and Linden plutons. [Linden results from Sims and others (1972). Chemical analyses in weight percent oxides. Cancrinite totals low due to abundance of volatiles. Table is continued on next page. Biotite Aegirine Coon Lake Linden Coon Lake Sphene Linden Coon Lake Linden SiO2 51.49 51.62 51.8 53 46.85 37.16 35.69 44 43 30.01 30.65 31 TiO2 0.58 0.54 0.48 0.5 2.13 2.66 3.13 0.5 0.5 37.17 34.62 31 Al2O3 1.35 1.34 1.38 1 2.5 13.96 12.34 11 13 0.55 0.57 3.5 FeO 25.43 25.36 25.74 14.8 16.97 21.02 18.42 18 14 2.06 2.34 3 MnO 0.33 0.41 0.44 0.31 1.13 1.11 0.01 0.06 MgO 1.96 2.00 1.85 8.6 8.17 9.32 10.03 0.02 0.02 CaO 6.74 6.92 6.29 18.8 17.85 0 0.01 26.51 25.44 29 Na2O 9.74 9.51 9.88 3.5 2.49 0.14 0.22 1 1 0.29 0.33 3.3 K2O 0.00 0 0.00 0.5 0.5 9.21 9.34 10.5 11.5 0.01 0.01 Cr203 0.00 0.02 0.00 nd 0.01 0 0.00 0 Total 97.62 97.70 97.85 99.84 94.60 90.28 96.62 94.04 100.80 100.7 Number of cations based on 6 oxygen 14 99.0 16.2 99.20 Number of cations based on 24 oxygen Si 2.1 2.1 2.1 2.02 1.88 6.31 6.33 6.96 6.71 4.89 5.11 4.91 Ti 0.02 0.02 0.01 0.01 0.06 0.34 0.42 0.06 0.06 4.55 4.34 3.69 Al 0.06 0.06 0.07 0.04 0.12 2.79 2.58 2.05 2.39 0.11 0.11 0.65 Fe 0.87 0.86 0.87 0.47 0.57 2.98 2.73 2.38 1.83 0.28 0.33 0.40 Mn 0.01 0.01 O.D2 0 0.01 0.16 0.17 0.00 0.01 Mg 0.12 0.12 0.11 0.49 0.49 2.36 2.65 3.3 3.77 0.01 0.01 Ca 0.29 0.3 0.27 0.77 0.77 0 0.00 0.00 0.00 4.63 4.54 4.92 Na 0.77 0.75 0.78 0.26 0.19 0.05 0.08 0.31 0.3 0.09 0.11 1.01 K 0 0 0 0.02 0,03 1.99 2.1I 2.12 2.29 0.00 0.00 Cr 0 0 0 0.00 0.00 0.00 0.00 0.00 0.00 118 �Trip 5 – Alkalic plutons Table 5-3 continued Exsolved albite Perthite Nepheline Coon Lake Linden Cancrinite Coon Lake Coon Lake SiO2 72.37 63.422 66.725 67.493 66.235 61.5 46.153 46.641 48.068 46.297 45.345 37.455 37.613 37.472 Al2O3 19.868 15.347 17.527 18.889 18.612 19 33.061 34.701 35.152 34.217 33.123 28.053 26.956 28.416 BaO 0.000 0.000 0.132 0.000 0.104 0.000 0.000 0.000 0.028 0.000 0.122 0 0 CaO 0.031 0.000 0.018 0.001 0.000 0.5 0.084 0.076 0.116 0.102 0.484 5.583 5.432 5.348 Na2O 10.58 0.646 3.055 3.785 0.716 1.5 15.947 14.848 13.198 15.917 13.574 18.776 18.468 18.201 K 2O 1.268 12.139 12.521 12.016 15.932 15.8 6.186 6.211 5.877 6.133 6.068 0.04 0.052 0.206 Total 104.118 91.553 99.977 102.185 101.598 99.8 101.43 102.476 102.412 102.693 98.594 90.029 88.52 89.643 Number of cations based on 8 oxygen Number of cations based on 32 oxygen No. of cations based on 12 O Si 3.03 3.13 3.04 3 3.01 2.92 8.67 8.62 8.79 8.59 8.70 3.02 3.08 3.02 Al 0.98 0.89 0.94 0.99 1 1.06 7.33 7.56 7.58 7.48 7.49 2.67 2.6 2.7 Ba 0 0 0.01 0 0.01 0 0 0 0.01 0 0.01 0 0 Ca 0 0 0 0 0 0.03 0.02 0.02 0.02 0.02 0.10 0.48 0.48 0.46 Na 0.86 0.06 0.27 0.33 0.06 0.14 5.81 5.32 4.68 5.72 5.05 2.94 ·2.93 2.85 K 0,07 0.76 0.73 0.68 0.92 0.96 1.48 1.46 1.37 1.45 1.49 0 0.01 0.02 119 �Trip 5 – Alkalic plutons Linden Pluton The Linden pluton [2681.00±0.29 Ma] and its smaller satellite to the east intrude mafic to felsic volcaniclastic and sedimentary rocks that are metamorphosed to the sillimanite grade at the north edge of the pluton and to chlorite grade at the southern edge. A narrow amphibolite-grade metamorphic aureole surrounds the pluton (Jirsa and others, 1992), as is evident in drill core LF-1 (Fig. 5-1). In this core, thin syenitic dikelets cut biotite-amphibole schist that has centimeter-thick green bands dominated by bright-green sodic amphibole and brown bands dominated by biotite. The country rock here is of sillimanite grade, and the aureole along the north edge of the Linden pluton may reflect retrogression. Sims and others (1972), who describe amphibolite-grade contact metamorphism of mafic volcanic rocks adjacent to the western margin of the pluton, suggest that the foliation at a high angle to the regional fabric is the result of forcible pluton emplacement. The Linden pluton is roughly 54 mi2 in size and elongate to the northwest, whereas the satellitic intrusion is about 4.5 mi2 in size and elongate to the northeast (Fig. 5-1). Exposures are limited to the pluton edges, but ten drill cores from the pluton were obtained by various private and governmental agencies. Records of these cores and the cores themselves are on file at the Minnesota Department of Natural Resources, Division of minerals in Hibbing. A summary of the cores is given in Table 5-4. The LP-series of drill cores were subsequently examined by Himmelberg (1973), and thin sections from these cores were briefly reexamined in conjunction with this report. The LP-series descriptions are directly from company logs, and the OB­ series descriptions are from the Minnesota Department of Natural Resources, Division of Minerals (Martin and others, 1988). The target of company drilling is not known, but complete metals analyses, together with Na, K, Al, Ca, Ba, and Sr abundances, were obtained by the explorationists. The satellite intrusion is not exposed, but one short drill core was obtained by the Minnesota Geological Survey (Meints and others, 1993). The main Linden pluton is generally uniform in composition within the exposed portions and in the drill cores. The typical phase consists of trachytic, variably porphyritic, salmon-pink and greenish-black, medium-to coarse-grained aegirine-augite syenite with conspicuous dark-brown sphene and centimeter-scale elliptical pyroxenite clots. As reported by Sims and others (1972), and confirmed here, the dominant minerals of the syenite are braidtextured perthite and dark-green aegirine-augite (Ac7Wo41En26Fs25, Table 5-3 and Fig. 5-7), with variable but lesser amounts of plagioclase, sphene, apatite, biotite, hornblende, magnetite, and epidote. Modal analyses and compositions of selected minerals are summarized in Tables 51 and 5-3. Complete descriptions of exposures are given in Sims and others (1972), and the drill cores are summarized in Table 5-4. Textures in the groundmass of drill holes CD-13 (Meints and others, 1993) and LF-2 are suggestive of cataclasis, yet other features in these cores, such as tabular plagioclase, blocky pseudomorphic biotite and epidote (presumably after pyroxene), and euhedral diamond-shaped sphene, show no evidence of brittle deformation. Thus the granoblastic fabric of the groundmass is most likely the result of plastic flow deformation of a viscous, mostly crystallized magma, in conjunction with late deuteric fluids. Drill hole CD-4 in the Linden satellite also contains zones of moderately sheared syenite characterized by rounded, rolled feldspar phenocrysts, suggestions of C­ S fabric, and streaky pink and gray segregations. Shear bands <I inch to 10 feet thick are foliated parallel to the trachytic fabric of unsheared portions, and the mineralogy of the sheared rock in thin section is identical to that of the undeformed portions (Table 5-1). As is the case in the Linden pluton proper, the annealed texture implies that deformation occurred in a hot, 120 �Trip 5 – Alkalic plutons semiplastic state, under near-magmatic temperatures. These submagmatic deformation features are also present in the Idington and Coon Lake plutons. Gheen Pluton The Gheen pluton, some 3 miles east of the Linden pluton, intrudes sillimanite-grade metasedimentary rocks. It is currently exposed at a high level, and its magnetic signature conforms to the long, sinuous, northeast­ elongate shape deduced from scattered outcrops along the length of the body. Local trachytic fabric defined by tabular perthite phenocrysts is steep and subconformable to the pluton contacts and the pluton shape at both map and outcrop scale is subcordant to the schistosity of the host supracrustal rocks. The pluton is chiefly mesocratic, pink and dark-green, porphyritic syenite and ranges to dark-greenish-black, coarse-grained pyroxenite and leucocratic, pink, coarse­ grained alkalifeldspar syenite. Conflicting internal intrusive relationships are common, with melanocratic phases occurring both as inclusions and as dikes in the porphyritic phase. However, pink leucosyenite dikelets cross all other phases. Mesocratic syenite phases contain 1- to 3-cm tabular perthite phenocrysts in a groundmass of fibrous amphibole, euhedral sphene, blocky oxides, stubby prismatic apatite, and minor calcite and epidote. Trace amounts of dark-green pyroxene are present, but most has been deuterically altered to bright-green fibrous amphibole. Calcite occurs both in irregular veinlets with amphibole and as magmatic, interstitial grains against sharp comers of feldspar phenocrysts. The melanocratic monzodiorite phase consists predominantly of relict pale-green pyroxene up to 2.5 mm across and lesser amounts of sericitized plagioclase, finegrained, euhedral sphene, and apatite. The pyroxene is variably replaced by euhedral, pale-green hornblende. Biotite occurs as brown books within hornblende, and is slightly altered to chlorite. Unaltered microperthite occupies a late anhedral interstitial position. Sills of pink, leucocratic, coarse-grained syenite up to 10 feet wide emanate from the Gheen pluton and cut adjacent metasedimentary rocks. This phase is characterized by irregular microcline phenocrysts in a seriate groundmass of macroscopically identified, pink microcline, fine-grained biotite, and minor white albite. Local planar miariolitic cavities lined with K-feldspar crystals are consistent with the interpretation that the pluton is exposed at a high level. The Gheen pluton differs from the other alkalic intrusions by its pervasive deuteric alteration and relatively abundant chalcopyrite. Modal analyses of the melanocratic and porphyritic mesocratic phases are listed in Table 5-1 and shown on Figure 5-3. 121 �Trip 5 – Alkalic plutons Table 5-4. Descriptions of cores from the Linden pluton; Dominant lithology in bold type. Drill Description Hole Pink, slightly porphyritic, medium-grained homogeneous leuco alkalifeldspar syenite. Trachytic foliation defined by aligned mafic minerals. LF-2 Feldspar varies from coarse blocky phenocrysts with recrystallized edges to granoblastic groundmass. LF-3 LF-4 LF-5 Dark-gray, slightly porphyritic, fine- to medium-grained heterogeneous poikilitic syenite to monzonite. Poikilitic feldspar encloses pyroxene, biotite, and sphene. Weak trachytic fabric defined by aligned pyroxenes and feldspar oikocrysts. Feldspathic dikelets and aegirine veinlets. Light-grayish-white, medium-grained, granular to hypidiomorphic hornblende monzonite with 1-cm mafic segregations. Deuteric pyroxene alteration. Light-pinkish-gray, medium- to coarse-grained, moderately porphyritic syenite; 2- to 4-cm mafic segregations of slightly porphyritic, euhedral aegirine in groundmass of feldspar, hornblende, sphene, etc. Mineralogy Perthite, biotite, muscovite, aegerine, melanite, sphene, calcite, epidote, oxides, pyrite. Microperthite that grades to antiperthite, pyroxene, biotite, sphene, apatite Perthite, antiperthite, hornblende-biotite-oxide clusters after pyroxene, epidote, sphene, apatite, calcite in brittle veinlets. Perthite, aegirine, sphene, apatite, oxides, hornblende, biotite, chlorite. Perthite, aegirine-augite, biotite, sphene, apatite, oxides. Red-stained perthite, fine granular plagioclase, stilpnomelane after biotite, chlorite after hornblende or pyroxene, leucoxene after sphene. Very coarse feldspar, 7-80% aegirine, sphene. Not described. OB-207 Pinkish-gray, coarse-grained, trachytic, aegirine syenite. OB-212 Coarse-grained, green and pink, trachytic syenite. Deuteric alteration of mafic minerals. LP-1 Pink, coarse-grained syenite with erratic distribution of mafic minerals. Contains a 1-foot-wide dike of melasyenite (biotite pyroxene-carbonate). LP-2 Upper 20 feet, leucosyenite with 75-90° dipping trachytic fabric; rest is pink and green mesocratic syenite with biotite segregations. Not described CD-13 Pink and green, coarse-grained, weakly trachytic syenite with annealed cataclastic texture. Microperthite phenocrysts, pyroxene altered to tabular clusters of biotite plus epidote. Foliated matrix of fine-grained plagioclase. Microperthite, plagioclase, biotite, epidote, melanite, sphene, sericite. CD--4 Linden Satellite Gray, coarse-grained, trachytic, porphyritic syenite with narrow pink and green, fine-grained shear bands. Granoblastic-recrystallized texture in shear bands grades into unsheared rock, contains rolled feldspar phenocrysts. Sheared portions of same mineralogy as unsheared. In unsheared portion, tabular perthite rimmed by granular plagioclase, zoned euhedral aegirine-augite rimmed by pale-green fibrous amphibole. Apatite, chlorite, sphene, allanite, oxides. 122 �Trip 5 – Alkalic plutons Baudette pluton, Lake of the Woods County The Baudette pluton is not exposed, but is seen in a drill core obtained by the Minnesota Geological Survey (hole 1986-CUSMAP-1; Mills and others, 1987). It is located 8 miles south of the town of Baudette, in Lake of the Woods County, and intrudes felsic schists. As judged from geophysical data, the pluton is approximately 1 mile long and half a mile wide, but the best resolution of available geophysics is only 1:250,000 (USGS data in Chandler, 1991). The core consists of coarse-grained, green and pink, porphyritic garnet-biotite syenite. Trachytoid phenocrysts of pink perthite up to 2 cm long with irregular granular borders are in a groundmass of green biotite, melanite garnet, plagioclase, lesser epidote, sphene, and aegirine­ augite, and accessory apatite and zircon or monazite. The brownish-yellow melanite varies from small euhedral crystals to large granular masses enclosed within coarser green biotite. The biotite varies greatly in grain size from fine-grained mats to medium-grained books with a decussate intergrown fabric. MONZODIORITE CLAN The monzodioritic group includes the Side Lake, Morcom, Idington, Lost Lake, and Cook plutons, as well as the Daisy Bay pluton (Sims and Mudrey, 1972), which is shown on Figure 5-1 but not discussed here. All are within the Wawa subprovince, adjacent to the north edge of the Shannon Lake granite phase of the Giants Range batholith (Jirsa and others, 1991). These plutons tend to be irregular in shape and elongate to the northeast, parallel to the regional D2 fabric (Jirsa and others, 1992) of the supracrustal country rocks. The rock is largely pink and green porphyritic monzodiorite, but varies erratically to dark-green pyroxenite and pink granite, granodiorite, and Na-rich trondhjemite. The pyroxenite tends to occur in small irregular pods and segregations, whereas the felsic differentiates occur in thin, straight dikes and in larger segregations. In addition, the monzodiorite group contains minor dark-green poikilitic phases in which antiperthite is grown over pyroxene, sphene, apatite, and hornblende. Porphyritic phases typically have strong trachytoid fabrics, defined by aligned feldspar phenocrysts that are subconformable to the borders but more erratic in the centers of the intrusions. The typical porphyritic phase is characterized by coarse, blocky, pink to gray phenocrysts of Na-rich antiperthite in a groundrnass of predominantly fine-grained euhedral aegirine-augite, along with sphene, perthite, polygonal plagioclase, lesser proportions of hornblende, biotite, apatite, epidote, chlorite, opaque oxides, and rare quartz. Feldspar phenocrysts are typically antiperthitic, but range from albitic plagioclase to strongly perthitic K-feldspar. The normative composition is commonly midway between the K-rich and Na-rich end members in contrast to the compositions implied by point counting, because of difficulties in properly quantifying modal abundances of the strongly exsolved feldspars (Fig. 5-3B). Idington Pluton The Idington pluton is located in west-central St. Louis County near the former village of Idington. Its irregular horseshoe shape, roughly 8 mi2 in size, is elongate to the northeast (Fig. 5-1). Trachytic foliations are predominantly northeast-oriented, subcordant to the pluton boundary, and dip generally more than 70°. However, exposures are limited to central parts of the pluton, where trachytic fabrics are less likely to conform to the pluton shape. The pluton has a rather irregular magnetic anomaly (Fig. 5-1), but the magnetic pattern has been somewhat obscured by a 150-foot-wide, strongly magnetic diabase dike and possibly by late north-trending brittle faults. No contact relationships with country rocks were observed, except at the southwestern edge of the pluton, where it is intruded by the 2,674-Ma Shannon Lake granite of the Giants Range batholith (Jirsa and others, 1991; Boerboom and Zartman, 1993). 123 �Trip 5 – Alkalic plutons Rock types in the Idington pluton are consistent in mineralogy but extremely erratic in modal proportions. Mesocratic, porphyritic pyroxene monzonite predominates, but dark-green aegirine-augite pyroxenite is common, and a small proportion of pink, sodic leucotrondhjemite occurs in aplopegmatite dikes 1 to 10 cm wide in the heart of the pluton. A mappable segregation of leucotrondhjemite exposed at the pluton's northeast corner contains minor flat-lying vuggy fractures lined with purple fluorite and a dark brown translucent tetragonal mineral tentatively identified as zircon or cassiterite. Modal analyses from four samples of the Idington pluton (porphyritic phase-C.551.X, pyroxenite phase­ C.650.A, and two felsic differentiates-C.552.B and C.561.A) are listed in Table 51. In the porphyritic phase, feldspar phenocrysts are mostly gray, rectangular, 1- to 4- cm crystals of coarsely exsolved antiperthite, but small intergrown polygonal grains of plagioclase and perthite also are abundant in the groundmass. Aegirine-augite crystals are prismatic, weakly pleochroic, and zoned with darker green rims. Apatite inclusions are common near the edges of pyroxene crystals. Euhedral, dark-brown sphene is prominent in hand sample and is microscopically associated with biotite. Melanocratic diorite is medium to coarse grained, with trachytically aligned aegirine-augite prisms in a groundmass dominated by fresh, zoned, anhedral plagioclase. Hornblende in this phase occurs as dark-green patches of secondary origin within pyroxene and as larger subpoikilitic grains with inclusions of apatite, pyroxene, and plagioclase. Green biotite forms clusters of euhedral blocky grains aligned parallel to the trachytic fabric. Sphene is mostly euhedral, but locally is subpoikilitic-anhedral and partially encloses pyroxene and other mafic minerals. Accessory minerals include allanite and secondary chlorite, calcite, and epidote. The leucocratic differentiate contains albitic feldspar as large as 40 cm across, which is characterized by graphic Intergrowths with quartz. Aplitic parts of the leuco-phase contain radial-plumose sheaves of albite as long as 3 cm. The erratic distribution between phases, which typifies exposures of this intrusion, can be documented on an outcrop scale to have formed by filter-pressing of a mafic liquid out of a feldsparphenocryst slurry. The groundmass in the porphyry is identical to the melanocratic pyroxenite, which occurs as irregular amoeboid to net-vein segregations as large as several feet. However, at the southwestern end of the pluton, cumulus modal layering is also present in the form of melanocratic layers tens of centimeters thick interlayered with mesocratic, porphyritic diorite. This diorite itself shows layering by changes in phenocryst size and abundance and trachytoid foliation parallel to layering. The layering and trachytic fabric have been drag-folded into widely spaced, crosscutting ductile shear bands, which lack cleavage or schistosity, but instead have an annealed, granoblastic habit similar to that in the Linden pluton. Pink granite pegmatite dikelets (Shannon Lake granite?) commonly occupy these shear planes. In these bands, elliptical deformed relict feldspar phenocrysts are recrystallized into granoblastic aggregates, and the pyroxenes have been replaced by bright-green hornblende. The annealed textures and lack of through-going fabric development imply that ductile deformation, recrystallization, and annealment, caused by self-induced strain during emplacement or the last vestiges of regional deformation, occurred while the rock was still hot. In a series of exposures along Highway 53 at the south edge of the Idington pluton a sharp line of demarcation exists whereby outcrops of Shannon Lake granite have inclusions only of Idington monzodiorite north of an east-west line, and those south of it have inclusions only of granodiorite derived from the Britt pluton, an early, D2-deformed intrusion (Jirsa and others, 1991, 1992). This implies that the earlier intrusive contact of the Idington pluton into the Britt granodiorite was overprinted but preserved by upward stoping of the Shannon Lake granite (Boerboom and Zartman, 1993). Side Lake Pluton The long, arcuate Side Lake pluton, roughly 27 mi2 in dimension, extends eastward from Side Lake in western St. Louis County. The pluton crops out only at its very western and eastern limits. Two exposed satellitic plugs off the eastern tip of the main pluton (Fig. 5-1) are identical in mineralogy to, and 124 �Trip 5 – Alkalic plutons conterminous with, the eastern end of the Side Lake pluton. These satellites have complex intrusive relationships with the country rocks, and are described in a separate section below. The main pluton intrudes metamorphosed basaltic volcanic and felsic sedimentary rocks along most of its length. It is just north of the Shannon Lake granite, but intrusive relationships with the granite are unknown because of lack of exposure. Crosscutting Proterozoic diabase dikes have lowered the magnetism along their length, in contrast to the Idington pluton, where the diabase dikes have enhanced the magnetism. Rocks from the exposures of the Side Lake pluton range from mesocratic biotite-hyperstheneclinopyroxene diorite on the west to hornblende monzodiorite on the east Trachytic fabrics in most outcrops are generally conformable to the margins of the pluton. Segregations of melanocratic pyroxenite to hornblendite occur throughout the intrusion, but are more abundant to the west, where they occur as irregular dikelets, segregations, and small inclusions in mesocratic diorite. In addition, dikes of diorite and pyroxenite up to 150 feet wide, which emanate from the western margin of the pluton, are parallel to the pluton boundary and intrude metabasaltic rocks. These dikes are clearly discordant to the regional metamorphic fabric in the intruded basalts; some of them contain wispy felsic stringers parallel to their walls produced by flow segregation. Thin, straight, late-stage pink granitic to syenitic dikelets are common within and adjacent to the pluton. Mesocratic, medium-grained, pinkish-gray diorite, which predominates at the western end, contains 5-10% tiny euhedral grains of pleochroic pale-pink to green hypersthene, and at least 20% euhedral, palegreen clinopyroxene with light-colored rims. These pyroxenes range in size from less than 1 to 3 mm; the hypersthene is generally finer grained, and the clinopyroxene is variably phenocrystic. Plagioclase is the predominant feldspar. A sample from the pluton 200 feet from the western contact contains strongly zoned, subhedral, trachytic plagioclase, with fuzzy grain boundaries. Another thin section 500 feet from the contact has fine-grained granoblastic plagioclase, orthoclase, and quartz between larger augite phenocrysts. Apatite is abundant as fine-grained euhedral prismatic grains included within pyroxene and feldspar. Oxides occur both within augite as wormy blebs of apparent secondary origin and as scattered blocky, fine-grained crystals. Sphene is rare or lacking at the western end. Brown biotite is a generally minor component at the western end of the pluton, but in some outcrops composes up to 5% of the rock. It locally forms vertically oriented poikilitic plates that are up to 2 cm long and oriented parallel to the trachytic fabric of the monzodiorite. The eastern outcrops consist of moderately heterogeneous, medium- to coarse-grained, pink and green pyroxene-hornblende monzodiorite. Here moderately developed trachytic foliation plunges 2030° to the souL'1west, down the axis of the pluton. Fine-grained, centimeter-sized, angular cognate xenoliths of mafic monzodiorite, in addition to dark-green mafic stringers, are common. Side Lake Pluton Satellites Two small plugs are exposed northeast of the Side Lake pluton. One, about 1/4 mile east of the Side Lake pluton, is round and 3/4 mile in diameter. It consists of medium- to coarse-grained, trachytic, weakly porphyritic, green and pink hornblende-pyroxene monzodiorite. Subhedral 5-mm orange-white antiperthite and scattered 2-mm aegirine-augite phenocrysts are set in a fine-grained groundmass of green prismatic pyroxene, fine-grained anhedral feldspar of unknown composition, minor hornblende, and brownish-green biotite mostly replaced by chlorite. Apatite, biotite, opaques (oxides and pyrite), and sphene each compose about 1% (Table 5-1). Clots and irregular segregations of pyroxenite are common, as are late dikelets of pink syenite up to 10 cm wide which cut across trachytic fabrics. The trachytic fabric is locally variable and inconsistent in orientation, but generally has a shallow southwest plunge toward the main Side Lake pluton. The next satellite is a horseshoe-shaped body, 0.75 mi2 in size, located 1.5 miles east of the Side Lake pluton (Fig. 5-1). Its linear trachytic fabrics are subconformable to the edges of the plug, and again plunge shallowly southwest toward the Side Lake pluton. This small intrusion is highly varied in 125 �Trip 5 – Alkalic plutons texture and mafic content, but dark-green poikilitic diorite and white leucocratic monzodiorite with small inclusions of pyroxenite predominate. The poikilitic diorite has 1-cm, oval-shaped antiperthitic to perthitic poikilitic feldspar with inclusions of aegirine-augite, biotite, sphene, and apatite. The long axes of the poikilitic feldspars and prismatic pyroxenes are parallel and define a primary trachytic fabric. The dark poikilitic phase is sharply cut by sills of the white monzodiorite. However at one location, the two rocks are commingled in a pillow-like fashion that suggests mixing of immiscible liquids. Thus the field relationships, as well as mineralogy, indicate that the dark poikilitic and leucocratic phases are comagmatic. Relationship of Satellitic Intrusions to the Side Lake Pluton Their similar lithological and textural attributes and aligned trachytic fabrics imply that the Side Lake pluton and the two satellites are derived from a common source at depth to the west. Further evidence of comagmatism is the presence of numerous thin anastamosing sills of white monzodiorite, identical to the white phase in the small plugs, which are parallel to the schistosity of the surrounding metasedimentary country rocks and intercalated with them. The intercalated monzodiorite and schist define a mappable unit along a discrete zone (dashed area on Fig. 5-1 that links the two satellitic plugs to the Side Lake pluton. This zone continues past the eastern satellite for at least 2.5 miles, where it merges back into a magnetic anomaly interpreted as another alkalic intrusion (Jirsa and others, 1991). At the intersection of Highway 73 and the Sturgeon River east of the Side Lake pluton, the intercalated monzodiorite and metasedirnentary rocks are transected by a late, north-trending brittle shear zone, which has minimal offset but has reduced the rocks to a fine-grained cataclasite. Morcom Pluton [Thin section CD-7] The Morcom pluton is just north of the Side Lake pluton and may be related to it at depth, because a large positive gravity anomaly underlies the area between the two. The Morcom pluton, which intrudes metasedimentary rocks, has a bulbous shape with a long narrow appendage to the east (Fig. 5-1). Scattered outcrops exist at the eastern limit of the pluton, and a drill core was obtained from the western end, near the north side. Trachytic foliation near the southeast edge dips 80°N, and at the eastern tip plunges 40°SW. In the drill core the foliation dips 40-45°, presumably toward the pluton center. Rock types are similar in the Morcom pluton, the eastern Side Lake pluton and its satellites, and the Idington pluton. However, the major-element geochemistry of the Morcom is very similar to that of the Linden pluton (Fig. 5-5). The drill core consists of multiphase, medium-grained, weakly porphyritic biotite-hornblende-pyroxene monzodiorite, with a trachytoid foliation defined by alignment of plagioclase phenocrysts and prismatic mafic minerals. Dark-green, poikilitic monzodiorite occurs in the core as 15-cm inclusions or layers; small miariolitic cavities lined with fine-grained crystalline biotite, pyroxene, and pyrite are also present. Late brittle slickensided faults and fractures, oriented obliquely to foliation, locally transect the core. Feldspar compositions vary from clean plagioclase with narrow twin lamellae to untwinned plagioclase, and from antiperthite to perthite, the latter confined to anhedral grains in the groundmass. Euhedral, light-green, aegirine-augite has hornblende rims and alteration patches; hornblende is also present as subhedral to prismatic, brownish- to bluish-green grains with patchy color zonation and rare deuteric overgrowths of colorless actinolite. Biotite is dark green and pleochroic, and is associated with hornblende. Accessory minerals include sphene, allanite, apatite, epidote, calcite, and minor secondary oxides within pyroxene. The nonpoikilitic and poikilitic phases have similar mineralogy (Table 5-1). Exposures at the eastern tip of the pluton consist of pink to gray, medium-grained monzodiorite with abundant 5- to 10-cm, elongate xenoliths of foliated felsic to pelitic schist, together with cognate xenoliths of fine-grained melanocratic monzodiorite. Some of the intrusive-breccia xenoliths are themselves an intrusive breccia. Pink monzonitic dikelets are abundant and cut all the earlier intrusive phases and xenoliths. The monzodiorite contains scattered sericitized plagioclase phenocrysts in a fine­ grained groundmass consisting of up to 5% quartz intergrown with granular K-feldspar and plagioclase, 126 �Trip 5 – Alkalic plutons along with hornblende, actinolite, sphene, chlorite, apatite, and minor oxides and secondary calcite. Melanocratic clots are of similar mineralogy but with a higher proportion of mafic minerals. Elsewhere at the eastern terminus, the rock lacks xenolithic inclusions but is still heterogeneous and cut by late, pink felsic differentiates. This inclusion­ free monzodiorite has a trachytic fabric defined by aligned feldspars and mafic clots, and is similar in mineralogy to the core from the western end of the pluton. Lost Lake Pluton (2675.1±0.5 Ma) The Lost Lake pluton, the "pluton southwest of Lost Lake" of Sims and Mudrey (1972), was described as a circular pluton composed of heterogeneous syenite with a local, conspicuously porphyritic facies, a pegmatitic facies with miariolitic cavities, and small bodies of pyroxene­ biotite lamprophyre. They noted that the borders of the pluton tend to be quartzose and contain small angular inclusions of metagraywacke and slate of the Lake Vermilion Formation. Based on detailed remapping, the authors have redefined the shape of the pluton as a long, sinuous and bulbous, northeast-trending body that is 1 mile or less wide but approximately 9 mi2 in size. The eastern tip of the pluton lies 1/4 mile south of Lost Lake, and the western terminus is just south of Angora on State Highway 53, about half a mile north of the Idington pluton (Fig. 5-1). The uniform magnetic signature of the pluton has been lowered locally by late, north-south, brittle faults which have minimal offset. The western end of the Lost Lake pluton is not exposed and its shape is inferred from aeromagnetic data, whereas the central portion is well exposed, and scattered outcrops exist over the eastern end, mainly adjacent to more resistant crosscutting Proterozoic diabase dikes. Two mappable intrusions of quartz monzonite 1/4 mile in diameter occur adjacent to the main body (Jirsa and others, 1991). These small plugs are similar in composition to leucocratic dikes within the main pluton, and are related to the pluton. Subvertical trachytic fabric, which is defined by both phenocrysts and elongate poikilitic feldspar, strikes generally northeast, subparallel to the length of the intrusion. The small, separate bodies of pink monzonite also possess a northeast-oriented trachytic fabric, defined by orbicular clots of biotite, disseminated biotite, or aligned feldspar crystals. The Lost Lake pluton is mineralogically similar to the Idington and eastern Side Lake plutons, but contains a higher proportion of pink leucocratic phases. The main rock types range from pink and green, porphyritic monzodiorite to dark-green, poikilitic biotite-pyroxene monzodiorite to pink monzonite, syenite, and quartz monzonite. Pyroxenite occurs in small segregations, in the same fashion as in the Idington pluton. In general, the poikilitic and porphyritic phases are earliest and are cut by the pink rock varieties. Small dikes of pink granitic pegmatite cut all other rock types, but it is unclear whether these dikes are related to the pluton or are from an external source, such as the Shannon Lake granite of the Giants Range batholith. The pink monzodiorite and syenite differentiates are medium grained, equigranular to weakly porphyritic, and commonly aplitic to pegmatitic, with pyroxene, hornblende, and biotite as the predominant mafic phases. The two small felsic plugs of quartz monzonite to granodiorite contain a mafic mineral assemblage of varied proportions of biotite, chlorite, and hornblende, and up to 30% quartz. The margins of these plugs contain abundant inclusions of felsic volcanic country rocks up to 15 feet across which were clearly deformed prior to incorporation, and small dikes emanating from these plugs cut across fold axes in the supracrustal rocks. One of the plugs contains a unique medium-grained, pink, orbicular granodiorite with trachytically aligned discs of black, concentrically foliated biotite that are as much as 0.5 cm thick and 5 cm long. The biotite orbs which contain intergrown sphene, apatite, plagioclase, quartz, and magnetite, compose as much as 15% of the rock. The orbicular rock grades into a non-orbicular phase with the same proportion of biotite, but as medium-grained, uniformly disseminated flakes. In addition to the typical phases, related rocks in the small plugs include coarse-grained, dark-green biotite-hornblende lamprophyre; green poikilitic monzodiorite; and pink monzonitic pegmatite. 127 �Trip 5 – Alkalic plutons Cook Pluton The Cook pluton (Cook Airport pluton on Southwick, 1993), 1 mile south of the town of Cook (Fig. 5-1), is inferred from aeromagnetic data to be 1.5 mi2 in size, elongate to the east. No outcrops of this pluton are known, but a drill hole in the western margin of the pluton recovered core of uniformly coarse-grained, peppery, dark-greenish-black and light-green, epidote-altered hornblende-biotite diorite. Strong trachytic foliation, which is defined by tabular plagioclase and mafic minerals, dips 50° from horizontal. One fine-grained cognate xenolith, 1 cm x 3 cm in size, is present near the bottom of the 10-foot core. The rock has a primary hypidiomorphic-granular texture, but pervasive, small euhedral crystals of secondary epidote are overprinted on all primary minerals, preferentially in the cores of plagioclase, and as fine-grained granular masses in biotite. Pale-green hornblende is rimmed by green biotite, and zoned plagioclase is clean and well-twinned, despite the pervasive epidote alteration. Accessory minerals include apatite, sphene, oxides, calcite, and interstitial orthoclase. Scattered late, brittle fractures which dip as much as 20° from horizontal are lined with coarse, lineated chlorite, pinkaltered feldspar, and a crust of epidote and white carbonate. The pristine trachytic igneous texture and lack of metamorphic fabric indicate that the pervasive epidotization is the result of deuteric alteration, rather than regional metamorphism. GRANITOID PLUTONS The granitoid group includes the Stingy Lake, Rice River, and Bello Lake plutons, all within the Wawa subprovince. These plutons tend to be oval in shape and elongate to the northeast. Rocks in this group are characterized by substantial quantities of quartz and are vaguely to strongly porphyritic and trachytic. Hornblende is the predominant mafic phase, along with biotite and rare pyroxene. These plutons are considered part of the alkalic group on the basis of their similarity to the other alkalic plutons in size, shape, high Ba and Sr content, magnetic signature, and trachytic fabric. Stingy Lake Pluton The Stingy Lake pluton is a 9 mi 2 circular pluton located 3 miles south of Sturgeon Lake, adjacent to the Giants Range granite, and is inferred to intrude mafic volcanic rocks (Fig. 5-1). Although unexposed, its round shape is well defined by its aeromagnetic anomaly (magnetic rim and nonmagnetic core). A 10-foot drill core was obtained from the northwest side of the pluton (Meints and others, 1993). The intrusion is cut by two Proterozoic diabase dikes. The rock in the core is uniformly coarse-grained, porphyritic, pink granodiorite to quartz monzodiorite. Tabular phenocrysts of string-and-braid microperthite up to 7 mm long, together with weakly zoned plagioclase up to 2 mm long having narrow twin lamellae and weakly sericitized cores, define the 45°-dipping trachytoid foliation. The perthite contains small blocky plagioclase inclusions, and the areas between abutting perthite grains are also stuffed with small blocky plagioclase grains. Lightgray anhedral interstitial quartz with shadowy extinction has been recrystallized into coarse polycrystalline aggregates. Hornblende is mostly altered to green biotite, epidote, and granular oxides; however, fresh, dark-green, euhedral hornblende is locally preserved within quartz and feldspar. Accessory minerals include blocky primary oxides, sphene, zircon, and apatite (Table 5-1). Late closely spaced, vertical brittle fractures lined with epidote and chlorite are pervasive in the 10-foot core. Rice River Pluton The Rice River pluton, about 5 miles west of Cook, intrudes metamorphosed sedimentary rocks. It is inferred to be approximately 15 mi2 in size, although its magnetic signature (Fig. 5-1B) of magnetic rim and nonmagnetic core is irregular and overprinted at the western edge by a north­ trending Proterozoic diabase dike. A drill hole in the magnetic eastern rim of the pluton recovered core of gray, coarse-grained, porphyritic quartz monzonite to monzodiorite. Steeply inclined trachytic foliation is 128 �Trip 5 – Alkalic plutons defined by euhedral, strongly zoned, 1- to 2-cm perthite phenocrysts and small elliptical melanocratic clots that are fine-grained cognate xenoliths. Groundmass to the phenocrysts consists of 3- to 6-mm subhedral microcline, zoned plagioclase, hornblende, and anhedral interstitial quartz; the phenocrysts are rimmed by fine-grained plagioclase and myrmekitic quartz-feldspar intergrowths. Plagioclase grains are heavily sericitized, preferentially in the cores. Hornblende is weakly zoned and slightly altered to biotite, chlorite, and opaques. Euhedral sphene, allanite rimmed by epidote, and secondary calcite occur in minor proportions. Scattered chlorite-pyrite veinlets dip 5-10° from horizontal and occupy brittle fractures; some have slickensides that dip shallowly in the fracture planes. Bello Lake Pluton The Bello Lake pluton (Jirsa, 1990; Jirsa and Boerboom, 1990) is just southwest of the Coon Lake pluton. It is approximately 60 mi2 in size, elongate to the northeast parallel to the regional strike of the mafic to felsic supracrustal rocks that it intrudes. The Bello Lake pluton is unexposed, but three 10-foot drill cores were obtained by the Minnesota Geological Survey, two near the western end, and one near the eastern end of the pluton (Fig. 5-1). The intrusion is magnetically quiet relative to the mafic volcanic rocks around it, but the pluton margins are strongly magnetic locally. As judged from the cores, the Bello Lake pluton is uniform in color and texture, but moderately variable in composition, ranging from pyroxene monzonite to hornblende granite. The pyroxenebearing phase occurs close to the pluton margin, whereas the hornblende monzonite occurs near the center of the pluton at its western end, and the hornblende granite is near the eastern end of the pluton. Data are insufficient to properly judge spatial variation of rock types, but the observed distribution suggests that the pluton may be zoned from a pyroxene-bearing phase at the rim, to a more differentiated, quartz-bearing phase near the center. Pyroxene monzodiorite near the pluton border (KIB-39; Table 5-1) is green and pink, medium grained, and seriate in texture with a strong trachytic fabric defined by rectangular, zoned plagioclase and subhedral, weakly uralitized augite crystals. Accessory oxides, sphene, hornblende, chlorite, biotite, and apatite all formed late in the crystallization sequence, and tend to occur together. Green and pink, medium-grained, slightly porphyritic hornblende monzonite (KIB-40; Table 5-1) on the western side of the pluton contains subhedral-prismatic hornblende and small phenocrysts of grayish-pink, blocky microperthite in an allotriomorphic-granular to weakly seriate groundmass of plagioclase, perthite, and minor quartz. This rock is similar to the pyroxene monzodiorite in hole KIB-39, except that hornblende occupies the position of pyroxene. Hornblende granite from the eastern part of the pluton (drill hole KIB-7; Table 5-1) is characterized by strongly zoned, blocky plagioclase with sericitized cores surrounded by poikilitic microperthite. Myrmekitic feldspar-quartz intergrowths occur along perthite-plagioclase grain boundaries. Quartz is coarse and anhedral interstitial, and hornblende forms dark-green, irregular grains with abundant tiny quartz inclusions near the edges and granular oxide inclusions in the cores. Accessory green biotite, epidote, and chlorite are associated with hornblende as alteration products, and blocky apatite crystals are associated with mafic phases. PETROGENETIC AND GEOCHRONOLOGICAL STUDIES Arth and Hanson (1975), using data on major, trace, and rare earth elements, and isotopic data from the Linden pluton, concluded that the magma formed from 5 to 10% partial melting of a mixed eclogite and garnet peridotite source at mantle depth. Stern and others (1989) believe that the Linden originated by partial melting of a LILE­ enriched mantle peridotite at shallow depths under hydrous conditions created by mantle metasomatism from rapid subduction of oceanic 129 �Trip 5 – Alkalic plutons lithosphere. However, they have lumped the Linden pluton in with the "sanukatoid suite," a very broad suite of rocks of variable size, timing, and associations throughout the Superior Province. The age of D2 deformation of t h e supracrustal rocks at the southern edge of the area from Cook to Side Lake (Jirsa and others, 1991) was bracketed by U-Pb zircon geochronology to between 2,685 and 2,669 Ma (Boerboom and Zartman, 1993). The alkalic plutons lack significant D 2 fabrics, and thus could not have been emplaced until approximately 2,669 Ma. [NOTE: This has been discredited by the new ages of 2681.00±0.29 Ma on the Linden and 2675.1±0.05 Ma on the Lost Lake plutons – Boerboom and others, 2022] The Idington pluton is intruded by a granite pegmatite inferred to have originated from the Shannon Lake granite, which was dated by Boerboom and Zartman (1993) at 2,674 ± 5, or a minimum of 2,669. Catanzaro and Hanson (1971) obtained a discordant Pb207/Pb206 age of 2,740 ± 10 Ma on sphene f r o m the Linden pluton. Prince and Hanson (1972) obtained a similar age of 2,740 Ma, based on a Rb/Sr isochron through apatite and two whole-rock samples. These older ages on the Linden pluton relative to the younger age implied for the ldington pluton indicate that the syenitic rocks may be slightly older than the monzodioritic group, or that pluton emplacement may have progressed from north to south. Clearly, modern high-precision U-Pb zircon dates on the alkalic plutons are needed. [NOTE: Recent age of 207Pb/206Pb 2681.00±0.29 Ma (Boerboom and others, 2022)] CONCLUSIONS The alkalic intrusions in northern Minnesota can be generally subdivided into a syenitic group, a monzodioritic group, and a granitoid group. The syenitic plutons are somewhat north of the monzodioritic intrusions, whereas the granitoid plutons are interspersed with the monzodiorites. Although these groups differ in mineralogy, they are all similar in terms of size, texture, map pattern, geochemistry (e.g., high Ba and Sr), aeromagnetic signature, and timing of emplacement All of the alkalic plutons have porphyritic textures, and the syenitic and monzodioritic plutons typically contain abrupt phase transitions from predominantly mesocratic, porphyritic rocks to dark-green pyroxenites and pink felsic differentiates. The granitoid plutons are more uniform in composition and texture. The plutons are eroded to various levels. The northeastward-elongation and en-echelon map pattern of the Gheen and Lost Lake plutons and the eastern Side Lake pluton and its satellites indicate exposure at high levels, whereas the broad, rounded map shapes of the Linden, Coon Lake, and Bello Lake plutons indicate a deeper level of erosion. The map patterns of the relatively well exposed Side Lake, Idington, and Lost Lake plutons indicate a similar style of emplacement, in which the plutons have penetrated the supracrustal rocks to different levels. The Side Lake pluton plunges to the west, as indicated by the deeper level of erosion at the western end of the pluton and the west-plunging linear trachytic fabrics in the Side Lake satellites. This westward plunge may be either a primary emplacement feature or the result of tilting of the pluton prior to unroofing. Several other plutons of alkalic affinity are suggested by the aeromagnetic data, but they are not exposed and their existence has not been verified by drilling. ACKNOWLEDGMENTS Field work and geochemical analyses for this project were funded by the Minerals Diversification Program administered by the Minerals .Coordinating Committee for the Minnesota Legislature. The Minerals Division of the Minnesota Natural Resources Research Institute (also supported by the Minerals Diversification Program) coordinated the analytical work for several of the geochemical samples. 130 �Trip 5 – Alkalic plutons FIELD TRIP STOPS These stop descriptions are brief – see introductory section for more detail about the individual plutons. Directions: To stop 1: Drive west on Highway 169 to Highway 5 at Chisholm • go ~15 miles north on Hwy. 5 to road 915/McCarthy Beach road the •go west on McCarthy Beach road between Side and Sturgeon Lakes, this turns into Link Lake trail—stay on for a total of ~7 miles to small trail (491988, 5283636) • walk SW on trail ~ 1000’ / 300m and go W-SW to ridge where there are some peeled outcrops. Work your way back NE along top of ridge for more outcrop back to road. Stop 1. (NAD83: 491723, 5283344) (47.70606°, -93.10680°) Side Lake pluton – multiphase Side Lake Pluton ultramafic to felsic. NEXT: Head back east on Link Lake Trail • at about one mile turn left on Beatrice Lake road at sharp bend in trail. • Take Beatrice Lake road ~1.7 miles to intersection with Snake Trail and turn right. • Follow Snake Trail ~2.5 miles to Hwy. 5. • Go right/south on Hwy. 5 for ~1.8 miles. • Go left on Hwy. 65/Perch Lake road for 1.7 miles then • take a left / north on Dean Forest Road (268). • Follow Dean Forest Road (276) ~4.6 miles through a series of jogs to a small road on the right / south (Mud Hole Road # 276). • Drive ~500 feet and park next to knob on the west side of the road (504781, 5284935), go up on outcrop knob to east. Multiple peels. Stop 2. (NAD83: 504781, 5284935) (47.71778°, -92.93625°) Roof zone of Side Lake Pluton. Many dikes of multiphase monzonitic rocks cut high-grade garnet-staurolite-sillimanite bearing metasedimentary rocks of the Lake Vermilion Formation. Some dikes may be unrelated tonalite. One smaller outcrop near road along south edge of knob has 3-5 x 7-15 cm mafic enclaves in quartz tonalite to monzonite. This lies in what is interpreted as the roof zone of the Side Lake Pluton. NEXT: Go back to road # 276 • turn right / east and drive ~3.4 miles to Highway 73. • Turn left / north on Hwy 73 for 4 miles to Hwy. 22 • turn left / west for 3 miles to road 931. • Turn left / south on 931 for 0.5 miles to crest of small hill, outcrop in the east side of road. Stop 3. ***Private Property please be respectful*** (NAD83: 504882, 5290921) (47.77164°, -92.93484°) Morcom Pluton – Monzdioritic intrusive breccia of widely variable grain size bearing many inclusions of different phases of itself that range from intermediate-porphyritic to ultramafic. One thin section was made from this outcrop and in it the mafic phase is dominantly hornblende, in contrast to samples from other parts of the pluton which contain abundant green Na-pyroxene in addition to hornblende. Sphene, magnetite, and apatite are also relatively abundant. This pluton is not well exposed with only a few outcrops in this vicinity on the east end and a drill hole on the west end; extent outlined via aeromag data. NEXT: Head back east to Hwy. 73 • turn left / north for 5 miles to Highway 1 • Turn left / west on Hwy. 1 for 3.4 miles to large outcrop ridge and find a safe place to park... Stop 4. Linden Pluton (NAD83: 504222, 5301045) (47.86273°, -92.94355°) (2681.00±0.29 Ma) Brownish-pink medium- to coarse-grained, strongly foliated, moderately porphyritic pyroxene syenite. Tabular crystals of gray perthite and prismatic dark green pyroxene phenocryst are surrounded by a pink groundmass composed mainly of fine granular albitic plagioclase. The foliation (and weak subvertical lineation), interpreted as magmatic, is defined by the phenocrysts of microcline and 131 �Trip 5 – Alkalic plutons pyroxene, and dips steeply to the northwest parallel to the pluton margin. The syenite also contains cm-scale ellipsoidal ultramafic pyroxenite enclaves that are flattened parallel to the main foliation. Dominantly composed of perthitic alkali feldspar, albitic plagioclase, and dark-green prismatic aegirine-augite (Ac7Wo41En26Fs25, Table 5-3 and Fig. 5-7), with lesser amounts of sphene, apatite, biotite, hornblende, magnetite, and epidote. Relatively coarse reddish-brown titanite/sphene is readily visible in hand sample. Plagioclase is generally restricted to the groundmass as very fine-grained granoblastic grains. In thin section the pyroxene is very fresh, bright green, sub-euhedral, and weakly zoned with roundish lighter green cores. Sphene forms small euhedral crystals, apatite forms thick irregular to subprismatic crystals, and minor proportions of biotite form strongly pleochroic light brown to deep brownish-green irregular books commonly intergrown with or included in pyroxene. Strongly aligned perthitic orthoclase forms blocky-rectangular crystals up to 7mm in length that are commonly Carlsbadtwinned. The groundmass matrix between the orthoclase and pyroxene is composed of fine-grained granoblastic feldspar that appears to have undergone brittle deformation; however within this granulated matrix are pristine pyroxene, sphene, and apatite crystals that show no evidence of shearing or rotation. This coupled with the apparent lack of shear bands on the outcrop implies that the granulation of the groundmass may have occurred during emplacement by semi-plastic deformation during upward flowage of the magma. Just north of the highway at the lowermost east end of this outcrop is an old adit that goes straight into the hillside; not sure as to when or why this was made. NEXT: Head back east on Hwy. 1 to Hwy. 73 • turn left / north for 5.2 miles to Highway 53. • Hang a right (go SW) on Hwy. 53 for 2 miles to where there are outcrops on the northeast side of the road. There is a driveway adjacent to this outcrop (on the north end) that would be a good place to park. Stop 5. ************WATCH OUT FOR TRAFFIC THIS IS A BUSY ROAD************ Gheen Pluton (NAD83: 515162, 5306034) (47.90745°, -92.79711°) The Gheen pluton is a spectacular example of multi-phase magma mingling textures. Strongly and coarsely porphyritic pyroxene syenite grades into, is cut by, and has inclusions of, medium-grained dark green hornblende gabbro to pyroxenite. Pink aplite and pegmatite forms the latest phase as small dikes that cross the other phases, and seems to have preferentially permeated the more mafic phases. Bluish chloritic slickenside surfaces look similar to those in the Linden pluton. The phenocrysts in the porphyritic phase at this stop are composed of braid-textured perthite to antiperthite in a matrix of bluish-green hornblende and prismatic actinolitic amphibole, abundant sphene, magnetite, and apatite, and anhedral to subpoikilitic saussuritized plagioclase. The aphyric mafic phases are composed varied combinations of pale green augite magmatic hornblende, secondary actinolitic amphibole, biotite, and accessory sphene, apatite, magnetite, and calcite. NEXT: Continue southeast on Hwy. 53 for about 14 miles, through the town of Cook, to County Road 467. • Turn left / east for 0.6 miles then veer right to stay on 467. Continue on 467 to railroad crossing; from there go another 0.75 miles to Forest Road 258D • Either drive or walk south on this road for 0.5 mile to an outcrop on the left / east in an overgrown clearcut, next to a logging trail that goes east. Stop 6. Idington Pluton (NAD83: 529108, 5287200) (47.73752°, -92.61175°) 132 �Trip 5 – Alkalic plutons The Idington (eye-ding-ton) pluton as not been dated, but it is intruded by the 2,674-Ma Shannon Lake granite along the southwestern margin of the pluton. This pluton is characterized by its coarsely porphyritic character as dramatically shown at this stop, and the phenocrysts are typically aligned by magmatic flow (trachytoid) in a crystal mush. In places the phenocrysts are randomly oriented, and in others oriented in a circular fashion that implies they were caught in an eddy. The mafic matrix is composed dominantly of aegirine or aegirine-augite and commonly has been squeezed out (i.e. filter-pressed) of the crystal mush to form small to large (10’s of meters) zones of an ultramafic phase. White ‘aplite’ dikes and some larger segregations cut the syenite and pyroxenite phases; these are interpreted as late residual differentiated melts that were squirted around through the semi-solid pluton. Very fine, delicate compositional zonation is visible in some of the phenocrysts here at this stop, where the rock is properly weathered. For a more thorough description of the pluton as a whole refer to the appropriate section of the introduction. The strikingly porphyritic syenite at this stop is typical of the Idington pluton although the phenocrysts are larger than normal. The phenocrysts are composed of perthite / antiperthite and NEXT: Go back north to the main road (467) and head east for 2.25 miles then follow road around bend to north (turns into County road 381) • Continue on 381 for 2.75 miles to Highway 1 • Turn right / east on Hwy. 1 for 3.25 miles to County Road 361 • Turn left / north and drive 1.5 miles to a small flat outcrop in the east ditch. There are also outcrops along the road on the way here one can stop at. Stop 7. Lost Lake Pluton (NAD83: 537263, 5293105) (47.79023°, -92.50248°) 2675.1±0.5 Ma The Lost Lake pluton is an irregularly-shaped intrusion that elongate to the east-northeast. The small outcrop in the road ditch shows mafic pyroxene-rich enclaves in pink syenitic phase. Outcrops nearby in the woods to the east demonstrate many different phases ranging from uniform pink to coarsely porphyritic to dark green and ultramafic. The ultramafic phases/enclaves in the road ditch outcrop are composed primarily of deep green (in thin section) aegirine as small equant to subprismatic crystals, a lesser proportion of larger blocky to subpoikitic biotite, and accessory sphene and apatite in a groundmass of poikilitic calcite (it fizzes) and minor sodic plagioclase. The pink portion is composed of allotriomorphic-granular mosaic of anhedral sodic plagioclase, perthite to antiperthite (commonly poikilitic), prismatic green aegirine, apatite, biotite, and interstitial calcite. At this stop we also will display some drill core from the Lost Lake pluton which demonstrates the multiple phases and diversity within this unit. NEXT: End of trip. Go back south to Highway 1 then east to Highway 169, then south back to Mountain Iron. Thank you for attending. 133 �Trip 5 – Alkalic plutons REFERENCES CITED Arth, J.G., and Hanson, G.N., 1975, Geochemistry and origin of the early Precambrian crust of northeastern Minnesota: Geochimica et Cosmochimica Acta, v. 39, p. 325-362. Barker, J.G., and Arth, J.G., 1976, Generation of trondhjemite-tonalite liquids and Archean bimodal trondhjemite-basalt suites, Geology, v. 4, p. 596-600. 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: Unpublished Ph.D. dissertation, University of Minnesota, Minneapolis. Boerboom, T.J., Jirsa, M.A., Southwick, D.L., Meints, J.P., and Campbell, F.K., 1989, Scientific core drilling in parts of Koochiching, Itasca, and Beltrami Counties, north-central Minnesota, 1987-1989: Summary of lithological, geochemical, and geophysical results: Minnesota Geological Survey Information Circular 26, 159 p. 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. Boerboom, T.J., Block, Amy Radakovich, Jirsa, M.A., Chandler, V.W., and Peterson, D.M., 2022, Bedrock Geology, pl. 2 of Jirsa, M.A., project manager, Geologic Atlas of Lake County, Minnesota: Minnesota Geological Survey County Atlas C-54, pt. A, 6 pls., scale 1:200,000 Card, K.D., and Ciesielski, A., 1986, DNAG #1. Subdivisions of the Superior Province of the Canadian Shield: Geoscience Canada, v. 13, p. 5-13. Catanzaro, E.J., and Hanson, G.N., 1971, U-Pb ages for sphene in northeastern Minnesota-northwestern Ontario: Canadian Journal of Earth Sciences, v. 8, p. 1319-1324. Chamberlain, K.R., Boerboom, T.J, and Bleeker, W., 2015: 2070 Ma dyke of southern Superior Province: a test of the radiating dyke model for the Kenora-Kabetogama/Fort Frances swarm, in Reconstruction of supercontinents back to 2.7 Ga using the large igneous province (LIP) record: with implications for mineral deposit targeting, hydrocarbon resource exploration, and earth system evolution; Supercontinent.org report number A194, 9 p. Chandler, V.W., 1991, Aeromagnetic map of Minnesota: Minnesota Geological Survey State Map Series S-17, scale 1:500,000. Deer, W.A., Howie, R.A., and Zussman, J, 1966, An introduction to the rock-forming minerals: London, Longman Group Limited, 528 p. Geldon, A.L., 1972, Petrology of the larnprophyre pluton near Dead River, in Sims, P.K., and Morey, G.B., eds., Geology of Minnesota: A centennial volume: Minnesota Geological Survey, p. 153-159. Himmelberg, G.R., 1973, Geologic descriptions of drill core from greenstone belts in northeastern Minnesota: Minnesota Geological Survey Open-File Report Irvine, T.N., and Baragar, W.R.A., 1971, A guide to the chemical classification of the common volcanic rocks: Canadian Journal of Earth Sciences, v. 8, p. 523-548. Jirsa, M.A., 1990, Bedrock geologic map of northeastern Itasca County, Minnesota: Minnesota Geological Survey Miscellaneous Map M-68, scale 1:48,000. Jirsa, M.A., and Boerboom, T.J., 1990, Bedrock geologic map of parts of Koochiching, Itasca, and Beltrami Counties, north-central Minnesota: Minnesota Geological Survey Miscellaneous Map series M-67, scale 1:250,000 / Jirsa, M.A., Boerboom, T.J., Chandler, V.W., and McSwiggen, P.L., 1991, Bedrock geologic map of the Cook to Side Lake area, St. Louis and Itasca Counties, Minnesota: Minnesota Geological Survey Miscellaneous Map Series M-75, scale 1:48,000. Jirsa, M.A., Southwick, D.L., and Boerboom, T.J., 1992, Structural evolution of Archean rocks in the western Wawa subprovince Minnesota: Refolding of pre­ cleavage nappes during D2 transpression: Canadian Journal of Earth Sciences, v. 29, p. 2146-2155. Martin, D.P., Meyer, G.N., Lawler, T.L., Chandler, V.W., and Malmquist, K.L., 1988, Regional survey of buried glacial drift geochemistry over Archean terrane in northern Minnesota: Minnesota Department of Natural Resources, Division of Minerals Report 252, V. 1, 74 p.; V. 2, 386 p. Meints, J.P., Jirsa, M.A., Chandler, V.W., and Miller, J.D., Jr., 1993, Scientific core drilling in parts of Itasca, St. Louis, and Lake Counties, northeastern Minnesota, 1989-1991: Summary of lithologic, geochemical, and geophysical results: Minnesota Geological Survey Information Circular 37, 159 p. 134 �Trip 5 – Alkalic plutons Mills, SJ., Southwick, D.L., and Meyer, G.N., 1987, Scientific core drilling in north-central Minnesota: Summary of 1986 lithologic and geochemical results: Minnesota Geological Survey Information Circular 24, 48 p. Prince, L.A., and Hanson, G.N., 1972, Rb-Sr isochron ages for the Giants Range granite, northeastern Minnesota: Geological Society of America Memoir 135, p. 217-225. Ruotsala, A.P., and Tufford, S.P., 1965, Chemical analyses of igneous rocks: Minnesota Geological Survey Information Circular 2, 87 p. Sage, R.P., 1988a, Geology of carbonatite-alkalic rock complexes in Ontario: ·Poohbah Lake alkalic rock complex, district of Rainy River: Ontario Geological Survey Study 48, 68 p. Sage, R.P., 1988b, Geology of carbonatite-alkalic rock complexes in Ontario: Sturgeon Narrows and Squaw Lake alkalic rock complexes, district of Thunder Bay: Ontario Geological Survey Study 49, 117 p. Sage, R.P., 1988c, Geology of carbonatite-alkalic rock complexes in Ontario: Wapikopa Lake alkalic rock complex, district of Kenora: Ontario Geological Survey Study 52, 63 p. 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 evolution of the southern Superior Province: Geological Society of America Bulletin, v. 118, p. 82-93. Sims, P.K., and Mudrey, M.G., Jr., 1972, Syenitic plutons and associated lamprophyres: in Sims, P.K., and Morey, G.B., eds., Geology of Minnesota: A centennial volume: Minnesota Geological Survey, p. 140-152. Sims, P.K., Morey, G.B., Ojakangas, R.W., and Viswanathan, S., 1970, Geologic map of Minnesota, Hibbing Sheet: Minnesota Geological Survey, scale 1:250,000. Sims, P.K., Sinclair, D., and Mudrey, M.G., Jr., 1972, Linden pluton: in Sims, P.K., and Morey, G.B., eds., Geology of Minnesota: A centennial volume: Minnesota Geological Survey, p. 160-162. Southwick, D.L., 1993, Geologic map of Archean bedrock, Soudan to Bigfork area, northern Minnesota: Minnesota Geological Survey Miscellaneous Map Series M-79, scale 1:100,000. Stern, R.A., Hanson, G.N., and Shirey, S.B., 1989, Petrogenesis of mantle-derived LILE-enriched Archean monzodiorites and trachyandesites (sanukitoids, in southwestern Superior Province: Canadian Journal of Earth Sciences, v. 26, p. 1688-1712. Streckeisen, A.L., 1973, Plutonic rocks: Classification and nomenclature recommended by the IUGS Subcommission on the Systematics of Igneous Rocks: Geotimes, v. 18, no. 10, p. 26-30. Wirth, K.R., Vervoort, J.D., and Heaman, L.M., 1995, Nd isotopic constraints on mantle and crustal contributions to 2.08 Ga diabase dykes of the southern Superior Province (abstract), Program & Abstracts for the Third International Dyke Conference, Sept. 4-8, 1995, Jerusalem, Israel, A. Agnon, G. Baer, 84, 1995. 135 �Trip 6 – Colvin Creek FIELD TRIP 6 Unique Keweenawan Inclusion (Colvin Creek) in the Duluth Complex Mark Severson (retired)1,2, Allison Severson3 and Laurie Severson (retired)4 1 (1988–2012) Natural Resources Research Institute, University of Minnesota, Duluth, 5013 Miller Trunk Hwy, Duluth, MN 55811 2 (2013–2018) Previously Teck American, then Teck Resources Unlimited, now NewRange (joint venture between Teck and PolyMet Mining Inc.) 3 Minnesota Geological Survey, College of Science and Engineering, University of Minnesota, 2609 Territorial Road, St. Paul, MN 55114 4 Earth Science Teacher, Woodland Middle School, ISD 709, Duluth, MN 55811 In memory of Richard Patelke 1957-2011 “Well, it ain’t my truck” 136 �Trip 6 – Colvin Creek INTRODUCTION Magnetic basalt inclusions within the Duluth Complex were first described by Bonnichsen (1974). Most of that description pertained to limited outcrops in what is now informally referred to as the South Colvin Creek Hornfels (Fig.6-1). Later, Tyson (1976) looked at four basalt inclusions including three nonmagnetic basalt inclusions in railroad cuts to the north, as well as the South Colvin Creek Hornfels. He concluded that the South Colvin Creek Hornfels was different and theorized that the magnetic basalts were derived from weathered and oxidized basalt flows, correlative with the North Shore Volcanic Group, and metamorphosed by the Duluth Complex. This field trip will visit the North Colvin Creek Hornfels (NCCH), shown in Figure 6-1, which is better exposed, contains several internal mappable units, and was first described by Severson and Hauck (1990). There they found a unique very fine-grained and cross-bedded unit, of gabbroic composition, within the inclusion. They initially theorized that the NCCH was formed as a result of magmatic currents (a concept they no longer support). It was also theorized that the cross-bedded unit represents a portion of a shear zone (Ojakangas and Holst, pers. com., sited in Patelke (1996)). This theory is also no longer deemed viable. Lastly, Patelke (1996) mapped and described the NCCH in more detail and proposed that it was an inclusion containing both metavolcanic and metasedimentary rocks that can be correlated with the North Shore Volcanic Group. This field trip will visit the NCCH which is referred to simply as the Colvin Creek Inclusion for the remainder of this guide. The thesis by Patelke (1996) is the source of almost all of this guide. GEOLOGIC SETTING The Colvin Creek inclusion is a large inclusion (2,500 X 800 meters), associated with a magnetic high, that has been rotated to a near vertical position and exhibits stratigraphic tops to the northwest as defined by pipe amygdules, sheeted amygdules, local convoluted flow bases and flow tops, and crossbedding. Patelke (1996) subdivided the inclusion into five major mappable units that include: two granoblastic, fine-grained metavolcanic units; two gabbroic sill units that bound the inclusion on the north and south; and a 350-meter-thick, cross-bedded, granoblastic, fine-grained metasedimentary unit of gabbroic composition. Overall, the inclusion strikes about N60°E with dips of 70-90° to the northwest. The entire inclusion has been metamorphosed to pyroxene grade facies and rotated to a subvertical position by the Duluth Complex. According to Miller and Severson (2005), the Colvin Creek inclusion is situated near the bottom of a “heterogeneous upper troctolitic cumulate” of the Partridge River intrusion (PRI). Geochemical work by Patelke (1996) indicate that metamorphism of the magnetic metabasalt units was isochemical and that they are probably equivalent to intermediate olivine tholeiites of the North Shore Volcanic Group (NSVG). The metasedimentary rocks are more problematic in that they are not analogous to any of the interflow sandstones of the NSVG as described by Jirsa (1980, 1984). At about 350 meters thick they are as thick as the total measured section of the NSVG interflow sedimentary rocks and show: no rock fragments; NO quartz, no conglomeratic horizons, and no intercalated volcanic rocks. Patelke (1996) suggested that the cross-bedded rocks were most likely deposited in a restricted basin as an eolian sediment that was derived from a strictly basaltic terrain – thus no quartz. Similar inclusions of crossbedded sediments with a gabbroic composition have been found at six locations within the Duluth Complex (i.e., geologic maps of the Babbitt SE and Babbitt SW quadrangles). Patelke (1996) thought that the informal “Phantom Lake sandstone,” an inclusion in the Whyte Quadrangle to the north of Two Harbors, was the most similar to the sediments in the Colvin Creek inclusion. While Patelke (1996) felt that these two units were similar, he concluded that neither of them can be strictly correlated with any other of the interflow sandstones in the Keweenawan system. 137 �Trip 6 – Colvin Creek Figure 6-1. Generalized geologic map of a portion of the Partridge River intrusion showing locations of the Colvin Creek Hornfels inclusions (gray) relative to the known Cu-Ni deposits. Base map from Miller and others (2001). GEOLOGY OF THE COLVIN CREEK INCLUSION Patelke (1996) mapped six major units associated with the Northern Colvin Creek Hornfels inclusion (exposed in Sections 27, 28, 33, and 34, T.59N., R.13W.). These units are briefly described below, and their distribution is shown in the geology map of Figure 6-2. The six units are, from south to north (also stratigraphically younging to the north) labeled as: MCC, AMG, AA, XBB, and GOG. These names are acronyms for field textures observed by Severson and Hauck (1990), and while these are not appropriate rock names, Patelke (1996) retained them in his thesis. MCC (Massive Colvin Creek unit) The MCC unit is a plagioclase-augite-oxide (titanomagnetite>ilmenite) rock, with local orthopyroxene and/or olivine, that under Phinney’s classification system (1972) is an oxide-bearing gabbro to augite troctolite. The MCC displays a massive, fine- to medium-grained texture similar to the units that stratigraphically overlie it but also shows primary decussate igneous texture. It is variably ophitic, and locally porphyritic. Granoblastic triple point junctions are reasonably well developed where the feldspar is 138 �Trip 6 – Colvin Creek Figure 6-2. Geology of the Northern Colvin Creek inclusion from Patelke, 1996. equant. Locally, there are ovoid clots of granular plagioclase that could be interpreted as amygdule infillings. The bottom contact of the MCC unit is not exposed. The upper contact with the AMG unit consists of rock types attributable to both MCC and AMG units within a 3-meter zone. For this reason, Patelke (1996) suggests that the MCC was injected sill-like and was mixed into the AMG while both were in a plastic state. Unfortunately, this particular exposure will not be visited during this trip. AMG (Amygdaloidal Gabbro unit) The AMG is stratigraphically above the MCC unit and is interpreted to be a subaerial metavolcanic unit with recrystallized amygdules. The rock is classed as an oxide melagabbro to augite troctolite. In the vast majority of the exposures, it is fine- to medium-grained and composed of plagioclase, augite, and oxide (titanomagnetite>ilmenite) with local orthopyroxene and poikilitic olivine. The AMG unit shows a persistent fine-grained, polygonal-granoblastic, sugary texture. In a few instances this texture is interrupted by clusters of plagioclase and by rounded to amoeboidal clot-like segregations of augite; both of which are lengthened parallel to the overall strike of bedding. These layers of pyroxene-rich segregations are interpreted to be recrystallized amygdules within relict volcanic flowtops. In areas of outcrop with common pyroxene-rich layers, the spacing of layers indicates flow thicknesses of 0.5 to 3.5 meters. The upper contact with the AA unit is exposed in only one outcrop (Fig. 6-3 - not visited this trip) wherein a black pyroxenemagnetite rich convoluted flowtop of the AMG is overlain by the base of a flow in the AA unit. 139 �Trip 6 – Colvin Creek Figure 6-3. Contact between the AMG (bottom, dark) and AA (top, light) units. The dark portion of the image is related to increased pyroxene and oxide content and interpreted to be a rubbly flow top that is abruptly overlain by the AA unit. MGC (Medium-Grained Gabbro unit) The MGC is a sill of very limited extent in the SW end of the Colvin Creek inclusion (Fig. 6-2). The rock is a medium- to coarse-grained orthopyroxene-bearing anorthositic gabbro according to the classification system of Phinney (1972). The sill crosscuts only the AMG unit, exhibits apparent chilled margins, and is estimated to be about two meters thick. Exposures of this unit will not be visited. AA (Amoeboidal Augite unit) The AA unit overlies the AMG unit and is also a fine- to medium-grained, massive, granoblastic magnetic basalt unit. At the outcrop scale, the AA is distinguished from the AMG by increased amounts pyroxene-filled avoids (amygdules) and by elongate pyroxene segregations that are interpreted as recrystallized pipe amygdules. Petrographically, the AA and AMG are very similar. Mineralogy consists of plagioclase, diopsidic augite, and titanomagnetite>ilmenite with local orthopyroxene (a major constituent in one outcrop). Plagioclase has a bimodal grain distribution consisting of fine-grained equant to stubby grains (0.25-1 mm) and patchy distributed laths (2-7 mm). The modal layering within this unit is defined by pyroxene stringers and ovoid clots that are interpreted to define both lava flow bases and amygdaloidal tops. Individual flows range from 0.5 to several meters thick. Elongate pyroxene masses lying perpendicular to strike are thought to be recrystallized pipe amygdules near the flow base (Fig. 6-4). 140 �Trip 6 – Colvin Creek Figure 6-4. Two flow units in the AA unit. Base of a single flow (6 inches to left of hammer head) with recrystallized, coalescing upward-trending pipe vesicles. Black wavy lines in extreme upper right of photo is the base of a third lava flow. The location of this outcrop is not documented. “Pyroxene interval” At the very top of the AA unit is a 0-2 meter thick, black, melagabbro unit, or “pyroxene interval” as mapped by Patelke (1996) in a few scattered outcrops. This unit consists of fine- to coarse-grained ferrosalite pyroxene and plagioclase. At one locality this unit contains: 1-10% brown garnet, 2-5% ilmenite>>titanomagnetite, and trace amounts of cordierite and hercynite. At one exposure (Fig. 6-5 – to be visited), there are several “veins” of potassium feldspar masses (up to 20-40 cm long by 1-10 cm wide), or tension gashes according to Patelke (1996). These “veins” are perpendicular to, and truncated by, the upper contact with the overlying XBB unit. The base of the XBB unit often exhibits a trough-like morphology downwards towards these feldspar masses. At several locations where this “pyroxene interval” is present, Patelke (1996) thought that there was some evidence of left-lateral tectonic movement. The tension gashes are one of his lines of evidence. Overall, Patelke (1996) thought that the “pyroxene interval” represents a deeply weathered flow top or soil developed on the AA unit and the effects of faulting are secondary. 141 �Trip 6 – Colvin Creek At another outcrop along the contact between the AA and XBB units (Stop 5 - to be visited), the “pyroxene interval” is absent. In its place are several sigmoidal-shaped pyroxene-rich lenses. Patelke (1996) thought that these lenses were developed along a bedding parallel fault. Figure 6-5. “Pyroxene Interval” (bottom 2/3rds of photo) between the AA and XBB units. Note convolute contact and k-spar-filled “tension gashes” as described by Patelke (1996). Figure 6-6. Typical cross-bedding exhibited by the XBB unit in a flat-laying outcrop. 142 �Trip 6 – Colvin Creek XBB (Cross-Bedded Belt) To the north of, and overlying the magnetic basalt units, is the Cross-Bedded Belt unit of roughly gabbroic composition. The rock is composed of fine-grained (1mm average) plagioclase-diopsideorthopyroxene-titanomagnetite>ilmenite with minor amount of orthopyroxene, hematite, hercynite, and geikielite. The rock exhibits beautiful bedding, cross-bedding (Fig. 6-6), density graded modal layering, and concave upward cross-beds along with scour and fill structures. Throughout the unit are localized minor biotite. There are several intervals, 0.5-3.0 meters thick, that are located near the base of the XBB that contain poikiloblastic pyroxene, up to several inches long (Fig. 6-7) that appear to have grown along bedding planes. The density graded modal layering consists of oxide- and pyroxene-rich basal layers grading upward into plagioclase-rich layers. Grain size for any individual mineral (1 mm) remains constant throughout the bed thickness. Angles of bedding and cross-bedding change over short distances in most outcrops. In some areas, the bedding exhibits a weak convolution or deformation (Fig. 6-7); possibly due to either soft sediment deformation and/or partial melting by the Duluth Complex. GOG (Gabbro-Olivine Gabbro unit) A gabbro to olivine gabbro unit, labelled as GOG, bounds the Colvin Creek inclusion at its upper (northwest) contact. The GOG was classed as a unit of the inclusion because it contains contact parallel layering (as does the inclusion) and shares strike-length and general tabular form with the other units of the inclusion. The GOG is medium to coarse-grained and composed of plagioclase (37-72%), augite (9-42%), olivine (0-23%), ilmenite (3-12%), and titanomagnetite (2-8%). The GOG unit is best exposed at the northwest end of the Colvin Creek inclusion where four contact-parallel zones were described by Severson and Hauck (1990) and by Patelke (1996). These zones (Fig. 6-2) are: A. weakly modally layered granular-textured augite troctolite zone with a plagioclase foliation B. phenocryst-rich gabbroic zone with anorthositic inclusions up to 5 inches across; C. a zone of heterogeneous gabbroic rocks with local inch-scale layering and cross-bedding (Fig. 6-8) indicative of magmatic currents; and D. a zone of anorthositic gabbro grading upward to gabbro with olivine gabbro interbeds. The trend of all of these zones parallel the contact trends of the underlying Colvin Creek inclusion. Unfortunately, this unit is too far away to bushwhack to Figure 6-7. Poikiloblastic pyroxene (black squares) in XBB unit. Note strange gain access and convolutions of bedding. visit during this trip. 143 �Trip 6 – Colvin Creek Common Characteristics of Colvin Creek Hornfels Listed below are characteristics common to all rock types of the Colvin Creek Hornfels: • All of the units are strongly magnetic and microscopically exhibit polygonal/granoblastic triple point junctions • Thin veins and later pods of pyroxene and/or massive magnetite are locally common. They are arranged in both parallel sets and discontinuous cross-cutting stringers • Thin, brown hornblende and/or orthopyroxene rims around titanomagnetite are commonly seen in thin-section • Sericitized plagioclase is seldom seen • Olivine, where present, is usually fresh and never serpentinized • Biotite is generally absent except in the XBB just above the “pyroxene interval” • The titanomagnetite is titanium-rich and the ilmenites are magnesium-rich. Figure 6-8. Rhythmic layering in GOG unit (subzone C) consisting of alternating olivine-rich and olivinepoor layers. Upper massive gabbro (hammer) truncates bed sets. FIELD TRIP STOPS Access starting from Mountain Iron will be heading south on Highway 53 through Virginia. Shortly after crossing the Tony Rukavina Bridge, over the Rouchleau Mine, turn and head east on Road 135 through the towns of Gilbert, Biwabik, and Aurora. Within Aurora, turn right at the stop sign and head south on CSAH 100, cross the railroad tracks, proceed to a stop sign and turn left on CSAH 110. Proceed to Hoyt Lakes on this highway. Within Hoyt Lakes continue straight through two stop signs and head out of town on Highway 110 (also called Skibo Vista Road). for about 4 miles. Just after passing the Bird Lake Recreation area, turn left onto road UT9235 (also called 569/Skibo Rd). Proceed down this road about 2.7 miles, cross the railroad tracks and continue east for another 1.9 miles. Turn left (north) on forest road 113 (yellow “share the road” sign at this intersection). Go 5.9 miles north on 113 to an unmarked logging road (another yellow ”share the road” sign at this intersection). Turn left on unmarked logging road (Figure 69) and head west about 1.5 miles depending on road conditions. Turn vehicles around, park as best as possible, and walk about 0.2 miles to the west (through an old beaver pond) to the first stop. Locations of the trip stops are shown in Figures 6-9 and 6-10. 144 �Trip 6 – Colvin Creek Figure 6-9. Access to Colvin Creek hornfels area and trip stops via remote logging road. Figure 6-10. Field trip stops (black dots) relative to mapped geology (modified from Patelke, 1996). 145 �Trip 6 – Colvin Creek Stop 1: MCC (Massive Colvin Creek unit) (NAD83: 577246E/5268002N) (47.56084°, -91.97315°) The MCC unit at this exposure is enigmatic. The rock is massive, lacks modal layering and regular concentrations of minerals. It is classed as a gabbro to augite troctolite composed of plagioclase, augite, orthopyroxene, olivine, and oxide (titanomagnetite>ilmenite). It is fine- to medium-grained with granoblastic triple point junctions. Locally, there are clots of granular plagioclase that could be interpreted as amygdule infillings, as in the overlying basaltic units. However, the MCC also displays primary decussate igneous textures, is variably ophitic, and locally porphyritic. For this reason, the distinction between the MCC and overlying AMG are often unclear. Patelke (1996) felt that the portions of the MCC were injected sill-like into the base of the inclusion while they were both in a plastic state. Directions: Continue down the road for about 2 minutes to a flagged trail off to the north. Follow the trail for another 5 minutes to Stop 2. Stop 2: AMG (Amygdaloidal Gabbro unit) (NAD83: 577230E/5268130N) (47.56199°, -91.97334°) At first glance, the AMG unit at this exposure is similar to the previous stop in that it consists mostly of massive, fine- to medium-grained “oxide gabbro.” However, within this exposure are several localized dark-gray, very fine-grained internal patches of basalt, that contain unquestionable plagioclasefilled amygdules. These patches exhibit gradational contacts with the surrounding medium-grained “oxide gabbro.” Thus, both fine-grained basalt and medium grained “gabbro” are present here (best seen after peeling a large area of the exposure). It is unknown whether these basalt patches represent true inclusions or are remnant unmetamorphosed patches in a rock that has undergone various degrees of partial melting to produce the “gabbroic” portions. Directions: Return to fork in flagged trail and continue north a few minutes to Stop 3. Stop 3: AA (Amoeboidal Augite unit) (NAD83: 577094E/5268083N) (47.56159°, -91.97515°) The AA unit overlies the AMG unit and is also a fine- to medium-grained, massive, granoblastic magnetic basalt unit. At the outcrop scale, the AA is similar to the AMG except for zones that contain common pyroxene-filled ovoids (recrystallized amygdules) and by pyroxene-rich horizons that are interpreted as sheeted amygdules and/or flow tops. Petrographically, the AA and AMG are very similar. Mineralogy consists of plagioclase, diopsidic augite, and titanomagnetite>ilmenite. Several basalt flows can be distinguished in portions of this outcrop based on massive flows grading upward (northward) into amygdule-rich basalt that in turn grades into pyroxene-rich flow tops. Individual flows range from over several meters to less than one meter thick. Note the presence of a cluster of coarse-grained pyroxene with minor K-spar (similar features will be seen at stop 6). Directions: Return to road and proceed further west for about 1-2 minutes to another flagged trail leading to the north. The Stop 4 exposure is about 100 feet north of the road on this trail. Stop 4: XBB Unit (Cross-Bedded Belt) (NAD83: 577065E/5268019N) (47.56101°, -91.97555°) To the north of, and overlying the magnetic basalt units, is the Cross-Bedded Belt unit of roughly gabbroic composition. This is the first of several exposures of the XBB unit that will be viewed during this trip. The rock is composed of fine-grained (1mm average) plagioclase-diopside-orthopyroxenetitanomagnetite>ilmenite. The rock exhibits beautiful bedding, cross-bedding, density graded modal layering, and concave upward cross-beds along with scour and fill structures. Note that NO quartz has ever been noted in this unit! 146 �Trip 6 – Colvin Creek There appears to be small-scale convolutions in the bedding trends that may be related to either soft-sediment slump or folding during intrusion of the Duluth Complex and subsequent rotation of the Colvin Creek inclusion. The location of this exposure along a curved mapped contact (Fig. 6-9) suggests that there is a small open fold between the AA and XBB units as suggested by Patelke (1996). Directions: Return to the road and head 3 minutes to the west to Stop 5 (about 50 feet north of the road). Stop 5: Contact of XBB and AA units (below photo) (NAD83: 576907E/5267935N) (47.56028°, 91.97767°) Both the AA and XBB units are present in this exposure. At the southern end of the exposure is a massive basalt unit that grades upward (northward) into a rock that contains abundant pyroxene-filled amygdules, which in turn, contains several pyroxene-rich lenses that represent sheeted amygdules and flow tops. Several flows are defined in the outcrop and the contact with the XBB unit is well defined (see Fig. 6-11). The overlying XBB unit consists of a fine-grained gabbroic rock with bedding planes similar to the previous stop but actual cross-bedding is not as striking. In regard to the contact between the two units, the intervening “pyroxene interval” is largely absent except for thin irregular pyroxene-rich lenses that display sigmoidal shapes. Patelke (1996) thought that sigmoidal-shaped pyroxene-rich lenses were developed along a bedding parallel fault with left-lateral movement. At the extreme north end of the exposure is an irregular, cross-cutting, massive oxide vein up to 3 inches wide. Figure 6-11. Contact between AA (left) and XBB (right) units with very poorly defined “pyroxene interval” in the contact zone. Note sigmoidal shapes of pyroxene layers at the contact. To the left of the contact (not in photo) are 2-3 trough-shaped zones (less than 2x3 feet) that contain bedded sediments similar to the XBB unit. Whether these zones are sedimentary interbeds or enfolded patches is unknown. 147 �Trip 6 – Colvin Creek Directions: Return to the road and proceed further west for about 10 minutes to Stop 6 on the southern edge of the road. On the way to Stop 6 there are numerous pavement road-crop exposures that consist mostly of massive magnetic basalt with local amygdules. Stop 6: AA (Amoeboidal Augite Unit) (NAD83: 576560E/5267549N) (47.55684°, -91.98234°) This outcrop is situated about 2,000 feet down the road from Stop 5 and serves more as a rest and regrouping stop. At this locale, the unit is massive and grades upwards (toward the road) into typical amygdaloidal basalt. Directions: Proceed down the road 350 feet and follow a flagged trail through the woods for about 840 feet westward (10 minutes). Stop 7: Contact of XBB & underlying AA unit (NAD83: 576259E/5267482N) (47.55628°, -91.98636°) This is the best exposure of “pyroxene interval” along the contact (see Figure 6-5 and description in text). Perpendicular to the contact, and wholly within the “pyroxene interval,” are at several “vein-like” potassium feldspar veins (up to 20-40 cm long by 1-4 cm wide) and a mass about 4 ft long by 1.5 ft wide. Patelke (1996) felt that the veins are tension gashes formed by lateral movement along the contact. The contact between the XBB and “pyroxene interval” exhibits some folding (soft-sediment?) with small-scale V-shaped troughs projecting downward into the “pyroxene interval.” Some of these troughs exhibit truncated bedding of the XBB against the “pyroxene interval” At one of the “V’s”, biotite, garnet and cordierite have been identified by Patelke (1996). At the extreme east end of the exposure, it appears that a bed of the XBB is folded(?) downward into the “pyroxene interval.” Patelke (1996) thought that the “pyroxene interval” represents a deeply weathered flow top or soil developed on the AA unit. Directions: At the west end of the Stop 7 exposure proceed northward for short distances (<100 feet) to several outstanding outcrops of the XBB unit of Stop 8. Stop 8: XBB Unit (Cross-Bedded Belt) (NAD83: 576248E/5267546N) (47.55685°, -91.98649°) Numerous exposures of beautifully cross-bedded XBB unit are present on the top of this hill. The rock is a very fine-grained granoblastic rock with a general modal composition of oxide-bearing anorthositic gabbro to gabbroic anorthosite. It is composed of plagioclase, diopsitic augite, and various iron-titanium-manganese oxides making up to 8-15% of the rock, NO quartz has ever been documented. As shown in Figures 6-6, 6-7 and 6-11, the rock is bedded and cross-bedded, exhibits density graded modal layering, concave upward cross beds, and scour and fill features. Some of the cross-beds show an unusually high angle of repose over very short distances possibly related to the environment of deposit (aeolian). 148 �Trip 6 – Colvin Creek Figure 6-12. Classic exposure of the XBB unit. Bedding tops to the north (right). Directions: Return to Stop 7 and proceed west for about 5 minutes to large exposures of the XBB and AA units. Note that between stops 8 and 9 is a glacial erratic of the GOG unit with stupendous inch-scale layering. This erratic is a good example of the GOG unit (otherwise inaccessible on this field trip). Stop 9: AA (Amoeboidal Augite Unit) (NAD83: 576167E/5267425N) (47.55578°, -91.98759°) After crossing over a large outcrop of the XBB unit, proceed southward a short distancer to a large tip over exposure (uprooted and wind-fallen tree) of the AA unit consisting of multiple basalt flows with ropey tops. This outcrop is present near the upper contact of the unit and small exposures of the XBB are present to the north and west. Pipe vesicles are present in one small area of the AA unit. Also present is a very small exposure of the “pyroxene interval.” Directions: Return to vehicles. Return to Mountain Iron Community Center (47.51869°, -92.58997°). References Bonnichsen, B., 1972, Southern Part of the Duluth Complex. In: Sims, P.K. and Morey, G.B. (eds), Geology of Minnesota – A Centennial Volume, Minnesota Geological Survey, p. 361-388. Jirsa, M.A., 1980, The Petrology and Tectonic Significance of Interflow Sediments in the North Shore Volcanic Group, Northeastern Minnesota, unpublished M.S. Thesis, University of Minnesota Duluth, 125 pages. Jirsa, M.A., 1984, Interflow Sedimentary Rocks in the Keweenawan North Shore Volcanic Group, Northeastern Minnesota: Minnesota Geological Survey, Report of Investigations 30, 20 p. 149 �Trip 6 – Colvin Creek 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. Miller, J.D., Jr. and Severson, M.J., 2004, Geology and Mineralization of the Western Contact of the Duluth Complex, Partridge River and South Kawishiwi intrusions, Northeastern Minnesota: Institute on Lake Superior Geology, 50th Annual Meeting, Duluth, MN, Part II: Field Trip Guidebook, p. 227-258. Miller, J.D., Jr. and Severson, M.J., 2005, Patelke, R.L., 1996, The Colvin Creek Body, A Metavolcanic and Metasedimentary Mafic Inclusion in the Keweenawan Duluth Complex, northeastern Minnesota: unpublished M.S. Thesis, University of Minnesota, 232 p. 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. 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/GMIN-TR-89-11, 236p. (with plates). Tyson, R. M., 1976, Hornfelsed Basalts in the Duluth Complex: unpublished M.S. Thesis, Cornell University, Ithaca, New York, 85 p. 150 �Trip 7 – Classic Outcrops FIELD TRIP 7 Classic Outcrops of Northeastern Minnesota Dean M. Peterson1 and George J. Hudak2,3 1 Big Rock Exploration, 2505 W. Superior St., Duluth, MN 55806 George Hudak Geosciences P.L.L.C., Duluth, MN 55804 3 Department of Earth and Environmental Sciences, University of Minnesota Twin Cities, 116 Church Street SE, Minneapolis, MN 55455 2 Introduction This field trip will investigate a wide variety of Neoarchean, Paleoproterozoic and Mesoproterozoic rocks that illustrate the diversity of Precambrian rocks in northeastern Minnesota. The field trip is an updated version of “Field Trip 5 – Classic Outcrops of Northeastern Minnesota” that was run during the 50th Annual Meeting of the Institute on Lake Superior Geology that took place in Duluth, Minnesota during May, 2004. As such, several of the field trip stop descriptions in this guidebook are derived from this earlier field trip guide, with updates based on recent geological studies. Generalized Stratigraphy of Northeastern Minnesota Neoarchean Vermilion District 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 WawaAbitibi 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 7-1. The Soudan belt (Figures 7-1 and 7-2) 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 7-1) include: (1) the Lower member, Soudan Iron-Formation member, and Upper member (Upper Ely) of the Ely Greenstone 151 �Trip 7 – Classic Outcrops Figure 7-1. 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. Figure 7-2. Generalized geology and geochronology of the Vermilion District in the vicinity of the Tower-Soudan anticline (modified after Peterson, 2001; Hudak et al., 2014). 152 �Trip 7 – Classic Outcrops 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, Peterson, 2001) and the 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). Regional chronostratigraphic correlations between the Wawa Greenstone (northwestern Ontario) and the Abitibi greenstone belt (eastern Ontario and Quebec) are indicated in Figure 7-3. Geochronological information for supracrustal and intrusive lithologies in the Vermilion District is relatively sparse (Figure 7-3). Peterson et al. (2001) obtained a U-Pb zircon age 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. Allerton et al. (2024a) obtained a crystallization age of 2708 ± 25 Ma for the Purvis Pluton, which intrudes the Eagles Nest Succession of the Lower Ely Member, and has been interpreted as a synvolcanic intrusion (Peterson, 2001). The age of the Upper Member of the Ely Greenstone formation is currently unknown. Jirsa (2016) obtained an age of 2715.74 ± 0.50 Ma for a felsic volcanic unit within the Newton Lake Formation (Boerboom, T. J., 2020). Lodge et al. (2013) obtained a U-Pb zircon age 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 ages ranging from 2680-2690 Ma from greywackes that comprise the Lake Vermilion formation. These dates confirm the source of the detritus in the Lake Vermilion Formation was derived locally from the volcaniclastic rocks comprising the Gafvert Lake Sequence. Figure 7-3. Regional chronostratigraphic correlations between the Vermilion district (Minnesota), the Wawa greenstone belt (northwestern Ontario), and the Abitibi greenstone belt (eastern Ontario and Quebec; after Ayer, 2010). Table 7-1. Lithostratigraphic units within the western Vermilion District (modified after Peterson and Jirsa, 1999; Peterson et al., 2009; Hudak et al., 2012). 153 �Trip 7 – Classic Outcrops Intrusive Rocks Late Intrusions 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). Vermilion Granitic Complex Granite, schist, amphibolite, and schist-rich migmatite Giants Range Batholith Granite, granodiorite, monzodiorite, and schist-rich migmatite. U-Pb zircon dates indicate a crystallization age ranging from 2640-2777Ma (Allerton et al., 2024a). Supracrustal Rocks Newton Belt Newton Lake Formation Tholeiitic and komatiitic basalt lava flows, intrusions, and clastic strata (deep subaqueous?) Bass Lake Sequence Tholeiitic basalt lava flows, iron-formation, and felsic porphyries (deep subaqueous) Soudan Belt Knife Lake Group Graywacke, slate, conglomerate, and sheared equivalents Lake Vermilion Formation Graywacke, slate, dacitic tuff, minor conglomerate. Detrital zircons from planar bedded, normal-graded resedimented volcaniclastic rocks have UPb age dates of 2680-2690 Ma (Lodge et al., 2013; subaerial to subaqueous) Gafvert Lake Sequence Dacitic to rhyodacitic tuff, lapilli-tuff, tuff-breccia, and iron-formation. Basal dacite tuff-breccia deposits in Lake Vermilion State Park have UPb age date of 2689.7 ± 0.8 Ma (Lodge et al., 2013; subaerial to subaqueous) Britt Sequence Tholeiitic basalt lava flows (deep subaqueous?) Upper Member – Ely Greenstone Tholeiitic basalt lava flows and iron-formation (deep subaqueous?) Soudan Member – Ely Greenstone Oxide-facies iron formation with intercalated basalt lava flows and felsic volcaniclastic rocks (deep subaqueous) Lower Member – Ely Greenstone Calc-alkaline and tholeiitic basalt-rhyolite lava flows, tuffs, epiclastic rocks, and minor iron-formation (shallow- to deep subaqueous) Central Basalt Sequence 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 (transition from shallow- to deep water environment) Fivemile Lake Sequence 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; interpreted as shallow subaqueous environment). Eagles Nest Sequence Algoma-type iron formation, basalt-andesite lava flows, hydrothermal exhalites, felsic tuffs. 154 �Trip 7 – Classic Outcrops 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. 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. 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. The age of the orebodies at the Soudan Mine were previously interpreted to be syn- or postdepositional to the precipitation of the Soudan Member of the Lower Ely Greenstone Formation (Gruner, 1926; Klinger, 1960; Thompson, 2015). Recent U/Pb and (U-Th)/He radiometric dating by Allerton (2024b) suggest the massive hematite orebodies at Soudan formed during Paleoproterozoic time (1640.8 ± 47.2 Ma – 1740.4 ± 72.5 Ma) and have been overprinted by a Mesoproterozoic hydrothermal event at approximately 1100 Ma (1093.1 ± 16.4 Ma). 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 Timiskamingtype 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 (Figure 7-2). 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. Paleoproterozoic Superior Type Iron Resources of the Mesabi Iron Range Superior type iron formation resources of Minnesota are exemplified by the long-standing mining of iron resources of the Biwabik Iron Formation along the length of the Mesabi Iron Range. The Mesabi Iron Range is largely located in St. Louis and Itasca counties and has been the most important iron ore district in the United States since ~1900. The Mesabi Iron Range is 120 miles long, averages one to two miles wide, and is comprised of rocks of the Paleoproterozoic Animikie Group. The Animikie Group on 155 �Trip 7 – Classic Outcrops the Mesabi Iron Range consists of three major conformable formations: Pokegama Formation at the base; Biwabik Iron Formation in the middle; and the overlying Virginia Formation. On the Mesabi Iron Range, these three formations generally dip gently to the southeast at angles of 3-15 degrees. Since the early 20th century, the Biwabik Iron Formation has been subdivided into four informal members referred to as (from bottom to top): Lower Cherty member, Lower Slaty member, Upper Cherty member, and Upper Slaty member (Wolff, 1917). The cherty members are typically characterized by a granular (sand-sized) texture and thick-bedding (beds ≥ several inches thick); whereas the slaty members are typically fine-grained (mud-sized) and thin-bedded (≤1 cm thick beds). The cherty members are largely composed of chert and iron oxides (with zones rich in iron silicate minerals), while the slaty members are composed of iron silicates and iron carbonates with local chert beds. Both cherty and slaty iron-formation types are interlayered at all scales, but one rock type or the other predominates in each of the four informal members, and they are so-named for this dominance Severson et. al. (2009). Leached and iron enriched direct ores (or natural ores) were the first materials mined, with the first shipments beginning in 1892, from strongly oxidized pockets along fault and fracture zones and the blanket oxidation of the iron formation at the surface. Taconite, which is the material that is mined today using magnetic separation methods, constitutes most of the iron formation and pertains to the hard, non-oxidized portions of the iron-formation. Production has been dominantly controlled by vertically integrated steelmakers since 1901, and therefore the mining and utilization of these ores have been dictated largely by US ironmaking capacity and demand. The taconite typically contains 30-35% iron and 40-50% SiO2, plus other components (Morey, 1992). The Biwabik Iron Formation is around 175-300 feet thick in the extreme eastern end of the Mesabi Iron Range at Dunka Pit, 730-780 feet thick in the central Mesabi Iron Range/Virginia Horn area near Eveleth, around 500 feet thick in the western Mesabi Iron Range near Coleraine, and eventually exhibits a “nebulous ending about 15 miles southwest of Grand Rapids” (Marsden et al., 1968) on the extreme western end of the Mesabi Iron Range. Maps of currently active taconite mining operations on the Mesabi Iron Range are presented in Figure 7-4. Mesoproterozoic 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 (Figure 7-5). 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 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. 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. 156 �Trip 7 – Classic Outcrops Figure 7-4. Bedrock geology and iron mining features of the Mesabi Iron Range. Felsic series—Massive granophyric granite and smaller amounts of intermediate rock that occur as a semicontinuous 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). 157 �Trip 7 – Classic Outcrops 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). South Kawishiwi Intrusion The South Kawishiwi intrusion (SKI), together with the similar sized Partridge River intrusion (PRI) immediately to the south, are most renowned for hosting the largest tonnage of Cu-Ni sulfide mineralization in the world (Naldrett, 1997). The realization that the SKI hosts vast quantities of Cu-Ni mineralization over 50 years ago has led to the publication of numerous geologic maps, (Green et al., 1966, Bonnichsen, 1974, Foose and Cooper, 1974, Miller et al., 2001, Peterson, 2002e, f, Peterson et al., 2004, Peterson, 2006b, Peterson et al., 2006), articles (Bonnichsen et al., 1980, Weiblen and Morey, 1980, Ripley, 1986, Chandler and Ferderer, 1989, Lee and Ripley, 1996, Hauck et al., 1997, Peterson, 2001b) theses (Weiblen, 1965, Vislova, 2003, Marma, 2003, Gal, 2008, White, 2010), and reports (Phinney, 1969, Phinney, 1972, Listerude and Meineke, 1977, Morey and Cooper, 1977, Foose, 1984, Dahlberg, 1987, Figure 7-5. Geologic map of northeastern Minnesota. 158 �Trip 7 – Classic Outcrops Dahlberg et al., 1989, Kuhns et al., 1990, Severson, 1994, Zanko et al., 1994, Hauck et al., 1997, Peterson, 1997, Peterson, 2001c, Miller et al., 2002, Peterson, 2002d, Patelke, 2003, Severson and Hauck, 2003). The SKI is shallow dipping (~20º to the east-southeast) sill-like intrusion dominantly composed of troctolitic cumulates that are exposed in an 8 x 32-km arcuate band along the northwestern margin of the Duluth Complex. Footwall rocks include the Paleoproterozoic Virginia Formation in the Serpentine and Dunka Pit deposits, the Paleoproterozoic Biwabik Iron Formation in the Dunka Pit and Birch Lake deposits, and the Archean Giants Range batholith from the northern Birch Lake deposit north to the Spruce Road deposit. The presence of shallow-dipping Biwabik Iron Formation inclusions 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 SKI, leaving the Giants Range batholith as the dominant footwall rock type. Alternately, the Virginia and Biwabik Iron Formations may simply have been largely eroded prior to the development of the Mid-Continent Rift. Also present as inclusions in the SKI are mafic volcanic hornfels (North Shore Volcanic Group), quartz sandstone hornfels (either the Puckwunge or Nopeming sandstones), and anorthosite (of the Anorthosite series). Anorthositic series rocks about the SKI on the northeast – and enclose an interpreted SKI feeder dike (the NLM) that extends farther northeast – the PRI forms the southern sidewall of the SKI, and the BEI and Anorthositic series rocks overlie the SKI to the east. On the regional Duluth Complex map of Miller et al. (2001), the SKI is subdivided into five major map units. These are, from the base upward, 1. Heterogeneous sulfide-bearing troctolite, gabbro, and norite with localized hornfels inclusions, 2. A thick unit of subophitic to ophitic augite troctolite, 3. Discontinuous and localized layers of poikilitic leucotroctolite, 4. A thick homogeneous sequence of ophitic troctolite, and 5. A thick uppermost sequence of homogeneous troctolite that contains numerous anorthositic layers. Severson (1994) and Zanko et al. (1994) further subdivided the SKI into 17 different lithostratigraphic units that are present in over 180 drill holes over a strike length of 31 kilometers. Sulfide mineralization is confined to the BH, BAN, UW, and U3 units near the base of the intrusion, and to a lesser extent the U1, U2, and PEG units. Major marker horizons that are correlated in drill holes include three horizons with abundant cyclic ultramafic layers (U1, U2, and U3 units) and a pegmatite-bearing unit (PEG unit) that was initially recognized by Foose (1984). The understanding of the significance of a large anorthositic inclusion, originally intersected in six deep drill holes east of the Maturi deposit, and its role in magma dynamics of the SKI has been a key feature in the development of an exploration model for Duluth Metals Limited’s Maturi Extension deposit (Peterson, 2001c). Terminology It is important to note the terminology utilized in this field trip guide for: 1) volcaniclastic rocks; 2) bedding characteristics; and 3) description and unit coding of outcrops in the Duluth Complex. 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); 159 �Trip 7 – Classic Outcrops • • 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 due to contact with water. Deposition of such deposits is is influenced by the continued emplacement of the lava in the presence of water, and the thicknesses of the hyaloclastite deposits can be dictated by the temperature of the magma, the effusion rate, and the distance from the volcanic vent (Cas and Wright, 1987; Gibson et al., 1999; Newkirk et al., 2001a, 2001b); and Peperite deposits, which are generated when magma intrudes into unconsolidated clastic material and mingles with (generally wet) debris to form a volcaniclastic deposit (McPhie et al., 1993). 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. The terminology for volcaniclastic rocks has historically been somewhat confusing because many different classification schemes have been developed (for example Fisher, 1961; Fisher 1966; Schmid, 1981; Cas and Wright, 1987; McPhie et al., 1993; White and Houghton, 2006), and different classification schemes are preferentially used in different parts of the world. As a result, the terminology relating to volcaniclastic rocks is commonly misused or misinterpreted. Four classification schemes that have been used most in the recent geological literature include: • • • • 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”. 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!” Also, classification is especially difficult in ancient volcaniclastic rocks because key aspects 160 �Trip 7 – Classic Outcrops 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). For this field trip guidebook, we will utilize Fisher’s (1966) classification (Figure 7-6) for volcaniclastic rocks. This classification scheme is based on the relative proportions of ash-sized material (< 2mm), lapilli-sized material (2-64mm), 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. Figure 7-6. 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. More recently, White and Houghton (2006) have developed a modified version of Fisher’s (1966) volcaniclastic classification scheme (Figure 7-6). 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). Specific terms for bedding thicknesses are also used in this guidebook. The terminology for bedding thickness has been adopted from McPhie et al. (1993) and includes: • • • • Laminated Very thinly bedded Thinly bedded Medium bedded <1 centimeters thick 1-3 centimeters thick 3-10 centimeters thick 10-30 centimeters thick 161 �Trip 7 – Classic Outcrops • • Thickly bedded Very thickly bedded 30-100 centimeters thick >100 centimeters thick Figure 7-7. The classification scheme used to describe and code mafic intrusive rocks within the Duluth Complex, modified after Phinney, 1972. Classification of outcrops and map units within the Duluth Complex have relied on the early work of William Phinney (Green et al., 1966; Phinney 1969, and Phinney, 1972) and is given in Figure 7-7. 162 �Trip 7 – Classic Outcrops FIELD TRIP STOPS Table 7-2. Simplified description of the twenty-field trip stops presented in this guidebook. Also included are the coordinates (UTM, Nad83, Zone 15N and Lat-Long), mileage along the route to the stop, and the age of features that will be observed and discussed on the outcrops. Stop 1: Laurentian Divide at Confusion Hill Longitude/Latitude: 47.51868699°N, -92.58996713E UTM NAD 83 Zone 15N: 530870E, 5262888N 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. Dikes of tonalite that cut the adjacent high-grade supracrustal rocks of the Minntac sequence contain metamorphic fabrics, yet little evidence of metamorphic origin can be seen in the interior of the body, implying it is syntectonic 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. 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 40-mile wide 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 Tower-Soudan area - making stratigraphic correlation between the two districts speculative. Return to bus. 163 �Trip 7 – Classic Outcrops Figure 7-7. Simplified bedrock geology map overlain by the field trip stops and traveled route. Stop 2: Pike River Dam Greywackes Longitude/Latitude: 47.57736062°N, -92.54367719E UTM NAD 83 Zone 15N: 534317E, 5269428N This glacially scoured outcrop exposes a nearly perfect cross-section of straight-bedded, variably graded, feldspathic graywacke and black slate. The feldspar-rich, dacitic composition of sandy textured beds is presumed to represent derivation from the Gafvert Lake felsic volcanic sequence exposed to the east in the Soudan area. Regionally, a series of outcrops from Gafvert Lake westward shows an irregular transition from proximal, possibly subaerial deposition on the east, to distal submarine turbiditic fan deposition to the west. The beds contain numerous "soft-sediment" deformation features including load structures, flames, intrafolial slump folds, and possibly some of the cross-stratal faulting. Bedding is nearly vertical, and graded beds indicate stratigraphic younging to the south. This topping direction, and the presence of a weak D2 cleavage that is left of bedding, indicate westward structural facing in the cleavage; consistent with a position on the south limb of a large, south-overturned, regional, D1 fold structure—the western extension of the Tower–Soudan Anticline. Northeast-trending kink bands, fault zones, and raised quartz veins traversing the outcrop. One of the truly classic outcrops of greywacke of the Lake Vermilion Formation is beautifully exposed at this stop. Prior to about the 1950s, no depositional mechanism could satisfactorily explain the coincidence in graywacke of; 1) coarse sand derived from a source many kilometers distant and having an altered clayey matrix; 2) interbedded black slate; and 3) the lack of evidence for reworking in shallow water (indicative of deposition below wave base). This was changed when the concept of turbidity currents was introduced to the geological profession by Kuenen and Migliorini (1950). Despite widespread publication on turbidites in more modern geologic settings through the 1950s and 1960s, the facies model was not 164 �Trip 7 – Classic Outcrops refined and applied to Archean and Proterozoic strata in the Lake Superior region until somewhat later (Morey, 1965; Ojakangas, 1966). Return to bus. Stop 3: Gafvert Lake Sequence Volcaniclastic Rocks Longitude/Latitude: 47.80135914°N, -92.28615141E UTM NAD 83 Zone 15N: 553454E, 5294469N This relatively new roadcut (approximately 15 years old) exposes rhyodacitic to dacitic composition Gafvert Lake Sequence lapilli tuffs and tuff breccias. The deposits have tentatively been interpreted to represent mass flow units produced by slumping of volcanic and volcaniclastic material from the Gafvert Lake volcano into an adjacent, probably submarine basin. Close inspection of the unit indicates the presence of a variety of lapilli and blocks including: 1) subrounded to subangular plagioclase ± quartz-phyric coherent dacite to rhyodacite; 2) subrounded to subangular pumice; 3) angular carbonate-rich fragments; 4) angular chert fragments; and 5) local subangular to angular massive sulfide fragments. Locally, abundant (up to 10%) <1mm euhedral pyrite cubes are disseminated in the matrix. The presence of both carbonate and massive sulfide fragments, as well as plagioclase- and quartz phyric coherent rhyodacite to dacite lapilli, may suggest the slumps are derived from a Gafvert Lake sequence subaqueous lava dome that was affected by local hydrothermal alteration and the deposition of chemical exhalates (e.g. carbonate, chert and massive sulfide fragments). Structurally, this outcrop occurs on the southern margin of an east-southeast – west-northwest trending D2-associated structure that extends from Pike Bay (northwest) to south of Putnam Lake (southeast). Here one can observe a strong, steeply dipping E-NE foliation and a well-developed lineation that plunges approximately 70° E. Return to bus. Stop 4: Soudan Member Banded Iron Formation Longitude/Latitude: 47.820074°N, -92.2365908E UTM NAD 83 Zone 15N: 557144E, 5296585N (NOTE: Modified from Peterson et al., 2009 and Hudak and Peterson, 2014.) This classic exposure of the Soudan iron-formation member of the Ely Greenstone Formation lies on the north limb of the Tower-Soudan anticline approximately 75 meters north of the stratigraphic top of the volcanic sequences known collectively as the Lower member of the Ely Greenstone. The outcrop displays two generations of tight folding in delicate laminae of chert (creamy white), chert-hematite jasper (red), and magnetite-chert (black to silver-colored). The second generation of folds (F2) is tectonic in origin, having subvertical axial surfaces that trend east, and steeply plunging axes. Most display Z-asymmetry. The earlier folds (F0-1) appear to have been sharply refolded to produce complex interference patterns. Lundy (1985) studied folding at this locality and concluded that some of the apparent interference structures are the product of early-formed sheath folds that did not involve refolding by D2. The F1 structures are predominantly intrafolial and exhibit a great variety of style and orientation; implying they formed by layerparallel, soft-sediment slumping (Fig. 7-8). Lundy’s mapping of this outcrop is an interesting demonstration of unraveling details at a single outcrop that led to recognition that D1 deformation was not systematic here, but likely soft sediment. Furthermore, it is a microcosm of regional-scale deformation. 165 �Trip 7 – Classic Outcrops It is interesting to observe the rhythmic microlaminae (1 mm or so thick) in various cherty beds exposed here and speculate about the paleoenvironment—that is, whether these represent daily heating/cooling, tidal, climatic, annual, or some other repetitive influence (e.g waxing/waning of a hydrothermal system) in the depositional environment. What is known about units of iron-formation in the Ely Greenstone, of which there are many, is that deposition occurred in deep water (below wave base) during periods of relative volcanic and tectonic quiescence by the slow subaqueous “rain” of chemical precipitates. The deep excavations in this area are the early workings of the Soudan iron mine, the first in Minnesota. The mine produced about 16 mt of high-grade hematite ore (60-63 percent iron converted to a park. Most of the production came from underground workings that began here in 1900, and which now can be visited on guided tours. The mine previously housed several underground physics research facilities. These include Soudan 1 (23rd level) which studied neutrino decay; 2) Soudan 2 (27th level), also to study neutrino decay; and 3) the MINOS (Main Injector Neutrino Oscillation Search) lab, which was built on the 27th level adjacent to Soudan 1 and studied the decay of neutrinos within the earth as they passed from Fermilab to Soudan (Peterson et al., 2009b). Figure 7-8. Outcrop map showing bedding trajectories and multiple generations of folds and faults (from Lundy, 1985). F1 folds are non-systematic and include both nappe- and sheath fold geometries. Return to bus. Stop 5: Murray Shear Zone Along Hwy 1/169 Longitude/Latitude: 47.81809993°N, -92.20694376E UTM NAD 83 Zone 15N: 559366E, 5296388N A series of roadcuts along Highway 169 expose a transect through the northern edge of the Murray Shear Zone, which is one of the most striking Neoarchean structural features in the Tower-Soudan area 166 �Trip 7 – Classic Outcrops (Peterson and Patelke, 2003). This series of outcrops perfectly display a classic feature of Neoarchean ductile (shear zone) structures - strain partitioning (Figure 7-9). A close look at these outcrops also gives one hints of broader economic geology implications via the presence of carbonate alteration of the chlorite schists. The carbonate (ankerite and/or ferro-dolomite) strain hardens the ductile deformed schistose rocks and allows for subsequent brittle deformation (and perhaps orogenic gold mineralization in cross-cutting quartz-ankerite-sulfide veins. On the larger scale, the D2 Murray shear zone transposes rocks 3-5 km eastwards in the zone bounded by its northern and southern strain partitioned boundaries. The overall geometry of this panel of Figure 7-9. Scanned image of the field sheet used to map outcrop OC-567. On the right are digital photographs of outcrop OC-567: A) the overall outcrop view looking WNW; B) view to the north of steeply east plunging, lineated and rod-shaped pillowed andesite (rock hammer 68cm for scale); and C) close-up view of rock sample S-604, taken from the west side of the outcrop (bright zone on the left side of picture A). Data from Peterson & Patelke, 2003. 167 �Trip 7 – Classic Outcrops rocks is like the geometry of “wedge-shaped shear zones” described in detail by Ramsey and Huber (1987). Peterson’s mapping and collection of structural data in the Murray shear zone panel is largely confined to a series of outcrops along the northern edge of the zone. Field observations of these outcrops indicate that the strain symmetry along this boundary is largely constrictional, with a dominant steeply east-dipping, elongate and rod-shaped structural fabric. A stereonet projection of planar and linear structural features within the Murray panel is shown in Figure 7-10. The mean value of the strike and dip of planar features is 282°/82°, and the trend and plunge of linear features has a mean orientation of 87°/71°. The overall mapscale internal geometry of the Murray panel clearly shows dextral asymmetry, with a strong sigmoidal wrapping of iron-formation (see field trip geologic maps) to the northeast. Figure 7-10. Stereonet projections of foliation, shear fabrics, and linear features from the Murray shear zone. An estimate of the amount of displacement of the rocks within the panel of rocks bounded by the Murray shear zone is given in Table 7-3. These values were calculated geometrically by using the average plunge of measured lineations (71°) and two measured lines of possible correlative stratigraphy offset by the bounding shear zones. The calculated total displacement values (net slip) are quite large (up to 13.8 km, or 43,000 feet of net slip), but the displaced rocks would still fall within the range of depth generally associated with greenschist facies metamorphism. Table 7-3. Calculated displacement along the Murray shear zone Strike Slip Lineation Plunge Dip Slip Net Slip 71° 4.5 13.1 13.8 71° 3.0 8.7 9.2 Return to bus. 168 �Trip 7 – Classic Outcrops Stop 6: Lower Ely Member (Central Basalt) Sheet and Pillowed Flows Longitude/Latitude: 47.8306566°N, -92.17157352E UTM NAD 83 Zone 15N: 561999E, 5297811N (NOTE: Modified from Field Trips of Hudak et al., 2004, 2014; Peterson et al., 2005, 2009). This classic outcrop has been visited during field trips associated with the 2004, 2009 and 2014 ILSG conferences (Hudak et al., 2004; Peterson et al., 2009; Hudak et al., 2014). 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 have been modified from Peterson et al. (2009) and Hudak et al. (2014). The Central Basalt sequence (Peterson and Patelke, 2003, Peterson, 2005) 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. (2009), 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 scoria-rich 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. At this stop, 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. 7-11). All three lava flows at this vicinity illustrate tholeiitic, MORB-like lithogeochemistries (Hudak et al., 2007). 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 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 bun- shaped 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 169 �Trip 7 – Classic Outcrops Figure 7-11. Detailed geological map of sheet flows, pillow lavas, and associated hyaloclastite deposits at Field Trip 3, Stop 1 (after Hudak et al., 2014; Hudak and Peterson, 2014). 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% subrounded 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 170 �Trip 7 – Classic Outcrops 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 psuedo-pillows that are up to 50 cm in diameter. Return to the bus by walking back down the hill. Stop 7: Mud Creek Shear Zone Longitude/Latitude: 47.87440908°N, -92.14025702E UTM NAD 83 Zone 15N: 585753E, 5309482N This outcrop shows highly strained rocks in the Mud Creek shear zone. The rock type is a quartziron carbonate-sericite schist, having quartz and tourmaline knots, abundant pyrite, and trace amounts of gold. Its protolith is unknown, because of the intense deformation, but could be any of several rock types in the region, including quartzofeldspathic porphyry, basaltic metavolcanic rocks, or graywacke. The shear fabric trends east-northeast, and lineations plunge at shallow angles to the east. Development of this shear zone, which occupies most of the valley of Mud Creek, is a product of largely dextral transpressive deformation that has been partitioned into discrete zones, presumably late in D2 deformation. It is generally believed that gold-bearing mineralization was introduced during these later deformation events, and the Mud Creek shear zone and environs have attracted considerable attention as a gold target (Peterson, 2001, Peterson and Patelke, 2004a, 2004b). The Mud Creek shear is a broad, anastomosing zone that forms the boundary between rocks of the Ely Greenstone and Lake Vermilion Formation on the south, and volcanic and iron formation-bearing rocks known informally as the Bass Lake sequence on the north. The Bass Lake rocks may be equivalent to parts of the Newton Lake Formation exposed north of Ely, but a complex series of faults in the intervening area makes this correlation speculative. Return to bus. Stop 8: Newton Lake Formation Variolitic Pillow Lavas and Hyaloclastite Longitude/Latitude: 47.93291301°N, -91.85191152E UTM NAD 83 Zone 15N: 585753E, 5309482N (NOTE: Modified from Field Trip Stop 5-16 (Jirsa et al., 2004), and Field Trip Stop ET-1 (Peterson et al., 2009)). The Neoarchean Newton Lake Formation is composed primarily of tholeiitic and komatiitic pillowed mafic lava, diabasic gabbro, differentiated mafic-ultramafic sills, intermediate-mafic pyroclastic rocks and siliceous marble with minor felsic-intermediate volcaniclastic rocks and lava flows. This formation is approximately 2,350 m thick. The unit overlies the Knife Lake Group in central part of the Vermilion district and the Lake Vermilion Formation in western part of Vermilion district (USGS National Geologic Map Database, https://ngmdb.usgs.gov/Geolex/UnitRefs/NewtonLakeRefs_9525.html). The Newton Lake Formation differs from the Ely Greenstone Formation in that the former contains a high proportion of mafic-ultramafic sills and lava flows, abundant diabasic sills and rare iron-formation. Lava flows in the Newton Lake Formation typically have larger MgO and incompatible element contents than those of the Ely Greenstone Formation, and some are classified as komatiites and komatiitic basalt (Schulz, 1980; Jirsa et al., 2004; Grotte and Hudak, 2014). The Newton Lake Formation (and possibly equivalent Bass Lake sequence) appears to be younger than the Lower Ely Member (~2723 MA; Peterson et al., 2001) with an age date of ~2715 MA (Jirsa, 2016) and was previously interpreted to be the youngest Archean supracrustal sequence in the Vermilion district until the Gafvert Lake Sequence was dated at 171 �Trip 7 – Classic Outcrops approximately 2689 MA (Lodge et al., 2013). Rocks having nearly identical composition and stratigraphic/structural setting occur in Itasca County some 80 kilometers to the west (Jirsa, 1990; Jirsa et al., 2004). A sequence of exceptionally well-preserved steeply-dipping, south-topping, lower greenschist facies metamorphosed Newton Lake Formation variolitic pillow lavas is exposed along a series of outcrops located on the west side of the road approximately one-half mile north of CR-88 on the Echo Trail. Variolites are defined as “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). According to Arndt and Fowler (2004), variolites result from either magma mingling or blotchy alteration, or they are a type of plagioclase spherulite. A generalized cross-section through these pillows from a detailed field and petrographic study of this outcrop (Grotte and Hudak, 2014) is presented in Figure 7-12A. Pillows vary from “bun-” to “mattress” shaped (Dimroth et al., 1978) and range from <1 to >2.5 meters in diameter. Pillow shapes, as well as the local occurrence of quartz-filled vacuoles within individual pillows, indicate younging directions to the south. Pillow cores tend to be dark green to pale yellow-green in color depending upon the abundance of secondary epidote alteration. The pillow cores commonly contain massive, globular variolites with local <1cm diameter spherical variolites., and are locally variolitic (Figure 7-12B). Pillow selveges are well preserved and commonly contain concentric zones globular to spherical Figure 7-12. Summary of field and petrographic observations of Newton Lake Formation variolitic pillow lavas at this location (from Grotte and Hudak, 2014). A) Generalized cross-section through a Newton Lake Formation pillow lava at this location. B) Outcrop photo of the margin of a pillow lava at this location noting the transition from well-preserved interpillow hyaloclastite into a variolitic pillow selvege. C) Thin section scan illustrating the well preserved cuspate, angular shards comprising the interpillow hyaloclastite. Dark spherical shapes on the right half of the photo are variolites. 172 �Trip 7 – Classic Outcrops variolites that mimic individual pillow shapes. Interpillow hyaloclastite is extremely well preserved and is composed of jigsaw-puzzle-fit angular cuspate shards that were originally glass but are now composed of fine-grained alteration minerals (Figure 7-12C). Petrographic observations (Grotte and Hudak, 2014) indicate that 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 needlelike to acicular skeletal plagioclase crystals and absence of phenocrysts suggest that the pillow lava flows at this location erupted at temperatures above the liquidus and experienced relatively large degrees of undercooling before undergoing rapid crystallization on the Neoarchean seafloor. As indicated in Jirsa et al. (2004), the Newton Lake is separated from the Ely Greenstone to the south by a complex zone of faulting (Shagawa Lake and Sibley faults) developed within sedimentary rocks of the Knife Lake Group. Although relatively undeformed conglomerate and sedimentary rocks of the Knife Lake Group are exposed just a few miles to the east, they are typically so sheared and altered in this area as to obscure lithologic and sedimentary interpretations. Return to bus Stop 9: Giants Range Batholith Longitude/Latitude: 47.81587746°N, -91.79083789E UTM NAD 83 Zone 15N: 590518E, 5296544N (NOTE: Modified from Hudak and Peterson, 2014). Footwall rocks to the northern part of the South Kawishiwi Intrusion are part of the Neoarchean Giants Range batholith (GRB). At this exposure along Highway 1, the GRB consists of porphyritic hornblende quartz monzonite that contains distinctive 1-2cm diameter potassium feldspar phenocrysts. One may also observe a distinctive foliation represented by alignment of black to dark green amphiboles and locally dark brown biotite. The massive nature of this unit creates an excellent footwall for Duluth Complex-associated intrusions and associated Cu-Ni-PGE deposits as the GRB lacks bedding and thus rare (if ever) gets incorporated into the mineralized zone as barren xenoliths. Additionally, melting of the GRB beneath longlived magma channels (Peterson and Boerst, 2013) at the base of the Maturi deposit has contaminated the South Kawishiwi intrusion, inducing additional sulfide immiscibility and the genesis of Ni- and Co-rich massive sulfide bodies. Return to bus Stop 10: Maturi SW Roadcuts of BH and U3 Units Longitude/Latitude: 47.78505228°N, -91.79056387E UTM NAD 83 Zone 15N: 590592E, 5293118N 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 173 �Trip 7 – Classic Outcrops 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. Return to bus Stop 11: SKI Magmatic Slurry Igneous Breccia Longitude/Latitude: 47.780584°N, -91.79273111E UTM NAD 83 Zone 15N: 590437E, 5292619N At this stop, we’ll examine perhaps the best exposure of Severson’s (1994) BH unit in the whole of the 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. We believe that all geologists who log drill core within the Cu-Ni-PGE 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 10 and 11. 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 had for many years its students complete a 1:5,000-scale bedrock geology map of this area. Return to bus Stop 12: Main AGT Longitude/Latitude: 47.81303067°N, -91.73468023E UTM NAD 83 Zone 15N: 594727E, 5296295N 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. Return to bus Stop 13: Spruce Road Bulk Sample Site/Discovery Burrow Pit Longitude/Latitude: 47.83271644°N, -91.67864227E UTM NAD 83 Zone 15N: 598885E, 5298553N 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 and visit several outcrops with fresh Cu-Ni sulfide minerals. This short stop will examine the bottom of an old borrow pit where participants can walk on and sample sulfide-bearing 174 �Trip 7 – Classic Outcrops 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. Return to bus Stop 14: Nickel Lake Macrodike Longitude/Latitude: 47.83079527°N, -91.63760896E UTM NAD 83 Zone 15N: 601959E, 5298393N The Nickel Lake Macrodike (NLM) is a northwest to southwest-trending, steeply dipping, asymmetric troctolitic and gabbroic intrusion interpreted to be a feeder dike for the northern portions of the SKI. The macrodike is interpreted to be located within a major rift-parallel normal fault (down to the southeast) now obscured by intrusion of NLM igneous rocks. Regional southward tilting (based on the deep level of erosion of the northern Bald Eagle Intrusion directly east of this area) leads to the interpretation that the southwest end of the NLM (near Omaday Lake) is structurally higher than the northeastern portion of the dike, and represents the location where magma flow changed from dike-like to sill-like, as it exited Figure 7-13. Bedrock geology map of the southwestern end of the Nickel Lake Macrodike. 175 �Trip 7 – Classic Outcrops the dike – thus the magma velocity slowed – and entered the growing SKI magma chamber. Excellent potential exists for Ni-Cu rich massive sulfide at the basal contact where the dike enters the SKI (Section 31, T62N, R10W). The 6.5km long by 1.0 km wide macrodike is composed of three main units: 1) inclusion-rich, locally sulfide-bearing, heterogeneous troctolite (unit Mpth) along the northwestern margin; 2) layered troctolite, melatroctolite, and dunite (unit Mltmt) along the southeastern margin; and 3) a late, cross-cutting, coarse-grained to pegmatitic oxide-rich, olivine-gabbro to melagabbro (unit Mxog) traversing generally through the center. Small (< 1m) to large (hundreds of meters long) xenoliths include Mesoproterozoic Anorthositic Series wall rocks (unit Mai) and North Shore Volcanic Group basalts (unit Mhb), and Paleoproterozoic Biwabik Iron Formation (unit Pifs) and Virginia Formation (unit Pvf). For this field trip we are simply going to take some walks in the bush, mostly along logging roads and snowmobile trails as time allows and look at numerous outcrops of the NML and adjacent rocks and discuss geology as we see it. Numerous detailed bedrock geology maps, reports, and presentations of the NLM and adjacent areas have been published over the last couple of decades (Peterson, 2002a, 2002b, 2002c, 2006a, 2006b, 2006c, 2008, Peterson and Albers, 2007) and a compilation of detailed geologic mapping data for the southwestern NLM is given in Figure 7-13. Return to bus Stop 15: Remnant Saprolite, Middle SKI Longitude/Latitude: 47.77089981°N, -91.66297244E UTM NAD 83 Zone 15N: 600176E, 5291703N A short field trip stop to examine locally layered troctolitic rocks of the Upper SKI of Peterson and Boerst (2013). This outcrops in this area epitomizes the “Sea of Troctolite” that occurs throughout the vast majority of the SKI (Middle and Upper SKI of Peterson and Boerst, 2013). Careful attention will be given to an outcrop next to the bus where spheroidal weathering of the troctolite is forming rounded core stones of troctolite, which we’ll see once again at stop 18. Return to bus Stop 16: Anorthosite Series Roadcut Longitude/Latitude:: 47.75915521°N, -91.64719916E UTM NAD 83 Zone 15N: 601381E, 5290418N Large, glacially polished roadside outcrop of the gabbroic and troctolitic anorthosites of the Anorthositic Series of the Duluth Complex. At this location these anorthositic rocks form the eastern sidewall of the SKI and are cut by a series of northeast-striking valleys. The valleys were interpreted by geologists of Duluth Metals Limited as steeply west-dipping reverse faults that were formed by emplacement of the SKI immediately to the west. Approximately 2.5 km to the southwest of this roadcut Cold Spring Granite Company quarries a large gabbroic anorthosite xenolith similar to this stop in their Mesabi Black quarry. Return to bus 176 �Trip 7 – Classic Outcrops Stop 17: Bald Eagle Intrusion Longitude/Latitude: 47.7385175°N, -91.6405279E UTM NAD 83 Zone 15N: 601921E, 5288133N A quick stop to observe a roadside outcrop of troctolite of the Bald Eagle Intrusion (BEI). The BEI is a large (4.5 to 16.5 km x 31 km) troctolitic to gabbroic body that was emplaced partially within Anorthositic series rocks, the SKI, and the Greenwood Lake Intrusion (see BEI on Figure 7-7). Weiblen (1965) mapped the well-exposed northern portion of the intrusion and showed that it is funnel-shaped and consists of an outer zone of troctolite and an inner zone of olivine gabbro. In the poorly exposed southwestern portions of the intrusion, field mapping by Green et al., (1966) and Foose and Cooper (1978) showed the BEI and SKI in direct conformable contact. Steep foliation and modal layering (Weiblen, 1965, Green et al., 1966) integrated with a distinct gravity anomaly over the northern BEI imply that the northern part of this intrusion is funnel shaped and necks down to a steep feeder dike. Weiblen and Morey (1980) interpreted the limited cryptic variation (Weiblen, 1965), the steep dip of lamination and layering, and adcumulate nature of the BEI as indicative of its being an open conduit to higher intrusions and perhaps volcanic flows. Petrologic observations and geophysical interpretations (Chandler, 1990, Chandler and Ferderer, 1989) suggest that the BEI and SKI were emplaced by successive overplating of magmas from a common feeder centered on the northern BEI and extending along the trace of the NLM that links the BEI and SKI. In a related analogy, Cartwright and Møller-Hansen (2006) have shown that interconnected sill complexes transect the middle to upper crust over a vertical distance of 8-12 km offshore of Norway. The geometry of the gravity and magnetic anomalies of the BEI, as well as the overall Midcontinent Rift is very similar to the pattern of the seismic reflections profiles of active ridge systems (Vislova, 2003). In detail, the geophysical expressions of the BEI have the same shape and dimensions as the “bulls’ eye” pattern of low velocity seismic reflection anomalies along the East Pacific Rise. These anomalies are interpreted to define regions of melt concentrations, i.e., active magma chambers. These data suggest that the BEI could be a “frozen” dynamic magma chamber (Weiblen et al., 2005, Peterson and Hauck, 2005). Eight exploration holes drilled by Duluth Metals Limited in 2011 revealed several new distinct features of the BEI. All of these holes encountered chromitite layers within horizontally layered troctolites with many of the chromitite horizons occurring as “rip up” clasts within troctolites. Duluth Metals Limited’s hole LOD-06, drilled 12km SSW of this field trip stop, encountered flowing gas at a depth of 1,778 feet. The gas was analyzed and found to contain >10% helium. This is the site where Pulsar is currently exploring with the aim of producing helium gas. Return to bus Stop 18: Vermilion Moraine Longitude/Latitude: 47.6918526°N, -91.81063993E UTM NAD 83 Zone 15N: 589247E,5282737N In common with most of the high latitude regions of North America, northeastern Minnesota was repeatedly glaciated during the ice ages of the Pleistocene Epoch. Glaciogenic sediments and landforms in this 2025 ILSG field trip area are associated with the Rainy Lobe of the Laurentide ice sheet. While there are a number of possible definitions of what constitutes the Rainy Lobe – sedimentological, textural, compositional, and association with particular geomorphic features – a definition rooted in glacial dynamics perhaps works best. In this sense, the Rainy Lobe refers to that portion of the Laurentide ice sheet lying northwest of Lake Superior (occupied by the Superior Lobe), and east of the Winnipeg Basin and Red River Valley (occupied by the Red River Lobe). In common, Rainy Lobe landforms and glaciogenic sediments 177 �Trip 7 – Classic Outcrops reflect a general northeast to southwest ice flow direction, and a Labradoran (northeastern) sediment provenance. In common with much the Canadian Shield, glacial erosion has nearly completely stripped preglacial regolith from bedrock north of the Laurentian Divide. However, preglacial saprolites are a common occurrence underlying glaciogenic sediments in central and western Minnesota; the nearest such occurrences are exposed in open pit mines of the Mesabi Range, on the south flank of the Giant’s Range. Approximately 12,400 years ago, the retreating Rainy Lobe made a last stand in northern Minnesota to form the West-Northwest to East-Southeast trending Vermilion Moraine. This stop includes a quick walk over the end of the Vermilion Moraine and a view to the south over a glacial lake plain (Fig. 7-14). Return to bus Figure 7-14. Annotated lidar digital elevation model showing glaciogenic landforms in the stop 18 area. Stop 19: Contaminated Basal SKI, Dunka Pit Area Longitude/Latitude: 47.69423099°N, -91.85803438E UTM NAD 83 Zone 15N: 585687E, 5282948N The recently permitted extension of Cliffs Natural Resources Northshore mine required the rerouting of St. Louis County Road 623 to the north. The building of the new road resulted in the exposure of rocks of the ~1.85 Ga. Biwabik Iron Formation and the 1.1 Ga. South Kawishiwi intrusion. This short stop will include the examination of three new roadside outcrop areas, including: 1) Metamorphosed Biwabik Iron Formation, 2) Sulfide-poor gabbroic rocks, and 3) Sulfide-rich (pyrrhotite-dominant) 178 �Trip 7 – Classic Outcrops contaminated noritic rocks. Geochemical analyses of rock samples from these three outcrop areas (completed by Duluth Metals in 2012) are given in Table 7-4. Table 7-4. Geochemical analyses of rock samples taken from the roadside outcrops of Stop 19. Sample ID DMR0446 DMR0447 DMR0164 DMR0165 DMR0448 DMR0449 DMR0450 DMR0451 Rock Type Iron Formation Iron Formation Olivine Gabbro Biotitic Gabbro Sulfidic Norite Sulfidic Norite Sulfidic Norite Sulfidic Norite Outcrop # 1 1 2 2 3 3 3 3 Cu (ppm) 4 0 357 218 3960 2630 4380 4270 Ni (ppm) 0 0 82 87 1230 761 1030 1120 Co (ppm) 2 13 44 53 191 117 168 183 Pt (ppb) 5 1 1 1 9 20 6 16 Pd (ppb) 1 3 1 2 53 40 58 60 Au (ppb) 1 1 1 1 18 15 23 23 S (%) -0.01 -0.01 0.02 -0.01 2.61 1.62 1.58 1.79 SiO2 (%) 51.35 38.71 48.14 47.28 42.67 51.62 41.25 43.53 Al2O3 (%) 0.25 0.45 15.21 15.06 15.02 17.56 13.81 13.71 Fe2O3 (%) 40.11 55.01 15.64 16.35 21.04 11.98 22.12 21.34 CaO (%) 6.07 3.33 7.87 8.00 7.51 5.39 6.29 6.90 MgO (%) 1.93 1.52 4.93 5.47 7.15 5.13 6.01 6.97 Na2O (%)t 0.06 0.06 2.90 2.62 2.25 2.72 1.99 2.34 K2O (%) 0.01 0.01 1.12 1.10 0.68 2.54 0.56 0.63 Cr2O3 (%) -0.01 -0.01 0.02 0.01 0.02 0.03 0.02 0.02 TiO2 (%) -0.01 0.05 3.12 3.47 1.77 0.76 1.90 2.31 MnO (%) 0.97 0.51 0.19 0.20 0.16 0.08 0.14 0.18 P2O5 (%) 0.06 0.07 0.43 0.34 0.23 0.05 0.26 0.26 -0.95 -1.37 0.05 -0.09 1.25 1.80 5.13 1.02 LOI (%) Photographs of the Biwabik Iron Formation (outcrop #1) and sulfidic norite of the South Kawishiwi intrusion (outcrop #3) are presented in Figure 7-15. 179 �Trip 7 – Classic Outcrops Figure 7-15. Field photographs of roadside outcrops of stop 5. (A) outcrop of the Biwabik Iron Formation, (B) closeup shot of bedding in granular iron formation (GIF), (C) rusty weathering and gossanous outcrop of the basal mineralized zone of the South Kawishiwi intrusion, and (D) pyrrhotite-rich norite. Return to bus Stop 20: Giants Range Batholith Migmatite/Pyroxenite-Lamprophyre Dike Longitude/Latitude: 47.68689429°N, -92.05199159E UTM NAD 83 Zone 15N: 571144E, 5281936N This field trip ends where we began, within the Neoarchean Giants Range Batholith. This roadside outcrop of the Embarrass tonalite, an early phase of the GRB that was first mapped by Griffin and Morey (1969) and later remapped by Terry Boerboom in 2015. The outcrop, as mapped by Terry Boerboom, consists of intermixed migmatitic biotite-schist and tonalitic gneiss crosscut by a lamprophyre/pyroxenite dike. Approximately 4-miles to the west of this outcrop the Embarrass Tonalite was originally dated by UPb zircon at 2718 ± 67 Ma by Peterman in Southwick (1994). This data has been superseded by a second U-Pb zircon age of 2687 ± 0.6 Ma by Jirsa (2016). RETURN TO MOUNTAIN IRON COMMUNITY CENTER 180 �Trip 7 – Classic Outcrops Acknowledgements Characterizing and evaluating the detailed geology of northeastern Minnesota has been a team effort involving former NRRI geologists, former and current Minnesota Geological Survey geologists and geophysicists, personnel from the Minnesota Department of Natural Resources and students and faculty from the Precambrian Research Center Field Camp, the University of Minnesota Duluth, the University of Minnesota Twin Cities, and the University of Wisconsin Eau Claire. Their efforts are appreciated. As well, permission to map private properties that was granted by local landowners and mineral exploration/mining companies is much appreciated. The authors would like to thank Jim Essig (Manager, Lake Vermilion / Soudan Underground Mine State Park) and James Pointer (Interpretive Supervisor, Lake Vermilion / Soudan Underground Mine State Park) from the MDNR for their support, assistance, and guidance while planning and conducting detailed geological mapping by the NRRI geologists during the DUSEL project and PRC students and faculty in Lake Vermilion State Park in 2010 and 2011. Funding from the Minerals Coordinating Committee, the University of Minnesota Permanent University Trust Fund, the National Science Foundation, the University of Minnesota Duluth Undergraduate Research Opportunities Program, the University of Minnesota Duluth Graduate School, The University of Wisconsin Oshkosh StudentFaculty Research Program, and many mineral exploration companies also enabled geological research in northeastern Minnesota. 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Hauck, S., Severson, M., Ripley, E., Goldberg, S., and Alapieti, T., 1997, Geology and Cr-PGE mineralization of the Birch Lake area, South Kawishiwi intrusion, Duluth Complex: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report, NRRI/TR-97/13, 32 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. Hooper, P., and Ojakangas, R., 1971, Multiple deformation in the Vermilion District, Minnesota: Canadian Journal of Earth Sciences, v. 8, p. 423-434. 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/TR-2002/03, 390 pages. Hudak, G. J., and Peterson, D. M., 2014, Non-Ferrous Mineralization Associated with the Wawa-Abitibi Terrane and Duluth Complex Cu-Ni-PGM Deposits, Northeastern Minnesota: Society of Economic Geologists, Guidebook Series, v. 47, 150 p. 182 �Trip 7 – Classic Outcrops 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. 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L., 1988, Transpression in an Archean greenstone belt, northern Minnesota: Canadian Journal of Earth Sciences, v. 25, p. 1060-1068. Hudleston, P.J., 1976, Early deformational history of Archean rocks in the Vermilion district, north-eastern Minnesota: Canadian Journal of Earth Sciences, v. 13, p. 579-592. Jansen, A. C., Hudak, G. J., Heine, J. J., and Peterson, D. M., 2009, 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., 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. Jirsa, M. A., Boerboom, T. J., and Peterson, D. 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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., Southwick, D.L., and Boerboom, T.J., 1992, Structural evolution of Archean rocks in the western Wawa subprovince, Minnesota: refolding of precleavage nappes during D2 transpression: Canadian Journal of Earth Sciences 29:2146-2155. Klinger, F. L., 1960, Geology and ore deposits of the Soudan Mine, St. Louis County, Minnesota: unpublished Ph. D. dissertation, University of Wisconsin, Madison, 96 p. Kuenen, P.H., and Migliorini, C., 1950, Turbidity currents as a cause of graded bedding: Journal of Geology, 58:91127. Kuhns, M.J., Hauck, S.A., and Barnes, R.J., 1990, Origin and occurrence of platinum group elements, gold and silver in the South Filson Creek copper-nickel mineral deposit, Lake County, Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/GMIN-TR-89-15, 60 p., 3 pls. Lee, I., and Ripley, E.M., 1996, Mineralogic and oxygen isotopic studies of open system magmatic processes in the South Kawishiwi intrusion, Spruce Road area, Duluth Complex, Minnesota: Journal of Petrology, v. 37, no. 6, p. 1437-1461. 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: Hibbing, Minnesota Department of Natural Resources, Division of Minerals, Report 93, 74 p. Lodge, R. W. D., Gibson, H. L., Stott, G. M., Hudak, G. J., Jirsa, M. A., and Hamilton, M. 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(ed.), Ore Deposits of the United States, 1933-1967: New York, American Institute of Mining, Metallurgical, and Petroleum Engineers, Inc., The Grafton-Sales Volume, v. 1, p. 518-537. 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. 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. Morey, G.B., 1965, The sedimentology of the Precambrian Rove Formation in northeastern Minnesota (abs.): Institute on Lake Superior Geology, 11th Annual Meeting, St. Paul, Minnesota, Proceedings p. 25-26. Morey, G.B., 1992, Chemical composition of the eastern Biwabik Iron Formation (Early Proterozoic), Mesabi Iron Range, Minnesota: Economic Geology, v. 87, p. 1649-1658. Morey, G.B., and Cooper, R.W., 1977, Bedrock geology of the Hoyt Lakes–Kawishiwi area, St. Louis and Lake Counties, northeastern Minnesota: Minnesota Geological Survey, Open-File Report, scale 1:48,000. 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. Naldrett, A.J., 1997, Key factors in the genesis of Noril’sk, Sudbury, Jinchuan, Voisey’s Bay and other world-class Ni-Cu-PGE deposits: Implications for exploration: Australian Journal of Earth Sciences, v. 44, no. 3, p. 283-315. Newkirk, T., Hudak, G. J., and Hauck, S. A., 2001a, Preliminary lava flow morphology studies at the Five Mile Lake VMS prospect, Vermilion District, NE Minnesota: Implications for volcanic processes, volcanic paleoenvironments, and VMS exploration: Institute on Lake Superior Geology, 47th Annual Meeting, Proceedings Volume 47, Part 1- Program and Abstracts, p. 69-70. Newkirk, T., Hudak, G. J., and Hauck, S. A., 2001b, Preliminary lava flow morphology studies at the Five Mile Lake VMS prospect, Vermilion District, NE Minnesota: Implications for volcanic processes, volcanic paleoenvironments, and VMS exploration: Geological Society of America Abstracts and Programs Volume 33, No. 6, p. A-398. Ojakangas, R.W., 1966, Precambrian stratigraphy and structure of the Tower, Minnesota quadrangle (abs): Institute on Lake Superior Geology Proceedings, 12th Annual Meeting, Sault Ste. Marie, Michigan, Proceedings p. 17. Patelke, R.L., 2003, Exploration drill hole lithology, geologic unit, copper-nickel assay, and location database for the Keweenawan Duluth Complex, northeastern Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report, NRRI/TR-2003/21, 97 pages, 1 CD. Peterson, D.M., 1997, Ore deposit modeling of the footwall mineralization of the Duluth Complex: Minnesota Department of Natural Resources, Division of Minerals, Project 317, 55 p., 46 pls. 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., 2001b, Development of a conceptual model of Cu-Ni-PGE mineralization in a portion of the South Kawishiwi intrusion, Duluth Complex, Minnesota: Society of Economic Geologists, Second Annual PGE Workshop, Sudbury, Ontario, 3 p. Peterson, D.M., 2001c, Copper-Nickel-PGE mineral potential of the eastward extension of the Maturi Cu-Ni deposit, Duluth Complex, Lake County, Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Confidential Report of Investigations NRRI/RI-2001-02, 29 pages, 15 plates, 1 CD. Peterson, D.M., 2002a, 3-Dimensional view through a mineralized system: the South Kawishiwi intrusion, Duluth Complex: Institute on Lake Superior Geology, 48th Annual Meeting, Thunder Bay, Ontario, v. 48. 184 �Trip 7 – Classic Outcrops Peterson, D.M., 2002b, Cu-Ni-PGE mineralization in the South Kawishiwi intrusion, northeastern Minnesota, Variation due to magmatic processes: Institute on Lake Superior Geology, 48th Annual Meeting, Thunder Bay, Ontario, v. 48. Peterson, D.M., 2002c, Variation in the Cu-Ni-PGE mineralization in the South Kawishiwi intrusion, Duluth Complex, northeastern Minnesota: 9th International Platinum Symposium, Billings, Montana, USA, July 21-25. Peterson, D.M., 2002d, Copper-Nickel grade maps for the Spruce Road deposit, South Kawishiwi intrusion, Duluth Complex: University of Minnesota Duluth, Natural Resources Research Institute, Report of Investigations NRRI/RI-2002/03, 99 p. Peterson, D.M., 2002e, Shaded relief map of the basal contact of the South Kawishiwi intrusion, Duluth Complex, northeastern, Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Map Series NRRI/MAP-2002-01, scale 1:75,000. Peterson, D.M., 2002f, Bedrock geology, sample location, and property position maps of the west Birch Lake area, South Kawishiwi intrusion, Duluth Complex, northeastern, Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Map Series NRRI/MAP-2002-02, scale 1:10,000. 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, Northeastern Minnesota: Unpublished geological map, North-Central Section of the Geological Society of America Meetings, Minneapolis, 1:20,000 scale. Peterson, D.M., 2006a, 3D Visualizations of mafic Intrusions in the Duluth Complex, northeastern Minnesota, Institute on Lake Superior Geology, 52nd Annual Meeting, Sault Ste Marie, Ontario, May 8-12, 2006, Minnesota, v. 52. Peterson, D.M., 2006b, Digital base for geological mapping within the northern South Kawishiwi intrusion: Lake and St. Louis Counties, northeastern Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Map Series NRRI/MAP-2006-01, scale 1:20,000. Peterson, D.M., 2006c, New ideas on mineralization in the Duluth Complex, Oral presentation and online pdf file to the Mesabi Range Geological Society, December 20, 66 pages. 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. 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., 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. 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. 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, 47th Annual Meeting, Madison, Wisconsin, Proceedings v. 47, Part 1, p. 77-78. Peterson, D.M., and Hauck, S.A., 2005, Visualization of "Frozen" dynamic magma chambers in the Duluth Complex, northeastern Minnesota: Eos Trans. AGU 86(52), Fall Meet. Suppl., Abstract V23A-0680. Peterson, D.M., and Jirsa, M.A., compilers, 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. Peterson, D.M., Jirsa, M. A., and Hudak, G. J., 2005. Field Trip 9: Architecture of an Archean Greenstone Belt: Stratigraphy, Structure and Mineralization: in Robinson, L., ed., 2005, Field Trip Guidebook for Selected Geology in Minnesota and Wisconsin: Minnesota Geological Survey Guidebook 21, p. 154-180. Peterson, D.M., Jirsa, M., and Hudak, G., 2009, 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. 185 �Trip 7 – Classic Outcrops Peterson, D.M., and Patelke, R. L., 2004a, Bedrock Geology and Lode Gold Prospect Data Map of the Mud Creek Road Area, Northern St. Louis County, Minnesota: Natural Resource Research Institute Geologic Map NRRI/MAP-2004/01, 1:12,00 scale, available for free download at http://www.nrri.umn.edu/egg/REPORTS/MAP200401/MAP200401.html. Peterson, D.M., and Patelke, R. L., 2004b, 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., Patelke, R.L., and Severson, M.J., 2004, Bedrock geology map and Cu-Ni mineralization data for the basal contact of the Duluth Complex west of Birch Lake, St. Louis and Lake Counties, northeastern Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Map Series NRRI/MAP2004-02, scale 1:10,000. 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. 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. Phinney, W.C., 1969, The Duluth Complex in the Gabbro Lake quadrangle, Minnesota: Minnesota Geological Survey, Report of Investigations 9, 20 p. 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. Phinney, W.C., 1972, Northwestern part of Duluth Complex, in Sims, P.K., and Morey, G.B., eds., Geology of Minnesota: A centennial volume: Minnesota Geological Survey, p. 335-345. Ramsey, J.G., and Huber, M.I., 1987, The Techniques of Modern Structural Geology, Academic Press Inc. (London) Ltd. 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. 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. Schulz, K.J., 1980, The magmatic evolution of the Vermilion greenstone belt, NE Minnesota: Precambrian Research 11:215-245. Severson, M.J., 1994, Igneous stratigraphy of the South Kawishiwi intrusion, Duluth Complex, northeastern Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR93/34, 210 p., 15 plates. 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., 1 CD. 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, Minnesota: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR- 2009/09, 173 p. + 37 plates. 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. Thompson, A., 2015, A hydrothermal model for metasomatism of Neoarchean Algoma-type banded iron formation to massive hematite ore at the Soudan Mine, NE Minnesota: unpublished M. S., thesis, University of Minnesota Duluth, 59 p. Vislova, T., 2003, Petrology of the Bald Eagle intrusion and associated rocks and its relevance to crystallization in dynamic magma chambers in the Midcontinent Rift: Unpublished Ph.D. Thesis, University of Minnesota. 186 �Trip 7 – Classic Outcrops Weiblen, P.W., 1965, A funnel-shaped, gabbro-troctolite intrusion in the Duluth Complex, Lake County Minnesota: Unpublished Ph.D. Thesis, University of Minnesota, 161 p. Weiblen, P.W., Morey, G. B., 1980, A summary of the stratigraphy, petrology, and structure of the Duluth Complex: American Journal of Science, vol. 280A, Part I, p 88-133. Weiblen, P.W., Peterson, D.M., and Vislova, T., 2005, Implications of Midcontinent Rift and oceanic ridges analogies and 3-D interpretations of the subsurface structure of the Bald Eagle intrusion in the Duluth Complex and the East Pacific Rise: Institute on Lake Superior Geology, 51st Annual Meeting, Sault Ste Marie, Ontario, v. 51, 3 p. White, C., 2010, The Nokomis Deposit, a Masters of Geology thesis: University of Minnesota, Duluth. White, J. D. L., and Houghton, B. F., 2006, Primary volcaniclastic rocks: Geology, v. 34, no. 8, p. 677-680. 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: University of Minnesota Duluth, Natural Resources Research Institute, Technical Report NRRI/TR93/52, 90 p., 3 pls. 187 �Trip 8 – Glacial FIELD TRIP 8 Glacial Lake Norwood and the Koochiching Lobe Phil Larson1, Andrew Breckinridge2, and Howard Mooers3 1 Vesterheim Geoscience PLC Natural Sciences Department, University of Wisconsin Superior, 202 Barstow Hall, Superior, WI 54880 3 Department of Earth and Environmental Sciences, University of Minnesota Duluth, 1114 Kirby Drive, Duluth, MN 55812 2 Introduction The region north of the Giants Range is draped by sediment deposited during the final retreat of the Laurentide ice sheet from northeastern Minnesota. These sediments record the retreat of the Rainy Lobe ice margin to the northeast, the formation of Glacial Lake Norwood (GLN), two successive advances of the Koochiching Lobe from the northwest, and the opening of a western outlet of GLN and its succession by Glacial Lake Agassiz, all over the span of a few thousand years. Historically, the Quaternary geology of this region has received scant attention. However, recent work integrating varve chronology, high resolution LiDAR digital terrain models, till geochemistry, rotasonic drilling, and mapping has resulted in substantially improved and nuanced understanding of the sedimentary processes active, and the sequence of events, during deglaciation. A key finding is that GLN was of significantly longer duration than previously believed, and consequently a stronger control on sediment and landform distribution in the region. Within the footprint of GLN, there is scant evidence for preservation of glacigenic sediments predating the Late Wisconsinan. Interbedded till and glaciolacustrine sediment thicknesses up to 70 m thick preserve evidence of extremely high sedimentation rates in a dynamic sediment system. High rates of sediment delivery by the Koochiching Lobe and analogues from the west served as the dominant sediment source, while intense reworking by wave action in GLN was a dominant control on sediment distribution. Historical Background 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 in the deglaciation sequence, although he did not define its entire length. 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 areaWinchell (1899) compiled Upham, Elftman, and his own observations into a map of large portions of northeastern Minnesota and description of the surficial deposits. Winchell (1900) described evidence for glacial lakes in Minnesota, including naming Glacial Lake Norwood. Notably, he did not recognize the full extent of Glacial Lake Norwood, assigning portions of the Norwood basin to other glacial lakes. 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. 188 �Trip 8 – Glacial Figure 8-1. Extent of Glacial Lake Norwood (light green-blue). The lack of modern lakes (blue) in the GLN basin highlight area of significant glaciolacustrine sediment thickness. Major Rainy lobe recessional moraines (brown lines). Proglacial Lake Northofnashwauk is the high-level (elev. >1500’) proglacial lake predating GLN dammed by the St. Louis sublobe. Modern understanding of the surficial geology of northeastern Minnesota began with Wright Wright (1956), who was the first to conduct systematic fieldwork in the area between the border lakes and Lake Superior. Wright and Watts (1969) reconstructed the postglacial vegetational history of northeastern Minnesota, and established the first regional deglaciation chronology, including use of radiocarbon dates to establish absolute ages to deglaciation. These early efforts were summarized in the comprehensive general glacial geologic framework of Minnesota Wright (1972). The United States Geological Survey conducted a comprehensive study of surficial geology and groundwater availability on the Mesabi Iron Range. An initial map (Cotter, Young, and Winter 1964) was later followed by additional publications on the glaciation sequence (Winter 1971) glacial sediment composition (Winter, Cotter, and Young 1973), and groundwater hydrology in glacial drift-hosted aquifers (Winter 1973). Hobbs (1983) provided the first comprehensive account of Glacial Lake Norwood’s extent and history. At that time, he rechristened GLN as Glacial Lake Koochiching, not recognizing the continuity with Winchell’s (1900) definition of GLN. In this respect he was hampered by the paucity of well-defined strandlines for the upper levels of GLN in the rocky meltout till underlying much of the southern portion of the basin. He also posited a late, lower elevation outlet for Glacial Lake Koochiching southward along the Prairie River; recent (2012) LiDAR elevation data (MNDNR 2012) combined with a better defined isostatic rebound reconstruction (Breckenridge 2015) suggest the existence of a southern outlet to GLN untenable. Significantly, Hobbs recognized that Glacial Lake Norwood expanded westward, ultimately establishing an outlet via the McIntosh Channel into Glacial Lake Climax. Continued retreat of the Red River lobe ice margin resulted in coalescence of Glacial Lake Climax and GLN into Glacial Lake Agassiz at the Herman level at 13.9±0.3 cal kyr BP (Lepper et al. 2007). 189 �Trip 8 – Glacial Björck (1990) expanded Wright and Watts (1969) pioneering work by collecting radiocarbon dates north of the Giants Range. He obtained basal radiocarbon dates from Sabin Lake (located in the outflow to GLN) of 10,230±230 and 10,320±170 14C kyr BP. Bjorck’s oldest date was 12,100±150 14C kyr BP from Heikkila Lake, located within the Big Rice moraine. Lowell et al. (2009) reported a radiocarbon date from north of the Vermilion moraine of 12,000±85 14C kyr BP, assigning a minimum age to the Vermilion phase and the moraines to the south of 13.9±0.2 cal kyr BP. These dates establish that Glacial Lake Norwood and drainage through the Embarrass Gap persisted long after the Laurentide ice sheet margin retreated from the Vermilion moraine. Johnson et al. (2016) assigned the glacigenic deposits in northeastern Minnesota to a formal statewide lithostratigraphic framework. Essentially all the aforementioned published work was opened to critical re-examination and revision upon release of 1m resolution LiDAR-derived digital terrain models in the spring of 2012 (MDNR 2012). This data provides resolution orders of magnitude greater than previous topographic models, allowing for vastly improved recognition of some classes of glacigenic landforms, and recognition for the first time of entire new classes of landforms. These advances allowed for significant refinement in mapping of glacial landforms, interpretation of sediment-landform relationships, and development of deglaciation process models. Breckenridge (2015) mapped glacial lake strandlines and developed an isostatic rebound model for much of northern Minnesota, including much of the Glacial Lake Norwood basin, demonstrating the previously unrecognized widespread extent of both Glacial Lake Norwood and the early, high levels of Glacial Lake Agassiz. Bauer et al. (2022) published a surficial geologic map and Quaternary stratigraphic interpretation of much of the GLN basin, relying primarily on the lithostratigraphic mapping approach favored by the Minnesota Geological Survey, but also incorporating landform interpretation based on the 2012 LiDAR data. Glacial History Northeastern Minnesota was continuously covered by ice from the earliest Late Wisconsin ice advance approximately 28 kyr bp until about 11 kyr bp by the Rainy lobe of the Laurentide ice sheet (Clayton and Moran 1982); (Mooers and Lehr 1997)). Although the Glacial Lake Norwood basin was subjected to multiple glacial cycles, the vast majority of glacigenic sediment was deposited during the last retreat of the Laurentide ice sheet during the Late Wisconsinan (<15 kyr bp). The Pleistocene stratigraphic record therefore principally reflects retreat of the ice sheet, and is composed of glacigenic sediment deposited at or near the ice margin. Bedrock Geology and Preglacial Regolith The GLN basin underlain by greenstone (metavolcanic and metasedimentary rocks) and granitoids of the ~2.7 Ga Wawa-Shebandowan Subprovince. The craton was intruded by mafic intrusives of the 2076 Ma Kenora-Kabetogama dike swarm (Southwick and Halls 1987; Buchan, Halls, and Mortensen 1996), while contact relationships indicate the Archean craton was peneplained by the time arenites, ironformation, greywacke, and argillite of ~1.85 Ga Animikie Basin were deposited to the south. Minor mafic dikes related to the 1.1 Ga Midcontinent Rift are known to intrude the Archean craton north of the Giants Range; it is probable that additional similar intrusives have yet to be recognized or mapped. Subsequent to cessation of the Midcontinent Rift, Precambrian bedrock in the GLN basin was subject to a nearly 1 billion year period of chemical weathering and saprolite formation. Saprolite formation was preferentially, but not necessarily, focused along joints, faults, and less weathering-resistant lithologies, forming deep linear weathering pendants beneath a more widespread blanket of saprolite. Commencement of glaciation at the beginning of the Pleistocene ~3 Ma subjected the Superior craton to significant physical erosion for the first time in nearly a billion years. Successive glacial cycles 190 �Trip 8 – Glacial preferentially eroded unconsolidated saprolite, removing first the extensive saprolite blanket, and then excavating saprolite from deep weathering pendants. To the southwest of the GLN basin (central Minnesota), the preglacial saprolite is largely intact beneath Pleistocene glacigenic sediment. To the northeast (northwestern Ontario), preglacial saprolite has been essentially completely removed. Here, the rugged ‘glacially sculpted’ shield terrain characteristic of this region is better explained as the unweathered bedrock surface forming the base of the preglacial saprolite; bedrock has undergone relatively little actual glacial weathering. In the GLN basin proper, preglacial saprolite removal by glacial erosion is incomplete, leaving patchy remnants of unconsolidated preglacial saprolite on the bedrock surface. Saprolite is occasionally intercepted in boreholes. Saprolite and incipient pendant weathering have been encountered associated with joints and fractures as deep as 100 m. Pre-Late Wisconsinan The overlying preglacial saprolite was removed by repeated cycles of erosion and deposition during the Pleistocene. Saprolite eroded as the ice sheet grew (relative early in a glacial cycle) was transported to the margin. The remnant saprolite was blanketed by glacigenic sediment as the ice margin receded (late in the glacial cycle. Subsequent glacial cycles removed both the older glacigenic sediment and additional saprolite. Winter (1971) and Winter, Cotter, and Young (1973) described a dark-colored, sandy-silty calcareous till in exposures in open pit mines on the Mesabi Iron Range. Since this till, where present, occurred immediately above bedrock, they referred to it as the “basal till”. Stark (1977) and (Lehr and Hobbs 1992) described occurrences of Winter’s basal till in exposures in the Dunka Mine. The matrix of Winter’s basal till is calcareous, and the pebble fraction contains carbonate clasts in addition to the granitic and metamorphic lithologies typical of Rainy lobe tills. A northeast-southwest pebble fabric in Winter’s basal till strongly supports a northeastern provenance for this till, indicating the carbonate in pebbles and till matrix is derived from Paleozoic carbonates in the Hudson Bay Lowlands (HBL). A distinctive greywacke lithology (Prest, Donaldson, and Mooers 2000) associated with carbonate-bearing tills has been recovered from glacigenic sediments north of the Giants Range (this author), indicating older carbonatebearing glacigenic sediment was actively reworked during the last retreat of the Laurentide ice sheet. Additional occurrences of this calcareous basal till have been intercepted in boreholes elsewhere in the GLN basin, indicating patchy remnants of preglacial saprolite and older (carbonate-bearing) glacigenic sediment are present beneath the relatively continuous blanket of glacigenic sediment deposited between ca. 15 kyr bp and 10 kyr bp during the last retreat of the Laurentide ice sheet. Post-Last Glacial Maximum – Rainy Lobe Recession of the Laurentide ice sheet margin following its last glacial maximum extent at ca. 20 kyr bp was characterized by rapid melting of ice during summer months followed by stabilization and minor re-advance during the winter. This process formed a series of small, annual recessional moraines, spaced 25-75 m apart, reflecting the long-term retreat rate of the ice sheet. To a significant degree, glacigenic sediment deposited by the Rainy lobe of the Laurentide ice sheet during retreat of its margin from the southwest to northeast is the oldest Pleistocene sediment preserved in the GLN basin. Post-LGM Rainy lobe sediments are typically comprised predominantly of sediment eroded locally from Archean greenstone and granitoid lithologies; this results in significant lithologic and geochemical compositional variability(Larson, 2004; Larson & Mooers, 2004). Lodgment tills are commonly ~2 m thick, while sand and gravel deposited in subaqueous recessional moraines commonly form sharp-crested ridges 5-40 m thick. Distal glaciolacustrine sand and silt commonly drapes older basal lodgment tills and recessional moraines. 191 �Trip 8 – Glacial This process of gradual ice margin retreat was punctuated by surges, episodes of major re-advance and stagnation. These surges resulted in deposition of moraines significantly broader and thicker than annual recessional moraines. The surges may not reflect re-advance of the ice sheet as a whole, but were likely restricted to sectors of the ice margin on the order of 100s of km. They may therefore not be a direct physical reflection of climate fluctuations, but rather reflect internal ice sheet dynamics. Most of the area exposed by ice margin retreat from the Giants Range was inundated in proglacial lakes, successively by Glacial Lake Nashwauk, Glacial Lake Norwood, and finally by Glacial Lake Agassiz. The extended interval between ice margin retreat, lake drainage, and establishment of terrestrial vegetation over most of this area significantly hinders the ability to establish a precise deglaciation chronology (compare Björck (1990)). Allen Phase The oldest major surge-stagnation moraine recognized in the GLN basin is the Allen moraine, which forms a WNW-ESE trending belt of stagnation topography (ice-walled lake plains, meltout tills, etc.), passing through the Embarrass Gap. Ice flow during the Allen phase was generally toward the SSW (bearing 190°). Ice margin retreat from the Allen moraine and opening of meltwater drainage through the Embarrass Gap was the event that by definition resulted in formation of Glacial Lake Norwood. Further ice margin recession and deposition of annual recessional moraines suggests about 150 years before the next major surge-stagnation event. Big Rice and Wahlsten Phases The second major surge-stagnation moraine recognized in the GLN basin is the Big Rice moraine, a W-E trending belt of thick meltout till and stagnation topography. Ice flow during the Big Rice phase was generally toward the SSW (bearing 190°). The third major moraine recognized in the TMM AOI is the Wahlsten moraine, an E-W trending belt of thick meltout till and stagnant ice topography. Ice flow during the Wampus phase was generally toward the S (bearing 180-190°), reflecting a significant reorientation in ice flow of the Laurentide ice sheet. Annual recessional moraines associated with the Wampus and Wahlsten phases consist of both subaqueous moraines composed of sand, gravel, and meltout tills deposited in Glacial Lake Norwood, and subaerial moraines predominantly composed of meltout tills. Vermilion Phase The fourth and final major moraine recognized in the GLN basin is the Vermilion moraine, a 40 m high, 1-2 km wide, WNW-ESE trending belt of thicker meltout till, stagnant ice topography, and subaqueous debris flow fans. Ice flow during the Vermilion phase was generally toward the SSW (bearing 195-205°). The Vermilion moraine truncates the eastern extent of the Wahlsten moraine, reflecting a further significant reorientation in ice flow of the Laurentide ice sheet. The next moraine formed by a major surgestagnation event lies >100 km to the northeast, suggesting an interval of >1000 years of gradual ice margin retreat after the Vermilion phase. Glacial Lake Norwood Retreat of the Rainy lobe margin north of the continental height of land at the Giants Ridge dramatically changed the character of sedimentation associated with the Laurentide ice sheet. South of the divide, meltwater generally flowed downslope away from the margin, depositing outwash in channels and as outwash plains with intervening rolling plains of subglacial lodgment till or moraines composed of hummocky supraglacial meltout till. Immediately upon marginal retreat north of the divide, ponding of meltwater against the ice sheet formed the first of a nearly continuous succession of proglacial lakes. Glacial 192 �Trip 8 – Glacial meltwater and other precipitation ponded against the ice sheet overflowed to the south through a series of successively lower outlets over the height of land, a process that continued until final collapse of the ice sheet in Hudson Bay. Initially, a series of ephemeral lakes formed in stagnant ice north of the divide. These lakes were dammed in part by the advance of the St. Louis Sublobe into the Glacial Lake Upham I basin (see Knaeble et al. (2005) and Larson et al. (2014)) Associated strandlines and meltwater channels are only poorly defined, and meltwater likely drained southward through stagnant Rainy lobe ice karst and St. Louis sublobe ice. Once the active Rainy lobe ice margin receded to the Allen moraine, a stable, relatively long-lived meltwater outlet was established through the Embarrass Gap. By definition, the first proglacial lake located north of the Laurentian Divide that drained through the Embarrass Gap is referred to as Glacial Lake Norwood. Three well-developed outlets to Glacial Lake Norwood are recognized, corresponding to relatively stable, long-lived lake levels. These are herein referred to as Glacial Lakes Norwood I, II, and III, corresponding to successively older and lower lake levels. The initial stable lake level (Glacial Lake Norwood I) was controlled by an outlet channel with a modern floor elevation of about 450 m amsl. This channel was bounded by the Giants Range ridge to the south, and the Allen moraine to the north. The Allen moraine at this location is a major recessional moraine, approximately 500 m wide with in excess of 15 m of vertical relief above the meltwater channel. The ice-cored Allen moraine formed an effective barrier to meltwater drainage blocking most of the Embarrass Gap until after the ice sheet margin retreated from the Vermilion moraine, a time interval of 100s to 1000s of years. Incision of the Glacial Lake Norwood I outlet was inhibited during this time interval in part because the channel was graded to its downstream inlet into Glacial Lake Upham II; only after drainage of this lake was further significant erosion and channel incision in the Embarrass Gap initiated (Larson and Mooers 2009). Gradual collapse of the Allen moraine due to ice melt led to resulted in an episode of collapse and downcutting of the moraine dam, and establishment of a second, lower stable outlet level for Glacial Lake Norwood II in the Embarrass Gap at a modern floor elevation of about 443 m amsl. Paleoislands of outwash and esker sediment located in Glacial Lake Norwood considerable distances north of the Vermilion moraine display well-developed shoreline features corresponding to this outlet, indicating that the downcutting episode occurred well after ice margin retreat from the Vermilion moraine, and that Glacial Lake Norwood II stood at this stable lake level for a relatively long time interval. A second collapse and downcutting episode through the Big Rice moraine led to establishment of the third, and final, lower stable outlet level corresponding to Glacial Lake Norwood III in the Embarrass Gap. An outlet with a modern floor elevation of about 433 m amsl corresponds to a second, lower welldeveloped strandline on esker and outwash paleoislands to the north. During its relatively long history, Glacial Lake Norwood expanded along the receding ice margin to form a lake that ultimately extended ~400 km E-W and in excess of 100 km N-S. The lake experienced two major ice re-advances into its western arm, evidenced by thick (>70 m) accumulations of glaciolacustrine sediment and till. The large fetch of the lake resulted in vigorous wave erosion along its shoreline and the considerable fraction of the lakebed situated above wave base. Final drainage of Glacial Lake Norwood III occurred when a western outlet (the McIntosh spillway) flowing into an early (Herman) level of Glacial Lake Agassiz formed in the vicinity of Trail, MN, 260 km to the west of the Embarrass Gap. 193 �Trip 8 – Glacial Figure 8-2. Outline of maximum extent of Glacial Lake Norwood in northern Minnesota. The lake extended over 350 km from east to west. The final outflow was westward into Glacial Lake Climax near Trail, MN. Glacial Lake Norwood sediments generally consist of gravels and sands in littoral (shallow) environments, reflecting local reworking of till and outwash by wave action, and silt and clay in benthic environments, reflecting settling of suspended fine-grained sediment from the water column. In general, Glacial Lake Norwood sediment sequences fine upward, reflecting diminished wave erosion and the increasing distance of the primary sediment source (the receding ice margin). Koochiching Lobe Subsequent to retreat of the Rainy lobe from the Vermilion moraine, Koochiching lobe (KL) ice re-advanced into the Glacial Lake Norwood basin, this time from the west and the Red River lobe (Meyer, 1993). In marked contrast to the sandy-textured till and glaciofluvial sediment associated with the Rainy lobe, KL diamicton is calcareous, and distinctly finer grained than Rainy lobe till; these sediments are placed in the Blackduck Formation in the MGS lithostratigraphic framework (Johnson et al., 2016). Based on rotosonic drilling, diamictons associated with at least two distinct advances into the GLN basin are present in the field trip area, separated by fine-grained glaciolacustrine sediment. The genesis of these diamictons – till or subaqueous debris flow – are enigmatic; fine-grained lacustrine sediment may grade upward into normally consolidated diamicton, which may grade upward into overconsolidated diamicton of similar composition. The Koochiching lobe advances overran older Rainy lobe landforms, including the Vermilion moraine. There is little evidence for erosion and entrainment of older glacigenic sediment by the KL, and no well-defined moraines or other landforms define the limits of the advances. Sediment was deposited from suspended sediment plumes or debris flows in the proglacial GLN, or as subglacial lodgment till. Although KL ice thickness is unknown, it was sufficiently thick relative to the depth of GLN to preclude development of a calving margin. 194 �Trip 8 – Glacial The majority of sediment shed by the advancing Koochiching lobe was deposited as glaciolacustrine sediment in GLN, and subject to a high degree of reworking in the lacustrine environment. Bedrock highs – shallow areas in GLN – are typically devoid of either older Rainy lobe or KL sediments. In places, a thin boulder lag containing limestone and dolomite clasts attests to the former presence of KL diamicton. In contrast, thicknesses of up to 70 m of till and glaciolacustrine sediment have been reported in intervening bedrock lows. Two distinct till compositions attesting to two distinct source areas have been reported in KL sediments. The younger, overlapping KL till bears greater similarity to calcareous Red River and Des Moines lobe tills elsewhere in Minnesota. In contrast, an older KL till is characterized by a distinctly higher Na2O content, similar to tills exposed at surface in Hubbard and Wadena Counties. Even as the Laurentide ice sheet margin was broadly retreating from Minnesota, both from the Red River Valley and from the Arrowhead, the advances of the Koochiching lobe into the GLN served to block development of meltwater outlets to the north and west. Ultimately, stagnation and wasting of the KL led to westward propagation of GLN until development of the McIntosh spillway. The massive sediment accumulations associated with the KL – up to 70 m in places as previously noted – were deposited over a time interval on the order of 1000 years. Description of Field Trip Stops Stop 1: Glacial Lake Norwood strandline 498550E/5283740N (UTM Zone 15, NAD83) (47.70704, -93.0193) Side Lake 7.5’ USGS Quadrangle This site is located on the uppermost relatively well-developed beach associated with Glacial Lake Norwood. A well-developed boulder lag and wave-cut notch attest to a relatively long-lived stable lake at this level characterized by energetic wave action. To the south, ice collapse pits in the subaqueous deposited Big Rice moraine evidence long lived stagnant ice along this moraine trend. Locally, the Big Rice and other moraines served as ice-cored dams preventing southern outflow. Stop 2: Gravel pit in minor Rainy lobe recessional moraine 498480E/5294980N (UTM Zone 15, NAD83) (47.80817, -93.0203014) Bear River 7.5’ USGS Quadrangle Here a small, sharp-crested subaqueous deposited recessional moraine has been developed into a gravel pit. The flanks of the moraine are draped by fine-grained glaciolacustrine sediment deposited in Glacial Lake Norwood. 195 �Trip 8 – Glacial Stop 3: Gravel pit in large Rainy lobe recessional moraine 495260E/5302290N (UTM Zone 15, NAD83) (47.8739285, -93.0633884) Bear River 7.5’ USGS Quadrangle This gravel pit is developed in a large subaqueous ice marginal fan(?) deposited at the margin of the retreating Rainy lobe. The fan was of sufficient height that its surface was above the GLN wave base, precluding deposition of finer-grained glaciolacustrine sediment. The presence of limestone and dolomite boulders on the fan surface indicate that this area was overrun by Koochiching lobe ice. Stop 4: Wave-washed bedrock high 492880E/5301180N (UTM Zone 15, NAD83) (47.8639194, -93.095198) Bear River 7.5’ USGS Quadrangle This wave-scoured bedrock high evidences the intensity of wave action in Glacial Lake Norwood. Rainy lobe sediment has been almost completely washed away, no Koochiching lobe sediment is preserved, and no glaciolacustrine sediment has been deposited. The very large boulder – a Rainy lobe erratic - attests to the ‘minimum’ particle size of this ‘boulder lag’. Stop 5: Glacial striae and grooves 489180E/5304010N (UTM Zone 15, NAD83) (47.8893303, -93.1447397) Rauch 7.5’ USGS Quadrangle Glacial striae and grooves on outcrop on either side of the road at this stop preserve evidence of ice flow directions for both the Rainy lobe (bearing 190° and 205°) and the later Koochiching lobe (bearing 140°). This indicates that Rainy lobe sediment was largely stripped from bedrock highs by wave action in Glacial Lake Norwood prior to advance of the Koochiching lobe from the west. 196 �Trip 8 – Glacial Stop 6: Borrow pit in reworked calcareous Koochiching lobe drift 490750E/5305460N (UTM Zone 15, NAD83) (47.9024009, -93.1237689) Silverdale 7.5’ USGS Quadrangle This small borrow pit on the margin of a wave-scoured bedrock high contains abundant carbonate pebbles and cobbles. These originated from calcareous Koochiching lobe till deposited on the bedrock high and later eroded by wave action. Stop 7: Slumping Koochiching lobe till and Glacial Lake Norwood 491460E/5311020N (UTM Zone 15, NAD83) (47.9524357, -93.114379) Silverdale 7.5’ USGS Quadrangle This site exposes a sequence of interbedded Koochiching lobe diamicton (till and debris flows(?)) and fine-grained glaciolacustrine sediment adjacent to the Littlefork River. The slope, already prone to slumping by stream erosion at the toe, was further destabilized by construction of the road. In the near vicinity to the southwest, an exploration rotosonic borehole intercepted around 70 m of such sediment. Stop 8: Samuelson Park 492580E/5310600N (UTM Zone 15, NAD83) (47.9459716, -93.0993661) Silverdale 7.5’ USGS Quadrangle Bedrock underlying the small waterfall in the Littlefork River has been striated by the Rainy lobe (bearing 196°). In the upstream direction, boulders eroded from the basal Rainy lobe lodgment till are visible in the stream bed and banks. Such bedrock and boulder lags serve as knickpoints defining the bed of the Littlefork River; steep and commonly slumping slopes adjacent to the river attest to the significant erosion of Koochiching lobe and Glacial Lake Norwood sediment during the Holocene. 197 �Trip 8 – Glacial Stop 9: Embarrass Gap 9A: 551980/5270340N (UTM Zone 15, NAD83) (47.583892, -92.3087077) 9B: 552150/5272700N (UTM Zone 15, NAD83) (47.6056089, -92.3061663) 9C: 552760/5272950N (UTM Zone 15, NAD83) (47.6078088, -92.2980211) Biwabik 7.5’ USGS Quadrangle These three stops are in the three successive major outlet channels for Glacial Lake Norwood. Stop 9A (elevation 450 m) is in a meltwater channel developed at the margin of the Rainy lobe, perhaps against an active ice margin. Stops 9B (elevation 443 m) and 9C (elevation 433 m) are two successively lower major outlets formed as the ice-cored Allen moraine collapsed over a time interval on the order of 1000 years. The outlet at 9C served as the stable outlet to Glacial Lake Norwood until opening of its final lower outlet to the west, through the McIntosh spillway in the vicinity of Trail, Minnesota. 198 �Trip 8 – Glacial REFERENCES Bauer, Emily J., Mark A. Jirsa, Amy Radakovich Block, Terrence J. Boerboom, Val W. Chandler, Dean M Peterson, Kaleb G. Wagner, Elizabeth L. McDonald, Jennifer M. Dengler, Gary N. Meyer, and Jacqueline D. Hamilton. 2022. “Geologic Atlas of St. Louis County, Minnesota.” Minnesota Geological Survey County Atlas Series C–51. Björck, Svante. 1990. “Late Wisconsin History North of the Giants Range, Northern Minnesota, Inferred from Complex Stratigraphy.” Quaternary Research 33:18–36. Breckenridge, Andrew J. 2015. “The Tintah-Campbell Gap and Implications for Glacial Lake Agassiz Drainage during the Younger Dryas Cold Interval.” Quaternary Science Reviews 117:124–34. https://doi.org/10.1016/j.quascirev.2015.04.009. Buchan, Kenneth L., Henry C. Halls, and James K. Mortensen. 1996. “Paleomagnetism, U-Pb Geochronology, and Geochemistry of Marathon Dykes, Superior Province, and Comparison with the Fort Frances Swarm.” Canadian Journal of Earth Sciences 33:1583–95. Clayton, Lee, and Stephen R. Moran. 1982. “Chronology of Late Wisconsinan Glaciation in Middle North America.” Quaternary Science Reviews 1:55–82. Cotter, Ralph D., H.L. Young, and Thomas C. Winter. 1964. “Preliminary Surficial Geologic Map of the Mesabi-Vermilion Iron Range Area, Minnesota.” USGS Miscellaneous Geologic Investigations Map I-403. Elftman, A.H. 1898. “The Geology of the Keweenawan Area in Northeastern Minnesota, Part I.” The American Geologist 21:90–109. Hobbs, Howard C. 1983. “Drainage Relationships of Glacial Lakes Aitkin and Upham and Early Lake Agassiz in Northeastern Minnesota.” Edited by James T. Teller and Lee Clayton. Geological Association of Canada Special Paper 26:245–59. Johnson, Mark D., Roberta S. Adams, Angela S. Gowan, Kenneth L. Harris, Howard C. Hobbs, Carrie E. Jennings, Alan R. Knaeble, Barbara A. Lusardi, and Gary N. Meyer. 2016. “Quaternary Lithostratigraphic Units of Minnesota.” Minnesota Geological Survey Report of Investigations 68:262. Knaeble, Alan R., Gary N. Meyer, Lisa M. Marlow, Phillip C. Larson, and Howard D. Mooers. 2005. “Deposits and Landforms in the Region Glaciated by the St. Louis Sublobe.” In Field Trip Guidebook for Selected Geology in Minnesota and Wisconsin, edited by Lori Robinson, Guidebook, 40–79. Minneapolis: Minnesota Geological Survey. Larson, Phillip C., Alan R. Knaeble, Howard D. Mooers, and Lisa M. Marlow. 2014. “The St. Louis Sublobe and Glacial Lake Upham.” Institute on Lake Superior Geology Field Trip Guidebook 60:102–18. Larson, Phillip C., and Howard D. Mooers. 2009. “Glacial Geology of the Vermilion Moraine.” Institute on Lake Superior Geology Field Trip Guidebook 55 (2): 81–99. Lehr, James D., and Howard C. Hobbs. 1992. “Glacial Geology of the Laurentian Divide Area, St. Louis and Lake Counties, Minnesota.” Minnesota Geological Survey Guidebook 18:82. Lepper, Kenneth, Timothy G. Fisher, Irka Hajdas, and Thomas V. Lowell. 2007. “Ages for the Big Stone Moraine and the Oldest Beaches of Glacial Lake Agassiz : Implications for Deglaciation Chronology.” Geology 35 (7): 667–70. https://doi.org/10.1130/G23665A.1. Leverett, Frank. 1932. “Quaternary Geology of Minnesota and Parts of Adjacent States.” USGS Professional Paper 161:149. 199 �Trip 8 – Glacial Lowell, Thomas V., Timothy G. Fisher, Irka Hajdas, K. Glover, Henry M. Loope, and T. Henry. 2009. “Radiocarbon Deglaciation Chronology of the Thunder Bay, Ontario Area and Implications for Ice Sheet Retreat Patterns.” Quaternary Science Reviews 28 (17–18): 1597–1607. https://doi.org/10.1016/j.quascirev.2009.02.025. MNDNR. 2012. “LiDAR Elevation, Arrowhead Region, NE Minnesota, 2011.” Minnesota Department of Natural Resources. ftp://ftp.gisdata.mn.gov/pub/gdrs/data/pub/us_mn_state_mngeo/elev_lidar_arrowhead2011/metadata/meta data.html. Mooers, Howard D., and James D. Lehr. 1997. “Terrestrial Record of Laurentide Ice Sheet Reorganization during Heinrich Events.” Geology, no. 11, 987–90. Prest, Victor K., J. Allan Donaldson, and Howard D. Mooers. 2000. “The Omar Story: The Role of Omars in Assessing Glacial History of West-Central North America.” Géographie Physique et Quaternaire 54:257–70. Southwick, David L., and Henry C. Halls. 1987. “Compositional Characteristics of the Kenora-Kabetogama Dyke Swarm (Early Proterozoic), Minnesota and Ontario.” Canadian Journal of Earth Sciences 24:2197– 2205. Stark, James R. 1977. “Surficial Geology and Ground-Water Geology of the Babbitt-Kawishiwi Area, Northeastern Minnesota with Planning Implications.” M.S. Thesis. M.S. Thesis, University of Wisconsin. Upham, Warren. 1894. “Preliminary Report of the Field Work during 1893 in Northeastern Minnesota, Chiefly Relating to the Glacial Drift.” Geological and Natural History Survey of Minnesota Annual Report 22:18–86. Winchell, Newton H. 1899. “The Geology of the North Part of St. Louis County.” Geological and Natural History Survey of Minnesota 4:222–65. ———. 1900. “Glacial Lakes of Minnesota.” Geological Society of America Bulletin 12:109–28. Winter, Thomas C. 1971. “Sequence of Glaciation in the Mesabi-Vermilion Iron Range Area, Northeastern Minnesota.” USGS Professional Paper 750–C:C82–88. ———. 1973. “Hydrogeology of Glacial Drift, Mesabi Iron Range, Northeastern Minnesota.” USGS Water Supply Paper 2029-A:31. Winter, Thomas C., Ralph D. Cotter, and H.L. Young. 1973. “Petrography and Stratigraphy of Glacial Drift, Iron Range Area, Northeastern Minnesota.” USGS Bulletin 1331–C:50. Wright, Herbert E. 1956. “Sequence of Glaciation in Eastern Minnesota.” Geological Society of America Guidebook 3:1–24. ———. 1972. “Quaternary History of Minnesota.” In Geology of Minnesota: A Centennial Volume, edited by Paul K. Sims and G.B. Morey, 515–47. St. Paul, Minnesota. Wright, Herbert E., and William A. Watts. 1969. “Glacial and Vegetational History of Northeastern Minnesota.” Minnesota Geological Survey Special Publication 11. 200 � 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, 2025 Description An account of the resource Institute on Lake Superior Geology. Mountain Iron, Minnesota. May 14-17, 2025. 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 2025-05-14 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/18fd51ee65642289cfb33fc14030102e.pdf 71db64055c86846756ee627afe76f091 PDF Text Text 70th Annual Meeting Institute on Lake Superior Geology Houghton, Michigan May 15-18, 2024 Proceedings Volume 70 Part 1 - Program and Abstracts �70th Annual Meeting Institute on Lake Superior Geology Houghton, Michigan May 15-18, 2024 Sponsored by: A. E. Seaman Mineral Museum Great Lakes Research Center Department of Geological and Mining Engineering and Sciences Michigan Technological University Meeting Co-Chairs Theodore J. Bornhorst, Erika C. Vye, Patrice Cobin, and James DeGraff Proceedings Volume 70 Part 1: Program and Abstracts Edited by Theodore J. Bornhorst and Erika C. Vye Cover Photo: The only known color photograph of in situ colorless calcite crystals with inclusions of native copper. Vug is about 15 cm across and 30 cm deep; located at the top of the Knowlton basalt lava flow at the 4 th level, 850 ft stope, of the Caledonia Mine, Michigan. Photo taken in 1994 soon after the vug was blasted open. Native copper in the calcite crystals has not been visibly altered despite being about 1 billion years old. Photograph by Theodore J. Bornhorst i �70th Institute on Lake Superior Geology Volume 70 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trip 1: Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan Trip 2: Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan Trip 3: Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty Trip 4: Keweenaw Fault System Geometry and Kinematics: Clues to Its Nature and Origin Trip 5: Geology and History of a Native Copper Mine: Adventure Mine, Ontonagon County, Michigan Trip 6: Southern Complex Granitoids, Gneisses, and Migmatites: New Data, Discoveries, and Perspectives Trip 7: Landslides on the Ontonagon River at Military Hill Reference to material in Part 2 should follow the example below: Authors, 2024, Field Trip title, 70th Institute on Lake Superior Geology, Abstracts and Proceedings, v. 70, Part 2, Field Trip Guidebook, p. xx-xx. Proceedings Volume 70, Part 1: Program and Abstracts and Part 2: Field Trip Guidebook are published by the 70th 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 black and white. 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 �Part 1: Program and Abstracts Table of Contents Institutes on Lake Superior Geology, 1955-2024 ............................................................. iv Sam Goldich and the Goldich Medal ............................................................................... vii Goldich Medal Guidelines ................................................................................................ ix Goldich Medalists ............................................................................................................. xi 2024 Goldich Medal Recipient ......................................................................................... xi Goldich Medal Committee ............................................................................................... xi Citation for 2024 Goldich Medal Recipient..................................................................... xii Honoring the Pioneers of Lake Superior Geology……………………………………….xiv Citation for 2024 Pioneer of Lake Superior Geology Recipient………………………...xv In Memoriam……………………………………………………………………………xix Eisenbrey Student Travel Awards ................................................................................... xx Joe Mancuso Student Research Award ........................................................................... xxi Doug Duskin Student Paper Awards and 2024 Student Paper Awards Committee ...... xxii Board of Directors and 2024 ILSG Meeting Volunteers .............................................. xxiii 2024 ILSG Meeting Volunteers and Session Chairs…………………………………..xxiv Field Trip Leaders and Guidebook Authors .................................................................. xxv Banquet Speaker Robert M Hazen ................................................................................ xxvi Report of the Chair of the 69th Annual Meeting ........................................................ xxvii Sponsors ........................................................................................................................ xxxi Technical Program ....................................................................................................... xxxii Abstracts ........................................................................................................................ xliii iii �Institutes on Lake Superior Geology, 1955-2024 # 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 iv �# 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 v 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 �# 70 Date Place 2024 Houghton, Michigan Chairs T.J. Bornhorst, E. Vye, P. Cobin, & J. DeGraff vi �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. Kalliokoski, 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 �INSTITUTE ON LAKE SUPERIOR GEOLOGY GOLDICH MEDAL viii �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 medalists 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 medalist, 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 �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 �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 2024 Suzanne W. Nicholson 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 2024 GOLDICH MEDAL RECIPIENT Suzanne W. Nicholson Goldich Medal Committee Serving through the meeting year shown in parentheses. Dorothy Campbell (2019-2024) Ontario Geological Survey, Government Member (Committee Chair) Dean Peterson (2022-2025) Big Rock Exploration, Industry Member Marcia Bjornerud (2023-2026) Lakehead University, Academic Member xi �Citation for the 2024 Goldich Medal Recipient Suzanne W. Nicholson It is a pleasure and honor to present the 2024 Goldich Medal to our close friend and colleague, Suzanne Nicholson, recently retired from a long and fruitful career at the U.S. Geological Survey. Suzanne began working for the USGS as a student field assistant in 1978, and then, after completing a master’s degree at the University of Massachusetts, was hired as a full-time employee in 1981. Suzanne’s interest in the geology of the Lake Superior region began with her dissertation work with Paul Weiblen at the University of Minnesota on felsic magmatism in the Portage Lake Volcanics in Michigan, part of the Mesoproterozoic Midcontinent Rift System (MRS). This study included detailed mapping and sampling (typically big samples that had to be carried long distances) followed by major and trace element whole rock analysis and determination of a suite of radiogenic isotopes (Sr, Nd, Pb). Using modern petrologic methods, her research documented the presence of two distinct felsic magma types, one derived by partial melting of felsic basement and the other related to rift basalts through partial melting and/or fractional crystallization. Suzanne also was one of the first to provide comprehensive radiogenic isotope analyses of host basalts, documenting their distinctive isotopic character and that of their mantle sources. Her 1990 seminal publication (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. 10,85110,868) described the unique geochemical character of rift magmatism around the Lake Superior region. Her on-going interest in MRS geochemistry culminated in the 1997 paper that established a rift-wide correlation of MRS basalts (Nicholson, S.W., Schulz, K.J., Shirey, S.B., and Green, J.B., 1997, Rift-wide correlation of 1.1 Ga Midcontinent Rift System basalts: multiple mantle sources during rift development: Canadian Journal of Earth Sciences, v. 34, p. 504-520), providing a foundation for future interpretations of MRS-related volcanic rock geochemistry. Suzanne was also a leader in advancing an understanding of the spatial-temporal evolution of MRS metallogeny (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), further refined in a 2020 paper (Woodruff, L.G., Schulz, K.J., Nicholson, S.W., and Dicken, C.L., 2020, Mineral deposits of the Mesoproterozoic Midcontinent Rift system in the Lake Superior region A space and time classification: Ore Geology Reviews, 103716). Along with colleagues from the USGS, Suzanne helped produce a series of 1:100,000-scale geologic maps for the MRS and adjacent rocks from the Keweenaw Peninsula, extending through Michigan into northern Wisconsin to the Minnesota state line. These maps summarized legacy mapping and, along with new fieldwork, resulted in interpretations and correlations that xii �are the current standard for understanding the distribution and origin of the MRS volcanic and intrusive rocks of that area. Suzanne also initiated and continues to lead an on-going cooperative government/academia effort to compile and digitize existing MRS geology, geochemistry, isotope data, and age dates that will promote and direct future research of the region. Throughout her career, Suzanne was a careful and meticulous researcher who held her own results to a very high standard for accuracy, completeness, and thoroughly documented interpretations. Through the years, Suzanne has been a strong supporter of the Institute on Lake Superior Geology. She was a first or co-author on 17 abstracts presented at ILSG meetings from 1990 through 2019, a co-leader for two ILSG field trips, and co-editor for the 1996 Proceedings, Part 1- Program and Abstracts volume. Suzanne also was always willing and able to help with anything needed at ILSG meetings (a common trait among ILSG participants), such as acting as a session chair or serving on the student paper committee. In 2015, Suzanne moved into increasingly responsible managerial positions within the USGS, which curtailed her direct involvement with research in the Lake Superior region. In 2020, she received the U.S. Department of Interior's second highest honorary award—the Meritorious Service Award— in recognition of her scientific leadership and noteworthy contributions to the USGS Mineral Resources Program. Suzanne retired from her position as Associate Program Coordinator for the USGS Mineral Resources Program in 2021 but was retained for 2 years as an annuitant to keep the Program on budgetary track during a time of transition. Her qualities as a scientist transferred to her administrative duties, demonstrating the same dedication and skills she brought to her research. Now that Suzanne’s service to the Program has ended, we look forward to her return to MRSrelated research as a USGS Emeritus scientist. Throughout her managerial tenure, Suzanne never lost her attachment to the Lake Superior region and was able to promote and maintain funding for ongoing regional project work for her USGS colleagues. This support resulted in many new and exciting discoveries, such as tracing the extent and nature of the Sudbury ejecta layer across Michigan and Wisconsin, and tackling legacy seismic data to help understand the tectonicmagmatic evolution of the MRS. Through her thoughtful discussions, critical reviews, cheerful field assistance, and friendship for the past 40-some years, Suzanne helped enrich the lives and careers of many people, including those of her fellow USGS MRS aficionados. We remain a convivial group and all of us look back fondly on the times we spent together in the field. Who could forget death marches across Isle Royale, or raccoons swiping rhyolite samples in the Porcupine Mountains, or six long weeks at the Hurley Holiday Inn, among our many other adventures? So now, the three of us, all former recipients of the Goldich Medal, are joined by Suzanne in that honor. In recognition of her decades of accomplishments and dedication to the geology of the Lake Superior region and to the Institute on Lake Superior Geology, it is our pleasure to present the 2024 Goldich Medal to its second female recipient, Suzanne Nicholson. Citation by: Laurel G. Woodruff, USGS, Goldich Medal Winner, 2014 Klaus J. Schulz, USGS, retired, Goldich Medal Winner, 2003 William F. Cannon, USGS, Emeritus, Goldich Medal Winner, 1992 xiii �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 the inception of the Institute in 1955. Biographical sketches of those pioneers will be presented at future annual meetings so that all may 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 summarize the contribution of the nominee. 2) The Organizing 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) 2024 Roland Duer Irving (1847-1888) xiv �2024 Citation for the Roland Duer Irving Pioneer of Lake Superior Geology It is my great honor to nominate and promote Roland Duer Irving (1847-1888) as the 2024 Pioneer of Lake Superior Geology. I suspect that many ILSG members are unfamiliar with Dr. Irving and his many truly pioneering contributions to our understanding of various aspects of Lake Superior geology. Were it not for his premature death at the age of 41, I have no doubt that his continued work on the Precambrian geology of the Lake Superior region would have ranked him as one of the greatest Lake Superior geologists of his time. As it stands, his nearly two decades of mapping, petrography, and geochemical studies and mentoring of students at the University of Wisconsin provided a firm and rational foundation for our further understanding of Lake Superior geology. Roland Duer Irving Roland Duer Irving was born in New York City on April 29th, 1847 (1847 – 1888) as grand-nephew to the classic American novelist-essayist Washington Irving and the New York State Supreme Court Justice, John Duer. John Wesley Powell (2nd USGS Director 1881-1894) noted in his memoriam of Dr. Irving (Powell, 1891) that in his youth, “Roland was subject to frequent and alarming attacks of illness, also to a weakness of sight, which proved to be his greatest obstacle through life” and as such “his early education was at home, his sisters and his father being his instructor”. Ultimately, he enrolled at Columbia College School of Mines in 1863, and with the continued help of his sisters, he graduated in 1868 with a degree in mining engineering. During and after his time at Columbia, he worked for coal mines and smelters in Pennsylvania and New Jersey. In 1870, he was offered a mining and metallurgy chair position at the University of Wisconsin. Irving’s arrival at the University of Wisconsin in 1870 marked the emergence of the “Wisconsin School of Precambrian Geology” (Dott, 2001). He quickly gained prominence within the university as a faculty leader and outside the university as a research investigator (Curti and Carstenson, 1949). Soon after his arrival, the Wisconsin Geological Survey was established by the legislature in 1873 with Irving, T.C. Chamberlain, and Moses Strong serving assistant geologists. By 1876, Chamberlin took the reins as chief geologist of the survey, a position he would hold until its legislative termination in 1879. In 1880, Clarence King, first director of the US Geological Survey, recruited both Chamberlin and Irving to join the USGS in an effort to develop a geologic map of the entire United States. In 1881, Chamberlin was appointed director of the Glacial Division of the USGS. In 1882, Irving was appointed to head the USGS’s Lake Superior Precambrian Division, all the while continuing as head of the Department of Mineralogy and Geology at the University of Wisconsin. xv �During Roland Irving’s teaching and research time with the University of Wisconsin, the Wisconsin Geological Survey and the USGS, he came to mentor and collaborate with several notable geologists who would make their own mark on Lake Superior geology (Dott, 2001). Charles Van Hise arrived at UW as a geology student under Irving’s supervision in 1874. He completed his BS in 1879, his MS in 1882, and his PhD in 1892 (1st PhD at UW). With the passing of Dr. Irving in 1888, Van Hise became not only the principal geologist for the USGS’s Lake Superior Division, but also the head of Wisconsin’s geology department. Another notable student of Roland Irving’s at Wisconsin was Florence Bascom. She conducted a petrographic study of the Mellen Complex under the supervision of Irving and Van Hise and received the second ever MS degree in geology from UW in 1887. She later earned her doctorate degree from Johns Hopkins in 1893, the first woman in the US to be awarded a PhD in geology. Irving’s work with the Wisconsin Geological Survey (1873-1879) involved many aspects of Wisconsin geology. In Volume 1 (actually published last in 1883), which was intended to be a general summary of the geology, natural history and economic geology of the state for the general education of the public, Irving contributed chapters on the minerals, rock types, and iron ores of the state. In Volume 2 (1877), Irving reported on the Precambrian, Paleozoic, and Quaternary geology of central Wisconsin. His descriptions of the general structure and lithologic attributes of the Baraboo Quartzite and its unconformable relationship with adjacent “Silurian” (Paleozoic) rocks is particularly noteworthy. Volume 3 (1880), which focussed on the geology of Northern Wisconsin, included Irving’s summary report on the general geology of the Lake Superior region (Part 1) and a more detailed report on the geology of the eastern Lake Superior District (Part III). This work, which was based on field studies conducted between 1875 and 1878, formed the basis of his subsequent USGS work detailing the overall geology and structure of the Keweenawan System in the Lake Superior region. It is noteworthy that the renown petrographer, Raphel Pumpelly, contributed a chapter on the petrography of Keweenawan rocks (Part II) collected by Irving and others. Irving relied heavily on petrographic examination of field samples in his subsequent USGS work. In the final volume (#4, 1882) devoted to the geology, paleontology, natural history, and glacial geology of the southern half of the state Irving’s contribution focussed on the field and petrographic attributes of crystalline rocks of the Wisconsin River valley. He recruited his MS student, Charles Van Hise, to carry out most of the petrographic descriptions. Joining the US Geological Survey in 1880 as head of the Lake Superior Precambrian Division, Irving took advantage of being able to explore beyond the confines of Wisconsin and immediately embarked on his long-standing desire to produce a “resume of the results obtained in the Lake Superior country by other geologists up to the present” (Geology of Wisconsin, Volume III (1880) Part 1, p. 3). Building on his own studies of the Keweenawan System in northern Wisconsin, he reviewed and, where appropriate, integrated all former geologic studies in the Lake Superior dating back to the Michigan surveys of Douglass Houghton (1831-1844), the surveys of Upper Canada starting with Logan (1846), and the work of Joseph Norwood in northeastern Minnesota as part of the D.D. Owen US Survey (1847-1852). Between July 1880 and March 1882, Irving conducted reconnaissance mapping, along with a crew of five assistant geologists, in several poorly understood areas throughout the Lake Superior basin. xvi �In 1880, the Minnesota Geological Survey, headed by N.H. Winchell, was in its 9th year of existence, but had only just begun to map the Precambrian geology of the state. As such, Irving decided to spend much of his mapping efforts on the north shore of Lake Superior between Duluth and Nipigon Bay to ascertain how it correlated with the south shore. This occurred at a time when many frontier states were developing their own geologic surveys with the expressed purpose of excluding the federal survey. The USGS already had a strong foothold in Michigan and after the ending of the Wisconsin survey in 1879, developed a strong presence there as well. Suffice it to say that Irving’s work in Minnesota was not well received or valued by the Winchell’s Minnesota Survey. Notwithstanding Winchell objections, Irving’s publication of USGS Monograph 5 - The Copperbearing Rocks of Lake Superior (Irving, 1883) proved to be a remarkably complete and accurate picture of the geology and structure of Keweenawan System (Figure 1). The many important observations and interpretations about Keweenawan geology put forth by Irving include: • formalizing the lithostratigraphy of the Keweenawan System • • defining the synclinal structure of the lavas in the Lake Superior area recognizing that eruptive rocks consist of basic, intermediate, and felsic types • noting no obvious relation of volcanic type to stratigraphic position • interpreting that basic lavas were erupted subaerially from fissures, not ash-generating volcanoes • recognizing that amygdaloidal zones capping basalts are themselves volcanic (not sedimentary) • accurately estimating the thickness of the North Shore Volcanics to be about 18,000’ • interpreting gabbroic and granitic rocks to be intrusive into the volcanic rocks (thus younger) and likely formed in staging chambers that fed surface eruptions Following on the publication of Monograph 5, Irving continued to apply his geologic and petrographic expertise to studies of other Precambrian systems (greenstones, quartzites, and iron formations) in collaboration with students and USGS colleagues. When Roland Irving unexpectedly died (from “paralysis”, perhaps a stroke) on May 30, 1888, he was engaged with Van Hise on another USGS monograph (#19) on the Gogebic Iron Range, which was published posthumously (Irvine and Van Hise, 1892). This monograph launched Charles Van Hise on a career path to becoming an internationally recognized expert on Lake Superior iron formations. While Van Hise will ultimately be recognized as pioneer of Lake Superior geology, it is fitting that we first acknowledge the remarkable accomplishments of his advisor and mentor, Roland Duer Irving. One can only imagine the professional stature he would have attained were he not struck down at the peak of his creativity and expertise. xvii �Figure 1: Plate 1 of USGS Monograph 5 by R.D. Irving, 1883 References Curti, M., and Carstenson, V., 1949, The University of Wisconsin, A History, 1848-1925 (v. 1). Madison, Wisconsin, University of Wisconsin Press, 739 p. Dott, Robert H., Jr., 2001, The remarkable legacy of the Wisconsin School of Precambrian Geology. Geoscience Wisconsin, v. 18, p. 27-40. Irving, R.D., 1883, The Copper-bearing Rocks of Lake Superior. USGS Monograph 5, 464p. Irving, R.D., and Van Hise, C.R., 1892, Penokee Iron-Bearing Series of Michigan and Wisconsin. USGS Monograph 19, 534p. Powell, J.W., 1891, Roland Duer Irving. Eleventh Annual report of the Director of the United States Geological Survey, Part 1- Geology: 1889-1890 p. 38-42. Citation by: James Miller University of Minnesota-Duluth xviii �In Memoriam Louis Mattson Obituary 12/31/2023 Louis A. Mattson, 89, of Pengilly, MN. passed away December 31, 2023, in Grand Rapids, MN. He was a long-time member of the Institute on Lake Superior Geology. The son of commercial fishermen, Lou was the last surviving member of the Mattson Tobin Harbor Fishery on Michigan’s Isle Royale. The fishery on Isle Royale, and family homesteads settled in the 1890’s at Larsmont, and the French River on Minnesota’s North Shore were part of Lou’s DNA. If you knew Lou, you knew about the family legacy on Lake Superior. The landscape of northern Minnesota inspired Lou to pursue a BS in Geology from the University of Minnesota Duluth and an MS in Geology from the University of Minnesota and the Colorado School of Mines. This education would lead to a 30-year career highlighted by travel around the world while working for M.A. Hanna’s Minerals Research Laboratory in Nashwauk. Lou’s sharp mind did not rest in retirement. He extensively researched family genealogy culminating in connections with relatives in Larsmo, Finland, enhanced his boat collection at the home he and Peggy built on Swan Lake, supported the Isle Royale Friends and Family Association (IRFFA). xix �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 end of the annual meeting. 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. xx �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 2023, the ILSG Board of Directors selected two students to be granted research funding of $500 each from the Joe Mancuso Student Research Fund. The awardees were: Adrian Perez Avila Lakehead University Braxton Murphy Michigan Technological University Department of Geological and Mining Engineering and Sciences TOPIC: Characterization of the host rocks to mineralization in the Shebandowan greenstone belt in the vicinity of the Moss Lake deposit, NW Ontario TOPIC: Determine the relative paleostress state and tectonic conditions that resulted in formation and movement of faults making up the Keweenaw fault system near Houghton, Michigan, USA. Zsuzsanna P. Allerton University of Minnesota- Twinn Cities TOPIC: Investigate the timing and genesis of massive and semi-massive hematite ore bodies located in the Neoarchean (~2.7 Ga) Lake Vermilion/Soudan Underground Mine State Park (SSP) Farhan Ahmed Bhuiyan University of Minnesota- Duluth TOPIC: Evaluating post-depositional mineral reactions in the 1.71 – 1.47 Ga Freedom Formation, Baraboo, WI xxi �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. 2024 Student Paper Awards Committee Stacy Saari – Minnesota Department of Natural Resources (Committee Chair) Paula Leier-Englehardt – HydroGeo Solutions LLC, Wisconsin Dan Hirvi – Consulting Geologist, Michigan Allison Severson – Minnesota Geological Survey xxii �Board of Directors Theodore J. Bornhorst, Chair (2024-2027) — Michigan Technological University Carysn Ames (2023-2026) — Wisconsin Geological and Natural History Survey Mike Easton (2022-2025) — Ontario Geological Survey Mark Smyk (2019-2024) — Lakehead University Peter Hollings Secretary (2019-2024) — Lakehead University Mark A. Jirsa Treasurer (2022-2025) — Minnesota Geological Survey Board member through the close of the meeting year shown in parentheses. xxiii �2024 ILSG Meeting Michigan Tech Volunteers Great Lakes Research Center Daniel J. Lizzadro-McPherson Student Volunteers: Affiliated with Michigan Tech Jhuleyssy Liesseth Sánchez Aguilar Gabriel Ahrendt Katherine Langfield Marie, Lansbery Braxton Murphy Abe Stone 2024 ILSG Meeting Session Chairs Allan Blaske, GEI Consultants Amy Radakovich Block, Minnesota Geological Survey Patty Cobin, A. E. Seaman Mineral Museum, Michigan Tech Mary Louise Hill, Lakehead University Allan MacTavish, AGC GeoConsulting Ashley Quigley, Michigan Geological Survey Bernie Saini-Eidukat, North Dakota State University Mark Smyk, Lakehead University xxiv �Field Trip Leaders and Guidebook Authors Field trips have been the mainstay of the ILSG since its inception 70 years ago. We give special thanks to the field trip leaders and guidebook authors who volunteered their time and talent in carrying that tradition forward. Trip 1: Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan Ted Bornhorst (Michigan Tech) Trip 2: Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan Tom Wright (Quincy Mine Hoist Association) Jim DeGraff and Ted Bornhorst (Michigan Tech) Trip 3: Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty Erika Vye, Charlie Kerfoot (Michigan Tech) Stephanie Swart (Michigan Department of Environmental Quality) Dione Price and Evelyn Ravindran (Keweenaw Bay Indian Community) Trip 4: Keweenaw Fault Geometry and Kinematics: Clues to Its Nature and Origin Jim DeGraff, Katherine Langfield, and Dan Lizzadro-McPherson (Michigan Tech) Trip 5: Adventure Mine, Ontonagon County, Michigan: Geology and History of a Native Copper Mine Matt Portfleet (Adventure Mining Company) Ted Bornhorst (Michigan Tech) Trip 6: Southern Complex Granitoids, Gneisses and Migmatites: New Data, Discoveries, and Perspectives Chad Deering (Michigan Tech) Trip 7: Landslides in the Glacial Lake Ontonagon Sediments Stan Vitton and Mohammad Sadeghi (Michigan Technological University) xxv �Mineral Informatics: A New Frontier in Understanding Earth Robert M. Hazen Banquet Speaker Senior Staff Scientist, Earth and Planets Laboratory Carnegie Institution for Science, Washington, DC 20015 Email: rhazen@carnegiescience.edu The story of Earth is a 4.5-billion-year saga of dramatic transformations, driven by physical, chemical, and biological processes. The co-evolution of life and rocks unfolded in an irreversible sequence of evolutionary stages. Each stage re-sculpted our planet’s surface, while introducing new planetary processes and phenomena. This grand and intertwined tale of Earth’s living and non-living spheres is coming into ever-sharper focus, thanks to advances in “mineral informatics” - a field that employs large and growing mineral data resources to tell the deep-time stories of our evolving planet. Minerals are remarkably information rich, holding dozens of trace and minor elements, scores of stable isotopes, solid and fluid inclusions, chemical zoning, twinning, exsolution, countless defects, and a host of optical, magnetic, electrical, and other properties. Every mineral specimen is a time capsule waiting to be opened—waiting to tell its story. This lecture will explore some of the advanced data analytical and visualization methods that are shining new light on the old field of mineralogy, while revealing in ever greater clarity the co-evolution of the geosphere and biosphere. A network diagram of all known minerals colored by the way the minerals form. For example, red indicates high-temperature igneous minerals, while green indicates minerals formed by life. xxvi �REPORT OF THE 69th ANNUAL MEETING OF THE INSTITUTE ON LAKE SUPERIOR GEOLOGY Robert Lodge (University of Wisconsin-Eau Claire), Esther Stewart, and Carsyn Ames (Wisconsin Geologic and Natural History Survey) hosted the 69th Annual Institute on Lake Superior Geology on April 23 – 26, 2023 at the Lismore Hotel and Conference Center in Eau Claire, Wisconsin. The meeting consisted of two days of technical sessions with pre- and posttechnical 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: Talon Metals, American Institute of Professional Geologists, Geological Society of Minnesota, Crystal Cave, and Visit Eau Claire. We also thank the Individual Contributors to the Student Travel Scholarship fund: Val Chandler, Jim DeGraff, Thomas Erickson, Tom Fitz, Dave Good, Bob Mahin, Gordon Medaris Jr., Jim Miller, Steven Pinta, Tod Roush, and Gerry White. The 2023 meeting was the first conference held in the US since the 2018 Iron Mountain meeting, the first in Wisconsin since the 2011 Ashland meeting, and the first meeting in Eau Claire since 1991. Total meeting registration was 126, including 19 students. Attendance from both Canadian and United States was excellent despite other conferences in the Lake Superior region in April and May (GAC-MAC, Sudbury; Northcentral GSA, Grand Rapids). The technical program was nevertheless excellent with a good array of topics from Archean and Paleoproterozoic geology, to Midcontinent Rift geology and mineralization, to Quaternary Geology, Geoscience Education and Geoheritage. In addition, a student-industry networking lunch was held at the Riverview Room in the Eau Claire Public Library and an evening social was held at Reboot Social. Proceedings Volume 69 was published in two parts. Part 1 – Program and Abstracts, compiled and edited by Carsyn Ames (WGNHS) contains 54 published abstracts for 34 oral and 19 poster presentations. Students presented 8 oral and 10 poster presentations. Part 2 – Field Trip Guidebooks, was compiled and edited by Robert Lodge (UWEC). It contains descriptions of two pre-meeting and two post-meeting field trips. Hard copies of the Abstract Volume and Field Trip Guidebooks for trip participants were printed by University Printing at the University of Wisconsin-Eau Claire. Both volumes are available for download from the Institute on Lake Superior Geology website. The 69th ILSG marked only the third time in the Institute’s long history that its annual meeting was held in Eau Claire, the last time being in 1991. Plans for another Wisconsin-based ILSG meeting had been discussed for a while. With recent work in the Paleoproterozoic geology and mineralization in the Penokean orogen in northwestern Wisconsin and the central location of Eau Claire, it seemed appropriate to host the meeting there. Eau Claire sits on the boundary between Precambrian Shield, Paleozoic Platform, and the terminus of the continental ice sheet and allowed organizers to host four field trips examining billions of years of geologic history. Two field trips focused on the Precambrian geology of the Penokean orogen exposed in the Chippewa and Eau Claire River Valleys. While the preconference field trip got to see historic flooding on the Chippewa River (there are not many days when a bunch of Precambrian geologists are xxvii �looking at the river rather than the rocks), waters receded for the post-conference field trip. Field stops on this trip were originally (in some cases, exclusively) described in previous ILSG meetings (Eau Claire, 1980; 1991) but benefitted from new research and analyses and new viewpoints on the tectonics and metallogeny of the region. One fieldtrip visited classic exposures Paleozoic stratigraphy around the Eau Claire and Menominee regions and enjoyed lunch in an ancient meteorite impact structure. One field trip visited Quaternary geology and fluvial geomorphology of the Chippewa River valley. All the field trips, and the meeting itself, were blessed with good weather. Total field trip participation was 116 (excluding leaders and volunteer drivers). A list of field trips is provided below: Pre-meeting field trips (and leaders) on Tuesday, April 23. 1) Precambrian Geology of the Chippewa River valley: A Transect through the Marshfield Terrane (Robert Lodge, Bob Hooper, UW-Eau Claire) 2) Wisconsin’s Paleozoic stratigraphy and tour of Crystal Cave (Carsyn Ames, Esther Stewart, Bill Batten, Eric Stewart, Ian Orland, WGNHS) Post-meeting field trips (and leaders) on Friday, April 26. 3) Precambrian Geology of the Eau Claire River valley: Re-discovering the Eau Claire Volcanic Complex (Robert Lodge, Evan Weber, Bob Hooper, UW-Eau Claire) 4) Quaternary Geology and Geomorphology of the Eau Claire Region (Doug Faulker, UW-Eau Claire; Elmo Rawling, WGNHS; Phil Larson, Minnesota State University, Mankato) Many registrants attended the welcoming reception on Tuesday evening. Furthermore, the vast majority of registrants and invited guests attended the annual ILSG banquet on Wednesday night. Although a Homer Award overview presentation was given, no “recipients” were identified during the 2023 annual meeting, or in the previous 4 years! As always, a highlight of the post-banquet activities was presentation of the 2023 Goldich Medal. This year’s very deserving recipient was Dr. Pete Hollings. The Goldich Medal citation was presented by Mark Smyk, his colleague for many years. Mark described Pete’s many contributions to the ILSG, to the greater understanding of Archean and Proterozoic geology of the Lake Superior region, and his commitment to students. Pete is indeed a worthy recipient of this prestigious award. The 69th ILSG continued the post-banquet guest speaker tradition. Curt Meine, a conservation biologist, historian, and writer from the Aldo Leopold Foundation and Center for Humans and Nature, gave a presentation entitled Imagining “Conservation Geology”: Lessons from the Driftless Area. His talk provided an insightful viewpoint of how geology and landscapes integrate with history and culture in the Driftless Area of central Wisconsin. In 2023, the student paper committee remarked on the high quality of student research across all participants and had a difficult time of selecting the best among the excellent oral and poster presentations. The committee awarded four prizes for the best oral and poster presentations by xxviii �both undergraduate and graduate students. The best graduate oral presentation was awarded to Justin Jonsson and his talk on “Petrogenesis of the mineralized horizons in the Offset and Creek zones, Lac des Iles Complex, N. Ontario”. The best undergraduate oral presentation was awarded to Blaize Briggs for his talk on “Quetico-Wabigoon Subprovince Boundary in the Superior Province north of Thunder Bay, Ontario, Canada”. The best graduate poster presentation was awarded to Fransisca Nunez Ferreira for her poster on “Morphometry and formation process of eskers developed under the Chippewa Lobe of the Laurentide Ice Sheet”. The best undergraduate poster presentation was awarded to Lillian Glodowski for her poster on “Characterizing volcanic host stratigraphy and syn-volcanic intrusions at the Lynne Zn-Pb-Cu deposit, Oneida Co., Wisconsin”. Eisenbrey Student Travel Grants were given to twelve students: Zsuzsanna Allerton, Ryan Barkley, Blaize Briggs, Tianna Groeneveld, Justin Jonsson, Daniel Lizzardo-McPherson, Francisca Nunez Ferreira, Jordan Peterzon, Sam Ghantous, Madeline Taylor, BJ Itai, and Katherine Langfield. The Institute’s Board of Directors met on Thursday, April 25, 2023, and a brief overview of the meeting notes is provided below: 1. Accepted report of the Chairs for the 68th ILSG, as published in the Proceedings volume, and minutes of last Board meeting, May, 2022 (Hollings) 2. Received, discussed, and accepted 2022-2023 ILSG Financial Summary (Jirsa). 3. Received, discussed, and accepted 2022-2023 report of the Secretary (Hollings). 4. Approved Carsyn Ames as on-going ILSG Board member 5. Discussed and approved Amy Radakovich Block as Assistant Treasurer in a non-voting role (end of term 2026). 6. Discussed and approved replacing Steve Kissin as the “member from academia” on Goldich Committee (end of term 2023) with Marcia Bjornerud 7. Approved Houghton as the site for the 70th annual ILSG meeting. The meeting will be hosted by Ted Bornhorst. 8. Discussed the role of the Michigan Tech archives as the host of hard copies of the publications of the Institute. A formal agreement has been signed with the Archives who will request financial support as needed rather than the previous model of providing a donation of $1 per attendee. 9. A number of future meeting locations were discussed. Peter Hinz has offered Kenora as a future site and Mark Jirsa is keen to host the Mountain Iron meeting that was cancelled due to the pandemic. Bernie Saini-Eidukat to be approached to see if he is still interested in organizing a meeting in St Cloud. 10. The cost of insurance was again discussed and it was agreed that the Board of Directors insurance should be maintained and that the costs would be included in the cost of each meeting. Given the high costs quoted for field trip insurance the Board to investigate field trip insurance options. Hollings to approach GAC. Amy to approach GSA and Carsyn to approach a risk advisor. The Board discussed embedding a Liability waiver in the registration process. 11. The Board discussed embedding a photo release in the meeting and/or field trip registration process such that anyone registering for the field trip is aware they are agreeing to be photographed and have their image used in ILSG publications/website etc. It was agreed that the Institute does not want to turn anyone away from the meeting/trips simply because xxix �they do not want to be photographed, and the result is that before any photos are uploaded to the website, someone on the Board (or a future social media position) will need to go through the photos and make sure that no one who has NOT signed a photo release is shown in a shot where they are identifiable. The 69th ILSG meeting was a great success, and we wish to thank all the people who contributed to that success, field trip leaders and drivers, UWEC student volunteers, and businesses and organizations in downtown Eau Claire that hosted and entertained visitors. Patty Cobin and Ted Bornhorst (A.E. Seaman Mineral Museum, Michigan Technological University) handled the premeeting registration and supplied the poster boards. Thanks also go to the staff at Lismore Hotel and Conference Center who helped the meeting and banquet run smoothly and providing lunches and snacks during the technical sessions. Thanks to Eau Claire Public Library for hosting our student-industry luncheon, Reboot Social for hosting our post-meeting evening social, Eau Claire Cheese and Deli for field trip lunches, and Eau Claire Student Transit for bus transportation for fieldtrips. Robert Lodge (UWEC), Carsyn Ames (WGNHS), and Esther Stewart (WGHNS) Co-Chairs, 69th Institute on Lake Superior Geology xxx �Donations to Support Student Participation at the Annual Meeting of the Institute on Lake Superior Geology A SPECIAL THANK YOU TO OUR INDIVIDUAL CONTRIBUTORS Roger Anderson Aaron Hirsch Wouter Bleeker Allan MacTavish Terry Boerboom Bob Mahin Ted Bornhorst Gordon Medaris Jr. Alex Brown Jim Miller Michael Carr Rick Sandri Val Chandler Isabel Serrano Kate Clover Mark Severson Abraham Drost Jim Small Thomas Erickson Gerry White Annia Fayon Graham Wilson Mary Louise Hill xxxi �TECHNICAL PROGRAM xxxii �Wednesday May 15, 2024 All field trips begin and end at the Michigan Tech Memorial Union Building Parking tickets are given between 7 am to 4 pm weekdays. Between ticketing hours all vehicles need a parking pass; these will be available from trip leaders. Pre-meeting Field Trips May 15, 2024 8:00 am - 5:00 pm PRE-MEETING FIELD TRIPS Trip 1: Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan Ted Bornhorst (Michigan Tech University) Trip 2: Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan Tom Wright (Quincy Mine Hoist Association), Jim DeGraff, Katherine Langfield, and Ted Bornhorst (Michigan Tech) Trip 3: Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty Erika Vye and Charlie Kerfoot (Michigan Tech) Stephanie Swart (Michigan Department of Environmental Quality) Dione Price and Evelyn Ravindran (Keweenaw Bay Indian Community) Wednesday evening May 15, 2024 4:00 pm - 8:00 pm Registration (2nd floor, Michigan Tech Memorial Union) 6:30 pm - 8:30 pm Poster Setup and Viewing (2nd floor, Michigan Tech Memorial Union) 6:30 pm - 8:30 pm Welcoming Reception (2nd floor, Michigan Tech Memorial Union) xxxiii �* 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 the paper if different than the first author. Thursday - May 16, 2024 All vehicles need a parking pass between 7am to 4pm weekdays; these will be available from registration 7:15 am - noon Registration (2nd floor, Michigan Tech Memorial Union) 8:15 am OPENING REMARKS (2nd floor, Michigan Tech Memorial Union) Ted Bornhorst, Erika Vye, Patty Cobin, and Jim DeGraff Co-Chairs, 2024 ILSG TECHNICAL SESSION I – ORAL PRESENTATIONS Session Chair: Mark Smyk 8:20 Jim MILLER Roland Duer Irving - Pioneer of Lake Superior geology 8:35 Graham WILSON, Charles BUTT, Robert GARRETT and Heather ROBINSON R.W. Boyle’s History of Geochemistry and Cosmochemistry 8:55 James DeGRAFF, Nolan GAMET, Katherine LANGFIELD, Daniel LIZZADROMcPHERSON, Sophie MUELLER, and Colin TYRRELL Transpressional Nature of the Keweenaw Fault System, Lake Superior Region, and Its Relationship to Grenville Orogenesis 9:15 Alex BROWN Re-interpretation of hydrothermal alteration, mineralization and host-rock oxidation to form the Keweenaw native copper lodes, northern Michigan 9:35 Esther STEWART, Michael TAPPA, Ann BAUER, Anthony PRAVE, and Latisha BRENGMAN Geochemical fingerprints from the late Mesoproterozoic epeiric seaway of the Nonesuch Formation, Wisconsin and Michigan, USA 9:55 END OF TECHNICAL SESSION I 9:55-10:10 COFFEE BREAK xxxiv �TECHNICAL SESSION II – ORAL PRESENTATIONS Session Chair: Ashley Quigley 10:10 Sarah GORDEE, Madison RIAN, Stacy SAARI, and Matthew CARTER Description and application of the Consolidated Minerals Database to support geological investigations: an example from the Cuyuna Range, central Minnesota 10:30 Stacy SAARI, Sarah GORDEE, Madison RIAN, and Matt CARTER Compiled historical drillhole and geochemical data from the Cuyuna Range, Minnesota, provides powerful new insights for geological and mineral potential investigations. 10:50 Bob MAHIN, Ashley QUIGLEY, John YELLICH, John ESCH, and Nolan GAMET Critical Mineral Systems in the Upper Peninsula of Michigan, A Cooperative Effort Between the USGS and the Michigan Geological Survey 11:10 Paul BEDROSIAN, Dana PETERSON, and Bennett HOOGENBOOM Geophysical imaging of the Paleoproterozoic Animikie basin in Minnesota 11:30 Dan HOLLIS Use of Ambient Noise Tomography for Mineral Exploration in the Lake Superior Region 11:50 END OF TECHNICAL SESSION II 11:50-1:10 LUNCH BREAK and ILSG BOARD OF DIRECTORS MEETING - lunches not provided to conference attendees- TECHNICAL SESSION III- ORAL PRESENTATIONS Session Chair: Mary Louise Hill 1:10 *Demily THIBODEAU-BELLO, Mary Louise HILL, Andrew CONLY, An evaluation of structural and mineralogical controls on gold mineralization on the GoldRich property in the Abbie Lake area, Wawa, Ontario 1:30 Dean PETERSON and Alex STEINER The geology and ore deposit model of the high-grade Emily Manganese Deposit, Cuyuna Range, Minnesota: Results from the 2023 drilling program 1:50 *Gabriel AHRENDT and Aleksey SMIRNOV Rock magnetic investigation of the Vulcan Iron Formation: Unveiling Paleoproterozoic Paleoenvironments 2:10 *Zsuzsanna ALLERTON, George HUDAK, Christian TEYSSIER, Annia FAYON, Martin DANIŠIK, Liam COURTNEY-DAVIES, and Phillip LARSON Geochronology and geochemistry of hematite ore in northeastern Minnesota xxxv �2:30 Joyashish THAKURTA and Beau HAAG Sulfur-isotope ratios in Paleoproterozoic Michigamme Formation at the Lake Superior Region: Implications on basin evolution and ambient seawater composition in the Greater Animikie Basin 2:50 *Jordan PETERZON, Noah PHILLIPS, Pete HOLLINGS, and Lionnel DJON Deformation conditions, micromechanics, and fault zone development in mafic protoliths at the Lac des Iles mine, northwestern Ontario 3:10 END OF TECHNICAL SESSION III 3:10-3:30 COFFEE BREAK TECHNICAL SESSION IV – POSTER PRESENTATIONS Session Chair: Allan Blaske and Patty Cobin 3:30-5:00 AUTHORS PRESENT AT THEIR POSTERS 5:00 END OF TECHNICAL SESSION IV Thursday evening May 15, 2024 6:00 pm RECEPTION AND CASH BAR (2nd floor, Michigan Tech Memorial Union) 7:00 pm ANNUAL BANQUET (2nd floor, Michigan Tech Memorial Union) 2024 Goldich Medal Recipient: Suzanne W. Nicholson Banquet Speaker: Robert M. Hazen, Carnegie Institution for Science “Mineral Informatics: A New Frontier in Understanding Earth” xxxvi �Friday - May 17, 2024 All vehicles need a parking pass available from co-chairs 8:15 INTRODUCTORY REMARKS AND UPDATES (2nd floor, Michigan Tech Memorial Union) Ted Bornhorst, Erika Vye, Patty Cobin, and Jim DeGraff Co-Chairs, 2024 ILSG TECHNICAL SESSION V – ORAL PRESENTATIONS Session Chair: Bernie Saini-Eidukat 8:20 *Farhan Ahmed BHUIYAN, Latisha BRENGMAN, and Esther STEWART Assessing depositional and post-depositional mineral associations in the <1.71 Ga Freedom Formation, Baraboo, WI, USA. 8:40 Jack MALONE, Ryan CLARK, Amira HARRIS-BOMMARITO, and David MALONE Baraboo Interval Quartzites in Iowa: Reassessing the Origin and Provenance of the Washington County Quartzite, SE Iowa 9:00 Gordon MEDARIS, Chloe BONAMICI, Phil BROWN, Laurel GOODWIN, Brian JICHA, Brad SINGER, Michael SPICUZZA and John VALLEY, The Evolution of Baraboo Interval Sedimentary Rocks: Deposition at 1.63 Ga and Metamorphism at 1.47 Ga 9:20 Amy Radakovich BLOCK, George HUDAK, and Kate SOUDERS Insights into the southwestern Superior Province: New igneous geochronology and geochemistry in northwestern Minnesota, USA 9:40 Ryan CLARK, David PEATE, Allison KUSICK, Kenny HORKLEY, and Chris MACFARLANE Baddeleyite age reveals timing of the Northeast Iowa Intrusive Complex (NEIIC) 10:00 END OF TECHNICAL SESSION V 10:00-10:20 COFFEE BREAK TECHNICAL SESSION VI – POSTER PRESENTATIONS Session Chair: Allan Blaske and Patty Cobin 10:20-11:40 AUTHORS PRESENT AT THEIR POSTERS 11:40 END OF TECHNICAL SESSION VI 11:40-1:00 LUNCH BREAK xxxvii �TECHNICAL SESSION VII – ORAL PRESENTATIONS Session Chair: Allan MacTavish 1:00 Jim MILLER and John GREEN Two decades of teaching the geologic heritage of Minnesota’s North Shore at the North House Folk School, Grand Marais 1:20 Erika VYE, Daniel LIZZADRO-MCPHERSON, and James JUIP The Keweenaw Geoheritage Summer Internship Experience 1:40 Eric NOWARIAK, S, Allison SEVERSON, and Amy Radakovich BLOCK Lithostratigraphy and Geochronology of the Lower Northeast Sequence of the North Shore Volcanic Group, Cook County, MN, USA 2:00 Pete HOLLINGS and Mark SMYK New Insights into the Geology and Geochemistry of the Osler Group and Related Rocks, Midcontinent Rift System, Northern Lake Superior, Ontario 2:20 David GOOD MCR Synthesis 1. Characterizing the MCR mantle plume 2:40 END OF TECHNICAL SESSION VII 2:40-3:00 COFFEE BREAK and TAKE DOWN POSTERS TECHNICAL SESSION VIII – ORAL PRESENTATIONS Session Chair: Amy Radakovich Block 3:00 Bill ROSE and James DeGRAFF Lidar Topography: Bright opportunity for reading Keweenaw Landscapes 3:20 Wouter BLEEKER, Natasha WODICKA, Sandra KAMO, Michael HAMILTON, Quinn EMON, and Jennifer SMITH The Lake Superior area “event layer”: Testing the connection with the Sudbury impact 3:40 Tien GRAUCH, S. HELLER, Laurel WOODRUFF, and Esther STEWART Revisiting geophysical interpretations of the Midcontinent Rift below Lake Superior— Insights from GLIMPCE seismic-reflection line C 4:00 Aaron HIRSCH Recent developments on the use of the Horizontal-to-Vertical Spectral Ratio (HVSR) passive seismic method to determine depth to bedrock in Minnesota 4:20 END OF TECHNICAL SESSION VIII xxxviii �4:20 Presentation of Student Awards Best Student Paper Awards – Stacy Saari Student Travel/Participation Awards – Ted Bornhorst 4:40 Concluding Remarks and Field Trips Ted Bornhorst, Erika Vye, Patty Cobin, and Jim DeGraff Co-Chairs, 2024 ILSG END OF TECHNICAL SESSIONS OF THE 70th ANNUAL MEETING Friday Evening May 17, 2024 7 pm ATDC Building across the parking lot from the A.E. Seaman Mineral Museum 2024 Edith D. and E. Wm. Heinrich Lecture “Mineral Evolution: A Case Study of a New Natural Law" by Robert M. Hazen Sponsored by the Edith D. and E. Wm. Heinrich Mineralogical Research Foundation and the A. E. Seaman Mineral Museum Saturday May 18, 2024 Field trips begin and end at the Michigan Tech Memorial Union Parking pass not needed on weekend. 8:00 am – 5:00 pm POST-MEETING FIELD TRIPS Trip 4: Keweenaw Fault Geometry and Kinematics: Clues to Its Nature and Origin Jim DeGraff, Katherine Langfield, and Dan Lizzadro-McPherson (Michigan Tech) Trip 5: Adventure Mine, Ontonagon County, Michigan: Geology and History of a Native Copper Mine Matt Portfleet (Adventure Mining Company), Ted Bornhorst (Michigan Tech) Trip 6: Southern Complex Granitoids, Gneisses and Migmatites: New Data, Discoveries, and Perspectives Chad Deering (Michigan Tech) Trip 7: Landslides in the Glacial Lake Ontonagon Sediments Stan Vitton (Michigan Tech) xxxix �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 + Denotes author that will present the paper if different than the first author. Numbered Posters and Abstracts are in sequential order Poster Number 1. withdrawn 2. Sheree HINZ GeologyOntario: a powerful search tool for Ontario explorationists Therese PETTIGREW, Robert CUNDARI, Rebecca PRICE, and Manuel DUGUET Identification of Fertile Parent Granitoid Units in the Superior Province of Ontario withdrawn 3. Mia MORSON and + Shannon ZUREVINSKI Quartz trace element chemistry: Exploring the link between a fertile parental granite and a mineralized pegmatite 4. *Kevin MEXIA, Pete HOLLINGS Geochemistry of Midcontinent Rift-related intrusive rocks of the Sunday Lake intrusion 5. Justin JONSSON, Paul MALEGUS, Sophie CHURCHLEY, and Rebecca PRICE Characterizing the geochemistry and nickel-copper-platinum group elements potential of mafic and ultramafic intrusions in northwestern Ontario withdrawn 6. *Andrea Paola CORREDOR BRAVO, Pete HOLLINGS, Matthew BRZOZOWSKI, and Geoff HEGGIE Magmatic and hydrothermal evolution of the Mesoproterozoic Current ultramafic PGECu-Ni deposit within the Thunder Bay North Intrusive Complex: insights from trace elements, Nd, Sr, O, and H isotopes 7. Max LAXER and + David GOOD Building a 3D model for Cu/Pd inflection points throughout the Marathon PGE-Cu Deposit 8. *Vlad SHESHNEV, Pete HOLLINGS, Noah PHILLIPS, Ryan WESTON, Matt DELLER, and Dana CAMPBELL Geochemistry and Petrology of the Eagle’s Nest Intrusion, McFaulds Lake Greenstone Belt, Northern Ontario xl �9. *Yiruo XU and Robert HOLDER Cooling of an Archean metamorphic terrane: garnet diffusion study of the Quetico Subprovince, Canada 10. *Clare BEAUDRY, Madelyn HESS, Cristian PEREIRA, and Bernhardt SAINI-EIDUKAT Petrology and geochemistry of Precambrian basement rocks in Walsh County, North Dakota 11. *Cristian PEREIRA, Timothy NESHEIM, Jeffrey D. VERVOORT, and Bernhardt SAINI-EIDUKAT Major element geochemistry and first zircon U-Pb age dates of Precambrian basement rocks in eastern North Dakota 12. Jamey JONES, Bill CANNON, Ben DRENTH, and Paul O’SULLIVAN Geologic and tectonic implications of detrital zircon U-Pb ages from the Dickinson Group in the western Upper Peninsula of Michigan, USA 13. Bill CANNON, A. SOUDERS, Ben DRENTH, and Robert AYUSO The Sacred Heart Orogeny in Michigan: Latest Archean Granites and the Great Lakes Tectonic Zone 14. A. SOUDERS, W. CANNON, B. DRENTH, R. SALERNO, J. THOMPSON, and P. SYLVESTER New LA-ICP-MS U-Pb geochronology of Archean rocks, central Upper Peninsula, Michigan, USA: a step toward refining the final assembly of the Superior craton 15. *Zsuzsanna ALLERTON, Annia FAYON, George HUDAK, Christian TEYSSIER, Liam COURTNEY-DAVIES, and Martin DANIŠIK Geochronology campaign in northeastern Minnesota 16. Ross SALERNO, Bill CANNON, Amanda SOUDERS, Jay THOMPSON Understanding the evolution of the upper Midwest Archean gneiss dome corridor using apatite, titanite, and monazite LA-ICP-MS U-Pb geochronology and microstructural analyses 17. *Trent EDIGER and Marcia BJØRNERUD Glimpses of a Paleoproterozoic landscape: Analysis of exhumed topography on Archean basement rocks northwest of Marquette, Michigan 18. Rebecca STOKES, Bill CANNON, and Ross SALERNO Characteristics of graphitization across a metamorphic gradient in the Michigamme Formation of the Marquette Trough and Baraga Basin, MI 19. Tom BUCHHOLZ, Alexander FALSTER, and Wm. SIMMONS Preliminary mineralogy of a pegmatite in the pyroxene syenites of the Stettin Complex, Wausau Complex, Marathon County, Wisconsin xli �20. *Katherine LANGFIELD, Nolan GAMET, James DeGRAFF Cross-sectional Geometry of the Keweenaw Fault System between Hancock and Mohawk, Upper Peninsula of Michigan 21. *Braxton MURPHY, Katherine LANGFIELD, and James DeGRAFF Geometry, Slip Kinematics, and Deformation along the Hancock Fault in the Quincy Mine Workings, Upper Peninsula of Michigan 22. *Kenz CARLTON, Basil TIKOFF, and Esther STEWART The Honey Creek Structure, Sauk County, Wisconsin: Asymmetric Faulting Associated with Seismic-Induced Fluid Escape 23. *Alex LAWRENCE, Adam VANDERKIN, and Robert LODGE Volcanic and Hydrothermal Reconstruction of the Paleoproterozoic Butler Zn-Cu occurrence, Clark County, Wisconsin 24. *Lyndsie VICKERS, and Robert LODGE Petrology and Geochemistry of Felsic Magmatism in the Paleoproterzoic Eau Claire Volcanic Complex, Northcentral Wisconsin 25. *Dan SHAKKED, Lucas ROBARGE, and Robert LODGE Analysis of deformation-related structures in the Eau Claire Volcanic Complex, Wisconsin 26. *Gwendolyn MARTIN and Marcia BJØRNERUD Investigating the origin of pervasive breccias in the Paleoproterozoic Saunders Formation in northern Wisconsin 27. Aaron HIRSCH, Emma SCHNEIDER Lithostratigraphic discrimination of Quaternary core in Minnesota using magnetic susceptibility 28. Bill ROSE and Erika VYE Michigan Coastal Path: A Social Commitment to Geoeducation 29. Bill ROSE and Erika VYE Jacobsville geoheritage is globally celebrated and locally loved 30. *Alice MARTIN, Zsuzsanna ALLERTON, Emma JOHNSON, Annia FAYON, Jim ESSIG, Sarah GUY-LEVAR,George HUDAK The Soudan Geology Trail Project: Let’s talk about rocks in northeastern Minnesota A tribute to Jean Peterman Kemp Zimmer and Jeanne Seaman Farnum by the A.E. Seaman Mineral Museum: Trailblazers for Women in Geology xlii �ABSTRACTS xliii ��Rock magnetic investigation of the Vulcan Iron Formation: Unveiling Paleoproterozoic Paleoenvironments AHRENDT, Gabriel1 and SMIRNOV, Aleksey1,2 1 Department of Geological Mining and Engineering Sciences, Michigan Technological University, 1400 Townsend Dr, Houghton, MI 49931 2 Department of Physics, Michigan Technological University, 1400 Townsend Dr, Houghton, MI 49931 The Paleoproterozoic (~1.88 Ga) Vulcan Iron Formation, located in the Southwestern Upper Peninsula of Michigan, is a significant Superior-type Banded Iron Formation, comprising four main members. The lower, Traders Member is characterized by banded ferruginous-siliciclastic layers with distinct alternating layers of ferric iron. The middle, Brier Member is a fissile slate with varying concentrations of magnetite from low to locally enriched. The upper, Curry Member, is an oolitic iron formation enriched with specular hematite and lacking noticeable banding. In some locations, the Curry Member is overlain by the ferric slate referred to as the Loretto Member. We conducted comprehensive rock magnetic investigations of three lower formation members, using thermal demagnetization of natural remanent magnetization, magnetic hysteresis and first-order reversal curve measurements, and thermomagnetic analyses. Our findings suggest that the members may be genetically distinct, reflecting shifts in depositional regimes that dramatically affected their texture and mineralogy. The Traders Member, characterized by abundant small paramagnetic grains, likely formed during a period of rapid subsidence and soluble transport of ferrous iron into a euxinic basin, followed by alternating periods of CO2 fixing and sulphide-oxidizing cyanobacteria. The subsequent transition to a shallow, foreshortened basin as the Pembine-Wasau terrane accreted led to the increased silica saturation and local concentration of superparamagnetic ferrous iron mud, forming the Brier slate. A further evolution due to the flooding of a shallow sea inducing high turbidity and increased oxygenation in the water column, resulted in the formation of the Curry Member, marked by a mix of magnetically hard minerals, including specular hematite. We speculate that, subsequently, a decrease in sea level, associated with the basin’s contraction, created conditions conducive to high silica input from continental margins and a change in the biotic regime which reduced the formation of granules and led to the creation of the Lorretto slate member. 1 �Geochronology campaign in northeastern Minnesota ALLERTON, Zsuzsanna1, FAYON, Annia1, HUDAK, George1, TEYSSIER, Christian1, COURTNEY-DAVIES, Liam2, and DANIŠIK, Martin3 1 Earth and Environmental Sciences, University of Minnesota, Minneapolis, MN 55455, USA Geological Sciences, University of Colorado, Thermochronology Research and Instrumentation Lab, Boulder, CO 80309-0399, USA 3 School of Earth and Planetary Sciences Department, John de Laeter Centre, Curtin University, Perth, WA 6845. Australia 2 Northeastern Minnesota is known from volcanic rocks along the north shore of Lake Superior and the intrusive suit of the Duluth Igneous Complex (DC) to its west with associated Ni-Cu-PGE mineralization (Leu, 2015; Miller, 2002). These lithologic units are part of the ~1100 Ma Midcontinent Rift System (MRS). During the rifting event, the hot (~1000°C) DC was emplaced into the “cold” Neoarchean (~2700 Ma) monzo-granitic Giants Range Batholith (GRB) footwall, resulting in contact metamorphism, sulfide mineralization (Benko et al., 2015), and hydrothermal activity. The focus of this project is to constrain the hydrothermal effect of DC emplacement and other regional events by tracking isotopic, chemical, and textural changes in accessory minerals zircon and apatite along a transect from the DC/GRB contact westward into the Archean basement. Samples in this study are collected from the GRB and the Purvis Lake tonalite. Six of the GRB samples were analyzed for U-Pb dating of in-situ apatite, and twelve GRB samples and one Purvis Lake tonalite sample were disaggregated to isolate individual zircon grains to acquire UPb radiometric dates (Figure 1). U-Pb data were obtained by Laser Ablation-Inductively Coupled Figure 1: Simplified geologic map of Minnesota's arrowhead region showing the study area (yellow inset). Samples are numbered within lithological units of the Giants Range Batholith (GRB) and Purvis Lake tonalite. Maps are modified from Griffin and Morey (1969) and Peterson and Jirsa (1999). 2 �Plasma Mass Spectrometry (LA-ICP-MS) at laboratories of University of Santa Barbara (in-situ apatite) and University of Colorado, Boulder (zircon separates). Results show bimodal dates signifying crystallization age and the timing of hydrothermal alteration. In-situ apatite U-Pb yields 2594.3 ± 30.8 Ma ages at ~ 2 km from the contact, and the rest of the apatite U-Pb dates suggest Pb-loss as a function of hydrothermal alteration from ~1067.24 ± 7.64 Ma to ~1084.89 ± 9.23 Ma within ~1 km of the contact. Zircons separated from twelve samples yield U-Pb crystallization ages ranging from ~2640 to 2700 Ma, and lower intercepts ranging from ~700 to 1200 Ma with uncertainties of ~4-200 Ma. The average lower intercept is ~1150 Ma, and we interpret these ages to record hydrothermal activity-induced Pbloss. Crystallization age and error increase, and the timing of hydrothermally driven Pb-loss becomes more elusive with distance from the contact. The tonalite ~20 km from the DC/GRB contact yields a crystallization age of 2708 ± 25 Ma and displays Pb-loss at 1140 ± 116 Ma. The large uncertainty associated with Pb-loss might be suggestive of multiple hydrothermal events. These data are consistent with hematite (U-Th)/He dates of the massive hematite ore bodies of Soudan Iron Mine, ~40 km (map distance) from the DC/GRB contact, that record hydrothermal alteration at ~ 1100 Ma (see Allerton at al., 2024 “Geochemistry and geochronology of hematite ore in northeastern Minnesota”, ILSG 2024). References Benko, Z., Mogessie, A., Molnar, F., Severson, M., Hauck, S., & Raic, S., 2015. Partial melting processes and Cu-Ni-PGE mineralization in the footwall of the South Kawishiwi Intrusion at the Spruce Road Deposit, Duluth Complex, Minnesota. Economic Geology and the Bulletin of the Society of Economic Geologists, 110(5), 1269-1293. Leu, A., 2016. Geology and Petrology of the Wilder Lake Intrusion, Duluth Complex, Northeastern Minnesota [thesis]. Miller, J., & Minnesota Geological Survey, 2002. Geology and mineral potential of the Duluth complex and related rocks of northeastern Minnesota. Report of investigations (Minnesota Geological Survey; 58). Saint Paul: University of Minnesota, Minnesota Geological Survey. 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. Griffin, W. L. and Morey, G. B., 1969. Geology of the Isaac Lake Quadrangle, St. Louis County, Minnesota. Published in Cooperation with the Minnesota Department of Iron Range Resources and Rehabilitation. Minnesota Geological Survey 5 P-8 Special Publication Series. University of Minnesota. 3 �Geochronology and geochemistry of hematite ore in northeastern Minnesota ALLERTON, Zsuzsanna1, HUDAK, George1, TEYSSIER, Christian1, FAYON, Annia1, DANIŠIK, Martin2, COURTNEY-DAVIES, Liam3, and LARSON, Phillip4 1 Earth & Environmental Sciences, University of Minnesota, Minneapolis, MN 55455, USA School of Earth and Planetary Sciences, John de Laeter Centre, Curtin University, Perth, WA 6845, Australia 3 Geological Sciences, University of Colorado Boulder, Thermochronology Research and Instrumentation Lab, Boulder, CO 80309-0399, USA 4 Earth and Environmental Sciences, University of Minnesota, Duluth, MN 55812, USA 2 The Neoarchean Lake Vermilion-Soudan Underground Mine State Park is known for its Algoma-type banded iron formation (BIF). The BIF encloses lens-shape high-grade (63-65% Fe) iron ore locally. The well-established view is that massive to semi-massive hematite ore bodies are the product of hydrothermal alteration of BIF, during which process hydrothermal fluids leached silica, resulting in volume reduction (production of vug spaces) and Fe-replacement (Gruner 1930; Klinger, 1960; Thompson, 2015). Studies have postulated that ore mineralization was syn- or post-depositional with BIF (Gruner, 1926; Thompson, 2015), but the absolute timing of hematite ore had not been established until now. Figure 1: The diagram illustrates hematite mineralization consistent with the timing of Yavapai and Mazatzal orogenies based on U-Pb ages, and Midcontinent Rift System signatures overprint mineralization ages with (U-Th)/He analysis. Presented is a novel technique based on coupled U-Pb and (U-Th)/He hematite radiometric dating (Courtney-Davies et al., 2022) to determine the formation age and thermal history recorded by hematite. Initial electron probe microanalyses (EPMA) allowed for hematite characterization (microcrystalline and microplaty) that helped locating inclusion-free mineral surfaces for radiometric age dating. U-Pb Laser AblationInductively Coupled-Plasma Mass Spectrometry results suggest Paleoproterozoic mineralization at 1740.4±72.5 Ma and 1640.8±47.2 Ma and (UTh)/He ages clustered at 1093.1±16.4 Ma, the latter indicating a hydrothermal overprint of the original mineralization event (Figure 1). We propose a regional scale model that describes the hydrothermal alteration of Archean BIF at ~1700-1600 Ma with the formation of hematite ore including the growth of microcrystalline then microplaty textures, followed 4 �by a thermal overprint at ~1100 Ma associated with the development of the Midcontinent Rift System (Figure 2). Figure 2: Schematic diagrams display S-N cross-section starting with A) pre-D2, showing Gafvert Lake sequence (GL) unconformably above the Soudan Member (SM) that is stratigraphically above the Lower Member (LM) of Ely Formation. B) D2 regional transpression resulting in right lateral shear zones within SM, constrained to 2685-2674 Ma from dating of regional metamorphic fabrics (Lodge et al., 2013). C) Orogenic magmatism—depicted by a mafic dike (MD) at 1700-1600 Ma—generates hydrothermal fluids, and shear zones are utilized for fluid flow and facilitate a 2-stage hematite ore mineralization (microcrystalline and microplaty). D) The Midcontinent Rift System at ~1100 Ma results in hydrothermal overprint of original mineralization recorded in hematite (U-Th)/He ages. Whole-rock — including major, trace and rare earth elements—lithogeochemical analysis has been performed on four iron formation and ore samples, and results are currently being processed. Klinger (1969) proposed a volume-to-volume replacement mineralization, while Thomson (2015) calculated 39% volume loss by silica leaching and 9% Fe mass gain with Fe replacement. Additionally, isocon analysis (Grant, 2005) is underway to better understand the relationship between mobile and immobile elements during ore mineralization and the paragenesis of hematite. References Courtney-Davies, L., et al., 2022. Hematite geochronology reveals a tectonic trigger for iron ore mineralization during Nuna breakup: Geology, v. 50, p. 1318-1323, doi: 10.1130/G50374.1. Grant, J.A., 2005, Isocon analysis: A brief review of the method and applications: Physics and Chemistry of the Earth, Parts A/B/C, v. 30, p. 997–1004, doi: 10.1016/j.pce.2004.11.003. Gruner, J. W., 1926. Hydrothermal alteration of iron ores of the Lake Superior type—a modified theory: Economic Geology, v. 32, p.121-130. Gruner, J.W., 1930. Hydrothermal oxidation and leaching experiments; their bearing on the origin of Lake Superior hematite-iron ores: Economic Geology, v. 25, p. 697-719. Klinger, F.L., 1960. Geology and ore deposits of the Soudan mine, St. Louis County, Minnesota [thesis]. Lodge, R.W.D., et al., 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. Schulz, K. J., 1982. The magmatic evolution of the Vermilion greenstone belt of Minnesota: Tectonophysics, v. 190, p. 233-268. Thompson, A., 2015. A hydrothermal model for metasomatism of Neoarchean Algoma-Type banded iron formation to massive hematite ore at the Soudan Mine, NE Minnesota [thesis]. 5 �Petrology and geochemistry of Precambrian basement rocks in Walsh County, North Dakota BEAUDRY, Clare1, HESS, Madelyn1, PEREIRA, Cristian1, and SAINI-EIDUKAT, Bernhardt1,2 1 Department of Earth, Environmental and Geospatial Sciences, 2Department of Chemistry and Biochemistry, North Dakota State University, Fargo, ND 58102, USA In 1977, thirty-two cores were drilled in eastern North Dakota and western Minnesota along the Red River, for the purpose of evaluating uranium potential (Figure 1). The project was funded by the Department of Energy and overseen by Bendix Corporation. A technical report (Moore, 1978), a M.S. thesis that focused on the weathered horizon at the top of the Precambrian bedrock (Kelley, 1980), and several ILSG abstracts were published. For this study, three cores from Walsh County, North Dakota were sampled at the North Dakota Geological Survey Drill Core Library (Grand Forks, ND). Samples were taken from RRVD #17, RRVD #18, and RRVD #19A to focus the study to Walsh County, ND. Figure 2 shows the lithology of the three cores and outlines sample locations. Figure 2: Stratigraphic column of RRVD drill core Precambrian layers. Black Xs indicate sample locations. Numbers to the left correspond to the XRF analysis in Table 1. Data taken from Moore (1979) and optical observations. Figure 1: Location map of Eastern North Dakota and Western Minnesota. Era of Precambrian Bedrock is outlined. Red River Valley Drill Cores are outlined. Petrography and whole rock geochemical analyses (Table 1) were carried out on Precambrian layers. Precambrian sediments are buried under younger layers in Eastern North Dakota, the sampled areas are underlain by Archean gneiss, (Klasner and King, 1986). RRVD #17 was characterized as quartz monzonite with heavier alterations of biotite and feldspar farther up in the core. The alterations may be due to stronger weathering agents on the paleoweathered horizon. RRVD #18 was characterized as a granodiorite with uniform foliation and mineral percentages throughout the drill core. RRVD #19A was characterized as a gneissic granite with 6 �higher foliation as the sample increases in depth. The top of the cores is bleached, likely an effect of paleoweathering processes. Analyses were plotted on AFM and TAS diagrams (Figure 3). Table 1: RRVD #17-785, 2: RRVD #18-645.5, 3: RRVD #18-655.5, 4: RRVD #19A-1284.5, 5: #19A1291, 6: #19A-1296.5. Chemical data from NDSU XRF analysis. wt% 1 2 SiO2 59.1 69.2 TiO2 0.58 Al2O3 Fe2O3 MnO 3 4 5 6 71 68.2 73.5 73.6 0.39 0.32 0.31 0.21 0.22 23.6 14.7 13.8 20.4 13.5 13.1 6.77 0.07 3.55 0.05 3.06 0.04 3.53 0.03 2.54 0.03 2.57 0.03 MgO 2.25 1.19 1.04 0.87 0.44 0.44 CaO 3.44 3.83 3.44 N.D. 3.12 2.8 Na2O 5.32 5.23 5.51 N.D. 4.82 4.32 K2O 1.84 1.46 1.32 6.4 1.36 2.47 P2O5 0.20 0.15 0.13 0.07 0.09 0.07 Total 103.1 99.75 99.66 99.81 99.61 99.62 Figure 3: Classification diagrams for measured samples. REFERENCES: Kelley, L.I., 1980, Kaolinitic weathering zone on Precambrian basement rocks, Red River Valley, eastern North Dakota and northwestern Minnesota. M.S. Thesis, University of North Dakota. 85 pp. Klasner, J.S. and E. R. King. 1986. Precambrian basement geology of North and South Dakota. Canadian Journal of Earth Sciences. 23(8): 1083-1102. https://doi.org/10.1139/e86-109 Moore, W. L., 1978, A preliminary report on the geology of the Red River Valley Drilling Project, eastern North Dakota and northwestern Minnesota: Bendix Field Engineering Company Subcontract H77-059-E, 292p. https://www.osti.gov/biblio/6538603 doi:10.2172/6538603. 7 �Geophysical imaging of the Paleoproterozoic Animikie basin in Minnesota BEDROSIAN, Paul A., PETERSON, Dana E. and HOOGENBOOM, Bennett E. U.S. Geological Survey, Bldg 20, MS 964, Denver Federal Center, Denver, CO 80225 The 1.88-1.83 Ga Penokean orogen is preserved as a discontinuous fold belt stretching nearly 1500 km from central Minnesota to eastern Ontario. In Minnesota, the supracrustal sequence occupies a NW-facing salient broadly divided into a southern fold-and-thrust belt and a northern tectonic foredeep. The former consists of volcanic and sedimentary rocks in several structural panels while the latter - the ‘main bowl’ Animikie basin (AB) - consists of thick sedimentary sequences and is one of the least deformed remnants of this former continentalmargin. Metasedimentary rocks of the AB include chemical sedimentary rocks (e.g., iron formation of the Mesabi iron range) and turbidites of the Virginia and Thomson Formations. The latter are an important source of sulfur for Ni-Cu mineralization within the intruding Duluth Complex and satellite intrusions. The USGS has been collecting geophysical data in the AB and surrounding areas, including airborne electromagnetic, broadband and nodal seismic, and magnetotelluric data. These data and resulting models reveal the main bowl AB to be more complex than suggested from surface geological mapping. High-electrical conductivity is mapped throughout the basin, including along its northern edge where it is linked to the gently dipping bedded-pyrrhotite-unit and along the steeply dipping SW edge of the Duluth Complex, where it may reflect a hornfels zone formed during contact metamorphism. At the basin scale, a discontinuous bowl-shaped high conductivity zone extends to ~5 km depth. This intra-basin conductor shows some relation to deformation boundaries, such as a demarcation between rocks exhibiting folding and cleavage and those that do not. Some deep (>5 km) geophysical variations are likely related to structural variations within the Archean basement or within thrust panels inferred to project some distance beneath the basin. Where exposed, strong conductors within some of the adjacent thrust belts suggest a correlation with metamorphic grade. Elevated conductivity can, in most cases, be related to a combination of metallic sulfides and graphite. A 9-20 km deep, steeply dipping conductive band is also imaged internal to the Duluth Complex and adjacent to modeled high-density bodies interpreted as magmatic feeder zones. We interpret this conductor as a remnant of AB metasediments preserved within the complex and speculate that the AB played an important control on magma emplacement. 8 �Assessing depositional and post-depositional mineral associations in the <1.71 Ga Freedom Formation, Baraboo, WI, USA. BHUIYAN, Farhan Ahmed 1, BRENGMAN, Latisha 1, and STEWART, Esther 2 1 University of Minnesota Duluth, Earth & Environmental Sciences Department, University of Minnesota Duluth, 1114 Kirby Drive, Heller Hall 229, Duluth, MN 55812. 2 Wisconsin Geological and Natural History Survey, University of Wisconsin-Madison Division of Extension, 3817 Mineral Point Rd, Madison, WI 53705. Some geochemical proxy records suggest that the period following Earth’s initial rise in atmospheric oxygen ~2.4 billion years ago was marked by low, fluctuating oxygen levels (Lyons et al., 2014; 2021). Such conditions would likely have made the ocean’s photic zone inhospitable to multicellular life forms that require oxygen (e.g. Krause et al., 2022). Interpreting the trajectory of Earth’s oxygenation is complicated due to uncertainties in the diagenetic effects on redox proxy records and limited preservation, especially across the late Paleoproterozoic and early Mesoproterozoic (e.g. Slotznick et al., 2022). To fill in critical knowledge gaps in surface redox conditions and geochemical characteristics of Mid-Proterozoic depositional environments, we investigated the Freedom Formation, a <1.71 Ga and >1.47 Ga iron-rich chemical sedimentary unit preserved in historic drill cores near Baraboo, WI, USA (Stewart et al., 2021). Our goal is to decipher primary redox information that links back to the depositing fluid. To accomplish this goal, and separate primary depositional signatures from post-depositional overprinting, we integrate core observations, mineral, petrographic, and geochemical datasets. The Freedom Formation includes a lower unit composed of thin-bedded, interlaminated, fine-grained clastic and chemical sediments and an upper unit composed of dolomite. We document a coarsening upward sequence in the lower Freedom Formation, accompanied by mineralogical changes in three drill cores (H122, H22, and H23). Two of the cores (H122 and H22) show a transition in mineralogy from a base assembly of chamosite, quartz, and magnetite to a Mn-carbonate and hematite-dominant assembly towards the top of the sections. This mineralogical transition is accompanied by an increase in the proportion of sand-sized material, a decrease in mud-sized material, and a noticeable transition to carbonate. Veining and disrupted beds occur throughout all the cores. Whole rock geochemical samples targeting carbonate-rich beds across the lower Freedom Formation indicate a decline in clastic contamination up section, marked by falling Al2O3 concentrations and reduced Zr/Hf ratios. Positive shale normalized Eu/Eu*SN anomalies indicate a role for hydrothermal fluids in precipitation of the authigenic mineral phases. Throughout the lower units, anoxic conditions are dominant, indicated by positive shale-normalized cerium (Ce/Ce*SN) anomalies. Interpreting the observed mineralogical transition in drill cores and the geochemical dataset requires detailed, systematic petrographic observation to distinguish the relative order of events and develop a paragenetic sequence. We identify texturally early minerals based on criteria outlined in LaBerge (1964) and separate those from post-depositional phases to interpret the history of the unit. To classify as a texturally-early phase, minerals must meet the following criteria: 1) be very fine-grained (where no grain size reduction can be attributed to metamorphism); 2) form even and consistent grain size distributions throughout the sample; 3) form the main component of granules or mud-sized particles in fine-grained layers characterized by a granular or particulate textural pattern; and 4) be associated with sedimentary features like 9 �bedding. From this work, we identified quartz and chamosite as the texturally earliest phases in the lowermost Freedom Formation. Towards the top of the lower Freedom Formation, carbonate phases were most commonly identified as texturally earliest. Additionally, the following key observations were made: (1) if quartz is not present, chamosite is the texturally earliest phase; (2) if chamosite is not present, then carbonate is the texturally earliest phase; (3) nano-scale hematite exists at boundaries of chamosite, quartz, carbonate, and stilpnomelane crystals, and within veinlets; (4) euhedral magnetite cross-cuts all other phases, and is often associated directly with chamosite; and (5) large hematite sometimes crosscuts small magnetite, or forms oxidized rims on euhedral magnetite crystals; (6) multiple generations of quartz and oxides exist; and (7) at least two types of carbonate are present (Fe- and Mn-rich and poor). Combining core and mineral datasets, we note that because multiple generations of oxides are present Figure 1: Reflected light photomicrograph of slide no: H122 538 (Fig. 1), post-formational fluid FF (50X) documenting oxidized hematite rims on magnetite flow may directly connect to crystals. observed redox changes in oxide phases. The most critical of these post-depositional observations is the oxidation of magnetite rims (Fig. 1). Overall, across all the cores, independent of redox changes observed in oxide phases, we note a transition in texturally early phases from reduced fine-grained, Fe2+- containing minerals to Mn-Carbonate. This mineralogical transition is marked by anoxic geochemical signatures and possibly indicates minor variations in oxygen conditions during the formation of the mineral phases preserved in the Freedom Formation. References: Krause, A. J., W. Mills, B. J., Merdith, A. S., Lenton, T. M., & Poulton, S. W. (2022). Extreme variability in atmospheric oxygen levels in the late Precambrian. Science Advances. https://doi.org/abm8191 LaBerge, G. L. (1964). Development of magnetite in iron formations of the Lake Superior region. Economic Geology, 59(7), 1313–1342. Lyons, T. W., Diamond, C. W., Planavsky, N. J., Reinhard, C. T., & Li, C. (2021). Oxygenation, life, and the planetary system during Earth’s middle history: An overview. Astrobiology, 21(8), 906–923. Lyons, T. W., Reinhard, C. T., & Planavsky, N. J. (2014). The rise of oxygen in Earth’s early ocean and atmosphere. Nature, 506(7488), 307–315. Slotznick, S. P., Johnson, J. E., Rasmussen, B., Raub, T. D., Webb, S. M., Zi, J. W., Kirschvink, J. L., & Fischer, W. W. (2022). Reexamination of 2.5-Ga “whiff” of oxygen interval points to anoxic ocean before GOE. Science Advances, 8(1), eabj7190. Stewart, E. K., Brengman, L. A., & Stewart, E. D. (2021). Revised Provenance, Depositional Environment, and Maximum Depositional Age for the Baraboo (< ca. 1714 Ma) and Dake (< ca. 1630 Ma) Quartzites, Baraboo Hills, Wisconsin. The Journal of Geology, 129(1), 1–31. 10 �The Lake Superior area “event layer”: Testing the connection with the Sudbury impact BLEEKER, Wouter1, WODICKA, Natasha1, KAMO, Sandra2, HAMILTON, Michael2, EMON, Quinn1, and SMITH, Jennifer1 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 2 Following the initial publication of Addison et al. (2005) [1], there has been growing recognition of a major “event layer” near the top of the Gunflint Formation in the Thunder Bay area [e.g. 2,3,4], and at equivalent stratigraphic levels in the Marquette Range Supergroup of Michigan and Wisconsin [5,6]. Despite some initial hesitation [7,8], a consensus quickly emerged that this event layer represents the local manifestations, in the Lake Superior area, of the 1850 Ma Sudbury impact event that, ~600 km to the east, formed a multiring impact crater centered on the Sudbury area (see [9,10] for a recent review). The deformed and partially preserved remnants of this crater are known as the “Sudbury Structure” and, with a reconstructed final crater diameter of ~300 km, it represents the largest terrestrial impact crater known in the geological record [9]. As such, it would have had far-reaching effects extending out to several crater diameters from “ground zero”, in addition to a global fall-out layer of impact material (cf. the global K-Pg ejecta layer). A curious question, raised in Bleeker & Kamo (2022) [10], is why then this event layer is not more widely recognized in Canada, in places where ca. 1850 Ma basinal stratigraphy is reasonably well preserved (e.g., Mistassini Basin, Fox River Belt, Belcher Islands, Labrador Trough etc.)? Some of these localities are not much farther away from a “ground zero” near Sudbury. This has prompted us to undertake further tests of the putative link between the Lake Superior area event layer and the Sudbury impact structure. Our first test is to more precisely date, by CA-ID-TIMS, felsic ash layers in the lowermost Rove Formation, i.e. the first well-defined and well-preserved tuff layers overlying the event layer. Currently our results suggest the oldest of these tuff layers is ca. 1842 Ma, thus tightening the permissible time interval for the event layer to 1856-1842 Ma [cf. 1,2]. Our second test (in progress) is to attempt a precise CA-ID-TIMS zircon date of the ca. 1850 Ma Peavy Pond Granodiorite (which currently has a SHRIMP age with ±11 Myr uncertainty, see [11]). This granodiorite is known to intrude the lower Michigamme Slates, Baraga Group, in Michigan (W. Cannon, pers. comm. 2023), thus constraining a minimum age for the event layer. Our third and potentially most definitive test is to identify “tracers” in the event layer of the Lake Superior area that can be uniquely tied to target rocks of the Sudbury area. One such tracer would be 2460 Ma zircons from the Copper Cliff Rhyolite and its subvolcanic intrusions (Creighton and Murray granites) that are unique to the area and represent the final felsic rift volcanism/magmatism of the lowermost Huronian Supergroup [12,9,10]. For this test we processed a large bulk sample (~7.5 kg) of the event layer, with its diagnostic grey accretionary lapilli, from the HWY 588 locality west of Thunder Bay. Zircons were separated and mounted for SHRIMP U-Pb analysis at the Geological Survey of Canada, Ottawa. Eighty three zircon grains were spot dated, of which 71 returned high-quality results (Figure 1). Figure 1: U-Pb concordia plot for spot dates by SHRIMP on 71 zircon grains from the Gunflint event layer, HWY 588 roadside outcrop. 11 �The age distribution shows several well-defined clusters with 2(3) of the dated zircons defining a small but discrete subpopulation at ca. 2460 Ma. One of these zircons shows possible shock features (PDFs, planar deformation features; see Figure 1 inset) identified during picking. Although this particular grain is likely ca. 2460 Ma in origin, its result shows considerable discordance and should thus be treated with caution. Nevertheless, we think the 2460 Ma subpopulation uniquely ties fall-out material in the event layer, including rare shocked quartz grains [e.g. 2], to the Sudbury crater and its target rocks. In addition to the conclusive result of the 2460 Ma zircons, the data also identify a distinct ca. 2310-2320 Ma subpopulation of zircon grains that are known to first show up (in a regional stratigraphic sense) in the Gordon Lake Formation of the upper Huronian Supergroup. These zircon grains could have been delivered to the Thunder Bay area either as 1) ejecta from the Sudbury impact event, or perhaps more likely 2) as reworked detrital zircons from widespread felsic ash material that was deposited across the wider Superior craton at 23102320 Ma. Finding shock features in these grains would favour the first scenario, whereas a total absence of shock features would favour the second scenario. To further constrain the impact event, we also subsampled the large sample from the HWY 588 event layer into 6 small slabs with varying abundances of 1–3 cm accretionary lapilli (i.e. from ~5 to ~95 vol% lapilli) and analyzed these for major and trace elements, and for low-level PGE abundances. Results show a negative correlation between siderophile elements such as Ir (also Ru, Ni, Cr etc.) and lapilli abundance, indicating that the lapilli consist largely of diluting material and are not the optimum target for identifying the nature of the impactor [e.g. 13]. The highest Ir content of 0.3–0.4 ppb, i.e. ~1–2 orders of magnitude above average crustal values, actually occurs in laminated, dark, fine sand- to silt-size sediments that overlie the lapilli-rich horizon fall-out material (~0.5–1.0 m above). Future work will entail more detailed sampling of this overlying stratigraphy to identify the Ir peak and define the detailed mineralogy and Ir deportment in this material. Incidentally, values of 0.3–0.5 ppb Ir are also the maximum recorded values in the upper Onaping Formation filling the Sudbury crater and overlying its melt sheet [14]. From our initial results it appears that a maximum of impactor material condensed relatively late and was least diluted in fine grained fall-out material near the top of the event layer, well above the accretionary lapilli. References [1] [2] [3] [4] [5] [6] [7] [8] [9] [10] [11] [12] [13] [14] 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. Geology, vol. 33(3), p. 193–196. Addison, W.D., Brumpton, G.R., Davis, D.W., Fralick, P.W., and Kissin, S.A., 2010. GSA Special Paper 465, p. 245–268. Jirsa, M.A., Fralick, P.W., Weiblen, P.W., and Anderson, J.L.B., 2011. GSA Field Guide 24, p. 147–169. Huber, M.S., McDonald, I., and Koeberl, C., 2014. Meteoritics & Planetary Science, vol. 49(10), p. 1749–1768. Pufahl, P.K., Hiatt, E.E., Stanley, C.R., Morrow, J.R., Nelson, G.J., and Edwards, C.T., 2007. Geology, vol. 35(9), p. 827–830. Cannon, W.F., Schulz, K.J., Horton Jr, J.W., and Kring, D.A., 2010. GSA Bulletin, vol. 122(1–2), p. 50–75. Kissin, S.A., and Fralick, P.W., 1997. Journal of the Royal Astronomical Society of Canada, vol. 91(5), p. 216. Kissin, S.A., Okamoto, M., Addison, W.D., and Brumpton, G.R., 2000. 46th Annual Meeting of the Institute on Lake Superior Geology, vol. 46, part 1, p. 31–32. Bleeker, W., and Kamo, S., 2022a. 68th Annual Meeting of the Institute on Lake Superior Geology, vol. 68, part 1, p. 5–6. Bleeker, W., and Kamo, S., 2022b. 68th Annual Meeting of the Institute on Lake Superior Geology, vol. 68, part part 2, Field Trip Guidebook, p. 4–57. Ayuso, R.A., Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., and Jackson, J., 2018. 64th Annual Meeting of Institute on Lake Superior Geology, vol. 64, part 1, p. 7–8. Bleeker, W., Kamo, S.L., Ames, D.E., and Davis, D., 2015. Geological Survey of Canada Open File 7856, p. 151–166. Mougel, B., Moynier, F., Göpel, C., and Koeberl, C., 2017. Earth and Planetary Science Letters, vol. 460, p. 105–111. Mungall, J.E., Ames, D.E., and Hanley, J.J., 2004. Nature, vol. 429(6991), p. 546–548. 12 �Insights into the southwestern Superior Province: New igneous geochronology and geochemistry in northwestern Minnesota, USA BLOCK, Amy Radakovich1, HUDAK, George J.2, SOUDERS, A. Kate3 1 Minnesota Geological Survey, 2609 Territorial Road, St. Paul, MN 55114 University of Minnesota –- Twin Cities, 116 Church Street SE, Minneapolis, MN 55455 3 U.S. Geological Survey, Denver, CO 80225 2 The U.S. Geological Survey Earth Mapping Resources Initiative (Earth MRI) program recently funded acquisition of new airborne geophysical (Allen Langhans and Drenth, 2023), geochronologic, and geochemical data in a part of the Superior Province in northwest Minnesota that is prospective for numerous Archean critical-mineral-producing systems. The study area comprises three subprovinces of the Archean Superior Province (Fig. 1). Previous work in the area (Jirsa et al., 2012; Jirsa et al., 1999) has been severely limited by an absence of outcrop, sparse drill hole data, and the existence of only one geochronologic age. These new ages and geochemical analyses, obtained through the Earth MRI program, represent the first highresolution geologic data in the Neoarchean subprovinces of the southwestern Superior Province. Seven new U-Pb zircon LA-ICP-MS magmatic ages (Souders, in review) establish the timing of intrusive activity in the southwestern extent of both the Wawa and Wabigoon subprovinces. A biotite-hornblende tonalite in the Red Lake Falls pluton (207Pb/206Pb weighted mean age of 2701 ± 4 Ma, 2s), a biotite tonalite in the Snake River batholith (207Pb/206Pb weighted mean age of 2738 ± 12 Ma, 2s), and a diorite in the Grygla pluton (207Pb/206Pb weighted mean age of 2771 Ma ± 8 Ma, 2s) define three distinct Neoarchean episodes of intermediate intrusive activity near the present-day southern margin of the Wabigoon subprovince. A preliminary magmatic age from a small hornblende monzodiorite stock in the Wabigoon indicates ca. 2727 intermediatemafic intrusive activity. In the Wawa subprovince, a 207Pb/206Pb weighted mean age of 2702 ± 6.5 Ma age (2s) from the Fertile pluton biotite granodiorite indicates Neoarchean intermediate intrusive activity at the northern margin of the Wawa subprovince coincident with similar activity in the Wabigoon. A small mafic body that intrudes a mafic volcanic sequence in the Wawa yields a preliminary age of ca. 2690 Ma. Finally, a combined 207Pb/206Pb weighted mean age from two closely spaced anorthosite samples confirms a Neoarchean (2737 ± 4.5 Ma, 2s) age for the Mentor Anorthosite Intrusive Complex (MAIC). High-precision CA-TIMS U-Pb zircon analyses provide age constraints on supracrustal rocks in both subprovinces. Two trachyandesite lapilli tuff samples from the Wabigoon subprovince yield 207 Pb/206Pb weighted mean ages of ca. 2730 Ma and ca. 2733 Ma (Block et al., in prep. b), >25 Ma older than the few other volcanic ages from the Wabigoon in Minnesota. Two feldspathic wackes in the Wawa subprovince are still being processed for ages. Newly dated intermediate intrusions are LREE enriched and have arc-like trace element patterns. Discrimination diagrams indicate that these intrusions are generally I-type, calc-alkaline, volcanic arc-granites. Samples from the MAIC exhibit complex REE patterns, and their interpretation is less straightforward. The newly dated intermediate volcanic samples from the Wabigoon are also calc-alkaline and exhibit arc-like signatures. In combination with ~130 additional geochemical analyses and detailed petrography, results presented here provide significant insight into the tectonic evolution of the southwestern Superior Province and invite comparison with well-studied rock packages in the Wabigoon and Abitibi provinces in Canada. 13 �Figure 1. Bedrock geology map of northwestern MN, USA. Units within the project area (Block et al., in prep. a) are shown in the legend. Units outside the map area are from Jirsa et al. (2012). Newly obtained LA-ICPMS U-Pb ages are shown as yellow stars, and newly obtained TIMS U-Pb ages are shown as green stars. 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. Block, Amy Radakovich, Drenth, Benjamin J., Souders, A. Kate, Hudak III, George J, Hirsch, Aaron C., and Saari, Stacy M., in prep. A, Geologic map of the Mentor Igneous Complex Focus Area, Northwest Minnesota: Minnesota Geological Survey, Miscellaneous Map Series M-200, scale: 1:100,000. Block, Amy Radakovich, Hudak III, George J, Souders, A. Kate, Drenth, Benjamin J., Schmitz, M., Hirsch, Aaron C., and Saari, Stacy M., in prep. b, Preliminary investigation of the geologic history and critical mineral potential of the Mentor Igneous Complex Focus Area, Northwest Minnesota: Minnesota Geological Survey, Report of Investigations 74. Jirsa, M. A., Boerboom, T. J., Chandler, V. W., 2012, Geologic Map of Minnesota, Precambrian Geology: Minnesota Geological Survey, Map S-22, 1:500,000. Jirsa, M. A., Chandler, V. W., and Runkel, A. C., 1999, Bedrock geologic map of northwestern Minnesota: Minnesota Geological Survey, Miscellaneous Map Series M-92, 1:200,000. 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 �Magmatic and hydrothermal evolution of the Mesoproterozoic Current ultramafic PGE-CuNi deposit within the Thunder Bay North Intrusive Complex: insights from trace elements, Nd, Sr, O, and H isotopes CORREDOR BRAVO, Andrea Paola1, HOLLINGS, Pete1, BRZOZOWSKI, Matthew1, and HEGGIE, Geoff2 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Clean Air Metals, 1004 Alloy Drive, Thunder Bay, ON P7B 6A5 Canada 2 The Mesoproterozoic PGE-Cu-Ni enriched Current intrusion, part of the Thunder Bay North Intrusive Complex, is located 50 km northeast of Thunder Bay, Ontario. The northwest-trending intrusion is a 3.4 km long conduit-type that intruded the rocks of the Quetico Basin during the early stages of the Midcontinent Rift System (MRS; Woodruff et al, 2020). The intrusion has four mineralized zones; the Current and Bridge Zone in the northwest are characterized by shallow and thin features; in the middle lies the BeaverCloud Zone, characterized by its substantial thickness, while the southeast is the deepest 437Southeast Anomaly (SEA) Zone (Kuntz et al., 2022). The intrusion exhibits a primitive mantle- Figure 1. Schematic model of the Current intrusion and the Quetico country rock. normalized pattern resembling ocean island Illustration compiled in Leapfrog using data basalt, characterized by LREE enrichment and provided by Clean Air Metals Inc. small positive anomalies in Nb, La, and Ce, consistent with minimal continental crust contamination. The La/Smn values in the Current intrusion samples, ranging from 1.8 to 2.6, align with previous studies, indicating a basaltic magma derived from an enriched mantle plume. The enriched nature of the magma in the Current intrusion is consistent with other mineralized and unmineralized intrusions associated with the MRS (Escape, Seagull, Lone Island intrusion, and Nipigon Embayment; Heggie, 2005; Hollings et al., 2007b; Caglioti, 2023; Yahia, 2023). The intrusion has slightly lower Sri (0.7021 to 0.7043) and εNd (-1.18 to -4.02) values compared to typical the mantle source at 1100 Ma. Therefore, it is suggested that the plume-derived magma interacted with an enriched subcontinental lithospheric mantle, which may have contributed to the slightly negative εNd values of the intrusion.The stable isotope analysis data from the Current intrusion indicates an interaction between magmatic mantle-derived fluids (δ2H from −40 to −80‰, δ18O from 5.5 to 7.0‰), meteoric fluids (δ2H <-80‰, δ18O <5.5‰), and devolatilization/ dehydration fluids of the Quetico country rocks (δ18O >7‰). Three distinct domains within the intrusion were identified based on alteration intensity and micro-textural observations and each showing varying secondary mineral assemblages Domain A consists of antigorite, tremolite, clinochlore, epidote, pyrite, cubanite, millerite, secondary pyrrhotite ± chamosite ± sericite, and ± secondary magnetite, Domain B consists of 15 �lizardite-chrysotile, tremolite, clinochlore, epidote, pyrite, cubanite, millerite ± sericite, and ± secondary magnetite Domain C consists of talc and carbonates. Domain A and B have characteristics of interaction with meteoric, mantle, and/or subcontinental lithospheric mantle -derived fluids, whereas Domain C is associated with fluids from devolatilization of the country rock and is overprinted on Domains A and B. The alteration processes in the different domains Figure 2. δ18O and δ2H values of bulk rock in the four involved two distinct fluid types at mineralized zones of the Current intrusion (Current, varying temperatures, Domain A likely Bridge, Beaver-Cloud, and 437-SEA) and the involved higher temperatures (>300°C) surrounding country rock of the Quetico basin. and fluids rich in H2O. In contrast, domain B was altered by fluids at lower temperatures (<300°C). Later CO2-bearing fluids of Domain C overprinted earlier alteration at temperatures below 50°C. The alteration of the intrusion also resulted in significant volume reduction of primary sulfides and oxides that have been replaced by secondary minerals, such as chalcopyrite and pyrrhotite were replaced by secondary magnetite and pyrite and primary magnetite was replaced by pyrite and chamosite. References Caglioti, C. (2023). PGE–Cu–Ni sulfide mineralization of the Mesoproterozoic Escape intrusion, northwestern Ontario (MSc). Lakehead University, Thunder Bay, Ontario. https://knowledgecommons.lakeheadu.ca/handle/2453/5195 Heggie, G. (2005). Whole rock geochemistry, mineral chemistry, petrology, and Pt, Pd mineralization of the Seagull Intrusion, Northwestern Ontario. Lakehead University, Thunder Bay, Ontario. https://knowledgecommons.lakeheadu.ca/handle/2453/689 Hollings, P., Richardson, A., Creaser, R. A., and Franklin, J. M. (2007b). Radiogenic isotope characteristics of the Mesoproterozoic intrusive rocks of the Nipigon Embayment, northwestern Ontario. Canadian Journal of Earth Sciences, 44(8), 1111-1129. https://doi.org/10.1139/e06-128 Kuntz, G., Wissent, B., Boyk, K., Harkonen, H., Jones, L., Muir, W., Buss, B., and Peacock, B. (2022). NI 43- 101 Technical report and preliminary economic assessment for the Thunder Bay North Project, Thunder Bay, Ontario Woodruff, L. G., Schulz, K. J., Nicholson, S. W., and Dicken, C. L. (2020). Mineral deposits of the Mesoproterozoic Midcontinent Rift system in the Lake Superior region–a space and time classification. Ore Geology Reviews, 126, 103716 Yahia, K. (2023). Geochemistry, petrography, geochronology, and radiogenic isotopes of the weakly mineralized intrusions in Thunder Bay North Igneous Complex (MSc). Lakehead University, Thunder Bay, Ontario. https://knowledgecommons.lakeheadu.ca/handle/2453/5283 16 �Re-interpretation of hydrothermal alteration, mineralization and host-rock oxidation to form the Keweenaw native copper lodes, northern Michigan BROWN, Alex C. 13250 rue Acadie, Pierrefonds, Quebec, Canada, H9A 1K9, acbrown@polymtl.ca The sandstone/conglomerate-hosted portions of the native copper ores of northern Michigan (e.g., the Calumet and Hecla Conglomerate ores) occur mostly in deeply reddish sediments. In the immediate vicinity of native copper ores, the reddish sediments appear to have been hydrothermally bleached to salmon-red colors (Butler and Burbank, 1929; Cornwall, 1956; White, 1968; Weege and Pollock, 1972). This communication notes that fine-grained salmoncolored clastic sediments may host fine-grained disseminations of native copper enclosed by salmon-red aureoles, not unlike grey reduction halos commonly found in red sandstones (Figs. 1, 2). If the interpreted origin and preservation of reduction halos in red sandstones is applied to the native copper-hosting aureoles in the Calumet and Hecla Conglomerate, the deep reddening of the conglomerates hosting native copper of the Keweenaw Peninsula may be interpreted as a post-ore event. Redbed sandstones commonly show centimeter-scale reduction spots and blotches, e.g., rift-hosted Carboniferous clastic sediments of eastern Canada (Poll and Sutherland, 1976), the Permian fluvial Abo Formation of New Mexico (Bensing et al., 2005), and the Jacobsville sandstones of northern Michigan. Petrographic and chemical analyses of reduction spots in the Abo Formation indicate that those reduction spots have never been reddened – ferrous clastic grains within the reduction spots are still ferrous while similar grains in the enclosing red sandstone are oxidized (Bensing et al., 2005). Interpretation: wood trash in the cores of reduction spots maintained elliptical reducing conditions in the immediate vicinity of wood trash (i.e., within the grey halos), while oxidizing post-sedimentary water reddened all other portions of the clastic sediments. Figure 1. Carboniferous redbeds of Dorchester Cape, New Brunswick, Canada, showing abundant centimetric-scale reduction spots with dark-greyish cores centered on fossilized organic matter. Figure 2. Close view of greyish reduction spots in redbeds of Figure 1 (tip of hammer for scale). Cores of reduction spots contain fossil wood debris and base-metal sulfides e.g., chalcocite, partially oxidized to malachite). 17 �Native copper in the Calumet and Hecla Conglomerate occurs as interstitial fillings in conglomerates and as fine-grained disseminations in finer sandy sediments. Curiously, very finegrained disseminations of native copper in fine-grained sediments are observed to be surrounded by elliptical salmon-red sediment (Fig. 3). A possible, chemically justified interpretation: native copper was deposited with salmon-red alteration, mostly within highly permeable conglomeratic portions of the sandstone-conglomerates and also as very fine-grains in associated sandy sediments. Subsequently, all sandstone-conglomerates were thoroughly oxidized to their classic deep-red color during post-ore circulations of oxygenated ground water, except in the finergrained sediments where local salmon-red alteration was preserved against reddening by reduction-inducing fine grains of native copper. Post-ore deep-red oxidation of copper in the fine-grained sediment was inhibited by the poor permeability of this fine-grained sediment to late deep-reddening groundwaters, but also by the reducing property of metallic copper. Figure 3. Cut and epoxy-ed sample of Calumet and Hecla Conglomerate native copper ore. Upper half: Deep red, coarse conglomeratic sediment with coarse-grained native copper. Lower half: Fine-grained clastic sediment containing fine-grained native copper enclosed by “bleached” salmon-red alteration halos. Bleached halos and core native copper are equated here to reduction spots with fossil wood debris common in redbed sandstones (see text for explanation). References Bensing, J.P., Mozley, P.S., and Dunbar, N.W., 2005. Importance of clay in iron transport and sediment reddening: evidence from reduction features of the Abo Formation, New Mexico, U.S.A. Sedimentary Research, 75: 562–571. Butler, B.S. and Burbank, W.S., 1929. The copper deposits of Michigan. US Geol. Surv. Prof. Paper 144, 238 p. Cornwall, H.R., 1956. A summary of ideas on the origin of native copper deposits. Economic Geology, 51: 615–631. Weege, R.J. and Pollock, J.P., 1972. The geology of two new mines in the native copper district of Michigan. Economic Geology, 67: 622–633. White, W.S., 1968. The native-copper deposits of northern Michigan, in Ridge, J.D., ed., Ore Deposits of the United States, 1933–1967 (Graton-Sales Volume 1), American Inst. Min. Metall. & Petrol. Eng., 303–326. 18 �Preliminary mineralogy of a pegmatite in the pyroxene syenites of the Stettin Complex, Wausau Complex, Marathon County, Wisconsin Buchholz, Thomas 1, Falster, Alexander 2, and Simmons, Wm 2 1 1140 12th Street North, Wisconsin Rapids, Wisconsin 54494 MP Research Group, Maine Mineral and Gem Museum, PO Box 500, 99 Main Street, Bethel, Maine 04217, USA 2 2 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 composed of various syenite phases. Recently an opportunity arose to examine the mineralogy of a pegmatite located in the pyroxene syenites of the Stettin Complex. Upper portions of the roughly horizontal pegmatite are below tilled soils but in-situ, are weathered, and fragments are coated with Fe-oxides and clays. Excavations over the last several years has exposed somewhat fresher material at depth and allowed better study. The pegmatite is zoned, but numerous included syenite screens complicate evaluation. Pegmatite-host syenite contacts are often sharp with no notable contact zones, suggesting relatively minor temperature contrast between the two phases, but thin 2-3 cm reaction zones are also common, with minor coarsening of feldspar and arfvedsonite compared to host syenite, small miaroles, scattered patches of abundant, tiny pink zircons and yet unidentified minerals, and in the freshest material, fluorite. The upper weathered portions of the dike, probably corresponding to border and intermediate zones, consist of anorthoclase (often showing “moonstone” visual effects; the recovery of these feldspars is the objective of those working the dike), highly altered former pyroxenes(?) with-sparse remnant hedenbergite, arfvedsonite, quartz, abundant clear pink to orange zircons and other accessory phases, including small miarolitic cavities. Per Medaris & Koellner (2010), pyroxenes in the Stettin complex range from Fe-rich diopside, to hedenbergite and aegirine. In this pegmatite pyroxenes appear to have been affected by late-stage oxidizing fluids, altering Fe2+-rich pyroxenes to Fe3+ rich smectite-group clays ± Fe-oxyhydroxides with sparse remnants of hedenbergite, while aegirine is absent from the dike. Ca released by pyroxene alteration may have contributed to the formation of various late-stage Ca-rich species. Deeper interior zones are mostly anorthoclase with arfvedsonite and other accessory minerals, with contact zones (or lack thereof) repeated around included syenite fragments. Graphic quartz-anorthoclase intergrowths “graphic syenite” are locally common, as are small miarolitic cavities. Several small-volume pegmatite units are: rare irregular patches of granular albite +- larger anorthoclase crystals, with abundant pyrochlore(?) crystals and possibly other species; and enigmatic thin 2-5 cm thick irregular veins or pods, mainly quartz, albite and anorthoclase: these are confined to the pegmatite and do not enter the host syenite, and includes sparse arfvedsonite, abundant cassiterite grains (≈2-4 μm), fergusonite, metamict zircons (some showing Hf-enrichment), sparse microlite (Ta-dominant pyrochlore group), sparse tantalite-(Mn) (Ta-Mn dominant columbite-group species), tiny grains of barite, and other yet-unidentified species. These anomalous pods or veins may be derived from highly fractionated late-stage differentiates. Perhaps high F activity supported unusual enrichment of HFSE in latecrystallizing melt. . 19 �Other accessory mineralogy of the main dike includes: Aeschynite-(Ce): Elongated, dark greyblack crystals in anorthoclase. Synchysite/Parisite: Small red to pinkish hexagonal crystals. High Ca contents suggest they are either synchysite or parisite. Bavenite(?): Small white bladed crystals included in clear quartz crystals; tentative ID based on their morphology and presence of Ca, Si and O (EDS). Calcite: White crusts and masses in vugs from lower portions of the excavation. Chevkinite-group(?): Dark grains with white borders; often heavily altered to soft, chalky, fine-grained niobian Ti-oxides with minor Th, Ca, LREE ± Si and Al. Columbite-group: Sparse columbite-(Fe) noted as inclusions in a porous fergusonite-(Y) grain. Fayalite: Uncommon gray radiating acicular crystals and glassy brown grains associated with arfvedsonite. Fergusonite-(Y): Small yellowish to reddish-brown tapering crystals in anorthoclase, arfvedsonite and miaroles. Ferro-anthophyllite: Uncommon, patches of white acicular crystals in smectite-rich altered pyroxenes. Fluorapatite: Sparse crystals in vugs with arfvedsonite. Fluorite: Late in vugs and isolated grains, likely often removed by weathering. Graphite: Sparse Thin hexagonal platy crystals, typically showing thin hexagonal overgrowths. Ilmenite: Common; thin black metallic plates with a pyrophanite component in anorthoclase and miaroles. Kainosite-(Y)(?): One off-white crystal in pegmatite, appears to be a Ca-Y-LREE silicate, may be kainosite-(Y). Magnetite: Common as irregular masses, rarely as well-formed octahedral crystals. Molybdenite: Sparse as thin soft hexagonal plates. Monazite-(Ce): Uncommon, small brick-red crystals in feldspar and in vugs. Niocalite(?): Yellow to pale yellow-brown elongated crystals in anorthoclase. Very sparse. Some compositions strongly suggest niocalite, others are yet-unidentified species. Pyrochlore: Rare: yellow-brown octahedral crystals. Quartz: Common, generally as a later-stage mineral. Siderite: now absent, but goethite pseudomorphs after probable siderite are common in small vugs. Thorite: Rare; red to red-black grains associated with fergusonite-(Y). Titanite: Uncommon, as brown to redbrown grains. Zircon: Abundant in upper intermediate zone, less so in coarse interior zones. As work continues, it is likely that additional phases will be identified, as there are a number of unknowns awaiting further work, and much material awaits cleaning and study. Thanks are due to Austin Gausmann, Bill Schoenfuss, and Trent and Shana Rebeck for access to the pegmatite. REFERENCES: Medaris, L. Gordon Jr., Koellner, Susan E., 2010. Ferromagnesian minerals in the Stettin Syenite Complex, Marathon County, Wisconsin: compositions and contrasts with the Wolf River Batholith (abstract): Institute on Lake Superior Geology Proceedings, 56th Annual Meeting, International Falls, MN, v. 56, part 1, p. 42-43. Van Wyck, N. 1994. The Wolf River A-type magmatic event in Wisconsin: U/Pb and Sm/Nd constraints on timing and petrogenesis (abstract): Institute on Lake Superior Geology, 40th Annual Meeting, Part 1, Program and Abstracts, p. 81-82. 20 �The Sacred Heart Orogeny in Michigan: Latest Archean Granites and the Great Lakes Tectonic Zone CANNON, W. F.1, SOUDERS, A. K2, DRENTH, Benjamin J.2, AYUSO, Robert A.1 1 U.S. Geological Survey, Reston, VA 20192, 2U.S. Geological Survey, Denver, CO 80225 The Sacred Heart Orogeny, as defined in Minnesota (Schmitz, et al., 2018), is the terminal Archean contractional and magmatic event during which the Minnesota River Valley Subprovince (MRVS) was sutured to the southern edge of the Superior craton along the Great Lakes Tectonic Zone (GLTZ), accompanied by voluminous granitic magmatism. The ages of those granites are tightly clustered from 2.58-2.6 Ga. In Michigan’s Upper Peninsula voluminous granites of the Southern Complex, the eastern extension of the MRVS, have ages very similar to those in Minnesota and were emplaced near and within the GLTZ during suturing. The granitic phase of the Bell Creek Gneiss (Cannon and Simmons, 1973), a coarse-grained, K-spar megacrystic granite, is well dated (Petryk, 2019; Barth, 2023). The average of ten dates is 2.56 Ga. An additional age from SHRIMP analyses (Ayuso, presented here) is 2584+5.3/-2.79 Ma, fully consistent with previous determinations. We have recently recognized and dated a batholithic-scale intrusion of medium-grained, massive, K-rich granite, informally the New Swanzy granite, that borders the granitic phase of the Bell Creek Gneiss to the east and is adjacent to and interacted with the GLTZ during suturing. A regionally extensive negative gravity anomaly south and southeast of the exposed granite suggests that the batholith is substantially larger than its exposed portion (see map below). Both granites intruded an older gneiss complex with ages ranging from 2.6-2.9 Ga. The Bell Creek granite has an internal foliation defined by alignment of K-spar megacrysts broadly conformable with trends of adjacent gneisses, and contains no angular xenoliths. In contrast, the New Swanzy granite, although compositionally uniform as judged in outcrop, has numerous pegmatitic segregations and dikes. It has sharp cross-cutting contacts with older gneisses and contains many xenoliths. These characteristics and several instances of dikes of New Swanzy-like granites cutting the Bell Creek, suggests that the New Swanzy granite post-dates the Bell Creek Gneiss, at least slightly, and has an intrusive contact with it. The New Swanzy granite has a complex interaction with sheared rocks of the GLTZ. The GLTZ is a 2.5 km-wide, SW-dipping zone of dextral shear-thrusting where rocks of the MRVS were thrust northward over the Superior craton (Sims, 1991). The age of suturing was suggested to be about 2.69 Ga. Close to the GLTZ, the New Swanzy granite has a shear foliation parallel to the GLTZ. Within the GLTZ areas of granite are enveloped in intensely sheared rocks whose protolith is uncertain. Many dikes and stringers of granite cut mylonitic foliation. These show varying intensity of deformation, but all were emplaced after the most intense shearing. An undeformed granite dike was intruded across shear foliation at 2559±19.5 Ma (U-Pb apatite). We consider this to set a minimum age for deformation in the GLTZ in this region. Thus, the New Swanzy seems best interpreted as a syntectonic granite emplaced adjacent to the active suture late in development of the GLTZ The data summarized here indicate the latest Archean tectonic and intrusive events in northern Michigan and Minnesota are correlative. Thus, we deem it appropriate to extend the Minnesota-derived term “Sacred Heart Orogeny” to the culminating phase of development of the Southern Complex in Michigan. Because rocks of the Southern Complex are exposed in direct contact with the GLTZ, they present a unique opportunity to study the interaction of intrusion and suturing of the MRVS to the Wawa-Abitibi terrane. 21 �A-Massive New Swanzy granite cutting country rock gneiss. B- GLTZ mylonite cut by granitic stringers with varying degrees of deformation. C- undeformed 2.56 Ga granite dike cutting foliation of GLTZ. References Barth, E. G., 2023, Age and chemistry of the Bell Creek Batholith: Michigan Technological University, M.S. Thesis https://doi.org/10.37099/mtu.dc.etdr/1589 Cannon, W.F., and Simmons, G C., 1973, Geology of part of the Southern Complex, Marquette District, Michigan: Journal of Research of the U.S. Geological Survey, v.1, n.2, p. 165-173. Petryk, B. 2019, The origin of an Archean batholith in Michigan’s Upper Peninsula: Michigan Technological University, M.S. Thesis. https://doi.org/10.37099/mtu.dc.etdr/932 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 p. Schmitz, M.D., Southwick, D.L., Bickford, M.E., Mueller, P.A., and Samson, S.D., 2018, Neoarchean and Paleoproterozoic events in the Minnesota River Valley subprovince, with implications for southern Superior craton evolution and correlation: Precambrian Research, v.316, p. 206-226. 22 �The Honey Creek Structure, Sauk County, Wisconsin: Asymmetric Faulting Associated with Seismic-Induced Fluid Escape Kenz CARLTON1, Basil TIKOFF1, & Esther K. STEWART2 1 2 University of Wisconsin–Madison, Department of Geoscience, 1215 West Dayton Street, Madison, Wisconsin 53706, USA Wisconsin Geological and Natural History Survey, UW-Madison Division of Extension, Madison, Wisconsin 53705, USA The Honey Creek structure occurs in the Balfanz quarry in Sauk County, Wisconsin. It is south of the Baraboo syncline and in the vicinity of the Denzer syncline. The Honey Creek structure is dominantly a N dipping fault (striking ~250 and dipping ~30 using right hand rule) that deforms the Ordovician Oneota Formation and underlying Cambrian Jordan Formation. The hanging wall contains units that are stratigraphically younger than those immediately adjacent across the fault; this geometry requires that the Honey Creek structure contains a normal fault. The footwall side of the fault was deformed, such that the beds were rotated to a nearly vertical orientation and occasionally overturned next to the fault. The folding in the footwall is consistent with drag on a reverse fault, which is opposite to the direction of inferred stratigraphic offset. Soft sediment deformation, interpreted as sand injection emanating from the Jordan Formation, is prevalent alongside the fault on the footwall side, although a sand lens crosscuts the fault in one place. A smaller-scale structure is located less than 30 m S of the Honey Creek structure. This feature has a similar strike and dip, normal sense of offset, and folding of the footwall. It differs from the Honey Creek structure insofar as there is brecciation but no sand injection along the fault. Finally, there appears to be a recumbent fold located less than 40 m S from the main Honey Creek structure, with a vergence away from the Honey Creek structure. We interpret this recumbent fold to have formed in the same deformational event. We consider two possible interpretations for this structure, both of which invoke significant ground shaking. First, the deformation could result from intracratonic seismicity. The timing of deformation (Early Ordovician) is broadly consistent with that of the Taconic orogeny, although the orientation of the fault is at a high angle to the inferred regional shortening direction (EW). Its location, directly south of the Baraboo syncline, could be consistent with reactivation of a Proterozoic fault. Second, the deformation could result from a distant (<200 km) meteor impact. The timing of deformation is consistent with a swarm of meteorites that occurred at the same time in the upper Midwest, and resulted in a number of craters (e.g., Decorah, Elm Creek, etc.). This interpretation is consistent with regional soft-sediment deformation in the Oneota Formation. 23 �Baddeleyite age reveals timing of the Northeast Iowa Intrusive Complex (NEIIC) CLARK, Ryan1, PEATE, David2, KUSICK, Allison2,3, HORKLEY, Kenny4, and MACFARLANE, Chris5 1 Iowa Geological Survey, University of Iowa, 300 Trowbridge Hall, Iowa City, IA, 52242, USA University of Iowa, Department of Earth & Environmental Sciences, 115 Trowbridge Hall, Iowa City, IA, 52242, USA 3 University of Wisconsin-Milwaukee, Department of Geosciences, 3209 N. Maryland Avenue, Milwaukee, WI, 53211, USA 4 University of Iowa, Materials Analysis, Testing and Fabrication Facility, 205 N. Madison Street, Iowa City, IA, 52242, USA 5 University of New Brunswick, Department of Earth Sciences, 2 Bailey Drive, Fredericton, New Brunswick, Canada 2 Recent geophysical surveys over portions of the entirely concealed Northeast Iowa Intrusive Complex (NEIIC) have provided a clearer picture of the region’s Precambrian basement geology (Drenth et al., 2015 and 2020). High amplitude magnetic and gravity anomalies remain the focus of further research into their mineral resource potential, due in part to the likelihood that the NEIIC is related to the ~1,100 Ma Midcontinent Rift System (MRS). A core drilled into the northeast-trending Osborne Anomaly in Clayton County, Iowa provides the only samples in the vicinity of the NEIIC. The Osborne core encountered 722 feet (220 m) of mafic-ultramafic rocks that has been previously described as olivine-plagioclase cumulate. Recent screening using a portable X-ray fluorescence (pXRF) spectrometer revealed elevated concentrations of zirconium (Figure 1) as well as aluminum and potassium in several discrete zones of late stage melt (Clark et al., 2019). Datable minerals in the form of zirconolite and baddeleyite have been identified in samples from these zones. Obtaining a reliable age from the Osborne core has been paramount to making the argument that the NEIIC is Keweenawan and thus possibly related to other magmatic intrusive terranes in the Lake Superior Region. A recent study (Drenth et al, 2020) obtained an age of ~1,170 Ma from LA-ICP-MS analyses of apatite crystals from the Osborne core. However, accurate U-Pb ages on apatites are often limited by the need for a precise correction for the common Pb component. Here, we present new U-Pb ages on baddeleyite crystals from a depth of 2,416.3 feet (736.5 m) that were analyzed by LA-ICP-MS at the University of New Brunswick. The U-Pb crystallization age of 1,148 ± 14 Ma (weighted average of six concordant baddeleyite analyses) stands as the first reliable date to come from the NEIIC region. This age is comparable with other intrusions outboard of the MRS, such as the Corson Diabase in eastern South Dakota (1,149 ± 7 Ma), the Great Abitibi dike (1,141 ± 2 Ma), and the Inspiration diabase (1,159 ± 33 Ma) (McCormick et al., 2017 and references therein), and indicate a wider regional magmatic event that pre-dated initiation of the MRS by ~50 Ma. The general age of these intrusions has been interpreted as early stage magmatism related to the onset of the MRS. The latest geophysical survey over the majority of the southern portion of the NEIIC shows that the Osborne Anomaly is cut by NEIIC intrusions (Drenth et al., 2020), thus providing a maximum emplacement age of ~1,150 Ma. 24 �Figure 1. Graph of Zr concentration by depth from two separate rounds of pXRF analyses illustrates two distinct zones of Zr-enrichment. Inset backscattered electron image shows an elongated zirconolite crystal (gray) with inter-grown baddeyelite crystal (white) from a sample at 2,614.3 feet depth. References Clark, R.J., Anderson, R.R., and Peate, D.W., 2019. The northeast Iowa intrusive complex: a magmatic conundrum related to the Midcontinent Rift System. Geological Society of America Abstracts with Programs, v. 51, no. 2. Drenth, B.J., Anderson, R.R., Schulz, K.J., Feinberg, J.M., Chandler, V.M., 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 Science, v. 52: 279-293. Drenth, B.J., Souders, A.K., Schulz, K.J., Feinberg, J.M., Anderson, R.R., Chandler, V.M., Cannon, W.F., and Clark, R.J., 2020. Evidence for a concealed Midcontinent Rift-related northeast Iowa intrusive complex. Precambrian Research, v. 347. McCormick, K.A., Chamberlain, K.R., and Paterson, C.J., 2018. U-Pb baddeleyite crystallization age for a Corson diabase intrusion: possible Midcontinent Rift magmatism in eastern South Dakota. Canadian Journal of Earth Science, v. 55: 111-117. 25 �Transpressional Nature of the Keweenaw Fault System, Lake Superior Region, and Its Relationship to Grenville Orogenesis DeGRAFF, James 1, GAMET, Nolan 2, LANGFIELD, Katherine 1, LIZZADROMcPHERSON, Daniel 1, MUELLER, Sophie 3, and TYRRELL, Colin 4 1 Michigan Technological University, Houghton, MI 49931 Michigan Geological Survey, Marquette, MI 49855 3 Nevada Gold Mines, Elko, NV 89801 4 Self-empoyed, Mass City, MI 49948 2 The Keweenaw fault system (KFS) is a connected set of faults that extends along the southern margin of the Midcontinent Rift System from northwest Wisconsin to near Keweenaw Point in Michigan. A component of reverse slip has thrust Portage Lake Volcanics (PLV, 1.1 Ga) southeastward over younger, mostly flat-lying Jacobsville Sandstone (JS) on some faults in the system (Fig. 1). This motion enhanced a regional northwesterly tilt to PLV strata, produced counter-regional tilt near major fault segments, and locally tilted footwall JS strata to vertical and overturned attitudes (1, 2). Regionally, the KFS azimuth changes by 65° from 35° near Houghton to 100° at Big Bay, as does the strike of PLV layers. Locally, the Keweenaw fault on published maps changes azimuth by up to 85° at unusual bends, some of which have been attributed to offsets on transverse faults. These changes in fault azimuth are important clues to the geometry of faults making up the KFS and to their individual and collective slip behavior. If opposing rock masses across the KFS are relatively rigid, a reasonable assumption, the fault system cannot be pure dip slip everywhere along its curved path, which inference also applies to its component faults. Mapping along the KFS since 2017 reveals that the sinuous, mostly single fault trace on published maps oversimplifies important structural relationships. In any part of the system, three directional fault sets are recognized: (1) a dominant set that defines the KFS trend and locally separates steeper dipping PLV layers to the northwest from shallower dipping PLV layers to the southeast; (2) splay faults angled 15-30° clockwise from set 1; and (3) connector faults angled 3575° counterclockwise from set 1 that join footwall splays to the main fault trend (Fig. 1). The three fault sets maintain these angular relationships as the curved KFS changes direction from near Houghton to the tip of the peninsula. Interconnections between faults define fault-bounded blocks with long dimensions roughly parallel to the local KFS trend. The fault-bounded blocks and footwall splay faults defining their southeastern and southern edges are arrayed in a left-stepping pattern along the KFS, suggesting a component of right-lateral strike slip. Analysis of fault-slip data, i.e. slickenlines and slip-sense indicators, from the fault population along the KFS indicates that the system’s slip characteristics change between Houghton and the tip of the Keweenaw Peninsula. Near Houghton where the KFS trends northeasterly, the ratio of strike slip to dip slip is about 1:1 and is bimodal, whereas the ratio is more than 2:1 and is unimodal near Bête Grise Bay where the KFS trends easterly. Geologic relationships across some faults are consistent with their northwest and north sides sliding to the right and upward relative to opposing sides. Inversion of fault-slip data indicates that a strike-slip regime existed near Keweenaw Point with an azimuth of maximum shortening of about 100°, which favors a component of right-lateral slip on the KFS. Folding within and adjacent to the fault-bounded blocks exhibits two styles. Multiple folds subparallel to shorter northeast and east ends of fault-bounded blocks (i.e. NE- to N-trending axes) formed by shortening across such boundaries. In contrast, single folds subparallel to longer sides of the blocks in footwall JS strata (i.e. NE- to ESE-trending axes) formed by drape 26 �of JS strata over steeply dipping, mostly strike-slip faults with little to no shortening across them. Based on this evidence, we infer that oblique slip on the KFS becomes mostly right-lateral strike slip near Keweenaw Point and that crustal shortening is along a line roughly perpendicular to the Grenville front about 550 kilometers to the east-southeast (Fig. 1). Acknowledgements: Funding was provided by the USGS EDMAP program, matched by MTU’s Department of Geology and Mining Engineering and Sciences, and supplemented by grants from the Michigan Space Grant Consortium, Keweenaw Community Forest Company, and the ILSG. We thank the Michigan Geological Survey for its sponsorship and G. Hubbell, I. Gannon, G. Hemmila, G. Ahrendt, J. Hawes, B. Murphy, B. Heusdens, and D. Breen for fieldwork assistance. Figure 1: Keweenaw fault system (black lines) north of Portage Lake, Michigan. Five largest fault-bounded blocks numbered 1 – 5. Black arrows show inferred maximum shortening direction. Inset map modified from Northwestern University maps online (https://www.earth.northwestern.edu/spree/Maps.html). References 1. 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. 2. 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, v. 135, no. 1/2, p. 449–466. 27 �Glimpses of a Paleoproterozoic landscape: Analysis of exhumed topography on Archean basement rocks northwest of Marquette, Michigan EDIGER, Trent and BJØRNERUD, Marcia Geosciences Department, Lawrence University, Appleton Wisconsin 54911 Background and purpose: Northwest of Marquette, MI, between the Yellow Dog River on the north, and Silver Lake and the Little Garlic River on the south, the modern land surface lies close to the nonconformity between an Archean granite-greenstone complex (the ~2.7 Ga Compeau Creek Gneiss and Mona Schist) and Paleoproterozoic metasedimentary rocks (the ~1.85 Ga Michigamme Formation). High areas are underlain by Archean rocks, while lower ones are underlain by the Michigamme Fm., suggesting that Michigamme sediments accumulated on an ancient land surface of hills and valleys with up to 70 m of relief. Although the Archean and Proterozoic rocks are locally in fault contact, good exposures of the nonconformity, together with systematic fining of grain size in the Michigamme with distance from Archean highs, support the interpretation that much of this area is an ancient exhumed landscape. Further evidence for a regional Paleoproterozoic landscape with significant relief comes from observations that the Sudbury ejecta layer in the region occurs at a wide range of stratigraphic heights above the base of the Michigamme Formation, and locally directly on Archean rocks (Cannon et al., 2010). The modern topography of this region is also qualitatively different - more rugged – than that seen on Archean rocks exposed only 25 km to the south, along the Black River east of Republic, even though the two areas share the same glacial history. In the study area, we suspect that the primary effect of glacial erosion was to remove the soft Michigamme Fm. from high spots, reexposing the sub-Michigamme surface. In the southern area, either the Michigamme Formation was never deposited or the pre-glacial landscape had been already been eroded to a level below the nonconformity. We believe, therefore, that the study area preserves a low-fidelity version of a Paleoproterozoic landscape and can provide insight into patterns of erosion and weathering at a time before land plants and modern atmospheric conditions. On an Earth with no vegetation and little to no soil, eroded sediments would have had a shorter residence time on landscapes and in river systems. Bedrock rivers would have been more common than they are today, and the primary mechanisms of landscape evolution would have been corrosion (chemical weathering) in a CO 2rich atmosphere, corrasion (abrasion of bedrock by entrained sediment), and cavitation (pitting of bedrock surfaces by bubble implosion in turbulent waters). Faults, joints, and other bedrock features would have been the primary influences on river channel location and potholes would have played an important role in channel development (Wohl, 1998). The goal of the study was to develop quantitative metrics to characterize the exhumed ancient landscape and contrast these with modern topography in areas with similar bedrock in order to gain a better understanding of geomorphologic processes in Paleoproterozoic time. Methods: The boundaries of the 165 km2 study area -- the extent of the exposed Archean nonconformity surface -- were drawn based on the provisional geologic map by Klasner et al. (1979) in combination with field observations and visual assessment of the topography. The rocks in the southern comparison area along the Black River are not a granite-greenstone complex like those in the study area, but they do include a mix of felsic and mafic lithologies (Archean gneisses and Paleoproterozoic dikes; Cannon, 1975) and thus serve as a reasonable analog. In order to understand the role of climate in generating relief in granite-greenstone complexes, we also analyzed the topography of two other granite-greenstone terranes: the Pilbara craton in the desert of northwestern Australia (ca. 3.5-3.2 Ga) and the Umburanas complex in the rainforest of Bahia 28 �Province, Brazil (ca. 3.4-3.1 Ga). In both areas, the bedrock lies close to the surface and the regions have been tectonically stable since at least Mesoproterozoic time. Digital Elevation Models (DEMs) for the study site and comparison site came from LiDAR data collected by the USGS 3D Elevation Program and were accessed via OpenTopography’s data map. DEMs for the Bahia province in Brazil and the Pilbara craton in Australia were generated by the Shuttle Radio Topography Mission and accessed through USGS EarthExplorer, and Geoscience Australia’s data map, respectively. When necessary, DEMs were merged into a single feature layer in Esri ArcGIS Pro 3.2.0 and clipped to the areas of interest. Roughness visualizations were calculated by determining the difference between the highest and lowest elevation cell in each 3x3 rectangular pixel neighborhood (Wilson et al. 2007). In ArcGIS Pro, focal statistics (statistical operations on each pixel based on specified neighboring pixels), were used to generate maximum and minimum elevation rasters. These intermediary rasters were then used to quantify roughness and create visualizations using the raster calculator tool. Results: By several measures, the topography of the study area is significantly more rugged than that of both the Black River comparison site and the Pilbara Craton. The roughest 90 m2 parcel in the study area has 73 m of relief compared with 27 m for the Black River and 46 m for the Pilbara. The study site also has a greater percent of land area in the highest roughness classes (>16 m of relief within 90 m2 parcels). The Black River area and Pilbara craton are surprisingly similar in the distribution of elevations, despite representing very different erosional conditions (glacial scouring and desert exposure, respectively). However, along the Black River, lithology seems to have little control on topography while in the Pilbara craton, contacts between Archean batholiths and volcanogenic sediments are the roughest areas. The Umbaranas site is much rougher than the other three, with up to 392 m of local relief where greenstones are exposed, possibly reflecting the wide range of volcanic and sedimentary lithologies in the Umburanas complex (Barbosa & Sabaté, 2002). In the study site NW of Marquette, topographic roughness is concentrated along linear zones – presumably erosion-enhanced faults or fractures in the Archean bedrock. These straight paleochannels differ from the meandering shapes typical of alluvial (sediment-dominated) river systems. Moreover, some of these channels have scalloped edges, a possible record of their evolution through linkage of bedrock potholes (Wohl, 1998). In summary, the topography of the study area differs not only from that of the nearby Black River site, which shared the same recent glacial history, but also from both the desert and rainforest sites. We suggest, therefore, that the region northwest of Marquette represents a Paleoproterozoic bedrock landscape that may have developed under warm, wet conditions in the absence of vegetation, a combination that does not occur on Earth today. References cited Barbosa, J., & Sabaté, P., 2002. Geological features and Paleoproterozoic collision of four crustal segments, Sao Francisco craton, Anais Da Academia Brasileira de Ciências, 74, 343–359. Cannon, W.F., Schulz, K., Horton, J., & Kring, D., 2010. The Sudbury ejecta layer in the Paleoproterozoic iron ranges of northern Michigan, USA. GSA Bulletin, 122, 50-75. Klasner, J., Cannon, W.F., & Brock, M., 1979. Bedrock geologic map of Baraga, Dead River and Clark Creek basins, Marquette County Michigan, USGS Open File Report 79-1305. Cannon, W.F., 1975. Bedrock geologic map of the Republic Quadrangle, Marquette County, Michigan. USGS Miscellaneous Investigations Series Map I-862. Wohl, E. 1998. Bedrock channel morphology. Rivers Over Rock. AGU Monograph 107, 133-149. Wilson, M., OConnell, B., Brown, C., Guinan, J., & Grehan, A, 2007. Multiscale terrain analysis of multibeam bathymetry data. Marine Geodesy, 30, 3–35. 29 �MCR Synthesis 1. Characterizing the MCR mantle plume GOOD, David Department of Earth Sciences, University of Western Ontario, 1151 Richmond Street, London, ON N6A 5B7, Canada The Midcontinent Rift (Keweenawan Large Igneous Province) contains the most diverse assemblage of mafic rock types for any LIP on earth with 9 distinct basalt groups and more than 15 major Ni-Cu-PGE occurrences or deposits. The main objective of the MCR synthesis is to build a coherent model to explain the vast array of observations, geochemical data and interpretations presented by numerous researchers over the past four decades. The project is subdivided into 4 related objectives: 1) Recognition of the key geochemical features of global rift/LIP settings that we should see in the MCR; 2) Build a classification scheme for all basalts, gabbros and ultramafic rocks using high-precision incompatible trace elements; 3) Apply analytical tools and modelling to unravel petrogenesis of recognized groupings; and 4) Represent the results in a model for the MCR that highlights spatial and temporal relationships in the rift. This project was inspired by several key events over the past decade, each of which indicate the project is feasible at this time. Proof of concept tests for objectives 2, 3 and 4 were presented in 2023 indicating a high degree of confidence for the success of this 4-to-5-year project. A few researchers have identified various stratigraphic units in the MCR to have originated by partial melting in the mantle plume and used their inherent isotopic or trace element compositions to model melt-crustal interaction and the petrogenesis of various intrusions that host Cu-Ni-PGE deposits. These plume-related basalt units include the lower Siemens Creek and Kallander Creek basalt in Michigan and the lower Osler and Mamainse Point basalt in Ontario. But these units each present slightly different trace element characteristics, so the question arises as to what criteria are useful for distinguishing mantle plume magmas from those generated in the upper mantle or at different depths within the plume. The main criteria for identifying plume magmas are based on ocean island basalt-like trace element characteristics. In this study, two MCR plume basalt types are identified (Groups 1 and 5) using the combination of λ1-λ2 REE coefficients, TiO2/Yb, and Gd/Yb diagrams. The differences between groups 1 and 5 are best explained by partial melting at different depths, based on differences between majorite (>~300 km) and pyrope garnet fractionation, respectively. Group 1 includes basalts from the Lower Osler and lower Kallander creek groups and the highly fractionated Devon volcanic unit. Group 5 includes basalts from the lower series A and B units at Mamainse Point, Central Osler Volcanic Group and the lower Siemens Creek basalt located at the Skinny, Bluff and Bond Falls sites in Michigan. A well-understood and fundamental characteristic of highly incompatible trace element ratios is their use to correlate basalt and intrusive rocks. These typically unique trace element signatures can be used in a manner like finger printing. However, in all cases, care must be taken, particularly for TiO2, to evaluate clinopyroxene or spinel fractionation, as is the case for basalt and intrusions in Group 1. Based on these comparisons, the Bovine, Current Lake, Disraeli, Haystack, Hele, Kitto, Riverdale Sill, Seagull, Shillabeer, and Thunder Intrusions belong to Group 1, whereas the McIntyre, Jackfish, and Logan sills belong to Group 5. 30 �Description and application of the Consolidated Minerals Database to support geological investigations: an example from the Cuyuna Range, central Minnesota GORDEE, Sarah 1, RIAN, Madison 1, SAARI, Stacy 1, and CARTER, Matthew1 1 Minnesota Department of Natural Resources, Division of Lands and Minerals, 1525 3rd Ave E, Hibbing, MN 55746 Over the past ~50 years, the Minnesota Department of Natural Resources (DNR) Lands and Minerals Division has amassed numerous collections of mineral exploration-related documents, amounting to well over 10,000 hardcopy materials containing geoscience and related land data. Curating these collections has proven to be challenging given the sheer volume of documents from multitudinous sources, necessitating a concerted solution to manage these materials. The Consolidated Minerals Database (CMD) is under development by the DNR to support the initiative to bring the agency’s collections of historical and contemporary documents into digital format and to make them readily available for public use. It is a database of unique collections containing cross-referenced documents with linkages to other internal and external databases, including the Hibbing Drill Core Library (DCL) database, and a web map, where geospatially linked documents can be retrieved from specific localities or regions. The various collections comprising the CMD are designated by project, company, or other relevant shared interest(s). Documents are individually entered into a particular collection and assigned a unique numerical identifier, which is used to cross-reference to different databases. Metadata (e.g., title, date, source) are recorded in a series of entries, and attributes of the document (e.g., scope, subject, content, methods, materials, discipline) are classified in a series of dropdown menus, enabling users to search and find documents meeting specific criteria relating to documents’ contents and origin. In the current initiative, the DNR utilizes the CMD intake application to produce digital records of documents from the Cuyuna Range in central Minnesota, where exploration and mining for iron and manganese ores was active throughout much of the early-middle 20th century. The objective is to curate mineral exploration documents and compile geological data from these records in a large database. A synthesis of these data will help to better understand the geological architecture and extent of historical exploration in the region, and the compiled datasets will help to evaluate the potential for additional iron, manganese, and other resources. Historical documents from the Cuyuna mining district are stored in the Hibbing Lands and Minerals office. Dozens of different exploration companies drilled at least 12,000 boreholes and created thousands of documents spanning multiple decades of mineral exploration and mining in this district. Relevant documents in this collection range from 1905 to the 1970s, and include geological maps, surface maps with drillhole collars and associated metadata, mine maps (surface, subsurface, infrastructure), tables with geochemical data, drillhole profiles with tabular geochemical data, geological information and interpretations, geological cross-sections, field notes, and notes and correspondences regarding property ownership, exploration results and resource estimates. Because of the number of companies involved and diversity in the presentation of data it is necessary to address certain challenges before curation into CMD. Documents are first sorted by company and locality in the Public Land Survey System, which allows for the identification and removal of duplicate maps and other documents shared 31 �among and across different companies. Once sorted, all relevant documents with clear datasets and sufficient metadata to identify the source and locality are curated into the CMD. Following curation, plan-view maps are converted to picture format and brought into ArcGIS, where they are spatially located using georeferencing methods. This method helps to resolve problematic drillhole locations, and to identify less obvious duplicate documents and datasets, including drillholes that were renamed over time as operators changed hands. Drillhole collar locations can then be added to a database of known drillholes in the region, and integrated and compared to cores from the DCL database. Once correctly positioned, individual geospatial datasets, such as geological logs, and geochemical and geophysical data, are extracted from each document. Extracting and compiling geologic data such as geologic logs and assays into tabular format has been a challenging endeavor. These data were hand-written or typed using a typewriter, and utilizing optical character recognition technology to extract text is not straightforward. However, transcribing the data by hand is a cumbersome and protracted process, and potentially introduces errors that must be checked for quality assurance. With the advance of artificial intelligence (AI) and machine learning, it is now possible to train an AI model to extract data into tables. Training the AI data extraction model is an efficient process. First, pages (10 minimum) containing example data listed in an internally consistent format are imported; then table(s) are delineated using the associated headers, labels, and rows per each page. Once the model is trained, numerous documents or pages of the same format are uploaded and the data are auto-extracted, and the output reflects the model’s specified number of tables, columns, and rows. Before the outputted data are extracted, they are enhanced within the model workflow to account for spelling errors, incorrect symbols, etc., so that it is unnecessary to resolve errors individually by hand. Once a model is trained for a specific table format, thousands of pages of tabular data can be extracted into a tabular database en masse in a matter of minutes. Using MircoMine modeling software, the tabulated data are visualized in a 3D model, where any remaining tabulation errors are identified and corrected. This process is ongoing, as many still-uncurated documents remain in the Cuyuna collection. To date, nearly 2,000 documents totalling over 20,000 pages have been scanned, georeferenced and lodged in the CMD, and the drillhole database contains over 4,000 individual drillholes with assay data totaling over 60,000 lines. Incompleteness notwithstanding, the current database is a growing and ever-refining, data compilation from thousands of geological investigations. Together, the newly compiled data are sufficiently expansive to make new observations and interpretations pertaining to the geology and distribution and style of iron and manganese resources, as well as the potential for other base and precious metal resources, in the Cuyuna Range. 32 �Revisiting geophysical interpretations of the Midcontinent Rift below Lake Superior— Insights from GLIMPCE seismic-reflection line C GRAUCH, V.J.S.1, HELLER, S.J.2, WOODRUFF, Laurel G.3, and STEWART, Esther K.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 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN 55112 4 Wisconsin Geological & Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 2 The 1.1 Ga Midcontinent Rift System (MRS) has been investigated in the Lake Superior region for more than a century, driven by mineral exploration, academic study, and, for a brief time, oil and gas exploration. Limited outcrops on land and the extent of MRS rocks under the lake have motivated many workers to use geophysical methods to investigate the nature and extent of the rift. The most influential geophysical data for modern paradigms has come from seismic-reflection profiles collected by the Great Lakes International Multidisciplinary Program on Crustal Evolution (GLIMPCE) in the late 1980s. Notably, many previous workers have used interpretations of GLIMPCE line C (Fig. 1, inset) to demonstrate the architecture of the MRS in western Lake Superior. A recurring theme from the previous work is that syn-extensional basalt flows accumulated in half-grabens bounded by normal growth faults, which then reactivated as reverse faults in response to later compression (e.g., Cannon et al., 1989; Hinze et al., 1992; Dickas and Mudrey, 1997; Stein et al., 2015). We are revisiting GLIMPCE line C by constructing a detailed velocity model for conversion of the seismic data measured in two-way travel time to a section plotted versus depth (Grauch et al., 2023). This approach allows for digital verification of the modeled velocities and more accurate depiction of thicknesses and dips of units to tie to geology onshore. We have constructed an analogous gravity model along line C that provides independent evaluation of our velocity model using velocity-density relations developed from analysis of region-wide rock property compilations (Grauch, 2023). Preliminary results from the velocity modeling, depth conversion, and ties to onshore geology have led to a significantly different view of Line C as primarily a sag basin rather than a half-graben, showing both syn- and post-magmatic subsidence (Fig. 1; Grauch et al., 2023). Narrow intervals of high velocities, which indicate a composition of gabbro (Grauch, 2023), emanate upwards along both sides of the sag basin from an inferred mantle bulge. The intervals are associated with strong linear reflections that truncate sub-horizontal layers in the sag basin and may obscure any minor faulting that occurred before or after intrusion. Cross-cutting mafic intrusions provide an alternate explanation for the termination of layers that was previously thought to indicate major faulting. This new view of line C implies that basin subsidence was the dominant process in the development of rift stage troughs rather than major half-graben structures. Other important interpretations include the following. • Portage Lake Volcanics show syn-magmatic basin subsidence • The Lower Oronto Group section shows post-magmatic basin subsidence • Onlap of Upper Oronto Group onto tilted Porcupine Volcanics suggest the deformation pre-dated deposition of Oronto Group sediments 33 �• Rocks of the lower northeast sequence of the North Shore Group may connect to rocks of similar age from the south shore that lie underneath the sag basin. Figure 1. Interpreted depth section for GLIMPCE Line C. No vertical exaggeration. NSVG, North Shore Volcanic Group. PLV, Portage Lake Volcanics. The new rendition of the Line C seismic data also raises several questions. • How do the sedimentary sections correlate from north to south? • What caused the truncation of volcanic layers at the volcanic-sedimentary contact in the middle of the seismic section and was reverse faulting involved? • What is the tectonic process that drove the syn-magmatic subsidence? • Where does the reverse Keweenaw fault extend into the section from the south shore and what was its influence? These and other questions can be addressed through the construction of velocity models and depth conversions of other seismic lines in the lake. Future insights will benefit from a three-dimensional view that these additional seismic lines will provide. Cannon, W.F., Green, A.C., 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.H., and Spencer, C., 1989, The North American Midcontinent rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332., https://doi.org/10.1029/TC008i002p00305 . Dickas, A.B., and Mudrey, M.G., Jr., 1997, Segmented structure of the Middle Proterozoic Midcontinent System, North America, in R.J. Ojakangas, A.B. Dickas, and J.C. Green (eds.), Middle Proterozoic to Cambrian Rifting, Central North America: Geological Society of America Special Paper 312, 37-46., https://doi.org/10.1130/08137-2312-4.37 . Grauch, V.J.S., 2023, Compressional wave seismic velocity, bulk density, and their empirical relations for geophysical modeling of the Midcontinent Rift System in the Lake Superior region: U.S. Geological Survey Scientific Investigations Report 2023-5061, 60 p., https://doi.org/10.3133/sir20235061. 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.), 69th Annual Institute on Lake Superior Geology Proceedings—Part 1, Program and Abstracts, p. 37-38. Hinze, W. J., Allen, D. J., Fox, A. J., Sunwood, D., Woelk, T., and Green, A. G., 1992, Geophysical investigations and crustal structure of the North American Midcontinent Rift system: Tectonophysics, v. 213, p. 17-32. 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, p. 1607-1616. 34 �GeologyOntario: a powerful search tool for Ontario explorationists HINZ, Sheree1 Ontario Geological Survey, 435 James Street South, Thunder Bay, ON, P7E 6S7 Canada W ith dr aw n In March of 2023, the Ministry of Mines released a new online platform to search and query Ontario Geological Survey data. The Ontario Geological Survey maintains and provides public access to a wealth of geological information including maps, publications, assessment files, mineral inventory points, miscellaneous data releases, geophysical data, abandoned mine site information, and more. Though this information has been available online for many years, the previous iteration of GeologyOntario had significant constraints, including a lack of spatial search abilities, and users experienced challenges in finding relevant information. The new GeologyOntario consists of separate text (Figure 1) and spatial search (Figures 2 and 3) tools which provides ample opportunities to discover information. Geological information is often dependent on spatial data, and the new spatial search tool runs on a powerful Esri-based system, allowing clients to build queries to focus on the types of data relevant to their interests, within the geographical areas they are working. The new GeologyOntario Search Hub is located at https://geology-ontario-en-mndm.hub.arcgis.com/. Figure 1. GeologyOntario text search page (https://www.geologyontario.mines.gov.on.ca/). 35 �aw n dr W ith Figure 2. GeologyOntario spatial search page showing regional geology, mineral inventory, and the results of a search for mineral inventory points listing lithium as a primary commodity in the area of the Separation Rapids pluton(https://mndm.maps.arcgis.com/apps/webappviewer/index.html?id=66ee0efe4d3c4816963737dbdb 890708). Figure 3. GeologyOntario spatial search page with regional geology, mineral inventory, assessment files, Resident Geologist Program (RGP) site visits, and exploration activity layers active. 36 �Recent developments on the use of the Horizontal-to-Vertical Spectral Ratio (HVSR) passive seismic method to determine depth to bedrock in Minnesota HIRSCH, Aaron C.1 1 Minnesota Geological Survey, University of Minnesota, 2609 Territorial Road, St. Paul MN 55114 Bedrock depth is an important dataset for water resource management, hydrological studies, mineral exploration, and general well planning. In Minnesota, bedrock depth is highly variable; thin to nonexistent in the northeast, up to 250m+ in areas to the west, and irregular elsewhere. In areas where bedrock depth is not known from existing water, exploration, or scientific drilling, various geophysical techniques can be used. One of these methods is the Horizontal-to-Vertical Spectral Ratio (HVSR) (Nogoshi and Igarashi, 1971; Nakamura, 1989) which utilizes horizontal ambient noise surface wave frequencies that are excited and amplified dependent on the depth to the basement bedrock below a less dense and seismically slower velocity upper layer (i.e. unconsolidated glacial sediments). The HVSR method has been utilized in Minnesota to estimate the depth to bedrock since the late 2000’s (Chandler and Lively, 2014) and has become a standard measurement in the MN County Geological Atlas program (e.g. Bauer et al., 2023; Mayer et al., 2023). The initial HVSR dataset used 1647 passive seismic measurements with 303 locations with a known bedrock depth, also known as control points, to develop parameters to accurately estimate the depth to bedrock across Minnesota (Chandler and Lively, 2016). The Minnesota Geological Survey has now collected a total of over 6000 HVSR measurements and 480 control points resulting in a new assessment from the larger and more geographically and geologically widespread dataset. Analyses has included a new quantitative data quality ranking using international HVSR guidelines (SESAME, 2004) and new control parameters have been investigated. The shape of the HVSR curve is now being captured in a passive seismic database due to its relationship with bedrock depth topography, bedrock weathering, and the underlying velocity structure. Ongoing evaluation of this database will help refine the HVSR depth to bedrock estimation and more accurately identify potential bedrock valleys while future work will include measuring densities and ultrasonic velocities of Quaternary cores to constrain the control point parameters more accurately. References Bauer, E. J., Cicha, J., Radakovich Block, A., Jirsa, M. A., Hirsch, A. C., Meyer, G. N., Scott, S. B., Lively, R. S.. (2023). C-55, Geologic Atlas of Otter Tail County, Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/256920. Chandler, V. W., Lively, R. S., 2016, Utility of the horizontal-to-vertical spectral ratio passive seismic method for estimating thickness of Quaternary sediments in Minnesota and adjacent parts of Wisconsin, Interpretation, Vol. 4, No. 3, p. SH71-SH90. http://dx.doi.org/10.1190/INT-20150212.1. Chandler, V.W., and Lively, R.S., 2014. OFR14-01, 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. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/162792. Mayer, J. A., Bradley, M. C., Retzler, A. J., Severson, A. R., Jirsa, M. A., Chandler, V.W., Conrad, D. R., Gowan, A. S., Radakovich Block, A., and Hamilton, J. D., 2023. C-58, Geologic Atlas of Lincoln 37 �County, Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/260212. Nakamura, Y., 1989, A method for dynamic characteristics estimation of subsurface using microtremor on the ground surface: Quarterly Report Railway Technical Research Institute, 25–30. Nogoshi, M., and Igarashi, T., 1971. On the amplitude characteristics of microtremor (part 2) (in Japanese with English abstract): Journal of the Seismological Society of Japan, 24, 26–40. SESAME, 2004. Guidelines for the implementation of the H/V spectral ratio technique on ambient vibrations. Measurements, processing, and interpretation: WP12 European commission — Research general directorate project no. EVG1-CT-2000-0026 SESAME, report D23.12, 62, http://www.gripweb.org/gripweb/sites/default/files/HV_User_Guidelines.pdf. 38 �Lithostratigraphic discrimination of Quaternary core in Minnesota using magnetic susceptibility HIRSCH, Aaron C.1, SCHNEIDER, Emma, L.1 1 Minnesota Geological Survey, University of Minnesota, 2609 Territorial Road, St. Paul MN 55114 Most of Minnesota is covered by Quaternary sediments deposited during multiple glaciation events. This Quaternary stratigraphy is highly complex due to multiple glacial events in which ice lobes emanated from differing locations north of Minnesota and deposited sediment (diamict, till) of variable thicknesses (up to 250m), provenance, and morphology across the state (Johnson et al., 2016). The Minnesota Geological Survey uses grain counts, color, sedimentary structures, and composition to establish Quaternary lithostratigraphic units that distinguish these deposits by lithology, stratigraphy, and geomorphology. Nine lithostratigraphic regions were identified using cuttings, outcrops, and rotary sonic core (Johnson et al., 2016). Magnetic susceptibility measurements were taken at 1-2 meter intervals from many of these rotary sonic cores during core analysis by applying a magnetic field to the core and recording the magnetic response. This study was conducted to determine if magnetic susceptibility measurements from cores can aid unit correlation across Minnesota regions as part of a USGS funded data preservation project. Over 11,000 measurements were recorded in a newly established Quaternary magnetic susceptibility database with over 7,000 measurements assigned to a lithostratigraphic formation and unit interpretation. Magnetic susceptibility logs were generated for each measured core and statistics calculated for each unit. Analysis of this database has identified lithostratigraphic units with distinctive magnetic susceptibility ranges as compared to nearby and similarly aged units. Due to these results, this newly established database functions as another tool for lithostratigraphic identification of Quaternary sediments and local and regional correlations. References Johnson, Mark D., Adams, Roberta S., Gowan, Angela S., Harris, Kenneth L., Hobbs, Howard C., Jennings, Carrie E., Knaeble, Alan R., Lusardi, Barbara A., and Meyer, Gary N., 2016. RI-68 Quaternary Lithostratigraphic Units of Minnesota. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, https://hdl.handle.net/11299/177675 39 �New Insights into the Geology and Geochemistry of the Osler Group and Related Rocks, Midcontinent Rift System, Northern Lake Superior, Ontario HOLLINGS, Pete and SMYK, Mark Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Ongoing geological reconnaissance and lithogeochemical sampling were undertaken on parts of the Black Bay Peninsula, St. Ignace Island and neighbouring islands in 2022 and 2023. New field and geochemical data have helped to both distinguish lithostratigraphic units and suggest common magmatic histories in developing a model for the Midcontinent Rift System (MRS)related Osler Volcanic Group and related intrusive rocks. The Osler Group (1108-1105 Ma), a ~3 km-thick succession of predominantly basaltic flows and clastic sedimentary rocks on the north shore of Lake Superior, represents some of the earliest MRS magmatism. Previous studies have largely focused on the paleomagnetism (e.g. SwansonHysell et al., 2019; Halls, 1974) and geochemistry (e.g. Hollings et al., 2007; Keays and Lightfoot, 2015) of the flood basalts in developing a stratigraphic sequence. However, only basic mapping and some initial studies had been conducted on felsic extrusive and intrusive rocks (e.g. St. Ignace Island Complex, Hollings et al., 2023, Smyk et al., 2006; geochronology, SwansonHysell et al., 2019 and references therein). Sampling efforts were most recently focused on these felsic rocks in order to determine their geochemical affinity and to suggest how they, and related mafic igneous rocks, may fit in with the provisional tectono-magmatic model for the Osler Group. Felsic rocks include the Agate Point rhyolite flows (1105.15 Ma); thin felsic fragmental units; aphanitic felsite; and massive, subvolcanic(?), quartz-feldspar-phyric rocks (aka “quartz-feldspar porphyry”/QFP). These red, brown or gray rocks occur predominantly on St. Ignace Island (including in the core of the St. Ignace Island Complex (SIC)) and on smaller islands, south to Agate Point. Rhyolites tend to be LREE-enriched (La/Smn= 5.20-5.47), have higher total REE than the QFPs and more pronounced negative Ti anomalies. The majority of QFPs, including those in the core of the SIC, tend to display a coherent, tightly grouped REE trend, characterized by moderate LREE enrichment (La/Smn= 2.34-4.71, averaging 3.50), relatively flat HREE patterns and pronounced negative Nb anomalies. This similarity in the REE distribution of both extrusive and subvolcanic felsic rocks suggests that they may share a common magmatic and fractionation history. In contrast, the basaltic flows into which felsic rocks have been emplaced have flatter REE distribution patterns (La/Smn= 1.61-3.51) than those of the felsic rocks, with less-pronounced negative Ti anomalies. Mafic and felsic flows situated above an unconformity/conglomerate at Bullers Bay, St. Ignace Island, display pronounced negative Nb anomalies whereas those below do not. This suggests that the lower flows are part of the more primitive Lower Formation of the Osler Group, while the flows above the unconformity resemble those of the more crustally contaminated Central Formation (cf. Keays and Lightfoot 2015; Hollings et al. 2007) as delineated on nearby Simpson Island. 40 �Gabbroic rocks occur at the margin of the SIC, in the Moss Lake Intrusion and as numerous diabase sills and dykes with various orientations which intrude the supracrustal rocks. SIC and Moss Lake gabbro samples display similar REE patterns, characterized by moderate LREE enrichment (La/Smn= 2.59-3.23), moderate negative Ti anomalies and pronounced negative Nb anomalies. By comparison, smaller, diabasic dykes and sills have relatively flat REE distribution and less-pronounced negative Ti anomalies. Prominent, regional-scale mafic dykes (i.e. McEachan, Shesheeb, Paps) display lower total REE and lack negative Sm anomalies. Hollings et al. (2023) suggested that the rocks of the SIC likely formed as the result of emplacement of a large mafic magma chamber at the base of the Osler volcanic pile that triggered partial melting to generate felsic end members which then ascended to shallower levels in the crust. The SIC QFPs are geochemically similar to both the massive, subvolcanic(?) QFPs elsewhere on St. Ignace Island and nearby islands, as well as to the rhyolites at Agate Point, suggesting a similar origin for all of these felsic rocks. REFERENCES Halls, H., 1974, A paleomagnetic reversal in the Osler Volcanic Group, northern Lake Superior: Canadian Journal of Earth Sciences, v. 11, p. 1200–1207, doi:10.1139/e74-113. Hollings, P., Fralick, P. and Cousens, B. 2007. 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, https://doi.org/10.1139/e06-084. Hollings, P., Hanley, J., Smyk, M., Heaman, L., Cousens, B., and Zajacz, Z. 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, Volume 64, Issue 6, June 2023, egad032, https://doi.org/10.1093/petrology/egad032. 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. Economic Geology, 110, 1235– 1267, https://doi.org/10.2113/econgeo.110.5.1235. Smyk, M.C., Hollings, P.N. and Heaman, L. 2006. Preliminary investigations of the petrology, geochemistry and geochronology of the St. Ignace Island Complex, Midcontinent Rift, northern Lake Superior, Ontario; 52nd Institute on Lake Superior Geology, Annual Meeting, Sault Ste. Marie, Ontario, May, 2006, Proceedings Volume 52, Part 1, p.61-62. 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 131 (5-6), 913-940. 41 �Use of Ambient Noise Tomography for Mineral Exploration in the Lake Superior Region HOLLIS, Dan1 1 Sisprobe SAS, 831 Pacific Street, #1A, Morro Bay, California, 93442 USA Abstract Ambient noise tomography (ANT) is a relatively new passive seismic tool used in mineral exploration. The method has been successfully used in the Lake Superior region in mapping subsurface structure and rock properties. This presentation will provide a brief introduction to the ANT method and review recent ANT work done in the Lake Superior region. Introduction Exploration for mineral resources uses a variety of geophysical methods to detect and delineate mineral deposits and systems in order to optimize core drilling programs: gravity, magnetics, active-source reflection seismic and electromagnetics methods to name a few. Ambient noise tomography is a relatively new seismic geophysical tool that has seen increasing use in the past couple of years for mineral resource exploration. ANT uses natural earth vibrations and human-generated seismic vibrations to image subsurface structure and map physical properties of the subsurface. Four ANT surveys for mineral resources in the Lake Superior region have been completed: three surveys within the Coldwell Complex near the town of Marathon, Ontario, and one in the Duluth Complex in northeastern Minnesota. Ambient Noise Tomography Method The Earth is constantly vibrating. For the ANT exploration method, useful vibrations, sometimes referred to as “ambient seismic noise”, are generated by hydrosphere-lithosphere interaction such as the oceanic microseism caused by swell (a similar microseism is caused by swell in Lake Superior and the other Great Lakes), and anthropogenic sources such as vehicular traffic, industrial sites, railroads, and other human activity. An ANT survey uses continuous passive seismic data collected with an array of nodal seismometers (“nodes”) and uses surface waves to image the subsurface. The data recording duration is usually between 1 to 4 weeks. Surface waves are dispersive with different frequencies propagating at different velocities related to the seismic velocities of subsurface lithology. Frequency-velocity dispersion curves are picked for all receiver pairs. These dispersion curves serve as input for a tomographic process resulting in an array of frequency-velocity points for each cell in a grid. All cells within the grid are inverted to produce depth-velocity profiles and the result is a 3D shear wave velocity (Vs) cube where velocity is the seismic velocity of the lithology and subsurface structure interpreted from the velocity model. Case Studies Coldwell Complex, Marathon Area, Ontario The first Marathon ANT survey was collected in October 2017 in the Marathon area over a VMS target. This first survey was intended as a noise test using 31one-component (vertical) 10 Hz nodes. The purpose of the noise test was to characterize the spectral power, temporal variation and azimuthal distribution of the local ambient noise. The collected data was also used to cross-correlate all receiver pairs to assess the signal-to-noise ratio of surface waves in the virtual source gathers and processed to produce a crude 3D velocity model which showed agreement with available core hole data. The positive results of the noise test led to the go-ahead for an expanded survey over the target. 42 �An 8.6 km2 expanded ANT survey over the Marathon target was collected in November-December 2017 using 91 one-component (vertical) 10 Hz nodes. The resulting 3D velocity volume had usable imaging down to 1500 meters and imaged the gabbro intrusion slab target. Details about the noise test and expanded ANT survey and its interpretation can be found in Hollis et al. In July-August 2018, the Sally ANT survey was conducted over an exploration target several kilometers to the northwest of the Marathon surveys. The Sally survey was collected using 196 three-component 5 Hz nodes. This survey demonstrated the effectiveness of using three-component data to produce a more accurate velocity model. Details of the Sally survey can be found in Lavoué et al. With the good results of the expanded Marathon survey, funding was obtained through the European Union Horizon 2020 program to conduct larger, higher resolution survey again over the Marathon target in order to test the limits of the ANT method and to test other potential ANT analyses and passive seismic methods. This survey was acquired in September-October 2018 using 983 one-component (vertical) 10 Hz nodes. This third Marathon ANT survey generated several publications on its results some of which are listed in the Reference section. Duluth Complex, Northeast Minnesota In September 2023, a 32 km2 survey was collected over a helium exploration target in Lake County, Minnesota. This survey used 183 three-component seismic nodes. Logging of a post-survey confirmation well has helium shows between 533 – 671 meters which agrees with the interpreted reservoir depth range from the ANT 3D data (Pulsar Helium). Conclusion Past work in the Lake Superior region has shown the ambient noise tomography is an effective tool for mineral exploration in the area. References Dales, P., L. Pinzon-Ricon, F. Brenguier, P. Boué, N. Arndt, J. McBride, F. Lavoué, C. J. Bean, S. Beaupretre, R. Fayjaloun, et al. (2020). Virtual Sources of Body Waves from Noise Correlations in a Mineral Exploration Context, Seismological Research Letters XX, 1–9, doi: 10.1785/0220200023. Hollis D., McBride J., Good D., Arndt N., et al (2019). Ambient noise surface wave tomography at the Marathon PGM-Cu deposit, Ontario, Canada, CSEG Recorder, June 2019. Lavoué A., Nicholas Arndt, John McBride, Aurélien Mordret, Florent Brenguier, Pierre Boué, Roméo Courbis, Sophie Beauprêtre, Charles Beard, Dan Hollis, and Richard Lynch, (2020), Ambient noise Rayleigh and Love wave tomography beneath the Sally Palladium-Copper Deposit (Ontario, Canada), SEG Technical Program Expanded Abstracts : 2075-2079. Pulsar Helium, https://files.elfsightcdn.com/eafe4a4d-3436-495d-b748-5bdce62d911d/2f2bca29-47e64d88-851e-d97bbca643b5/Pulsar_corp_deck_20Mar24x_FINAL-compressed.pdf. Accessed 3/29/2024. Sharma H., Molnar S., Hollis D. and McBride J. (2018). Application of ambient-noise analysis and velocity modeling in mineral exploration. SEG Technical Program, Expanded Abstracts, 3072–3076. Teodor, Daniela & Beard, Charles & Pinzon-Rincon, Laura & Mordret, Aurelien & Lavoué, François & Beaupretre, Sophie & Boué, Pierre & Brenguier, Florent. (2021). High-frequency ambient noise surface wave tomography at the Marathon PGE-Cu deposit (Ontario, Canada). 10.5194/egusphere-egu21-13152. 43 �Geologic and tectonic implications of detrital zircon U-Pb ages from the Dickinson Group in the western Upper Peninsula of Michigan, USA JONES, James V.1, CANNON, William F.2, DRENTH, Benjamin J.3 and O’SULLIVAN, Paul4 1 U.S. Geological Survey, Anchorage, AK 99508, USA jvjones@usgs.gov U.S. Geological Survey, Reston, VA 20192, USA 3 U.S. Geological Survey, Denver, CO 80225, USA 4 GeoSep Services LLC, Moscow, ID 83843, USA 2 In the Lake Superior region of the northern United States and southern Canada, Paleoproterozoic metasedimentary successions record the breakup of southern Superia (in present coordinates) that began ca. 2.3 Ga and the eventual transition to long-lived accretionary orogenesis along the southern Laurentia margin ca. 1.90–1.85 Ga. These successions are difficult to correlate for reasons that include contrasts in thickness and facies at multiple scales, similarities in depositional environment through hundreds of millions of years of sedimentation, and polyphase tectonism that variably produced intense deformational and metamorphic overprints. Detrital zircon U-Pb geochronology is useful for correlating siliciclastic strata that are widespread throughout the successions and for identifying provenance patterns in space and time. We present new data for samples collected from ca. 2.3–1.8 Ga strata across the western Upper Peninsula of Michigan and northern Wisconsin that provide a baseline for regional geologic mapping and correlations with similar strata regionally to globally. Our findings provide new insights into stratigraphic relationships of the ca. 2.1 Ga Dickinson Group and require revision of the depositional history, tectonic evolution, and regional significance of the succession. The Dickinson Group is a distinctive succession of metasedimentary and metavolcanic rocks exposed only in the Felch trough area of the western Upper Peninsula. The strata are bounded by the Randville Dolomite of the Chocolay Group to the north and Archean banded gneiss to the south. These bounding contacts are mostly interpreted to be structural. The lowermost unit of the Dickinson Group is the East Branch Arkose, a coarse cobble to boulder conglomerate that contains rounded clasts of granite and quartzite in a matrix of feldspathic to lithic wacke. The conglomerate is moderately sorted and generally matrix-supported. At a few localities, the East Branch appears in unconformable contact with coarse-grained granite, interpreted as one of the ca. 2.6 Ga batholiths that are common in the southern Superior Province. At these basal localities, cobbles are strongly flattened and the entire unit contains a well-developed foliation defined by the flattened cobbles and aligned biotite in the sedimentary matrix. The East Branch Arkose is overlain by the Solberg Schist, the lower part of which contains fine-grained mafic schist and amphibolite together with discontinuous calc-silicate horizons up to 15 cm thick. Compositional layering in the Solberg is isoclinally folded with a consistent foliation defined by fine-grained chlorite and amphibole. The middle Solberg contains a ~100-ft-thick bed of iron-formation called the Skunk Creek Member that includes biotitehornblende schist and thinly bedded metachert with magnetite layers (James, 1958). The upper Solberg is made up of interlayered biotite quartzite, massive gray quartz-mica schist, and staurolite-biotite schist. The Solberg Schist is overlain by the Six-Mile Lake Amphibolite, which is made up of fine- to medium-grained amphibolite with a strong tectonic foliation defined by hornblende. As originally mapped, the Dickinson Group defines a subvertical, south-facing homocline and was previously thought to be Archean based on an inferred gradational contact between the Six-Mile Lake Amphibolite and the Archean banded gneiss (James, 1958; James et 44 �al., 1961). However, detrital zircon data published by Craddock et al. (2013) showed that the East Branch arkose was deposited ca. 2.1 Ga or later, thus implying a Paleoproterozoic age for the entire succession. Our detrital zircon U-Pb data from the East Branch Arkose match previously published data from Craddock et al. (2013) and are dominated by ca. 2.6 Ga grains interpreted to reflect local granitic sources that are also observed as cobbles. Rare, but statistically significant ca. 2.1 Ga populations, confirm the Paleoproterozoic maximum depositional age. Mafic Solberg Schist that overlies the arkose does not contain abundant zircon, although some small grains that were recovered show a mix of ages ranging from ca. 3.1 to 2.6 Ga and a small ca. 2.1 Ga population. Detrital zircon age spectra from upper Solberg exposures are distinctly different, though. Samples of biotite quartzite and staurolite schist both contain prominent ca. 1.86–1.84 Ga age populations together with more minor ca 2.5 and 2.3 Ga age populations. The upper Solberg age spectra closely match samples of the Michigamme Formation from throughout the surrounding region, suggesting that the upper siliciclastic component of the Solberg schist should, instead, be mapped as Baraga Group. This revised interpretation raises the possibility that the mafic volcanic rocks and iron formation of the lower and middle Solberg Schist could also correlate with the lower Baraga and(or) Menominee Groups, though additional data are required to test these possibilities. Furthermore, it raises questions about the age of the Six-Mile Lake Amphibolite, the uppermost unit of James’ (1958) Dickinson Group. We suggest that the Six-Mile Lake may be Archean as previously inferred by James (1958), in which case its concealed contact with the upper Solberg or Michigamme Formation would be tectonic rather than depositional. We are presently working to test this revised hypothesis through new geochronology and 40Ar/39Ar thermochronology across the contact. The actual depositional age of the East Branch Arkose remains uncertain, as it can be younger than the ca. 2.1 Ga detrital zircon age populations that it contains. This age population overlaps with the ca. 2.1 Ga porphyritic red granite that crops out among the western exposures of Dickinson Group strata, though cross-cutting relationships between the Dickinson and porphyritic red granite are not observed. A ca. 2.1 Ga depositional age for the arkose would require rapid unroofing of the coeval granite in a manner not presently observed elsewhere in the region. In summary, prior interpretations of a continuous ca. 2.1–2.0 Ga Dickinson Group succession in the western Upper Peninsula of Michigan are not consistent with new detrital zircon ages from siliciclastic strata previously mapped as the upper Solberg Schist. These units correlate with the Michigamme Formation instead and raise new questions about the age, setting, and tectonic evolution of multiple Archean and Paleoproterozoic units in the region. References cited 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, https://doi.org/10.1086/673265. Drenth, B.J., Cannon, W.F., Schulz, K.J., and Ayuso, R.A., 2021, Geophysical insights into Paleoproterozoic tectonics of the Superior Province, central Upper Peninsula, Michigan, USA: Precambrian Research, v. 359, https://doi.org/10.1016/j.precamres.2021.106205. James, H.L., 1958, Stratigraphy of pre-Keweenawan rocks in parts of northern Michigan: U.S. Geological Survey Professional Paper 314-C, 44 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. 45 �Characterizing the geochemistry and nickel-copper-platinum group elements potential of mafic and ultramafic intrusions in northwestern Ontario JONSSON, Justin1, MALEGUS, Paul1, CHURCHLEY, Sophie1, PRICE, Rebecca1 1 aw n Resident Geologist Program, Ontario Geological Survey, Ministry of Mines, Suite B002, 435 James Street South, Thunder Bay, ON P7E 6S7 Canada Globally, magmatic sulphide deposits host significant resources of nickel, copper, cobalt and platinum group elements (PGE). These deposits occur as concentrations of sulphide minerals hosted within mafic to ultramafic intrusive rocks and are widespread across Ontario, occurring in every Precambrian geologic terrane. Ontario is home to 10 operating mines in magmatic sulphide deposits: 9 within the Paleoproterozoic Sudbury Igneous Complex and one within the Neoarchean Lac des Iles Complex. W ith dr In 1999, Operation Treasure Hunt was initiated by the Ontario Government to stimulate mineral exploration by acquiring new airborne geophysical data, surficial and bedrock geochemical data, and development of new methods. In 2003, following completion of the Operation Treasure Hunt project, the Ontario Geological Survey published a report (Vaillancourt et al. 2003) that assessed 109 mafic to ultramafic intrusions across Ontario. The purpose of this part of Operation Treasure Hunt was to characterize and publish data for intrusions known to be prospective for PGE-dominated magmatic sulphide mineralization. Many of the intrusions studied during Operation Treasure Hunt were host to significant known mineralization, including current and past-producing mines, and several of these intrusions are the focus of ongoing mineral exploration. Despite the work by Vaillancourt et al. (2003), there are hundreds of mafic to ultramafic intrusions in Ontario that have not been systematically assessed for magmatic sulphide mineralization potential. Many of these intrusions have favourable characteristics for potentially containing magmatic sulphide deposits, including geophysical anomalies (e.g., magnetic, conductivity), overburden geochemical anomalies and known sulphide mineralization. In 2023, the Resident Geologist Program of the Ontario Geological Survey initiated a project to systematically characterize geochemistry of a subset of mafic-ultramafic intrusions in northwestern Ontario that largely have not been subject to significant historical evaluation by academic researchers, government surveys, or mineral exploration companies. Evaluating the geochemistry of mafic to ultramafic intrusions can provide insight into the magma history, tectonic setting and potential for economic metal endowment. Factors that may influence metal endowment, that can be determined from the examination of geochemical data, include determining magma source characteristics, the timing of sulphur saturation and the degree of interaction of the magma(s) with their country rocks. Careful evaluation of physical characteristics and whole-rock geochemistry can inform future mineral exploration and/or the development of models for the emplacement of mafic to ultramafic intrusions and any hosted mineralization. 46 �W ith dr aw n Initial sample collection and analytical work took place during 2023. Areas of interest are shown in Figure 1, and include the Red Lake, Onaman–Tashota, and Heaven Lake greenstone belts. In this display, we provide examples of preliminary results and interpretations from areas targeted in the first year of field work, including the Trout Bay intrusion (Red Lake greenstone belt), Westwood intrusion (northeast of the Lumby Lake greenstone belt), and the Big Ghee Lake intrusion (south of the Shebandowan greenstone belt). Figure 38.1. Simplified bedrock geology map of a portion of northwestern Ontario, showing project target areas: Red Lake greenstone belt (outlined in blue); Heaven Lake greenstone belt (outlined in black); and Onaman–Tashota greenstone belt (outlined in white). Regional geology modified from Ontario Geological Survey (2011). References Ontario Geological Survey 2011. 1:250 000 scale bedrock geology of Ontario; Ontario Geological Survey, Miscellaneous Release—Data 126 – Revision 1. Vaillancourt, C., Sproule, R.A., MacDonald, C.A. and Lesher, C.M. 2003. Investigation of mafic-ultramafic intrusions in Ontario and implications for platinum group element mineralization: Operation Treasure Hunt; Ontario Geological Survey, Open File Report 6102, 335p. 47 �Cross-sectional Geometry of the Keweenaw Fault System between Hancock and Mohawk, Upper Peninsula of Michigan LANGFIELD, Katherine1, GAMET, Nolan2, DeGRAFF, James1 1 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, Houghton, MI, USA 2 Michigan Geological Survey, Marquette, MI, USA The Keweenaw fault system (KFS) is a major compressional feature along the Keweenaw Peninsula near the southern edge of the Midcontinent Rift System (MRS). The MRS formed in the Mesoproterozoic when a major extensional event split the ancient North American continent across the Upper Midwest, yielding large volumes of basaltic lava such as the Portage Lake Volcanics (1.1 Ga, PLV). The PLV strata were thrust southeastward over the Jacobsville Sandstone (JS) along the KFS during post-rift compression by the Grenville Orogeny (1), and some have postulated an earlier origin by normal faulting during rifting (2,3). A recent interpretation based partly on cross-section modeling is that faults making up the KFS are parts of a detached thrust system that formed during the Grenville Orogeny (4). Faults of the KFS have been interpreted to have dip slip – recent reverse slip and possibly earlier normal slip. Since 2017, bedrock mapping and analysis of fault-slip indicators have revealed a significant component of right-lateral strike slip on the KFS, which at its northeast end is inferred to have twice the magnitude of north-side-up reverse slip (5). The collective oblique motion across the KFS is accommodated on fault segments with three distinct orientations that overlap and intersect: (1) major segments parallel to the KFS trend, (2) splay faults striking clockwise to major segments by less than 35°, and (3) shorter connector faults striking counterclockwise to major segments and splays by up to 75° (Fig. 1). The ratio of dip slip to strike slip should vary among faults with such a range of orientations, as should the style of deformation in their hanging walls and footwalls. To help understand these relationships, cross-sections were constructed across various fault components of the KFS using recent mapping data, heritage data from published maps, and drill hole data. Cross-section work employed the dip-domain-bisector method and principles of detached thrust systems and conservation of volume. The new cross-sections attempt to model the subsurface geometry of the segmented KFS and to build on previous work in the area (4). Important unknowns are the JS thickness in the footwall and how JS strata deform adjacent to major faults. A minimum JS thickness of 800 meters was assumed, based on Mayflower drill hole #41 that crosses the Keweenaw fault at 476 meters below sea level (Fig. 2). Ductile deformation of a poorly indurated, mud-prone section near the base of JS was the method used to accommodate flexural slip in the overlying section, but other mechanisms remain to be investigated. A common feature of cross-sections transverse to the KFS trend is a thrust sheet with shallowly dipping PLV strata between a major fault segment and a splay fault. The cross-sections are adding to our understanding of deformation within the KFS and to the tectonic forces that created it. Acknowledgements This project was funded by the USGS EDMAP program (Award G21AC10681), Department of Geological and Mining Engineering and Science of Michigan Tech, ILSG Student Research Fund, Michigan Space Grant Consortium, and sponsored by the Michigan Geological Survey. We thank Tom 48 �Wright for access to Quincy Mine; Ian Gannon, Breeanne Heusdens, Jack Hawes, Braxton Murphy, and Dillon Breen for field assistance; and Dan Lizzadro-McPherson for ArcGIS assistance. References 1. Cannon, W.F., 1994, Closing of the Midcontinent rift ‒ A far-field effect of Grenvillian compression: Geology, v. 22, p.155-158. 2. Cannon, W.F., Green, A.G., Hutchinson, D.R. and nine others, 1989, The North American Midcontinent Rift beneath Lake Superior from GLIMPCE seismic reflection profiling: Tectonics, v. 8, p. 305-332. 3. 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, no. 2, p. 303-310. 4. 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, v. 135, no. 1/2, p. 449–466. 5. Lizzadro-McPherson., D.J., 2023, Structural Analysis and Slip Kinematics of the Keweenaw Fault System between Bête Grise Bay and Gratiot Lake, Keweenaw County, Michigan: Michigan Technological University M.S. thesis, 140 p. Figure 1: Updated bedrock geologic map and legend of study area, with fault segments labelled: KF – Keweenaw Fault, HFHancock Fault, AGF – Allouez Gap Fault Figure 2: Crosssection showing Keweenaw (KF) and Hancock Faults (HF) at Douglass-Houghton Falls. Main units: JS – Jacobsville Sandstone, PLV - Portage Lake Volcanics 49 �Volcanic and Hydrothermal Reconstruction of the Paleoproterozoic Butler Zn-Cu occurrence, Clark County, Wisconsin LAWRENCE, Alex1, VANDERKIN, Adam1, and LODGE, Robert, W.D.1 1 Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA The Butler Zn-Cu occurrence is located in western Clark County, northcentral Wisconsin, and is an example of a volcanogenic massive sulfide (VMS) deposit. These VMS deposits are mined globally for numerous metals including Zn, Cu, Pb, Ag, and Au, and are formed at or near the seafloor in extensional submarine volcanic environments through the discharge of hot, metalrich hydrothermal fluids (e.g. Franklin et al., 2005). The Butler occurrence is hosted in the Paleoproterozoic Eau Claire Volcanic Complex (1.8-1.9 Ga) within the Marshfield terrane of the Penokean Orogen (Shultz and Cannon, 2007; DeMatties, 2022). Historically, the interpreted setting for Penokean volcanism within the Marshfield terrane was a “continental’ setting with younger magmas emplaced within older Archean crust. New data from the Eau Claire Volcanic Complex suggests the absence of Archean crust during Penokean volcanism (Weber et al., 2023). The goal of this project is to interpret the volcanic and hydrothermal setting of the Butler Zn-Cu deposit and test whether the lithostratigraphic and petrochemical associations fit an oceanic or continental model. Extensional environments that form VMS deposits can exist in both oceanic and continental settings. This imparts unique lithostratigraphic (Franklin et al. 2005) and petrochemical (Piercey, 2011) characteristics on the host stratigraphy and alteration styles. The lithostratigraphy of continental-associated VMS tends to be more felsic in nature with a higher abundance of siliciclastic rocks relative to oceanic-settings. Mafic rocks are much more abundant in ocean environments whereas are rare and largely intrusive in continental settings. Chemically, felsic rocks in continental settings are HFSE- and REE-enriched while mafic rocks are typically alkalic- to MORB-affinities commonly found in continental rift settings. Approximately 2700 linear feet of drill core from the Butler occurrence were re-logged and sampled for petrographic and geochemical characterization of the host volcanic and hydrothermally-altered rocks. Petrography divided the host strata into three main units: 1) felsic volcanic rocks, 2) amphibolite, and 3) metapelite. The felsic volcanic rocks are fine-grained, foliated quartzofeldspathic schists (Figure 1A) that have layered volcaniclastic textures and local stretched and flattened lapilli fragments. The amphibolite units (Figure 1B) are fine- to mediumgrained, homogenous, and largely unaltered suggesting an intrusive origin that post-dates the main VMS-forming event. The metapelite units (Figure 1C) are made up of a micaceous matrix composed of muscovite, chlorite, and biotite. The metapelite units are characterized by large porphyroblasts of garnet, staurolite, and/or cordierite. Hydrothermally-altered rocks that host sulfide mineralization are metamorphosed to biotite±chlorite±talc schists and calc-silicate mineral assemblages. The sulfide mineralization is primarily pyrite with variable amounts of chalcopyrite, and sphalerite. Massive sulfides (Figure 1D) are weakly banded with chloritic gangue while semi-massive vein-type mineralization is found throughout altered rocks. The relative abundance of the felsic and amphibolite units coupled with an intrusive origin for amphibolite, suggests a bimodal-felsic type VMS, described in the paper Volcanogenic Massive Sulfide Deposits that is typical of continental magmatism (Franklin et al., 2005). Mafic units are interpreted to be island arc to MORB-type based on Ti vs. V discrimination plots. Felsic volcanic rocks have an FII-type affinity on Zr/Y vs. Y discrimination diagrams and have within plate affinities on Nb vs. Y discrimination diagrams. Geochemical abundances of the host rocks support a continental petrochemical association. 50 �Figure 1. Photographs of core samples from the Butler deposit featuring the host rocks of the VMS. White scale bar equals about 1 cm. (A) Foliated felsic volcanic rock that is the main host rock of the Butler formation. (B) Amphibolite unit, image displays how homogenous the matrix is. (C) Metapelite unit, tan staurolite porphyroblasts along with large purple to grey cordierite porphyroblasts found throughout the matrix. (D) Massive sulfide unit containing pyrite and chalcopyrite. 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. Franklin, J. M., Gibson, H. L., Jonasson, I. R., and Galley, A. G., 2005, Volcanogenic massive sulfide deposits, in Hedenquist, J. F. H., Goldfarb, R. J., and Richards, J. P., eds., Economic Geology, 100th Anniversary Volume, p. 523-560. Piercey, S. J., 2011, The setting, style, and role of magmatism in the formation of volcanogenic massive sulfide deposits: Mineralium Deposita, v. 46, p. 449-471. Schulz, K.J., and Cannon, W.F., 2007. The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157: 4-25. Weber, E.M., Lodge, R.W.D., Marsh, J.H. (2023). U/Pb geochronology and zircon petrochronology of Paleoproterozoic magmas from the Marshfield terrane, Penokean Orogen, Wisconsin. Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 1Program and Abstracts, p. 97-98. 51 �Building a 3D model for Cu/Pd inflection points throughout the Marathon PGE-Cu Deposit LAXER Max1, GOOD David1 1 Department of Earth Sciences, University of Western Ontario, 1151 Richmond Street, London, ON N6A 5B7, Canada The Marathon PGE-Cu deposit is hosted in the North American Midcontinent rift system, a failed continental rift (Good et al., 2015). Magmatic activity in the area created an optimal environment for the formation of economically significant sulphides (Smith et al., 2022) bearing copper (Cu), and platinum group elements (PGE) at the Marathon deposit. The deposit lies in the Two Duck Lake Gabbro, a subophitic, coarse-grained intrusion located at the eastern margin of the Coldwell Complex. This study explores how Cu/Pd varies in 3D space at the deposit scale and aims to use it as a vectoring tool to guide exploration. The ratio of Cu/Pd is a useful marker of the enrichment of Pd relative to the mantle. Low Pd relative to Cu indicates previous Pd depletion due to the early formation of sulphides in the intruding magma that formed the deposit, whereas a relatively higher Pd concentration implies Pd enrichment (Barnes et al., 1993). The 3D model helps to visualize the positions of the abrupt shifts (inflection points) in Cu/Pd ratio throughout the Marathon deposit. Identifying and modelling Cu/Pd inflection points facilitated the search for trends in mineralization. To create the model, a data filtration process was employed to define wide mineralization intervals containing at least 80 ppm Cu and 0.15 ppm Pd. Zones of continuous mineralization of at least 16 m in length were identified. To identify inflection points the difference in Cu/Pd ratio was evaluated at 10 m intervals. The mineralized zones were searched for points that surpassed the thresholds found to constitute trend reversals in Cu/Pd (ΔCu/Pd >5000 or <-5000). Approximately 1150 inflection points have been identified in 404 drillholes from a dataset of 997 drillholes and 61960 assays. Figure 1. Graphs showing the trends of concentration Cu and Pd in ppm, the ratio of Cu/Pd, and the 10 m difference calculations used to identify inflection points, down drill hole M-20-541 (depth in m) at the Marathon PGE-Cu deposit. The dashed lines indicate filtration cut-offs for Cu (80 ppm) and Pd (0.15 ppm) and the inflection point thresholds in the Cu/Pd graph (at -5000 and 5000). 52 �Three zones of interest were identified within the model of the deposit with distinct trends in the occurrence Cu/Pd ratio inflection points (Fig. 2). Area 1 included zones of high grade Pd mineralization occurring independently of any high-grade copper or any inflection points. In Area 2 the arrangement of inflection points suggests a boundary which aligns with the paleosurface at the contact between the Footwall and the Main Zone. A fault runs through Area 3 (Good et al., 2015), along which there are no Cu/Pd inflection points, indicating that there may be a link between faulting and the consistency of Cu/Pd ratio. The most prominent pattern observed in the 3D model of the inflection points was that Cu/Pd correlates better with lithological changes than with shifts in Cu and Pd grade. Area 1 Area 2 Area 3 Figure 2. Views of the Leapfrog model of the Marathon deposit, including all Cu and Pd assays and all inflection points. Showing plan views of the whole deposit, Areas 1 and 3 and a cross-section of Area 2. References Barnes, S.-J., Couture, J., Sawyer, E., & Bouchaib, C., 1993. Nickel-copper occurrences in the BelleterreAngliers Belt of the Pontiac subprovince and the use of Cu-PD ratios in interpreting platinumgroup element distributions. Economic Geology, 88(6), 1402–1418. 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(4), 983–1008. Smith, J. M., Ripley, E. M., Li, C., Shirey, S. B., & Benson, E. K. (2022). Magmatic origin for the massive sulfide ores in the sedimentary country rocks of mafic–ultramafic intrusions in the midcontinent rift system. Mineralium Deposita, 57(7), 1189–1210. 53 �Critical Mineral Systems in the Upper Peninsula of Michigan, A Cooperative Effort Between the USGS and the Michigan Geological Survey MAHIN, Robert2, QUIGLEY, Ashley2 YELLICH, John1, ESCH, John1, and GAMET, Nolan2, 1 Michigan Geological Survey, Western Michigan University, Kalamazoo MI 49008-5241 2 Michigan Geological Survey, Western Michigan University, Gwinn MI 49841 In 2018, the U.S. Geologic Survey (USGS) released a list of critical minerals defined as “non-fuel mineral or mineral material essential to the economic or national security of the U.S., and which has a supply chain vulnerable to disruption” and updated it in 2022 to a total of 50 critical minerals (Burton, 2022). Since 2021, President Biden has made the domestic supply of critical minerals a national priority. With federal funding, the USGS Earth Mapping Resource Initiative (EMRI) is collaborating with State geological surveys on geologic mapping and critical mineral assessments, as well as inventorying and characterizing mine wastes. The USGS has identified broad focus areas within the United States to target critical minerals (Hammarstrom and others, 2023). These focus areas are based on known mineral occurrences and favorable geologic settings. The Precambrian of the Upper Peninsula (UP) figures in 17 mineral systems. The Michigan Geological Survey (MGS) has narrowed the list to nine systems in the UP to focus our future work (Table 1). With the support of the USGS, forthcoming mapping and geochemical reconnaissance programs by the MGS over the next few years will assess these systems. Name of focus area Midcontinent Rift magmatic sulfide Ni-Cu-PGE Manganese (Mn) in ironformations Graphite in black shales Mineral system Deposit type(s) Critical minerals in Critical minerals the deposit types Identified Mafic magmatic Nickel-copper-PGE sulfide Co, Ni, PGE, Te Nickel, Co, PGE Marine chemocline Iron-manganese Co, Mn Manganese Metamorphic Graphite (carbonaceous sed) Humboldt Granite Porphyry Sn (granite-related) Porphyry/skarn Humboldt Granite Magmatic REE Southern Complex pegmatites Porphyry Sn (granite-related) Pegmatite LCT Mesoproterozoic Phosphate Marine chemocline Peavey Pond Complex IOA-IOCG Western Upper Peninsula, IOCG IOA-IOCG Peralkaline syenite/granite/rhyolite/ alaskite/pegmatites Graphite Be, Nb, Sc, Sn, Ta, W Be, Fl, Hf, Nb, REE, Ta, Te, V, Zr Be, Ce, Li, Nb, Sc, Ta, Sn Co, REE Phosphate Iron oxide apatite; Iron oxide copper Co, REE gold Iron oxide apatite; Iron oxide copper Co, REE gold Trace Trace Trace Phosphate Unknown Unknown Table 1: USGS-MGS Critical Mineral Focus Areas for the UP (modified from Hammarstrom and others, 2023) The existence of some critical minerals is well-established in the UP, such as magmatic sulfide Ni-Cu-Pt-Pd-Co. Others, such as graphite, manganese, and phosphate have been documented in small occurrences or as accessory minerals in larger deposits (Cannon and Klasner, 1976: Hwang and others, 1986; James and others, 1968; Peterman and others, 1987, 54 �Mancuso, 1975). Evidence for critical minerals such as rare-earth elements, beryllium, and fluorspar in pegmatites, granites, and iron-oxide-copper-gold/oxide apatite deposits (IOCG/IOA) is sparse. A limited number of studies of UP pegmatites and the Humboldt granite have identified trace REE, Be, and Ta minerals (e.g. Buchholz and others, 2014; Johnson and others, 2015; Moss, 1975, Schulz and others, 1988). Mineralization directly tied to IOCG/IOA has not been identified, although the tectonic history and metal endowment suggests the UP is a permissive, if not prospective region for them. The USGS has also mounted a mine waste characterization program intended to identify potentially recoverable critical minerals in historical mine stockpiles, waste piles and tailings and prioritized by size, potential mineral resources. As part of the effort, the MGS has identified over 80 mine sites in the UP within EMRI critical mineral focus areas that have published references to possible critical mineral content. Future assessments will involve representative sampling of mine waste features and geochemical evaluations. References Buchholz, T. W., Simmons, W. B., and Falster, A.U., 2014: Accessory mineralogy of the Black River Pegmatite and Humboldt granite, Marquette County, Michigan. In Fortieth Rochester Mineralogical Symposium: Contributed Papers in Specimen Mineralogy, Part 1, Rocks & Min., 89:4, 370-374. Burton, J., 2022. U.S. Geological Survey Releases 2022 List of Critical Minerals: https://www.usgs.gov/news/national-news-release/us-geological-survey-releases-2022-list-criticalminerals. Cannon, W.F. & Klasner, J.S., 1976. Phosphorite and other apatite-bearing sedimentary rocks in the Precambrian of Northern Michigan: US Geological Survey Circular, 746, 6 p. Hammarstrom, J.M., Woodruff, L.G., and Dicken, C.L., 2023, Critical mineral deposits of the United States: U.S. Geological Survey data release, https://doi.org/10.5066/P9K1HBNT Hwang, J. Y., Carlson, D. H., Johnson, A. M., and Van Alstine, J., 1986. Preliminary investigation of graphite resources in Michigan, in Process Mineralogy VI: Applications to Precious Mineral Deposits, Industrial Minerals, Coal, Liberation, Mineral Processing, Agglomeration, Metallurgical Products and Refractories, with Special Emphasis on Cathodoluminescence Microscopy, (Ed. by, Hagni, R. D.), p. 315- 327. 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, 127 p. Johnson, Christopher M. and Van Daalen, Christopher M., 2015. Mineralogy and geochemistry of Late Archean and Paleoproterozoic granites and pegmatites in the Northern Penokean terrane of Marquette and Dickinson Counties, Michigan. University of New Orleans Theses and Dissertations. 2088. Mancuso, J.J., 1975. Carbonate-apatite in Precambrian cherty iron formation, Baraga County, Michigan. Economic Geology, 70, p. 583-586. Moss, Michael J., 1975. Some pegmatites near Gwinn, Michigan. Master's Thesis 2451, Western Michigan University. Peterman, J.F., Johnson, A.M. and Van Alstine, J., 1987. Geological and geophysical investigation of graphite resources in Upper Michigan, in Institute on Lake Superior Geology, 33rd Annual Meeting, Proceedings and Abstracts, p. 53-54. Schulz, K.J., P.K. Sims, Z.E. Peterman. 1988. A post-tectonic rare-metal-rich granite in the southern complex, Upper Peninsula, Michigan. in Institute on Lake Superior Geology, 34th Annual Meeting, Proceedings and Abstracts, p. 34. 55 �Baraboo Interval Quartzites in Iowa: Reassessing the Origin and Provenance of the Washington County Quartzite, SE Iowa MALONE, Jack1, CLARK, Ryan1, HARRIS-BOMMARITO, Amira2, and MALONE, David2 1 Iowa Geological Survey, 300 Trowbridge Hall, University of Iowa, Iowa City, IA 52242 Geography-Geology, Felmley Hall of Science, Illinois State University, Normal, Illinois 61790 2 The Washington County Quartzite (WCQ) in southeastern Iowa is the southernmost occurrence of the Baraboo Interval quartzites in the midcontinent region (Figure 1). Two drill holes encountered poorly sorted quartzite and phyllite at a depth greater than 2,300 feet, likely deposited in a braided fluvial or deltaic environment near the Laurentian continental margin on the Columbia supercontinent. One hundred new LA-ICPMS detrital zircon ages from the WCQ show a prominent 1.78 Ga age peak, representing local Yavapai-aged basement, a secondary peak at 1.8-1.9 Ga representing a distal Penokean source, and minor <2.5 Ga peak derived from distal sources in the Superior Province. Multidimensional scaling of other Baraboo Interval quartzites and potential sources show that the WCQ is indistinguishable from the lower interval of the Baraboo quartzite (Figure 2). Cumulative and stacked probability plots also reflect principal source areas locally derived by erosion of underlying Yavapai-aged crust and distally derived Penokean and older sources from the Pembine-Wausau Terrane or southern Superior Province. The WCQ likely serves as the down slope equivalent during initial Baraboo deposition. Figure 1. Geological map of Precambrian basement rocks in the northern midcontinent (modified from Medaris et al., 2021). Baraboo Interval Inliers are B = Barron, F = Flambeau, M = McCaslin, T = Thunder Mountain, R = Rib Mountain, and N = Necedah. SLTZ = Spirit Lake Tectonic Zone, ECMP = East-Central Minnesota Batholith, GLTZ= Great Lakes Tectonic Zone, and NFZ = Niagara Fault Zone. 56 �Figure 2. Three-dimensional multi-dimensional scaling plot of Baraboo Interval strata in the northern midcontinent. Included data compiled from Malone et al. (2022), Van Wyck and Norman (2004), Stewart et al. (2018), Stewart et al. (2021), and Medaris et al. (2021). The red circle is the WCQ from this study. The yellow cluster includes samples from the lower members of the Baraboo Quartzite in the Baraboo Hills that are dominated by Yavapai-age zircons. The blue cluster includes the middle members of the Baraboo Quartzite in the Baraboo hills, other quartzites north of the SLTZ, and Necedah reflects a complex sedimentary provenance that includes proximal and distally derived zircons from the Penokean Province, Superior Province, and Trans-Hudson belt. The purple cluster includes the Waterloo Quartzite and the upper members of the Baraboo Quartzite in the Baraboo Hills that are dominated by southerly derived Mazatzal-age zircons. References Malone, D.H., Craddock, J.P., Holm, D., Krieger, A., and Baumann, S.J., 2022. Continent‐scale sediment dispersal for the Proterozoic Baraboo Interval quartzites in the Laurentian midcontinent. Terra Nova, 34(6): 503-511. Medaris, L.G., Jr., Singer, B.S., Jicha, B.R., Malone, D.H., Schwartz, J.J., Stewart, E.K., Van Lankvelt, A., Williams, M.L., and Reiners, P.W., 2021. Early Mesoproterozoic evolution of midcontinental Laurentia: Defining the geon 14 Baraboo orogeny. Geoscience Frontiers, 12: 101174. Stewart, E.K., Brengman, L.A., and Stewart, E.D., 2021. Revised Provenance, Depositional Environment, and Maximum Depositional Age for the Baraboo (<ca. 1714 Ma) and Dake (<ca. 1630 Ma) Quartzites, Baraboo Hills, Wisconsin. Journal of Geology, 129: 1-31 . Stewart, E.D., Stewart, E.K., Walker, A., and Zambito, J.J., IV., 2018. Revisiting the Paleoproterozoic Baraboo interval in southern Wisconsin: Evidence for syn-depositional tectonism along the south-central margin of Laurentia. Precambrian Research, 314: 221-239. Van Wyck, N., and Norman, M., 2004. Detrital zircon ages from early Proterozoic quartzites, Wisconsin, support rapid weathering and deposition of mature quartz arenites. Journal of Geology, 112: 305-315. 57 �The Soudan Geology Trail Project: Let’s talk about rocks in northeastern Minnesota MARTIN, Alice1, ALLERTON, Zsuzsanna1, JOHNSON, Emma1, FAYON, Annia1, ESSIG, Jim2, GUY-LEVAR, Sarah2, HUDAK, G. H. 1 1 Earth and Environmental Science Department, University of Minnesota, 150 John Tate Hall, 116 Church St. SE, Minneapolis, MN 55455, USA 2 Minnesota Department of Natural Resources, 1302 McKinley Park Rd, Soudan, MN 55782, USA As part of a larger research endeavor within the Archean terrane of northern MN, we are developing educational outreach content created to engage the public with portions of the important geology of the region. Rocks exposed in the Lake Vermilion-Soudan Underground Mine State Park, located near Tower, MN, record glimpses of environmental and tectonic conditions from 2.7 billion years ago to the present, including mysteries of the early Earth, complexities of modern history, and possibilities of the future. The planned content will be designed to follow outcrops located along a paved trail that runs through the park (Figure 1). Figure 1: This figure shows recently exposed units along the proposed trail (bold line in center of the figure). Starting at the banded iron formation (A), the trail leads north then east (clockwise) to the next units (B) consisting of pillow basalts and basaltic lava flows, followed by felsic tuff (C) and lastly a large, exposed outcrop along the paved road is chlorite schist with intertwined with banded iron formation (D). 58 �The chosen outcrops consist of well-preserved greenschist-facies metamorphosed igneous, sedimentary, and sheared rocks. Selected rock units along the trail include Neoarchean basaltic lava flows, pillow basalts, felsic tuff, gabbro, oxide-facies banded iron formation, and chloriteand sericite-dominant schists. These rock units represent an ancient submarine volcanic and hydrothermal environment that was subsequently regionally deformed (Hudak et al., 2016). The work being done will be included in educational materials designed to communicate the scientific content in digestible ways. Plain language writing, paired with visuals, modern analogues and analogies will present the information in a variety of ways with the intention of supporting a range of learning styles. The materials will be available in physical forms (on paper and/or trail signs) and with QR codes which will link to additional online content. Visual illustrations will be designed in collaboration with a Minnesota high school student. Combining educational material with the hands-on outdoor experience of visiting the trail and highlighted outcrops aims to facilitate cognitive development and understanding of the long history and importance of the regional geology. This initiative is a collaboration among park officials Jim Essig (Park Manager) and Sarah Guy-Levar (Interpretive Supervisor) from the MN Department of Natural Resources, the University of Minnesota Department of Earth and Environmental Sciences, and local and state educators. References Hudak, G.J., Peterson, D.M., Radakovich, A., Pignotta, G., Schwierske, K., and Students from the 2010-2013 Precambrian Research Center Geology Field Camp, 2016, Bedrock geologic map of Lake Vermilion/Soudan Underground Mine State Park – Report to the Minnesota Department of Natural Resources: Natural Resources Research Institute, University of Minnesota Duluth, Technical Report NRRI/TR-2016/20, 23 p. 59 �Investigating the origin of pervasive breccias in the Paleoproterozoic Saunders Formation in northern Wisconsin MARTIN, Gwendolyn and BJØRNERUD, Marcia Department of Geosciences, Lawrence University, 711 E Boldt Way, Appleton WI 54911 The Paleoproterozoic Saunders Formation is an enigmatic unit with limited exposure along the Brule River, which forms the border between northern Wisconsin and the Upper Peninsula of Michigan. The unit occurs just north of the Niagara Fault zone, the Penokean-age (ca. 1.88 Ga) tectonic suture between the Superior Craton and the Wisconsin Magmatic Terranes (Schulz & Cannon, 2007). Variously described as a “massive dolomite” (Allen 1910), a “silicified dolomite” (Sims, 1992), and a “silica rock” (Cannon, 1986), the Saunders Fm. is thought to be part of the lower Chocolay Group, correlative with the Randville, Bad River, and Kona Dolomites, and possibly also the quartzites underlying these units (Sturgeon River, Sunday, and Mesnard Fms.). Each of these carbonate formations is overlain by a major unconformity, at ca. 2.1 Ga., representing at least 100 million years of erosion. Every published description of the Saunders Fm. mentions that it tends to be brecciated, yet the nature of these breccias has not been explored in detail. Dutton & Linebaugh (1967) suggested that the Saunders Fm. represents a condensed section of basal Chocolay quartzite and dolomite, related to the formation of the regional unconformity. James et al. (1968) similarly hypothesized that “silcretes” within the Saunders had formed by Proterozoic weathering but also pointed out that none of the other Chocolay Group carbonate units displays evidence of such deep weathering. They speculated, therefore, that the Saunders breccias could have had a tectonic origin but did not pursue that hypothesis further. The purpose of this study was to characterize and interpret Saunders breccias in outcrops at Brule River Cliffs State Natural Area in Wisconsin. At this site, outcrops of the Saunders Formation are of two types: 1) beige to orange-colored dolostone with a ‘gritty’ but otherwise massive (unveined, unlaminated) texture; and 2) dramatically fragmented dolostone with extensive ‘stockwork” quartz veins that constitute most of the rock mass. Immediately southwest of the Natural Area boundary, large boulders of dolomite-matrix breccias with angular chert fragments are common. Although these are not in situ, we suspect they come from the Saunders Fm., and were transported ca. 12 km by glacial ice. If so, these chert breccias represent a third, distinct textural type within the Saunders. The gritty dolomite, which is typically unveined, has a distinctive diamictic, granular texture, with scattered mm-sized, rounded grains set in a much finer matrix. In thin section, the matrix also appears granular, unlike the crystalline texture typical of most carbonate rocks. Similar gritty/ granular dolomites have been observed along a major upper crustal fault zone in Namibia. Rowe et al. (2012) interpreted these unusual textures as records of decarbonation and fluidized granular flow caused by rapid frictional heating during seismic slip in rocks that had been at ambient crustal temperatures of around 200°C. (Carbonate rocks typically devolatilize, rather than melt, during seismic slip, so pseudotachylyte is rare along faults cutting through dolostone). The absence of talc or other calc-silicate metamorphic minerals in the Saunders Formation points to subgreenschist temperatures in an upper crustal setting comparable to the Namibian case. The veined breccias have an ‘exploded’ look, with quartz veins in multiple orientations that appear to have increased the volume of the rock mass more than 100%. The isolated fragments of host dolostone have narrow, slab-like shapes that suggest fragmentation occurred partly along bedding planes. In thin section, most of the veins have a coarse, blocky texture with no preferred orientation of crystals. Fluid inclusions arrays are common, particularly in the interiors of the crystals. Some of the vein quartz shows slight undulose extinction. The chert breccia boulders 60 �found southwest (in the down-ice direction) of the Saunders outcrops have the texture of cataclasites. The angular fragments of chert within these breccias appear to represent thin silicified stromatolitic layers that were fractured and dismembered. We interpret these three textural types of the Saunders Fm. as distinct areas within a major Penokean-age fault zone. The chert breccias may represent the outer part of the fault zone, dominated by non-seismic cataclasis. The gritty dolomite, bearing evidence of co-seismic heating, would have been closer to the fault core, together with the heavily veined dolostone, whose ‘exploded’ nature points to extreme dilational strain and forceful fluid influx with little cataclasis or grinding. The large amounts of vein material relative to the host rock, as well the blocky texture of the veins, are consistent with the introduction of large volumes of overpressured, silicasupersaturated fluids into the shallow crust during the propagation of a fault rupture upward from depth. Such fluid influx can happen when co-seismic slip breaks the barrier between deep crustal, low-permeability rocks in which fluids are at lithostatic pressures and overlying high-permeability rocks with fluids at hydrostatic pressure (Cox and Munroe, 2016). Silica-rich fluids traveling upwards from below this barrier would be far from chemical equilibrium in the shallow crust, and they would rapidly precipitate their dissolved silica, easily overcoming kinetic quartz growth limits that exist under equilibrium conditions (Williams and Fagereng, 2022). This interpretation of the Saunders breccias is supported by oxygen isotope analyses of seven vein quartz samples, all of which yielded 18O values between 15.94 to 17.46 VSMOW. Although brecciated textures described in previous studies of the Saunders formation may be related to deep weathering and the post-Saunders unconformity, the breccias exposed in the Brule River Cliffs Natural Area are clearly tectonic -- and probably coseismic -- in origin. The Saunders Formation thus provides further evidence for great earthquakes at various crustal depths along major fault zones during the Penokean orogeny (Larson & Bjørnerud, 2017; Taylor & Bjørnerud, 2023). References cited Allen, R., 1910. The Iron River iron-bearing district. Mich. Geol. Biol. Survey Pub. 3, Ser. 2, 151 p. Cannon, W., 1986. Bedrock geologic map of the Iron River 1º x 2º quadrangle. USGS Map I-1360-B. Cox, S. & Munroe, S., 2016. Breccia formation by particle fluidization in fault zones. Am. J. Science, 316, 241-278. Dutton, C. & Linebaugh, R., 1967. Map of Precambrian geology of Menominee district, USGS Map I-466. James, H., et al., 1968. Geology & ore deposits of Iron River-Crystal Falls District. USGS Prof. Paper 570. Larson, M. & Bjørnerud, M., 2017. Seismic slip, mylonitization and fluid flow along the Penokean TwelveFoot Falls shear zone, Marinette County, NE Wisconsin. Proc. Inst. Lake Superior Geol., 63, 56-57. Rowe, C., Fagereng, Å., Miller, J. & Mapani, B., 2012. Signature of coseismic decarbonation in dolomitic fault rocks of the Naukluft Thrust, Namibia. Earth & Planetary Science Letters, 333, 200-210. Schulz, K. & Cannon, W., 2007. Penokean orogeny in the Lake Superior Region. Precam. Res. 157, 4-25. Sims, P.K., 1992. Geologic map of Precambrian rocks, southern Lake Superior region, USGS Map I-2185 Taylor, M., and Bjørnerud, M., 2023. Deciphering the metamorphic and deformational history of the Hardwood Gneiss, Felch District, Michigan. Proc. Inst. Lake Superior Geol., 69, 89-90. Williams, R. and Fagereng, Å., 2022. The role of quartz in the seismic cycle. Rev. Geophysics, 60, 2021RG000768. 61 �The Evolution of Baraboo Interval Sedimentary Rocks: Deposition at 1.63 Ga and Metamorphism at 1.47 Ga MEDARIS, Gordon Jr., BONAMICI, Chloe, BROWN, Phil, GOODWIN, Laurel, JICHA, Brian, SINGER, Brad, SPICUZZA, Michael, VALLEY, John Department of Geoscience. University of Wisconsin–Madison, Madison, Wisconsin 53706 Supermature siliciclastic sedimentary rocks of the Baraboo Interval (Dott, 1983) were deposited in the southern Lake Superior region following the 1.65-1.63 Ga Mazatzal orogeny and subsequently experienced 1.47 Ga fluid-rock interactions related to the trans-Laurentian Pinware-Baraboo-Picuris orogeny (Daniel et al., 2022). Metamorphic fluid-rock interactions include dehydration and metasomatic varieties, the latter having been promoted by regional-scale advective flow of brines along permeable channels in the various Baraboo Interval occurrences. In the Baraboo Range, south-central Wisconsin, the supermature sedimentary rocks are composed of five oxides, viz. SiO2, Al2O3, Fe2O3, TiO2, and H2O, with CaO, Na2O, and K2O having been largely removed during weathering of the igneous basement. During metamorphism, the original sedimentary mineral assemblage of kaolinite + quartz + hematite + rutile was transformed to one of pyrophyllite + quartz + hematite + rutile through the dehydration reaction, Al2Si2O5(OH)4 (kln) + 2SiO2 (qtz) = Al2Si4O10(OH)2 (prl) + H2O (fluid) Note that in shale, which consisted mostly of kaolinite, the appearance of pyrophyllite is accompanied by diaspore, as expressed by the dehydration reaction, 2Al2Si2O5(OH)4 (kln) = Al2Si4O10(OH)2 (prl) + 2AlO(OH) (dsp) + 2H2O (fluid) Figure 1. Schematic cross-section of the Baraboo Range, indicating the various metasomatic mineral assemblages; the numbers specify 40Ar/39Ar plateau ages for muscovite. Folding and metamorphism in the Baraboo Range were accompanied by advective flow of brines and potassium metasomatism along the base of the quartzite and in the overlying slate (Fig. 1). At the base of the quartzite, kaolinite was replaced by muscovite in paleosol, kaolinite was replaced by pyrophyllite and accompanied by precipitation of muscovite in metapelite (pipestone), and thin diaspore hydrothermal veins (bordered by muscovite) intruded quartzite 62 �above the metapelite. Slate that overlies the quartzite consists of muscovite, chlorite, quartz, hematite, and rutile and contains 4.7% to 6.4% K2O, due to potassium metasomatism, compared to 3.5% K2O in average shale. The paleosol at the base of the quartzite in Baxter Hollow (Fig. 1), which is 796 cm thick, experienced a total flux of 0.46 mol cm-2 K2O during metasomatism. Five additional Proterozoic paleosols in the southern Lake Superior region, ranging in thickness from 300 cm to 950 cm, also experienced potassium metasomatism, with K2O fluxes between 0.22 and 0.73 mol cm-2, respectively; for all six paleosols taken together, K2Oflux = 0.00066  thicknesscm + 0.06, for which R2 = 0.82. SiO2 was mobilized high in the quartzite section at the base of the metasiltstone and metapelite horizon in the south limb of the syncline, where quartz was precipitated in several bedding-parallel slickenfiber layers, 3 mm to 8 mm thick (Fig. 1). Individual slickenfibers are cylindrical, having a:b:c fabric ratios of 8:1:1, in which the a-dimension is up to 4 mm in length. Slickenfibers plunge down-dip approximately perpendicular to the fold axis of the syncline, and slickenfiber steps record top-to-the-south shear. SiO2 was also mobilized above the metasiltstone/metapelite horizon, where quartz was precipitated in bedding-parallel quartzite breccia zones up to 100 m thick (Fig. 1). The breccia zones consist of angular red quartzite fragments cemented by a stockwork of white quartz veins that consist predominantly of coarse-grained quartz and small amounts of specular hematite and locally, coarse-grained muscovite. Euhedral quartz crystals occur in late-stage vugs, some of which are partly to completely filled by kaolinite. Values of 18O (-2‰ to +31‰ VSMOW, SIMS) in euhedral quartz correlate with complex patterns of growth zoning and healed fractures (SEM-CL) to reveal multiple fluid events, including high-T (~300 oC) and low-T (50-100 oC) exchange with hydrothermal and meteoric fluids (Schranz et al., 2017). The pyrophyllite + diaspore mineral assemblage in metapelite constrains the temperature of Baraboo recrystallization to between 315 oC and 360 oC at a pressure of 2.0 kbar. An isochor for fluid inclusions in quartz in a folded quartz vein in metapelite, combined with phase equilibrium considerations, yields T-P conditions between 320 oC, 2.7 kbar, and 385 oC, 4.0 kbar, corresponding to a thermal gradient of ~30 oC/km. Expressed another way, the thermal gradient for metamorphism of the Baraboo quartzite was ~1700 oC/GPa, which places Baraboo metamorphism in the high T/P type of metamorphism (775 oC/GPa < T/P < 2000 oC/GPa), as defined by Brown and Johnson (2019). 40 Ar/39Ar plateau ages for muscovite in paleosol (1467 ± 11 Ma), hydrothermal veins (1478 ± 12 Ma), quartzite breccia (1472 ± 3 Ma), and four samples of slate (between 1493 ± 3 and 1473 ± 3 Ma) demonstrate that recrystallization and K-metasomatism in the Baraboo Range were contemporaneous with emplacement of the 1476-1470 Ma Wolf River A-type ferroan granitic batholith in Wisconsin. Such metamorphism and magmatism in Wisconsin represent the local expression of the continental-scale geon 14 Pinware-Baraboo-Picuris orogeny, which is characterized by high T/P metamorphism and A-type ferroan granitic batholiths. References Brown, M. and Johnson, T., 2019. Metamorphism and the evolution of subduction on Earth. Am. Mineral., 104: 1065-1082; Dott, R.H. Jr., 1983. The Proterozoic red quartzite enigma in the north-central U.S. – resolved by plate collision? Geol. Soc. Am. Mem., 160: 129-141; Daniel, C.G., et al., 2023. Linking the Pinware, Baraboo, and Picuris orogens: Recognition of a trans-Laurentian ca. 1520-1340 Ma orogenic belt. Geol Soc. Am. Mem., 220: 175-190; Schranz, L., et al., 2017, Stable oxygen isotopes, fluid inclusions, and microstructures in Baraboo Quartzite breccia. Proc. ILSG, v. 63/1: 83-84. 63 �Geochemistry of Midcontinent Rift-related intrusive rocks of the Sunday Lake intrusion MEXIA, Kevin1 and HOLLINGS, Pete1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, On P7B 1J4, Canada. The Sunday Lake intrusion is located 25 km north of Thunder Bay, Ontario and hosts Ni-CuPGE mineralization. It has an age of 1109.0±1.3 (Bleeker et al., 2020), and is related to the ~1115 to 1106 Ma magmatic event of the Midcontinent Rift System (MRS; Heaman et al. 2007). The intrusion is emplaced in Archean rocks of the Quetico Basin along the Crock Fault (Flank, 2017). It is not exposed at surface but was first identified from airborne magnetic surveys. In 2008, platinum and palladium mineralization was discovered by HTX Minerals Corp. In 2017 Impala Canada Ltd. Partnered on a joint venture with more than 30 holes drilled to date. The intrusion is funnel-shaped with a width of up to 1.5 kilometers and is 3 kilometers in length. It varies from 350 meters to 1000 meters in thickness. The intrusion consists of maficultramafic layers divided into three series: the Upper Gabbro Series, the Lower Gabbro Series, and the Ultramafic Series (Flank, 2017). Reef-style mineralization present in the lower zones of the intrusion consists of disseminated to blebby chalcopyrite-pyrrhotite-pyrite bearing olivine melagabbro (Fig. 1; Miller, 2020). One hole intersected the basal zones of the intrusion with more than 20 meters of mineralization at 2.11 g/t Pt, 0.95 g/t Pd, 0.16g/t Au, 0.26% Cu, and 0.11% Ni (Flank, 2017). The objectives of this project are to characterize the paragenesis of the Sunday Lake intrusion and the Ni-Cu-PGE mineralization, investigate the effects of crustal contamination on mineralization within the Sunday Lake Intrusion, and to place the Sunday Lake intrusion within the evolution of the MRS. This project utilizes two representative drill holes from which a total of 71 samples were collected. A total of thirty polished thin sections were generated for petrographic studies. Rocks were classified based on relative proportions of olivine, clinopyroxene, and plagioclase with modal rock names such as melagabbro, olivine melagabbro, and wehrlites. Fifty-five samples were analyzed for major and trace elements. Spider diagrams show different compositions within the layered intrusion, with primitive samples having trends consistent with a plume-like composition (Fig. 2A) while others suggest interaction with and contamination by host rocks (Fig. 2B). Variation in the behavior of trace elements suggest contamination, assimilation, and fractional crystallization processes were involved in the magmatic evolution of the intrusion. Sixteen samples have been sent for Sm-Nd and Rb-Sr isotope studies. The results of this study will be used to assess the source of mineralization and extent of contamination of the Sunday Lake Intrusion. 64 �A SL23KM41 Cpy Po B 5 mm Gangue Figure 1. Photomicrograph in reflected natural light (PPL) of a gabbroic sample containing pyrrhotite and chalcopyrite. Polished thin section scanned using a Zeiss microscope. References Figure 2. Primitive mantle normalized REE spider diagram of two samples. A: Sample showing a plume-like trend. B: Sample suggesting an interaction with the host rock. Normalising values from Sun and McDonough (1989). 9 Bleeker, W., et al. "The Midcontinent Rift and its mineral systems: Overview and temporal constraints of Ni-Cu-PGE mineralized intrusions." Targeted Geoscience Initiative 5 (2020): 7-35. Flank, S. (2017). The Petrography, Geochemistry and Stratigraphy of the Sunday Lake Intrusion, Jacques Township, Ontario. School of graduate studies. Heaman, L. M., Easton, R. M., Hart, T. R., MacDonald, C. A., Hollings, P., & Smyk, M. (2007). Further refinement to the timing of Mesoproterozoic magmatism, Lake Nipigon region, Ontario. Canadian Journal of Earth Sciences, 44(8), 1055-1086. Miller, J.D., Green, J.C., and Severson, M.J. (2002). Terminology, nomenclature, and classification of Keweenawan igneous rocks of northeastern Minnesota. 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. 5-20. Miller, J.D. (2020). Report on the Petrography, Geochemistry, and Lithostratigraphy of DDH SL10-026 from the Southern Sunday Lake Intrusion. JDM GeoConsulting. 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, 42(1), 313-345. Wold, R.J., Hinze, W.J. (1982). Geology and tectonics of the Lake Superior basin. Geol. Soc. Am. Mem. 156, 280. 65 �TWO DECADES OF TEACHING THE GEOLOGIC HERITAGE OF MINNESOTA’S NORTH SHORE AT THE NORTH HOUSE FOLK SCHOOL, GRAND MARAIS MILLER, Jim1 and GREEN, John2 1 2 Department of Earth and Environmental Science, UMD – retired; current residence: Shuniah, ON Department of Earth and Environmental Science, UMD – retired; current residence: Duluth, MN Since 1997, the North House Folk School, located in Grand Marais, Minnesota, has been promoting lifelong learning in the traditional arts and crafts and in knowledge about our northern culture and environment - present and past. Starting with a dozen courses at its inception, North House currently offers over 350 classes per year to over 3,000 students. From 2004 to 2010, John Green lent his expertise to North House by offering weekend lectures and field courses each year on basic geology and the geology of the North Shore and the Gunflint Trail. As a bonus, he also compiled lists of the native plants seen on these field excursions. In 2013, Jim Miller revived the course and has come to offer 2-3 courses per year that explore the Midcontinent Rift geology of the North Shore (What’s This Rock series) and the diverse geology at the end of the Gunflint Trail (Geology up the Trail series). By 2022, three different North Shore classes are offered (two per year, rotating in May and August) that explore different segments of the shore: What’s this Rock? – Grand Portage to Grand Marais, What’s this Rock Too? - Grand Marais to Tettegouche State Park, and What’s this Rock 3? - Tettegouche to Two Harbors (Fig. 1). From 2014 to 2019, a mid-October weekend field trip at the head of the Gunflint Trail was run out of the Gunflint. We hope to reinstate this course this fall or next, but will run it out of Grand Marais. Due to the Covid pandemic, no in-person field courses were permitted during 2020 and most of 2021 (one WTR course was run in October 2021, but participants drove their own vehicles). North House opted to host on-line webinars during the winter of 2021-22. Jim presented three webinar series. In January 2021, three lectures were offered on North Shore geology which was virtually attended by 104 students. During March 2022, three lectures were presented on the geology of Minnesota State Parks and Waysides with 83 people logged in. Then, in January 2023, a two-lecture webinar on the geology of the Gunflint Trail was viewed by 50 people. With the lifting of all Covid restrictions in the winter of 2022-23, North House reverted to only offering in-person classes. As currently taught, the three weekend WTR courses start with an introductory meeting on Friday evening on the North House campus in Grand Marais to discuss trip logistics and provide a geologic overview. Saturday is devoted entirely to a field trip that visits various classic geological exposures along the North Shore. Travel has typically involved carpooling with personal vehicles, but starting in August, 2023, the field trips have used a mini-bus. This or a 15passenger van will be the preferred method of transport going forward. In the evening, participants have the option to gather for informal discussions or a special lecture on various topics, especially at wood-fired pizza party held either Friday or Saturday evenings at the North 66 �House campus. The weekend concludes with a half-day field trip on Sunday morning, after which the group gathers either on a cobbled beach on Lake Superior to practice their newfound rock identification skills. Each field course is limited to about 15 registrants. A total of over 400 students have attended the field courses we have taught at North House over the past 20 years. Participants have ranged in age from 10 to 80 and come from all over the US and Canada, though most are from Minnesota, especially the Twin Cities. Their backgrounds range from those who have never heard of plate tectonics, to those who have had a few geology courses in their past. The common denominator among all participants is that they can all be characterized as being “rock curious”. Figure 1: Geology of northeastern Minnesota showing the general locations of geology field courses currently taught at North House Folk School - three “What’s this Rock?” courses and a Gunflint Trail (GFT) course. 67 �Quartz trace element chemistry: Exploring the link between a fertile parental granite and a mineralized pegmatite MORSON, Mia, and ZUREVINSKI, Shannon Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada Recent studies have proposed the use of pegmatitic quartz trace element chemistry as an indicator to lithium mineralization of potential economic pegmatite deposits (Müller et al., 2021). Using Laser Ablation- Inductively Coupled Mass Spectrometry (LA-ICP MS), the trace elements are determined in situ within a single quartz grain. Trace elements such as Al, Ge, Ti, P, B, and Fe3+, can substitute for Si4+ at very low concentrations, and the elements H, Li, Na, K, Fe2+ can enter the quartz crystal lattice via interstitial lattice positions (Götze et al., 2004). In this study, quartz from the 2685 Ma fertile Ghost Lake granitic batholith (GLB) and related Mavis Lake pegmatites group (Dryden, Ontario) were used to compare trends in quartz trace element concentrations within a single system (Figure 1). The focus of this study was to test the applicability of using quartz to (1) help identify fertile parental granitoid plutons and (2) help decipher any internal fractionation trends in mineralized pegmatites. If trace element concentrations in quartz from the granitic parent show any relationship to the mineralized pegmatites, this would prove beneficial in other exploration programs where potentially enriched Li pegmatites have not yet been identified. It could allow for an early assessment of an S-type granitoid and show indication of a potential nearby mineralized pegmatite, essentially the technique could be deemed a ‘fertility indicator’. Furthermore, if the quartz trace element compositions in the pegmatite zones show enrichment trends similar to the proposed model of Černý (1991), the technique would prove powerful in defining potential areas for further investigation (i.e. following the trend of enrichment). Figure 1: Study sample location through the zones of fractionation and enrichment, from the Ghost Lake Batholith to the Mavis Pegmatites (modified from Breaks and Selway, 1991). The results show correlations with increasing Li, Al, and Ge, and decreasing Ti from the GLB fertile parent granite to mineralized Mavis Lake pegmatites (Figure 2). The quartz trace elements can be used to form a simple fractionation model, similar to the pegmatite fractionation model of Černý (1991) to show elemental enrichments in a fertile LCT pegmatite system upon evolution (Figure 3). In summary, the quartz chemistry indicates fractionation and enrichment 68 �trends can be identified across the pegmatite zones. It is still unclear whether or not the technique could be applied to granites in order to assess fertility, as more work is required to understand the key differences in the quartz chemistry between a barren granite and a fertile parent granite. Figure 2: Trends in quartz trace element concentration in the GLB granites, intermediate beryl columbite zone, and mineralized Mavis Lake pegmatites. (a) Al vs Li bivariate plot. (b) Ti vs Ge bivariate plot. Figure 3: Fractionation model showing quartz trace element trends within the GLB-Mavis Lake system (modified from Černý, 1991). Breaks F.W., Selway, J.B., Tindle, A.G., 2005. Fertile peraluminous granites and related rare-element pegmatites, Superior Province of Ontario. Rare-Eelement Geochemistry and Mineral Deposits: Geological Association of Canada (GAC) Short Coarse Notes 17: 87-125. Černý, P., 1991. Rare-element granitic pegmatites. Part 1: Anatomy and internal evolution of pegmatite deposits. Part 2: Regional to global environments and petrogenesis. Geoscience Canada 18:49– 81. Götze, J., Plötze, M., Graupner, T., Hallbauer, D.K., Bray, C. J., 2004. Trace element incorporation into quartz: A combined study by ICP-MS, electron spin resonance, cathodoluminescence, capillary ion analysis, and gas chromatography. Geochimica et Cosmochimica Acta, 68(18): 3741–3759. Müller A., Keyser W., Simmons W. B., Webber K., Wise M., Beurlen H., Garate-Olave I., Roda-Robles E., Galliski M. A., 2021. Quartz chemistry of granitic pegmatites; implications for classification, genesis and exploration. Chemical Geology, 584:1-17. 69 �Geometry, Slip Kinematics, and Deformation along the Hancock Fault in the Quincy Mine Workings, Upper Peninsula of Michigan MURPHY, Braxton, LANGFIELD, Katherine, DeGRAFF, James Department of Geological and Mining Engineering and Sciences, Michigan Technological University, Houghton, MI, USA 49931 The Hancock fault is one of several major compressional features along the southern edge of the Midcontinent Rift System. It forms part of the Keweenaw fault system (KFS) whose connected segments follow the spine of Michigan’s Keweenaw Peninsula. The Hancock fault is a splay in the hanging wall of the KFS that intersects the main Keweenaw fault zone at an acute angle to define a thrust slice (Fig. 1). It extends along an azimuth of 55° for 17 kilometers from a point west of Houghton to its intersection with the main Keweenaw fault zone between Calumet and Lake Linden. Volcanic and sedimentary layers of the Portage Lake Volcanics (PLV, 1.1 Ga) are shown on bedrock geology maps with left-lateral offset across the Hancock fault (1, 2). Figure 1: Hancock (HF) and Keweenaw (KF) faults shown on USGS bedrock geology maps (1, 2). Major layers: pb-Bohemia conglomerate, psc-Scales Creek flow, pkKearsarge flow, pgGreenstone flow, chc-Copper Harbor conglomerate (base). Like other faults of the KFS, the Hancock fault has a major component of reverse slip that occurred during compression related to the Grenville Orogeny (3). Recent mapping and fault-slip measurements on the population of faults associated with the KFS reveal that the fault system has a right-lateral component of strike slip. This raises the question of whether the leftlateral offset of units across the Hancock fault in map view can be reconciled with net reverse and right-lateral slip along the KFS. The work reported here builds on previously reported work on the Hancock fault in the Quincy Mine workings to address this and other questions about its geometry, slip kinematics, and deformation (4). An adit in east Hancock provides access to the 7th level of the historic Quincy Mine, whose workings run along and across the Hancock fault at four locations (Figs. 1 and 2). The current phase of the project focused on acquiring data from the fault southwest of the adit, which data 70 �were combined with data previously collected from the adit toward the northeast. Fault-slip measurements were made on all accessible fault surfaces and consisted of fault attitude, slickenline rake, and slip sense where possible. The Hancock fault’s strike and dip were measured at well-exposed locations, and the thickness of its gouge and breccia were measured where possible. Orientations were measured using the FieldMove Clino app as well as a Brunton compass. Stereonet and FaultKin freeware were used to plot and analyze the orientation data. Figure 2: Geology along the Hancock fault in the Quincy adit and connected mine workings. The project is ongoing but some results have emerged. The Hancock fault cuts upward across stratigraphy in the direction of thrusting at angles between 4° and 18°, similar to cut-off angles for the Keweenaw fault (5). The low cut-off angles of both faults, which have significant reverse slip, are consistent with the properties of a detached thrust system. Both faults also have components of strike slip as indicated by the population of nearby faults, but right-lateral slip is slightly dominant over left-lateral slip. This implies that the significant left-lateral offset of PLV layers seen in map view is the result of mostly reverse slip on the Hancock fault as it cuts slightly clockwise to strike of the layers. Acknowledgements Funding for this work was provided by the ILSG Student Research Fund and is gratefully acknowledged. References 1. Cornwall, H.R. and Wright, J.C., 1956a, Geologic Map of the Hancock Quadrangle, Michigan: U.S. Geological Survey, Mineral Investigations Field Studies Map MF-46, scale 1:24,000. 2. Cornwall, H.R. and Wright, J.C., 1956b, Geologic Map of the Laurium Quadrangle, Michigan: U.S. Geological Survey, Mineral Investigations Field Studies Map MF-47, scale 1:24,000. 3. Cannon, W.F., 1994, Closing of the Midcontinent Rift: a far-field effect of Grenvillian compression: Geology, v. 22, pp. 155-158. 4. Langfield, K.M., DeGraff, J.M., and Gamet, N.G., 2023, Slip kinematics of the Keweenaw and Hancock faults within the Midcontinent Rift System, Upper Peninsula of Michigan: Inst. on Lake Superior Geology, 69th An. Meeting, Eau Claire, WI, Part 1 – Program and Abstracts, v. 69, p. 50-51. 5. 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. 135, no. 1/2, p. 449–466. 71 �Lithostratigraphy and Geochronology of the Lower Northeast Sequence of the North Shore Volcanic Group, Cook County, MN, USA NOWARIAK, Eric S., SEVERSON, Allison R., BLOCK, Amy Radakovich Minnesota Geological Survey, Department of Earth and Environmental Sciences, University of Minnesota-Twin Cities, MN, USA The Lower northeast sequence (LNE) of the North Shore Volcanic Group (NSVG) in northeastern Minnesota represents some of the earliest known sedimentary and volcanic rocks of the Midcontinent Rift System (MRS) including the Puckwunge Formation (PF), Grand Portage Lavas (GPL), Esther Lake Lavas (ELL), and the Hovland Lavas (HL) (Miller and others, 2002). New 1:24,000 field mapping in northern Cook County, paired with U/Pb TIMS zircon geochronology have established an updated lithostratigraphic sequence of the LNE and understanding of its relationships with surrounding intrusive rocks, providing new insights on magmatic evolution and spreading rates during the Plateau Stage of MRS magmatism. The LNE is a shallow to moderately south dipping bimodal volcanic sequence that youngs southward and is segregated from the overlying Upper northeast sequence by the cross-cutting Brule-Hovland Gabbro, a complex of texturally and mineralogically varied gabbroic and diabasic intrusions (Fig. 1). To the north, the LNE is bounded predominantly by cross-cutting Early Stage Duluth Complex granophyric intrusions and locally by underlying Paleoproterozoic metasedimentary sequences. The arenitic sandstones and conglomerates of the PF, which forms the base of the LNE, lie unconformably on top of Paleoproterozoic bedrock and forms the base of the LNE on to which the NSVG rocks were erupted. GPL volcanics were erupted disconformably atop the PF, and consist of geochemically primitive basalts, as compared to overlying LNE volcanics (Mattis, 1972). Basal units of the GPL are defined by pillowed and fragmental olivine basalts, transitioning to thick flows of massive to sparsely amygdaloidal basalts with rubbly tops. The boundary between the GPL and the overlying ELL is marked by a distinct change from olivine tholeiite basalts to thick, pilotaxitic flows of amygdaloidal, oxide-rich ferroandesites and andesitic basalts which exhibit steeper REE 72 Figure 1. Schematic Stratigraphic Section of the LNE showing locations of geochronologic samples. See text for details. �profiles than GPL basalts. U/Pb zircon geochronologic analysis of a thin rhyolite flow interlayered with the massive andesitic basalts near the base of the ELL returned an age of ca. 1105.4 Ma (Fig. 1). Bimodal volcanics of the HL sequence above the ELL are typified by feldspar-phyric to glomeroporphyritic flows of basalt, basaltic andesite, trachyandesite, and rhyolites. Plagioclase phenocrysts within the mafic and intermediate lithologies are fractured, locally resorbed, contain inclusions of olivine and clinopyroxene, and have compositions estimated to be more calcic than the groundmass feldspars. Mafic and intermediate lithologies are locally pillowed, though most flows do not show evidence of subaqueous eruption. Rhyolites constitute a large volume of the upper half of the HL sequence. These rhyolites exhibit abundant autobreccia textures, flow foliation, pumice fragments, and phenocrysts of feldspar and quartz. Microscopically, HL rhyolites are texturally diverse, containing microlites, perlitic structures, fiamme, and spherules interpreted to be a product of volatile-rich lavas and pyroclastic flows. U/Pb zircon analyses of upper HL rhyolites returned ages of ca. 1105.69 Ma and 1106.0 Ma (Fig. 1). Basalt flows intercalated with the rhyolite units are thick, structured flows with basal breccias, columnar jointed massive flow centers, and amygdaloidal flow tops. Subvolcanic, plagioclase ultraphyric diabase dikes and sills of Burnell’s (1976) Brule Lake Porphyry (BLP) are abundant throughout the HL. The BLP contains 20-70% plagioclase phenocrysts hosted within a compositionally varied matrix, and although they have not been dated, a synvolcanic interpretation has been applied to these intrusions based on their local amygdaloidal nature and textural similarities to the HL. New ages from rhyolites in the basal portion of the ELL and upper portion of the HL at ca. 1106 Ma suggest these lavas are volcanic expressions of the coeval Early Stage Duluth Complex Cucumber Lake and Misquah Hills granophyres exposed to the north of the LNE (Vervoort and others, 2007). The Early Stage granophyres are interpreted to be a product of assimilation of continental crust (Vervoort and others, 2007). Such a change in magma composition is reflected in the lithologic and geochemical character of the ELL and HL as compared to the underlying GPL. The overlap in ages presented here suggests rapid eruption rates at ca. 1106 Ma, which indicate the entire ELL and HL sequence erupted within ~1 Ma. Phenocryst-rich volcanic and hypabyssal rocks throughout the HL and BLP may have been derived from anorthositic cumulates that formed in the roof of a mid-crustal staging chamber during a period of slow rift spreading, and subsequently remobilized during magma recharge and venting events at 1106 Ma. References Burnell, J.R., Jr., 1976, Petrology and structural relations of the Brule Lake intrusions, Cook County, Minnesota: Minneapolis, University of Minnesota, M.S. thesis, 105 p., 1 pl. Mattis, A. F., 1972. The Petrology and Sedimentation of the Basal Keweenawan Sandstones of the North and South Shores of Lake Superior. University of Minnesota – Duluth, M.S. thesis. Miller, James 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. 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(1-4), 235-268. 73 �Major element geochemistry and first zircon U-Pb age dates of Precambrian basement rocks in eastern North Dakota PEREIRA, Cristian1, NESHEIM, Timothy2, VERVOORT, Jeffrey D. 3, and SAINIEIDUKAT, Bernhardt1,4 1 Department of Earth, Environmental and Geospatial Sciences, North Dakota State University, Fargo, ND 58102, USA 2 North Dakota Geological Survey, 2835 Campus Rd., Grand Forks, ND 58202 USA School of the Environment: Earth Sciences, Washington State University, Pullman, WA 99164 USA 3 4 Dept of Chemistry and Biochemistry, North Dakota State University, Fargo, ND 58102, USA We are re-examining cores of the 1977 Red River Valley Drilling Project (Moore, 1978). Other previous work includes study of the paleoweathered horizon on the Precambrian bedrock (Kelley, 1980), Klasner and King (1986), Sims et al. (1991) and various ILSG abstracts. We sampled the Precambrian portions of three of these cores (RRVD #5, #8, #11) from eastern North Dakota, and a core cut by Kennecott Exploration Company in 2010 (10NDV001; Nesheim, 2013) (Fig. 1; Table 1). Samples were analyzed at Washington State University (WSU) for U-Pb zircon age dates. Kelley (1980) reported major element analyses and new analyses were carried out at WSU and North Dakota State University using XRF (Table 2). Figure 1. Precambrian geology map of North Dakota after Nesheim (2012) and Sims et al. (1991), with location map and core locations (stars) for this study Table 1. Summary of samples, lithology, and zircon age dates Core / depth Lat / Long lithology Zircon age (MSWD) RRVD 5 46.225514 Fine to medium grained 2715.0 +/- 18.1 Ma (3.4) 379 ft (115.5 m) -96.932504 quartz monzonite RRVD 8A 46.897502 Fine to medium grained 2782.7±9.3 Ma (0.72) 600 ft (182.9 m) -97.370662 chlorite gneiss RRVD 11 47.614926 Medium grained biotite 2 populations: younger 693 ft (211.2 m) -97.291738 granitoid 2671±23.2 Ma (3.0) 10NDV001 48.61706 Medium grained 2694.5±13.6 Ma (1.19) 837 ft (255.1 m) -97.316902 magnetite-rich granitic gneiss 74 �Table 2. Whole rock major element analyses wt.% 1 2 3 4 SiO2 74.00 68.91 64.30 65.80 TiO2 0.16 0.18 0.13 0.42 Al2O3 13.7 10.43 17.40 15.20 Fe2O3 1.36 2.04 4.67 FeO 1.65 MnO 0.01 0.31 0.02 0.10 MgO 0.04 0.44 0.71 0.00 CaO 1.49 4.60 0.91 1.65 Na2O 4.60 0.87 8.16 6.00 K2O 4.22 6.82 5.96 5.90 P2O5 0.09 0.03 0.06 0.05 SO3 0.04 LOI 5.44 sum 99.67 99.72 99.68 99.79 1: RRVD 5-383.5''; 2: RRVD 8A-602'; 3: RRVD 11-695'; 4: 10NDV001-836' Figure 2. U-Pb concordia diagrams for zircons from the Precambrian core samples with weighted mean 207 Pb/206Pb ages. Error ellipses represent 2SE uncertainties. Open ellipses with thick grey lines depict outlier U-Pb zircon analyses removed from final age determinations. The analyzed rocks contain 64-74 wt.% SiO2, with RRVD 11-695' showing high total alkalis (Na2O+K2O = 14.12 wt. %). All show Neoarchean zircon ages (2.7 –2.8 Ga) with the granitoids showing slightly younger ages than the gneisses (Table 1; Fig. 2). Sample RRVD 11-693 appears to be a 2 component rock with two zircon populations. These chemical results and measured ages are consistent with those measured in other areas of the Superior Craton (cf. Li et al., 2020). REFERENCES: Kelley, L.I., 1980, Kaolinitic weathering zone on Precambrian basement rocks, Red River Valley, eastern North Dakota and northwestern Minnesota. M.S. Thesis, University of North Dakota. 85 pp. Klasner, J.S. and E. R. King. 1986a. Precambrian basement geology of North and South Dakota. Canadian Journal of Earth Sciences. 23(8): 1083-1102. https://doi.org/10.1139/e86-109 Li, D., Hollings, P., Chen, H., Sun, X., Tan, C., and Zurevinski, S., 2020, Zircon U–Pb and Lu–Hf systematics of the major terranes of the Western Superior Craton, Canada: Mantle-crust interaction and mechanism(s) of craton formation, Gondwana Research, v. 78, p. 261-277. Moore, W. L., 1978, A preliminary report on the geology of the Red River Valley Drilling Project, eastern North Dakota and northwestern Minnesota: Bendix Field Engineering Company Subcontract H77-059-E, 292p. https://www.osti.gov/biblio/6538603 doi:10.2172/6538603 Nesheim, T., 2012, Review of Radiometric Ages from North Dakota’s Precambrian Basement. North Dakota Geological Survey Geologic Investigations No. 160. Nesheim, T., 2013, Recent Diamond Exploration in Eastern North Dakota. NDGS GeoNews, p. 5-7. Sims, P.K., Peterman, Z.E., Hildenbrand, T.G., and Mahan, S., 1991, Precambrian Basement Map of the Trans-Hudson Orogen and adjacent terranes, northern Great Plains, U.S.A.: USGS Miscellaneous Investigations Series Map, I-2214. 75 �The geology and ore deposit model of the high-grade Emily Manganese Deposit, Cuyuna Range, Minnesota: Results from the 2023 drilling program PETERSON, Dean1, and STEINER, Alex1 1 Big Rock Exploration, 2505 West Superior Street, Duluth, MN 55806. The Emily deposit is the highest-grade manganese resource in the USA. The deposit is located along the western margin of the Paleoproterozoic Animikie Basin (Southwick and Morey, 1991) and is hosted by the Emily Iron Formation, a shallow water Superior type iron formation. Recent work by Big Rock Exploration on a drilling program for Electric Metals has identified coherent zones of high-grade mineralization (30 to ≥40 wt.% Mn) over a 1.25-kilometer strike length. An ore deposit model has been developed that incorporates the deposition of primary thin-bedded manganese-iron carbonates (Fig. 1) and later conversion into massive manganese oxide through early folding (Fig. 2) and prolonged periods of weathering, oxidation, and erosion. Figure 1. Model of the primary depositional setting of the Paleoproterozoic Superior-type iron formations of the western Animikie basin. Figure 2. Schematic model for the formation of the Emily District thrust-front folds as related to the Penokean fold & thrust belt of the Cuyuna North and South iron ranges. 76 �Historic exploration and drilling in the 1940’s and 1950’s by Pickands Mather and US Steel identified iron and manganese-bearing mineralization within the Emily Iron Formation (Strond, 1959). US Steel developed but did not implement a preliminary mine plan for mining of the Emily Deposit. Following approximately 50 years of inactivity, Cooperative Mineral Resources (subsidiary of Crow Wing Power) pursued a pilot mining operation using pressurized water that ultimately proved unsuccessful. As a follow up investigation into the outcomes of the pilot mining, a small-scale drill program was completed in 2010-2012. An Emily deposit drilling program was designed and executed by Big Rock Exploration, LLC, in 2022-2023. A total of 29 drill holes were completed to extend mineralization and refine the previous resource estimates. A total of 13,107 feet of drilling was completed for this program. Data collected for this project includes lithological, structural, geotechnical, and geochemical data from drill cores as well as geophysical data from selected drill holes. Through interpretation of legacy, recent and new drilling data, Big Rock Exploration has redefined the stratigraphy of the Emily Iron Formation and developed an ore deposit model for the high-grade manganese oxides of the Emily deposit (Fig. 3). This ore deposit model and associated geological model have been used to support an updated and expanded mineral resource estimate for the Emily Deposit, to be completed by Forte Dynamics. Figure 3. Integrated stratigraphy, permeability, texture, and Mn-grade diagram for the Emily deposit. REFERENCES Southwick, D.L. and Morey, G.B., 1991, Tectonic imbrication and foredeep development in the Penokean orogen, east-central Minnesota; an interpretation based on regional geophysics and results of test drilling, U.S. Geological Survey Bulletin 1904-C, pp. C1–C17. Strong, R., 1959, Report on Geological Investigation of the Cuyuna District, Minnesota, 1949-1959, US Steel Internal Report, 318 pages. 77 �Deformation conditions, micromechanics, and fault zone development in mafic protoliths at the Lac des Iles mine, northwestern Ontario PETERZON, Jordan1, PHILLIPS, Noah1, HOLLINGS, Pete1, and DJON, Lionnel2 1 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 2 Faults and their associated damage zones are important geologic structures that serve as permeable pathways through the upper crust; however, the effect of host lithology on fault core development and damage zone structure remains poorly constrained. The development of fault cores and damage zones is typically controlled by the strength and composition of the protolith, conditions of deformation, and fluid chemistry (Caine et al., 1996). Fault zones are characterized by a variably developed fault core composed of unconsolidated gouge or silicified breccias, outward into a highly fractured damage zone and then a relatively unaltered protolith. Trapped mineralization may be offset or remobilized by later faulting. Faults may 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; Faulkner et al., 2010). Permeability is typically 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. This study examines deformation conditions and fluid-rock interaction of fault zones within the Lac des Iles Complex. The Lac des Iles Complex is a series of mafic-ultramafic intrusive bodies occurring within the Marmion terrane of the Superior Province (Figure 1). The complex has been dated at 2689 ± 1.0 Ma and was emplaced into a ~3.01 to ~2.68 Ga granitegreenstone terrane (Djon et al., 2018). Extensive Ni-Cu-PGE mineralization has been offset by two late reverse faults in the high-grade zones (>4 g/t Pd) called the Camp Lake fault and the Offset fault. A depletion in Pd is observed within the damage zone of each fault, approximately 145 – 180 meters from the actual fault. This depletion is likely due to late fluid flow within the damage zones. Fault cores in tonalite mainly are composed of breccias with calcite to quartz-rich matrix, while fault cores in gabbro are composed of chlorite-rich gouges (Figure 2). Fracture densities in felsic protoliths have a higher fracture density than mafic protoliths suggesting that fluid flow would be 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. We hypothesize that the development of a frictionally weak, chlorite-rich fault core impeded the development of a more fracture-dense damage zone in the gabbros. Electron microprobe analyses on chlorite grains reveal three generations of chlorite growth have occurred: pre-faulting at ~350°C, syn-faulting at ~150 – 200°C, and post-faulting at ~150°C (Figure 3). Elemental gains and losses from unaltered protolith to fault core were examined to understand the interactions between alteration and fault zones. Within fault cores and damage zones, there is an observable gain in Mg and Fe in mafic protoliths, due to the precipitation of new chlorite within the fault zone. In mafic protoliths, fluid-rock interactions play an important role in the development of fault core and damage zone structures. 78 �Figure 2 Variations in drill core with proximity to faulting. Figure 1 Local geology of the Lac des Iles Intrusion with faults of study. Modified from Djon et al., (2018). Figure 3 Results of chlorite thermometry from electron microprobe analyses. 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, and fluid flow properties of fault zones. Journal of Structural Geology, 32 (11): 1557-1575. 79 �Identification of Fertile Parent Granitoid Units in the Superior Province of Ontario PETTIGREW, Therese1, CUNDARI, Robert1, PRICE, Rebecca2, and DUGUET, Manuel3 1 Ontario Geological Survey, 435 James St. South, Thunder Bay, ON P7E 6S7 Ontario Geological Survey, 227 Howey St, Red Lake, ON P0V 2M0 3 Ontario Geological Survey, 933 Ramsey Lake Rd, Sudbury, ON P3E 6B5 Canada aw n 2 Peraluminous granites are widely distributed throughout the Superior Province of Ontario, most notably within and adjacent to the metasedimentary rocks of the English River and Quetico subprovinces from which they were derived by partial melting. There has been a significant amount of work that proposes a direct genetic relationship between peraluminous, S-type granitoids (i.e., fertile parent granites) and rare-element pegmatites of the lithium-cesiumtantalum (LCT) group across the world (see for instance Černý, 1989, 1991; Wise, Müller and Simmons, 2022, and references therein). W ith dr A fertile granite is the parental granite to rare-element pegmatite intrusions. Many granitic melts have the capability to generate fertile granite plutons that will, in turn, produce even more fractionated melts enriched in incompatible elements. In the case of the LCT group pegmatites, the residual melt may percolate into the surrounding host rock and crystallize rare-element pegmatites (Breaks, Selway and Tindle, 2003). Identifying fertile parent granites is an important step in the exploration for rare-element pegmatites as it greatly reduces the search area on a regional scale (Breaks and Tindle, 1997). A significant amount of work was performed by the Ontario Geological Survey (OGS) in the late 1990s and early 2000s to improve our understanding of rare-element pegmatites and their parent granitoid units in the Superior Province, with a focus on northwestern Ontario (e.g., Breaks, Selway and Tindle, 2003). In 2022, ten areas of the Superior Province in Ontario were identified for study as part of the fertile granite project (Figure 1). Locations for sampling were selected to complement the existing granitoid geochemical databases acquired by the OGS, as well as to provide coverage in areas previously not investigated for the presence of fertile granitoid rocks and associated LCT group pegmatites. The 2022 field season was intended as a preliminary investigation of the selected areas. A total of 100 samples were collected (Figure 1) and analyzed for major, trace element and rare earth element geochemistry at the Geoscience Laboratories (Sudbury) to identify potential fertile parent granite bodies. Due to several staffing changes during the spring and summer of 2023, focus on the project was delayed and did not resume until the winter of 2023-24. Further work in support of the fertile granite project will include preliminary evaluation of the geochemical data set generated during the summer of 2022, compilation of geochemical data from previous studies and planning for additional sampling during the 2024 field season. The primary deliverable of the project will be an MRD compiling previously released whole-rock geochemical data related to fertile granites. This compilation will be supplemented with the new whole-rock geochemical data acquired during this study. Additionally, several articles will be generated and released during the course of the project in both the Resident Geologist Program Recommendations for Exploration (released annually in January) and the Report of Activities (published annually in the spring). 80 �aw n dr W ith Figure 1. Locations of fertile granite project target areas (outlined in black) and sample locations (green dots) collected in 2022. Regional geology from Ontario Geological Survey (2011, see publication for a detailed geological legend). Subprovince boundaries are based on Stott (2011) and are outlined in blue (from Cundari, 2022). References 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, 179 p. Breaks, F.W. and Tindle, A.G., 1997. Rare-element exploration potential of the Separation Lake area: An emerging target for Bikita-type mineralization in the Superior Province of northwest Ontario; in Summary of Field Work and Other Activities, 1997, Ontario Geological Survey, Miscellaneous Paper 168, p. 72-88. Černý, P., 1989. Exploration strategy and methods for pegmatite deposits of tantalum; in Lanthanides, tantalum and niobium, Springer-Verlag, New York, p. 274-302. ——— 1991. Rare-element granitic pegmatites. Part 1: Anatomy and internal evolution of pegmatite deposits. Part 2: Regional to global environments and petrogenesis; Geoscience Canada, v.18, p. 4981. Cundari, R.M., 2022. Identification of Fertile Parent Granitoid Units in the Superior Province of Ontario: Project Description; in Summary of Field Work and Other Activities, 2022, Ontario Geological Survey, Open File Report 6390, p. 30-1 to 30-5. Ontario Geological Survey, 2011. 1:250 000 scale bedrock of Ontario; Ontario Geological Survey, Miscellaneous Release—Data 126 – Revision 1. Stott, G.M., 2011. A revised terrane subdivision of the Superior Province in Ontario; Ontario Geological Survey, Miscellaneous Release—Data 278. Wise, M.A., Müller, A. and Simmons, W.B., 2022. A proposed new mineralogical classification system for granitic pegmatites; The Canadian Mineralogist, v.60, p. 229-248. 81 �Lidar Topography: Bright opportunity for Reading Keweenaw Landscapes ROSE, Bill1 and DeGRAFF, James1 1Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 USA A GeoAtlas for Keweenaw, Houghton, and Baraga counties will soon be publicly accessible as an exciting new tool for understanding landscapes through lidar surveys and convenient geospatial display tools. In this presentation, we discuss hypotheses that emerged from a “first look” at this remarkable data which reveals greater detail than standard topographic maps. The purpose of this discussion is to stimulate robust investigation to build new geological awareness of geomorphology. The work may be a useful element for geoeducation because of the improved resolution. 82 �Figure 1: Lidar topography may reveal in situ differentiation within thick lavas of the Portage Lake Volcanics. Here, lidar data from the Keweenaw GeoAtlas is used with field mapping data from Cornwall (1951) and Longo (1984). These thick lavas reveal cooling of about 1000 years where the interior texture of the basalt is pegmatoidal, contrasting with ophitic textures near the bottom and top of the flow where solidification was faster. Arrows show the same crossections on lidar and geologic maps. Figure 2: Geology at Traverse Island in Keweenaw Bay. Left -aerial image from Google Earth with offshore tracing of Jacobsville Sandstone strata (yellow) and fractures (red). Onshore features are from Denning (1949). Right – onshore tracing of sandstone strata from 2-m resolution Lidar data. White outline is a caprock of “quartzite” with a possible channel form at its western edge. References Cornwall, H.R., 1951, Differentiation in Magmas of the Keweenawan Series, J Geol, v. 59, pp. 151-172. Denning, R.M., 1949, The Petrology of the Jacobsville Sandstone, Lake Superior: Michigan College of Mining and Technology [MTU], M.S. thesis, 71 p. 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. 83 �Michigan Coastal Path: A Social Commitment to Geoeducation ROSE, Bill1 and VYE, Erika2 1 Geological and Mining Engineering & Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 USA 2 Great Lakes Research Center, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 USA The geology of the Midcontinent Rift is beautifully exposed for researchers, teachers, students, and geotourists in the Keweenaw Peninsula and Isle Royale (http://carnegiekeweenaw.org/social-post/keweenaw-shorelines-bill-rose). Michigan holds title to the surrounding bottomlands of the Great Lakes under the Public Trust Doctrine in addition to a public trust interest in the shorelands up to the ordinary high water mark (OHWM). Michigan confronts the challenge of discerning the boundaries between public trust interests and private property rights at the shore (Norton et al, 2013). In 1968, the Michigan Legislature adopted an elevation-based approach for discerning the ordinary high water mark (OHWM). In 2005, the Michigan Supreme Court reaffirmed that Michigan's public trust interest extends up to the OHWM, but it left unresolved questions of exactly how the two methods of marking ordinary high water relate to one another, and precisely how far up the shore the state has authority to regulate private shoreline development extends. (Norton et al, 2011). Here we describe the definitions of high water lines for geologic discussion. How may both landowner and hiker amicably agree on the high water mark when we meet along the shore? Figure 1: Shoreline exposures reveal the rock details clearly - veins of calcite (near the Copper Harbor Light) and native copper (Washington Island). 84 �Figure 2: Shoreline of Copper Harbor Conglomerate on Manitou Island. The zonation and succession of shorelines may be seen and used to define the high water mark. References Norton, R. K., Meadows, G.A., and Meadows, L.A. (2013). The deceptively complicated “elevation ordinary high-water mark” and the problem with using it on a Laurentian Great Lakes shore. Journal of Great Lakes Research, Volume 39, Issue 4, pp 527-535. Norton, R.K. & Meadows, G.A. (2014). Land and water governance on the shores of the Laurentian Great Lakes. Water International 39:6, pages 901-920. 85 �Jacobsville Geoheritage is Globally Celebrated and Locally Loved ROSE, Bill1 and VYE, Erika2 1 Geological and Mining Engineering & Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 USA 2 Great Lakes Research Center, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 USA The Jacobsville Sandstone is a well-known red bed sandstone of Neoproterozoic age from Upper Michigan, USA (Cannon and Nicholson, 2001) and is part of the Keweenaw Supergroup related to the Midcontinent Rift System. The rift formed ~1100 Ma and is a ~3000 km long feature in North America, centered on the Lake Superior area. The Jacobsville is the youngest of the area’s Precambrian rocks and was deposited during the Rigolet Phase of the Grenvillian Orogeny (1010-980 Ma) (Hodgin et al 2022). Cliff exposures show crossbedding and channels interpreted as fluvial deposits. Jacobsville Sandstone was a fashionable building stone in much of Eastern North America. From 1885 to 1920, it was used in hundreds of prominent buildings including the famous Astoria Hotel in New York City (Eckert, 2000). It was mined from several quarry sites near Jacobsville, Michigan. The location is part of a significant geoheritage location where native copper has also been mined, valued, and utilized for thousands of years. The development of copper mining drove extensive immigration of Europeans to Upper Michigan. The Jacobsville quarries offered an alternative to underground employment in the booming mining industry of the Keweenaw. Since quarrying has ceased, Jacobsville quarries have been overgrown and are often overlooked. Highlighting the significance of these places and increasing access offers an opportunity to teach locals and visitors about Earth's history and natural/cultural resources. It connects people to a significant element of Keweenaw geoheritage often eclipsed by the history of copper mining. In recent years we have been building awareness of the geohistory and geoheritage of Jacobsville quarrying. This awareness is building educational outreach focused on this remarkable rock formation which features in many local towns. The International Union of Geological Sciences (IUGS) and UNESCO’s International Geoscience Programme (IGP) have announced that the Jacobsville Sandstone - a rock formation named for Jacobsville, Michigan - is now one of only 15 Global Heritage Stone Resources (GHSR) in the world and the first in the United States (Rose et al, 2017). Global Heritage Stone Resources (GHSR) are scientific designations created and managed by the Heritage Stone Subcommission – HSS (IUGS/IAEG) to enhance the geological knowledge, use, and conservation of natural stones of historical importance worldwide. Highlights of recent Jacobsville geoheritage efforts include boat tours that explore the rock exposures in spectacular cliff views and Michigan historic signage in Houghton and other towns. References WF Cannon and SW Nicholson, 2001, Geologic Map of the Keweenaw and Adjacent Area Michigan: U.S. Geological Survey Map I-2696, scale = 1:100,000. WI Rose, EC Vye, CA Stein, DH Malone, JP Craddock and S Stein (2017) Jacobsville Sandstone: A candidate for nomination for Global Heritage Stone Resource, Michigan, USA. Episodes 40 (3), 213-219 86 �K.B. Eckert, (2000). The Sandstone Architecture of the Lake Superior Region. Wayne State University Press, Detroit, USA, 344 p. EB Hodgin, NL Swanson-Hysell, JM DeGraff, ARC Kylander-Clark, MD Schmitz, AC Turner, Y Zhang, DA Stolper (2022). Final inversion of the Midcontinent Rift during the Rigolet Phase of the Grenvillian Orogeny. Geology 2022; 50 (5): 547–551 Fig 1: Red Jacket firehouse in Calumet (built in 1898–99) – National Register of Historic Places. Use through Creative Commons by Andrew Jameson. Fig 2: One of many cliff exposures of the Neoproterozoic Jacobsville Sandstone, here about 1 km N of the town of Jacobsville, 19 km SE of Houghton, in Michigan’s Keweenaw Peninsula. Cliff exposures are found in dozens of locations within Keweenaw Bay. Photo by Steve Brimm. Fig 3: Jacobsville Quarry near Portage Entry, in operation, about 1895 (MTU Neg 03965, Michigan Technological University Archives and Copper Country Historical Collections, Houghton, Michigan Michigan Technological University Archives). 87 �Compiled historical drillhole and geochemical data from the Cuyuna Range, Minnesota, provides powerful new insights for geological and mineral potential investigations. SAARI, Stacy1, GORDEE, Sarah1, RIAN, Madison1, and CARTER, Matt1 1 Minnesota Department of Natural Resources, Lands and Minerals, 1525 Third Ave. East, Hibbing, MN 55746 The Minnesota Department of Natural Resources (DNR) staff of Lands and Minerals (LAM) compiled a large tabular dataset of drill hole locations, geological logs, and geochemical data from the Cuyuna Iron Range of central Minnesota as part of the federally funded Earth Mapping Resources Initiative (Earth MRI). These data will help to determine where potential resources of iron and manganese, as well as other critical minerals, may exist in the underlying bedrock. Iron in the Cuyuna Range was discovered in 1903 by Cuyler Adams and was actively mined until 1984. Prior to the end of WW1, there were 37 active mines on the Cuyuna Range. The lack of outcrop hindered the geological understanding and definition of the resources, and over 12,000 exploration holes were drilled over this period (Morey et al., 1977). Some of the historical explorers and mining operators in this area included: Orelands Mining Company, Evergreen Mining, Pittsburg-Pacific, Pickands-Mather, Oliver Iron Mining (division of US Steel), Inland Steel, Rogers-Brown Company, and Zontelli Brothers (Sutherland, 2016). Since the end of active mining in this district, the DNR LAM office in Hibbing has accumulated thousands of mining and mineral exploration documents from various companies. Current estimates suggest that the Cuyuna Range is among the top three largest manganese occurrences in the United States, justifying continued interest in its resource potential (Strong, 1959; Beltrame et al., 1981; Kilgore and Thomas, 1982; Cannon et al., 2017). From the 1940s to 1960s, the US Bureau of Mines (USBM) assembled, coded, and entered location and geochemical data for about 40% of the 12,833 drill holes available from the USBM Minnesota Mineral Development Atlas. USBM used criteria such as depth of overburden, past mining activity, availability of manganese data, and the availability of the sample material to determine which data to prioritize. All data were entered from public land survey (PLS) sections if there were fewer than 80 drillholes. For PLS sections exceeding 80 drillholes, then only 5 drillholes were entered for each quarter-quarter. This compilation resulted in data for 5,045 drillholes across the entire Cuyuna Range (Morey et al., 1977). Further, in the early 1990s, the Minnesota Geological Survey (MGS) created several databases to compile the assays and geologic logs from the USBMs dataset. Their database contains around 12,000 holes which have limited assay data and generalized geologic logs. Many of these drillhole sites were also entered in the Minnesota Well Index. As part of the Earth MRI project the MGS transferred thousands of maps and drill logs to the DNR. Compiling and managing 10,000s of documents and drill hole locations is a complicated and enormous undertaking, however, the use of geospatial software (i.e., ArcGIS), artificial intelligence and machine learning, and MicroMine 3D modelling software has accelerated the process. Without these technological resources, it would be difficult and time consuming to track 88 �duplication among the various exploration documents as well as the rebranding of drillhole names by successive explorers. The DNR merged the USBM and MGS databases and used the MicroMine software to help identify and resolve missing intervals, overlapping intervals, missing or incorrect azimuth and inclination, drillholes without analytical data, missing total depth, values beyond the end of the drillhole, and coincident drillhole collars. It would be nearly impossible to uncover these errors by hand. DNR staff curated and compiled a collection of US Steel data from the 1950s that was not included in its entirety in either the MGS or USBM compilations. These and other historic exploration data added drill logs and assays for hundreds of holes to the DNR’s drillhole compilation, mainly from the Emily area. Intervals that were assayed from these holes were often missing geological information, likely because the alteration and mineralization made interpretation difficult for the earliest explorers. Any missing collar elevations were obtained from 2012 1-m Lidar, knowing that inaccuracies may exist from previously mined areas or existing stockpiles post-exploration. DNR staff consolidated lithology types by limiting modifiers related to alteration or mineralization and extracted all assay data to create a working 3D model of the Emily district. Not only will future users have access to this model, but they will also be able to easily search by key words, sort based on geochemical results, and view the geographical location of the data. The data compilation is the first of a three-phase project which will be followed by ground and airborne geophysics and later supported by petrographical, lithochemical and geochronological analyses. Subsequent geologic mapping, mineral potential evaluation, and a geological interpretation for the rest of the Cuyuna district will complete this project. The project will culminate in an Earth MRI data release and report in 2026. REFERENCES: Beltrame, R.J., Holtzman, R.C., and Wahl, T.E., 1981, Manganese resources of the Cuyuna Range, eastcentral Minnesota: Saint Paul, Minn., Minnesota Geological Survey Report of Investigations 24, 22 p. Cannon, W.F., Kimball, B.E., and Corathers, L.A., 2017, Manganese, in chap. L of Schulz, K.J., DeYoung, J.H., Jr., Seal, R.R., II, and Bradley, D.C., eds., Critical mineral resources of the United States—Economic and environmental geology and prospects for future supply: USGS Professional Paper 1802, p. L1–L28. Jirsa, M. A., Boerboom, T. J., Chandler, V. W., 2012, Geologic Map of Minnesota, Precambrian Geology, Minnesota Geological Survey Map S-22, 1:500,000. Kilgore, C.C., and Thomas, P.R., 1982, Manganese availability - Domestic: U.S. Bureau of Mines Information Circular 8889, 14 p. Morey, G.B., Broberg, J., Beltrame, R.J., and Holtzman, R.C., 1977, Manganese-Bearing Ores of the Cuyuna Iron Range, East-Central Minnesota, MGS Report of Investigation for Grant U.S.D.I., Bureau of Mines G0264002, 185 p. Strong, R., 1959, Report on Geological Investigation of the Cuyuna District, Minnesota, 1949-1959, U.S. Steel - Oliver Iron Mining Division, 301 p., 6 plates. Sutherland, Frederick E., 2016, The Cuyuna Range: Legacy of a 20th Century Industrial Community. Ph.D. thesis, Michigan Technological University, 271 p. 89 �Understanding the evolution of the upper Midwest Archean gneiss dome corridor using apatite, titanite, and monazite LA-ICP-MS U-Pb geochronology and microstructural analyses SALERNO, Ross1, CANNON, William1, SOUDERS, Amanda2, and THOMPSON, Jay2 1 U.S. Geological Survey, Reston, VA 20192, 2U.S. Geological Survey, Denver, CO 80225 The origin of the Archean gneiss dome corridor stretching across Minnesota, Wisconsin, and northern Michigan is an important question for understanding the Paleoproterozoic tectonic evolution of the upper Midwest. The formation of these gneiss domes was originally attributed to orogenic collapse during the Penokean orogeny (1.86-1.83 Ga) (Schneider et al., 2004). More but more recent work, however, indicates their exhumation may be more closely linked to a suite of younger structures and metamorphism which are broadly concurrent with the Yavapai event between 1.78-1.75 Ga (e.g., Tinkham and Marshak, 2004; Schulz and Cannon, 2007). In this study, we leverage new LA-ICP-MS U-Pb geochronology and microstructural observations using EBSD (electron backscatter diffraction) to shed light on the timing of metamorphism and deformation related to the exhumation of these domes to understand the broader tectonic framework in which they formed. We investigate a suite of metamorphosed and deformed rocks collected from both inside and adjacent to these gneiss domes in northern Michigan (Figure 1). We show that these rocks have metamorphic U-Pb ages ranging from Neoarchean to Mesoproterozoic, reflecting the prolonged tectonic history of the southern margin of Laurentia (Figure 2). In the Paleoarchean Watersmeet gneiss, titanite grains have U-Pb intercept ages of 2550±46 (2σ, n=36) Ma, concurrent with the Sacred Heart orogeny. The U-Pb concordia ages of apatite in the Watersmeet gneiss at 1869±32 (2σ, n=27) Ma, and monazite U-Pb ages in the Hardwood gneiss at 1826±21 (2σ, n=36) Ma, reflect metamorphism of these rocks during the Penokean orogeny. Several samples have apatite U-Pb concorida ages that indicate heating continued for tens of millions of years after the end of the Penokean orogeny at about 1830 Ma: at 1815±32 Ma (2σ, n=17) in the Republic trough, at 1803±29 Ma (2σ, n=49) in the Hardwood gneiss, and at 1796±29 (2σ, n=51)Ma in the Michigamme Formation directly adjacent to the Watersmeet dome. Our dataset documents the influence of post-Penokean orogenic events on the rocks of the gneiss dome corridor in northern Michigan. In the Neoarchean Carney Lake gneiss, migmatitic rocks have titanite U-Pb ages of 1752±71 Ma (2σ, n=53), indicating reactivation during the Yavapai orogeny. In the Solberg schist, in the Felch trough, titanite grains have recrystallized into aggregates of subgrains, likely formed during deformation. These titanite grains have U-Pb ages of 1713±32 Ma (2σ, n=82), which we interpret to reflect the timing of deformation-induced recrystallization. The U-Pb ages of apatite in the Solberg schist are markedly younger at 1588±28 Ma (2σ, n=46) and align with the timing of the Mazatzal orogeny. Together, these new U-Pb data add to a growing body of evidence that the present architecture of the gneiss dome corridor in the upper Midwest is at least in part due to post-Penokean orogenic events. 90 �Figure 1. Generalized geologic map of the study area (modified from Tinkham and Marshack, 2004) showing the locations of sample sites with yellow stars. Black dots show the position of towns W: Watersmeet, R: Republic, H: Hardwood, and M: Marquette. Figure 2. The LA-ICP-MS U-Pb ages of apatite, titanite, and monazite for the samples in this study. The vertical bars represent the timing of major orogenic events along the southern margin of the Superior craton and Laurentia. References Schneider. S., Holm. D., O’Boyle. C., Hamilton. M., Jercinovic. M., 2004, Paleoproterozoic development of a gneiss dome corridor in the southern Lake Superior region, USA: GSA Special Paper 380, 339357. Schulz. K., Cannon. W., 2007, The Penokean orogeny in the Lake Superior Region: Precambrian Research, 157, 4-5. Tinkham. D., Marshak. S., 2004, Precambrian dome and keel structure in the Penokean orogenic belt of northern Michigan, USA: GSA Special Paper 380, 321-338. 91 �Analysis of deformation-related structures in the Eau Claire Volcanic Complex, Wisconsin SHAKKED, Daniel, L.1, ROBARGE, Lucas, C.1 and LODGE, Robert W.D. 1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701, USA This study is focused on the Penokean-aged (1.8-1.9 Ga) deformation fabrics and microstructures in the Big Falls region of the Eau Claire Volcanic Complex (ECVC), Northwestern Wisconsin. Varying intensities of deformation and metamorphism within the Penokean Orogeny are extensive and well documented, particularly in the External Domain in the northern parts of the orogen. However, the southernmost regions are more poorly studied because outcrops are present as rare erosional outliers in river channels. Structural interpretation, and by association terrane boundaries, have largely been inferred from geophysical data. This project focused on describing the structural and metamorphic fabric development at two field areas: one studying the origin of the gneissic banding within the amphibolitic banded gneiss (Figure 1A), while the other being a study on the genesis of the migmatitic fabrics (Figure 1B). The goal of this research is to petrographically and geochemically interpret the deformation mechanisms of the ECVC and improve its tectonic context to the rest of the Penokean orogeny. There are two main volcanic terranes within the Penokean Orogeny and are sutured together by the Eau Pleine Shear Zone. The Pembine-Wausau terrane is characterized as a juvenile arc-system formed through subduction and was accreted onto the southern edge of the Superior Craton. This was followed by the accretion of the Marshfield Terrane; an Archean microcontinent overprinted by Paleoproterozoic magmatism. Historical interpretations of the ECVC suggested these rocks were deposited on the Marshfield Terrane, but recent geochemical and petrochronological studies show a mantle-derived, oceanic affinity (Lodge et al, 2023; Weber et al., 2023). Therefore, revisiting and reinterpreting the tectonic context of the ECVC to the Marshfield Terrane is warranted. The previous interpretation of the bedrock at Big Falls County Park suggests gneissic banding was inherited from igneous layering from a layered mafic intrusion (Cummings, 1984). This study describes field and petrographic observations that indicate intense ductile fabric development during shearing such as asymmetric inclusions, pressure shadows, and feldspar grain-boundary migration within the gneiss (Figure 1A, 1C). Comparison of these textures with other sheared amphibolites and amphibolitic gneisses support the interpretation that banding is, at least in part, caused by intense shearing (e.g. Bozkurt et al., 1997). Geochemical analysis of the Big Falls Region shows a hydrated, mantle-derived signature closely related to an E-MORB oceanic-arc system, indicating that these magmas are not derived from an Archean, continental fragment (Weber et al. 2023). Structural analysis of these rocks shows extensive fabric shearing and deformation, supporting the theory that these rocks are structurally emplaced. Petrographic and outcrop analysis of the tonalite intrusion and “lensoidal amphibolite” in the Little Falls region indicates the gneissic fabrics and banding are migmatites. Two possible theories for the genesis of the migmatites are suggested: anataxis of the amphibolite and granulite facies metamorphism (Ashworth 2011) or melt injection from the tonalite intrusion during the Penokean deformation. Previous interpretations of the Little Falls region indicated three episodes of metamorphism (Cummings 1984). Evidence of a weak foliation and recrystallization in the tonalite intrusion (Figure 1D) containing xenoliths of banded amphibolite gneiss suggests at least two metamorphic events. However, its relationship to the thermal event that formed the migmatites is uncertain. 92 �A B C D Figure 1: A: Outcrop photo of amphibolitic bands and sheared garnet-hornblende porphyroblasts in banded gneiss, Big Falls County Park, Wisconsin. B: Outcrop photo of migmatite at Little Falls County Park, Wisconsin. The neosome consists of quartz, plagioclase, and feldspars while the paleosome consists of hornblende, biotite, and chlorite in the picture. C: Photomicrograph (XPL, 10x) of strained feldspar crystal with grain boundary migration, and an amphibole band being deflected via shearing above the feldspar grain. D: Photomicrograph (XPL, 10x) of the tonalite intrusion at Little Falls which contains quartz, plagioclase, biotite, hornblende, and chlorite. References Ashworth, J.R., 2011, Migmatites. New York, NY, Springer, 371 pp. Bozkurt, E., and Park, R.G., 1997, Microstructures of deformed grains in the augen gneisses of southern Menderes Massif (western Turkey) and their tectonic significance: Geologische Rundschau: Zeitschrift für allgemeine Geologie, v. 86, p. 103–119. 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. Lodge, R.W.D., Weber, E.M., and Hooper, R.L., 2023, Precambrian Geology of the Eau Claire River Valley: Re-discovering the Eau Claire Volcanic Complex, in Lodge, R.W.D. ed., Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 2 – Field Trip Guidebooks, p.47-70. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian research, v. 157, p. 4–25. Weber, E.M., Lodge, R.W.D., and Marsh, J.H., 2023, U/Pb geochronology and zircon petrochronology of Paleoproterozoic magmas from the Marshfield terrane, Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 1-Program and Abstracts, p. 9798. 93 �Geochemistry and Petrology of the Eagle’s Nest Intrusion, McFaulds Lake Greenstone Belt, Northern Ontario SHESHNEV, Vlad1, HOLLINGS, Peter1, PHILLIPS, Noah1, WESTON, Ryan2, DELLER, Matt2, CAMPBELL, Dana2 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, Canada Wyloo Metals, 1-1127 Premier Way, Thunder Bay, ON, P7B 0A3, Canada 2 The Eagle’s Nest intrusion hosts economically significant orthomagmatic Ni-Cu-(PGE) mineralization located in the McFaulds Lake greenstone belt within the northern regions of the Archean Superior province, approximately 500km northeast of Thunder Bay, Ontario. The Eagle’s Nest intrusion is part of the 2734 Ma ultramafic-dominated Koper Lake subsuite of the larger Ring of Fire intrusive suite (Metsaranta et al., 2015; Houlé et al., 2020). The mineralized ore body of the Eagle’s Nest intrusion is zoned, with massive sulfide mineralization at its northwestern extent that gradationally becomes semi-massive, net-textured and disseminated to the southeast (Zucceralli et al., 2022). Mineralization consists of 11.1 million tonnes of proven and probable mineral reserves containing 1.68% Ni, 0.87% Cu, 0.89g/t Pt, 3.09g/t Pd and 0.18g/t Au (Burgess et al., 2012). The intrusion was emplaced along a sub-horizontal conduit, forming a blade-shaped dike (Barnes and Mungall, 2018). Mineralization is consistent with gravitational sulfide segregation at the basal, northwestern contact of the intrusion. Subsequent deformation rotated the intrusion into its present day, subvertical orientation, with a width of ~500m, thickness of ~150m and vertical extent >1600m. Parental magma composition of the Eagle’s Nest intrusion has been determined on a number of occasions but with contrasting outcomes. A low MgO komatiitic magma composition with ~22% MgO and ~12% total FeO was proposed by Mungall et al. (2010). In contrast, Zuccarelli et al. (2022) reported olivine compositions of Fo82-86 consistent with picritic parental magmas containing moderate MgO (10-20%) and high total FeO (>12%). The contrasting results means that parental magma composition of the Eagle’s Nest intrusion needs to be further constrained. The objectives of this study are to petrographically and geochemically characterize the (1) unmineralized portions of the Eagle’s Nest intrusion and (2) associated offshoot dikes in the vicinity, and (3) constrain the parental magma characteristics by using geochemical, petrographic, mineral chemistry, and radiogenic isotope techniques. A total of 136 samples were collected for this study. Forty-four samples came from offshoot dikes consisting of fine- to medium-grained mafic to ultramafic units. Eighty-seven samples were collected from within the intrusion comprising peridotite (Fig. 1), gabbro, and chilled margin samples. Lastly, five samples were collected from the host wall-rock of the intrusion which consists of tonalite. One-hundred and twenty-one samples were analyzed for major and trace elements using Inductively Coupled Plasma Atomic Emission Spectroscopy (ICP-AES) and Inductively Coupled Plasma Mass Spectrometry (ICP-MS), respectively. An initial batch of 30 thin sections was prepared consisting of 15 offshoot dikes, eight contacts, and seven peridotite samples. Twenty samples from the intrusion were selected for Sm-Nd isotopes. To constrain the parental magma composition, three approaches will be considered: (1) examination of preserved chilled margins along the length of the chonolith, (2) examination of chilled margins preserved in the magmatic breccia matrix within the stratigraphic hanging-wall 94 �of the intrusion, and (3) mineral chemistry of fresh olivine preserved within the peridotite horizons of the intrusion. The parental magma composition obtained from these three methods will be further compared to ensure consistency between the methods. The intrusion’s magmatic history will be constrained using the obtained Sm-Nd isotope data, which will provide further insights into the mantle source and contamination history of the melts that formed the Eagle’s Nest intrusion. Figure 1. Photomicrograph in crossed-polarized light (XPL) of a peridotite sample containing serpentinized cumulus olivine and poikilitic orthopyroxene with fresh olivine within the oikocryst. References Barnes, S.J. and Mungall, J.E. 2018 Blade-shaped dikes and nickel sulfide deposits: A model for the emplacement of ore-bearing small intrusions: Economic Geology, v. 113, p. 789 – 798. Burgess, H., Gowans, R., Jacobs, C., Murahwi, C. and Damjanović, B. 2012. Noront Resources Ltd.—NI 43–101 technical report feasibility study—McFaulds Lake Property, Eagle’s Nest Project, James Bay Lowlands, Ontario, Canada: Micon International Ltd., 197p. Houlé, M.G., Lesher, C.M., Metsaranta, R.T., Sappin, A.-A., Carson, H.J.E., Schetselaar, E.M., McNicoll, V., and Laudadio, A., 2020. Magmatic architecture of the Esker intrusive complex in the Ring of Fire intrusive suite, McFaulds Lake greenstone belt, Superior Province, Ontario: Implications for the genesis of Cr and Ni-Cu-(PGE) mineralization in an inflationary dyke-chonolith-sill complex: Geological Survey of Canada, Open File 8722, p. 141–163. Metsaranta, R.T., Houlé, M.G., McNicoll, V.J., and Kamo, S.L., 2015. Revised geological framework for the McFaulds Lake greenstone belt, Ontario: Geological Survey of Canada, Open File 7856, p. 61–73. Mungall, J.E., Harvey, J.D., Balch, S.J., Azar, B., Atkinson, J., and Hamilton, M.A., 2010, Eagle’s Nest: A magmatic Ni-sulfide deposit in the James Bay lowlands, Ontario, Canada: Society of Economic Geologists Special Publication, v. 15, p. 539–557. Zuccarelli, N., Lesher, C.M., Houlé, M.G., Weston, R. and Barnes, S.J. 2022. The diversity of nettextured sulfides in Magmatic Sulfide Deposits: Insights from the Eagle’s Nest Ni-Cu-(PGE) Deposit, McFaulds Lake greenstone belt, Superior Province, Canada: Economic Geology, v. 117 (8), p. 1731 – 1759. 95 �New LA-ICP-MS U-Pb geochronology of Archean rocks, central Upper Peninsula, Michigan, USA: a step toward refining the final assembly of the Superior craton SOUDERS, A.K.1, CANNON, W.F.2, DRENTH, B.J.1, SALERNO, R.A.2, THOMPSON, J.M.1, SYLVESTER, P.J.3 1 U.S. Geological Survey, Denver, CO 80225 USA (asouders@usgs.gov) U.S. Geological Survey, Reston, VA 20192 USA 3 Texas Tech University, Lubbock, TX 79409 USA 2 The central Upper Peninsula, Michigan consists of two contrasting Archean terranes of the Superior Province: the granite-greenstone terrane of the Wawa-Abitibi Subprovince in the north (Northern Complex) and the gneisses of the Minnesota River Valley Subprovince (MRVS) in the south (Southern Complex). The two terranes are separated by the Great Lakes Tectonic Zone (GLTZ). The suturing of the MRVS to the southern margin of the Superior Province and development of the GLTZ, long interpreted to have occurred at about 2.69 Ga, has more recently been suggested to be related to the ca. 2.58 - 2.6 Ga Sacred Heart Orogeny (Schmitz et al. 2018; Cannon et al. 2024, this volume), a proposal that we are still evaluating. In this study we present new LA-ICP-MS U-Pb geochronology for Archean crystalline rocks from both the Northern Complex and Southern Complex, across the GLTZ. This age characterization is essential to define/refine the regional geochronologic framework of ‘basement’ rocks in the central Upper Peninsula. This is essential to understand subsequent geologic processes. Heavy mineral separates were produced using Electro Pulse Dissagregation (EPD) followed by heavy liquid separation at Zirchron (AZ, USA). Individual zircon grains were hand-picked and mounted in 25 mm epoxy resin mounts and polished to a 1 µm finish. All samples were imaged via cathodoluminescence in the Denver Microbeam Lab (USGS) using the JEOL 5800 LV SEM. LA-ICP-MS analyses were made using a Nu AttoM sector field ICP-MS coupled to a NWR 193 ArF excimer laser system in the Mineral Isotope Laser Laboratory (MILL) at Texas Tech University. Typical laser ablation conditions during all analytical runs were a fluence of 3 J/cm2, 8 Hz, and 240 laser pulses using a 15 µm laser spot. Data was reduced using Iolite v.4 (Paton et al. 2011) and final age calculations were made using IsoplotR (Vermeesch, 2018). We are presently working on zircon LA-MC-ICP-MS Hf isotope analyses to characterize the source components of Archean crystalline rocks from the Northern Complex and Southern Complex. Zircon grains from nine rocks in the Southern Complex and six rocks in the Northern Complex were targeted for LA-ICP-MS U-Pb analysis. Examples of samples analyzed from Southern Complex granitic gneisses are shown in Figure 1. For all samples, the age spectrum of concordant grains is complicated with multiple inherited populations within a single sample. This observation is like that presented by Ayuso et al. (2018) for the ca. 2700 Ma Carney Lake gneiss and ca. 2750 Ma Hardwood gneiss, south of our current study area. Examples of crystalline samples analyzed from the Northern Complex are shown in Figure 2. A single age population for zircon grains analyzed with little to no evidence of Early or Middle Archean inheritance is common for Northern Complex basement samples. These data support the fundamental difference between the primitive volcanic/plutonic terrane of the Northern Complex and an older Archean continental crustal component in the Southern Complex. 96 �Figure 1. Examples of a subset of Archean granitic rocks sampled from the Southern Complex, MI Figure 2. Examples of a subset of Archean granitic rocks sampled from the Northern Complex, MI References Ayuso, R.A, Schulz, K.J., Cannon, W.F., Woodruff, L.G., Vazquez, J.A., Foley, N.K., 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. ILSG, Proceedings 64th Annual Meeting. Cannon, W.F., Souders, A.K., Drenth, B.J., Ayuso, R.A. (2024) The Sacred Hearth Orogeny in Michigan: Latest Archean Granites and the Great Lakes Tectonic Zone. ILSG, Proceedings 70th Annual Meeting. Paton, C., Hellstrom, J., Paul, B.,Woodhead, J. and Hergt, J. (2011) Iolite: Freeware for the visualisation and processing of mass spectrometric data. JAAS. doi:10.1039/c1ja10172b. Schmitz, M.D., Southwick, D.L., Bickford, M.E., Mueller, P.A., Samson, S.D. (2018) Neoarchean and Paleoproterozoic events in the Minnesota River Valley subprovince, with implications for southern Superior craton evolution and correlation. Precambrian Research, v.316, p. 206-226. Vermeesch, P. (2018) IsoplotR: a free and open toolbox for geochronology. Geoscience Frontiers. doi: 10.1016/j.gsf.2018.04.001. 97 �Geochemical fingerprints from the late Mesoproterozoic epeiric seaway of the Nonesuch Formation, Wisconsin and Michigan, USA STEWART, Esther K.1,2, TAPPA, Michael1, BAUER, Ann1, PRAVE, Anthony3, and BRENGMAN, Latisha4 1 Department of Geoscience, University of Wisconsin-Madison, Madison, Wisconsin 53706, USA Wisconsin Geological and Natural History Survey, UW-Madison Division of Extension, Madison, Wisconsin 53705, USA 3 School of Earth and Environmental Sciences, University of St. Andrews KY16 9TS, Scotland/UK 4 Department of Earth and Environmental Sciences, University of Minnesota-Duluth, Duluth, Minnesota 55812, USA 2 The Oronto Group (Copper Harbor Conglomerate, Nonesuch Formation, and Freda Formation) of the southern Lake Superior Region preserves an exceptional record of late Mesoproterozoic environments and associated microfossils (e.g., Cumming et al., 2013; Fedorchuk et al., 2016; Strother and Wellman, 2021). A lacustrine rift basin is often cited as the most plausible depositional setting for the ca. 1080 Ma Nonesuch Formation because of its association with alluvial deposits of the underlying Copper Harbor Conglomerate and overlying Freda Formation and because of its location within interior Laurentia (Elmore et al., 1989; Slotznick et al., 2023). Our recent sedimentologic and stratigraphic evidence demonstrates deposition of the lower Oronto Group within a tide- and wave-influenced estuary (Stewart et al., 2023, see also Hieshima and Pratt, 1991; Jones et al., 2020). New geochemical results from Nonesuch Formation carbonates (Figure 1), including strontium (Sr), carbon (C), and oxygen (O) isotope compositions, rare earth element - yttrium (REY) patterns, and trace element ratios complement and add new dimension to this environmental interpretation. Strontium isotope compositions refine the Precambrian marine 87Sr/86Sr curve (Chen et al., 2021), with the Nonesuch recording relatively radiogenic compositions at ca. 1080 Ma between previously reported values of 0.706600 at ca. 1109 Ma and 0.705965 at ca. 1058 Ma. Most shale-normalized REY patterns from Nonesuch Formation carbonates are characterized by positive lanthanum anomalies and elevated yttrium: holmium (Y/Ho) ratios. Many of the same samples are also enriched in heavy REE, while others record light REE enrichment. These patterns indicate Nonesuch Formation carbonates precipitated from brackish water, consistent with REY patterns observed in modern estuaries (Lawrence and Kamber, 2006). One sample has a flat shale-normalized REY distribution and likely precipitated within part of the estuary dominated by fluvial input. The combined geochemical evidence suggests Nonesuch Formation carbonates were minimally altered by diagenesis, and diagenetic alteration was dependent on sedimentary facies. While C and O isotopes are uncorrelated, initial Sr isotope compositions correlate positively with O isotopes, and C isotopes group by sedimentary facies. Minor diagenetic alteration thus resulted in less radiogenic Sr isotope compositions and did not impact C isotope compositions, which instead reflect facies-dependent incorporation of remineralized, isotopically light organic carbon during deposition or early diagenesis. Although ƩREY correlates with initial Sr isotope composition, there is no covariation between initial Sr isotope composition and lanthanide anomalies or Y/Ho ratios. This implies that carbonates likely precipitated in shallow pore waters where ƩREY was modified by contribution from surrounding detrital material, and REY profiles were determined by original pore water chemistry in connection with the overlying water body. 98 �Figure 1. Example of carbonate sampled for geochemistry from the Nonesuch Formation. (a) core scan showing fine-grained siliciclastic sediment (dark gray) and carbonate (light cray). Note molar tooth crack cross-cutting laminae. Core is 1 inch (2.54 cm) wide. (b) photomicrograph showing carbonate spar (light color) infilling molar tooth crack. Plain polars, scale bar is 1000 µm. Modified from Stewart et al. (2023). Chen, X., Zhou, Y., Shields, G.A., 2022. Progress towards an improved Precambrian seawater 87Sr/86Sr curve. Earth-Science Reviews 224, 103869. Cumming, V.M., Poulton, S.W., Rooney, A.D., Selby, D., 2013. Anoxia in the terrestrial environment during the late Mesoproterozoic. Geology 41(5), 583-586. 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 43(3), 191-213. Fedorchuk, N.D., Dornbos, S.Q., Corsetti, F.A., Isbell, J.L., Petryshyn, V.A., Bowles, J.A., Wilmeth, D.T., 2016. Early non-marine life: evaluating the biogenicity of Mesoproterozoic fluvial-lacustrine stromatolites. Precambrian Research 275, 105-118. Hieshima, G., Pratt, L., 1991. Sulfur/carbon ratios and extractable organic matter of the middle Proterozoic Nonesuch Formation, North American Midcontinent rift. Precambrian research 54(1), 65-79. Jones, S., Prave, A., Raub, T., Cloutier, J., Stüeken, E., Rose, C., Linnekogel, S., Nazarov, K., 2020. A marine origin for the late Mesoproterozoic Copper Harbor and Nonesuch Formations of the Midcontinent Rift of Laurentia. Precambrian Research 336, 105510. Slotznick, S.P., Swanson-Hysell, N.L., Zhang, Y., Clayton, K.E., Wellman, C.H., Tosca, N.J., Strother, P.K., 2024. Reconstructing the paleoenvironment of an oxygenated Mesoproterozoic shoreline and its record of life. Bulletin 136(3-4), 1628-1650. Stewart, E.K., Bauer, A.M., Prave, A.R., 2023. End-Mesoproterozoic (ca. 1.08 Ga) epeiric seaway of the Nonesuch Formation, Wisconsin and Michigan, USA. Geological Society of America Bulletin. Strother, P.K., Wellman, C.H., 2021. The Nonesuch Formation Lagerstätte: a rare window into freshwater life one billion years ago. Journal of the Geological Society 178(2). 99 �Characteristics of graphitization across a metamorphic gradient in the Michigamme Formation of the Marquette Trough and Baraga Basin, MI STOKES, Rebecca1, CANNON, William1, SALERNO, Ross1 1 U.S. Geological Survey, Geology, Energy and Minerals Science Center, Reston, VA 20192 Graphitization of carbonaceous material (CM) occurs by the progressive aromatization of carbon, expulsion of heteroatoms, and three-dimensional stacking of graphene layers — a process that alters the structure, isotopic, and trace element chemistry of the residual CM. Graphitization is generally considered thermally driven and has been extensively studied in the context of predictable changes in crystallinity as measured by Raman spectroscopy or X-ray diffraction, ultimately yielding the development of several graphite geothermometers (Henry et al., 2019). More broadly, crystalline graphite is an industrial mineral used in lithium-ion batteries and critical for the energy transition away from fossil fuels. The efficacy of graphite in the anode of batteries is directly related to its physicochemical properties which are a function of its geologic origin. The variably metamorphosed and deformed black shales and slates of the Michigamme Formation provide a natural laboratory to revisit our understanding of the graphitization process and the associated changes in structure and chemistry of CM in the context of technological applications. The Michigamme Formation is a Paleoproterozoic metasedimentary and metavolcanic sequence that is widespread in the Upper Peninsula of Michigan. The lower part of the formation is finer-grained black shale and siltstone, typically with a prominent slaty cleavage. In the Marquette Trough and Baraga Basin, the focus areas of our study, this fine-grained sequence has been mapped and formally designated the Lower Slate Member of the Michigamme Formation. Highly carbonaceous units are ubiquitous near the middle of this member and vary from ~70 meters on average along the Marquette Trough to 150 meters or more in the Baraga Basin. Along the Marquette Trough, the outcrop trace of the Lower Slate transects a regional metamorphic gradient from chlorite to staurolite grade whereas the metamorphic grade in the Baraga Basin is uniformly low, within the chlorite zone. In the Marquette Trough, the graphitic beds sampled for this study are along the north limb of this complex syncline, mostly dip steeply southward, and have a well-developed slaty cleavage that is axial planar to the larger structure of the trough. Deformational features diminish to the north and the northernmost of our samples, in the Baraga Basin, are from nearly flat lying beds with no penetrative structural features. An important and widely accepted distinction of the Michigamme Formation is that the development of the penetrative regional cleavage predates the final metamorphic event. Thus, graphitization of the CM occurred under both stressed (tectonic) and static conditions. A suite of thirteen core samples from the Upper Peninsula Geological Repository and three outcrop samples, all from carbonaceous sections of the Michigamme Formation, were selected for detailed evaluation (Figure 1). In all samples, CM occurs as disseminated and elongated fine-grained (<20 µm) particles that tend to be concentrated in bands parallel to the metamorphic fabric defined by phyllosilicates and quartz. Analysis of total carbon on decarbonated samples yielded values ranging from 1 to 24 wt.% C, with an average value of 7 100 �wt.%. Carbon isotopic analysis from decarbonated samples yielded a general trend towards heavier δ13C values with increasing metamorphic grade, ranging from -32.05‰ (Sample 4) to 21.85‰ (Sample 13). Raman spectroscopic analysis of CM yielded a similar trend with metamorphic grade across the sample suite. The R2 ratio, which is one parameter used to evaluate metamorphic grade, decreases from 0.64 in Sample 15 to 0.35 in Sample 8. These R2 values correspond to a temperature range from ~325°C to ~500°C using the empirical geothermometer from Aoya et al. (2010). Notably, Samples 11 and 13 from the garnet zone are significant outliers in both the Raman and C isotope datasets. Combining these results with additional data from scanning electron microscope imaging, X-ray diffraction mineralogical analysis, and laser ablation ICP-MS analysis of CM concentrates will yield a more detailed look at the evolution of CM with metamorphism and deformation. These results will help refine our understanding of the geologic processes that lead to economic graphite deposits and fine-tune graphite deposit models with an application focus. Figure 1. Geologic map and metamorphic isograds (red lines) of the field area in the Upper Peninsula region of Michigan. Map is generalized from Cannon and Ottke (1999). Core samples are noted with white dots, and blue dots for outcrop samples. References Aoya, M., Kouketsu, Y., Endo, S., Shimizu, H., Mizukami, T., Nakamura, D., Wallis, S., 2010. Extending the applicability of the Raman carbonaceous material geothermometer using data from contact metamorphic rocks. Journal of Metamorphic Geology, 28: 895–914. 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. Henry, D. G., Jarvis, I., Gillmore, G., & Stephenson, M. (2019). Raman spectroscopy as a tool to determine the thermal maturity of organic matter: Application to sedimentary, metamorphic and structural geology. Earth-Science Reviews, 198: 102936. 101 �Sulfur-isotope ratios in Paleoproterozoic Michigamme Formation at the Lake Superior Region: Implications on basin evolution and ambient seawater composition in the Greater Animikie Basin THAKURTA, Joyashish 1, and HAAG, Beau 2 1 Natural Resources Research Institute, University of Minnesota Duluth, 5013 Miller Trunk Highway, Hermantown, MN 55811, USA 2 Niblack Project LLC, 136 River Street, Elko, Nevada 89801, USA A considerable variation in δ34S-ratios of sulfide minerals has been found in a sulfide-mineralrich succession of slate, metasiltstone, and metagreywacke in the Paleoproterozoic Michigamme Formation located within the Baraga Basin at the eastern edges of the Greater Animikie Basin in the Lake Superior Region (Ojakangas, Morey, and Southwick, 2001). Sulfide minerals such as pyrite, chalcopyrite and pyrrhotite display δ34S-values ranging between 2 and 40‰ (V-CDT) and appear to systematically vary with respect to the stratigraphic intervals (Figure 1). This variability is gradational and devoid of any anomalous spike. It is predominantly a function of stratigraphic location and it shows no relationship with the type of sulfide mineral or the textural mode of occurrence. This observation rejects the possibility that the δ34S values were influenced by selective infiltration of externally derived sulfur-rich fluids along the stratigraphic layers of the Michigamme Formation. It also overrules the possibility that the observed δ34S values were caused by low-grade metamorphism and recrystallization of the sedimentary rocks of the Michigamme Formation. The values are consistent with primary δ34S ratios in sulfide minerals which were precipitated from intergranular fluids within a sequence of clastic sediments in a basin shortly after deposition. Consequently, the measured δ34S ratios in the stratigraphic horizons represent Sisotopic signatures inherited from the S-reservoir in the ambient basin, as well as changes introduced by basin-evolution and diagenesis of the siliciclastic sediments in a marine foreland depositional environment along the eastern portion of the Greater Animikie Basin. While the observed general trend of gradual increase in δ34S-values in the lower and upper members of the sequence can be explained by a systematic chronological trend in the global seawater composition (Paiste et al., 2020), a significant rise and fall in δ34S-values in the Lower Slate and Upper Greywacke Members can be attributed to a limited-term separation of the Baraga Basin from an open-ocean circulation to an isolated basin environment in response to structural adjustments caused by the formation of a fold and thrust belt along the southern shore line of the foreland basin during the waning stages of the Penokean Orogeny (Schulz and Cannon, 2007). In this period of isolation, the sulfur-isotope composition was primarily controlled by intrabasinal bacterial fractionation leading to a significant rise in the δ34S values up to 40‰ in the observed sediments. Upon subsequent erosional removal of the thrust sequence, and associated structural readjustments, the connection of the Baraga Basin to an open ocean was restored and the δ34S-values in the new sedimentary rocks mimic values that are consistent with deposition in a larger open ocean setting. 102 �Figure 1: Observed variation in δ34S-ratios. Stratigraphy adapted from Rossell and Coombes (2005) References: 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 141-142: 319-341. Paiste, K., Lepland, A., Zerkle, A.L., Kirsimae, K, Kreitsmann, T, Mand, K., Romashin, A.E., Rychanchik, D.V. and Prave, A.R., 2020. Identifying global vs. basinal controls on Paleoproterozoic organic carbon and sulfur isotope records. Earth-Science Reviews 207: 01320. Rossell, D. and Coombes, S., 2005. The Geology of the Eagle Nickel-Copper Deposit Michigan, USA. Report for Kennecott Exploration, dated April 29, 2005, 35 p. Schulz, K.J. and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, 157: 4-25. 103 �An evaluation of structural and mineralogical controls on gold mineralization on the GoldRich property in the Abbie Lake area, Wawa, Ontario THIBODEAU-BELLO, Demily, HILL, Mary Louise, CONLY, Andrew Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The GoldRich property in the Abbie Lake area is an active gold prospect in the Michipicoten greenstone belt, within the Wawa subprovince of the Archean Superior province of the Canadian shield. The property is located 30 km northeast of Wesdome’s Eagle River mine and 10 km northeast of Wesdome’s Mishi property in northern Ontario. The Main Shear trench on the GoldRich property is an area of regional metamorphism dominated by ductile deformation hosting orogenic gold. The Main Shear trench hosts mylonites of felsic and intermediate composition with varying strain intensities; all lithologies strike east-west. The mylonite of intermediate composition is characterized by having en-echelon quartz veins perpendicular to the foliation. Gold occurrences have been found in zones of high strain, in both felsic and intermediate mylonite lithologies. This HBSc thesis project relies on detailed trench mapping, microstructural and petrographic analyses, and geochemical methods to discover how gold is hosted on this property. Understanding the controls on gold mineralization will guide future exploration. Gold mineralization in the Main Shear trench is related to deformation. Based on geochemical and petrographic analysis there is no correlation observed between alteration mineralogy and gold mineralization. Microstructural analysis revealed pervasive subgrain-rotation quartz recrystallization in both mylonitic lithologies indicating that the Main Shear trench has undergone deformation at the upper greenschist to lower amphibolite facies regional metamorphic conditions. Gold is associated with deformation, microstructurally related to recrystallized quartz grain boundaries, boudinaged veins, and shear bands. Based on these results, it is recommended that further exploration on this property should be focused on locating zones of similar structure and strain intensity, including areas of boudinage, regardless of lithology, to continue building the gold prospect. 104 �Petrology and Geochemistry of Felsic Magmatism in the Paleoproterzoic Eau Claire Volcanic Complex, Northcentral Wisconsin VICKERS, Lyndsie A. 1, LODGE, Robert W.D.1 1 Department of Geology and Environmental Sciences, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA The 1.8-1.9 Ga Eau Claire Volcanic Complex (ECVC) (Figure 1) is the type locality for Penokean-age magmatism formed on an Archean crustal block (~2.6-3.0 Ga) called the Marshfield Terrane during the Penokean Orogen (Sims et al., 1989; Schulz & Cannon, 2007) and has influenced historic tectonic models and terrane-boundary maps. The other volcanic terrane within the Penokean Orogen, the Pembine-Wausau terrane (PWT), is interpreted to have formed with minimal influence of older crust and hosts about 150 million tonnes of volcanogenic massive sulfide (VMS) ores. The ‘continental’ setting of the Marshfield Terrane assumes a different metallogenetic system than the ‘oceanic’ setting of PWT and may be less prospective for the same VMS mineralization. However, recent U/Pb isotopic and other geochemical data (Lodge et al., 2023; Weber et al, 2023) indicates parts of the ECVC were mantle-derived and not contaminated by older Archean crust and challenges this ‘continental’ model. The ECVC is challenging to study because of a lack of mineral exploration (and drilling) coupled with rare outcrop exposure due to glacial/fluvial sediment and Paleozoic rock cover. This project studies remote, inaccessible outcrops along the Eau Claire River to refine the tectonic model and terrane boundaries of the southern Penokean Orogen. Samples obtained from mapping were petrographically characterized and analyzed for major and trace element geochemistry. Thirteen samples were analyzed for major and trace elements via WD-XRF and compiled with other geochemical datasets from the region. Trace element geochemical data are used to determine magmatic and tectonic settings of these rocks and improve regional tectonic models for the ECVC. Felsic magmatism in the region consists of fine-grained quartz-muscovite schists (Figure 1A) and banded quartzofeldspathic gneisses interpreted to be felsic volcanism and mediumgrained massive granodiorite (Figure 1B). Fine-grained quartz-muscovite schist are characterized by approximately 5% quartz porphyroclasts that are 1-3mm in size. The matrix is very finegrained quartz, feldspar, and muscovite that can have variable amounts of chlorite. Banded quartzofeldspathic gneisses have fine grained biotite and amphibole in quartz and feldspar-rich matrix that may define volcaniclastic textures at the outcrop-scale. Medium-grained, massive granodiorite is 15% biotite/hornblende and is characterized by weak to minimal foliation. These granodiorites are interpreted to be intruding felsic volcanic rocks, amphibolites, and metasedimentary units (Figure 1C). At one outcrop, the contact region is exposed revealing the formation of a megabreccia matrix, incorporating gneiss fragments that can reach up to 1 meter in size. These gneissic metasedimentary rocks are fine-grained with alternating bands of light and dark layers, ranging from straight to intensely folded. Samples from the ECVC on Zr/Ti vs. Nb/Y classification diagrams reveal a bimodal magmatic suite which is commonly associated with extensional tectonic settings. Felsic volcanic and intrusive rocks on Nb vs. Y discrimination plots suggest that felsic magmatism was likely formed in syn-collisional or volcanic arc settings. Rhyolite fertility discrimination diagrams (Zr/Y vs. Y) show that both felsic volcanic and intrusive suites from the ECVC are FII-type rhyolites, typical of upper-crustal melting in rift zones. Therefore, the ECVC region may be prospective for VMS mineralization and the tectonic setting should be further evaluated. 105 �Figure 1. Bedrock geologic map of the Eau Claire River region in northcentral Wisconsin showing extent of Precambrian bedrock. Map from Mudrey & Brown (1982). (A) Poorly exposed quartz-phyric micaceous schist near Rock Dam, WI interpreted as a metarhyolite. (B) Weathered surface of granodiorite intrusion along bank of Eau Claire River, (C) Contact region between granodiorite and dark gneiss or metasedimentary rocks. References Lodge, RWD, Weber, EM, Hooper, RL (2023), Precambrian Geology of the Eau Claire River Valley: Re-discovering the Eau Claire Volcanic Complex. in Lodge, RWD (Ed.), Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 2 – Field Trip Guidebooks. v.69, part 2, p.47-70. Mudrey, M.G., Jr., Brown, B. A., Greenberg, J. K. "Bedrock Geologic Map of Wisconsin." Wisconsin Geological and Natural History Survey, 1982. 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. 2145-2158. Weber, EM, Lodge, RWD, Marsh, JH (2023). U/Pb geochronology and zircon petrochronology of Paleoproterozoic magmas from the Marshfield terrane, Penokean Orogen, Wisconsin. Institute on Lake Superior Geology Proceedings, 69th Annual Meeting, Eau Claire, Wisconsin, Part 1-Program and Abstracts, p. 97-98. 106 �The Keweenaw Geoheritage Summer Internship Experience VYE, Erika1, LIZZADRO-MCPHERSON, Daniel2, and JUIP, James2 1 Great Lakes Research Center, Michigan Technological University, 1400 Townsend Drive, Houghton, MI, 49931 2 Geospatial Research Facility, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 Earth science education benefits from holistic interpretation of geologic features, processes, and landscapes through multiple ways of knowing (Deloria & Wildcat, 2001; Morton & Gawboy, 2003; Ricci & Riggs, 2019; Semken, 2005). Geoheritage is an evolving field that emphasizes the importance of the varied personal values people have for geologic features and explores the wide-ranging relationships we have with landscapes; as such it is an excellent place-based education tool to explore connections to our underpinning geology (Semken et al, 2017; Tormey, 2019). In this project setting, the Keweenaw region of Michigan’s Upper Peninsula on Lake Superior, we demonstrate a regional, place-based approach to help deepen participant understanding of the billion-year-old geologic processes at the heart of the Midcontinent Rift system. These processes created both the Lake Superior basin and the largest known native copper deposit on Earth in a region further defined as the ancestral and contemporary homelands and waters of the Keweenaw Bay Indian Community (KBIC) on Lake Superior. The Keweenaw Geoheritage Summer Internship Experience was created in partnership with the Keweenaw Bay Indian Community's Natural Resources Department (KBIC NRD) and Tribal Historic Preservation Office (THPO), and Michigan Tech’s Great Lakes Research Center (GLRC) and Geospatial Research Facility (GRF). The experience was created to support intergenerational and multicultural learning about the Keweenaw landscape, its stories, and geology. Tribal and non-tribal high school student interns, community partners, and knowledge holders spent time together reading the landscape and sharing reflections on our varied relationships with land and water. Week 1 entailed a 5-day field experience visiting valued sites of the Keweenaw Bay Indian Community that also teach us how geology impacts land, life, and culture in our place asking “what gifts does geology offer us? what are the relationships with land and water in this place?”. During the field experience, youth interns collected and documented local knowledge by engaging in multi-sensory and multimedia documentation strategies to record their experiences (photos, audio recordings, drawing, 360 virtual reality images, etc.). Sacred Anishinaabe knowledge was not sought or shared in these experiences. Data collected during the field experience was then used as the foundation for a 5-day geospatial workshop following the field experience. The goal of the workshop was for youth interns to design, create, and publish ARC GIS StoryMaps depicting their personal reflections of the diverse relationships with local landscape and understanding of its formation. ARC GIS StoryMaps is a story-authoring, web-based application that enables sharing of maps in the context of narrative text and other multimedia content. Workshop participants worked with the data they collected during the field experience in combination with supporting data layers and 107 �maps specifically created for the experience. Interns were mentored by the geospatial research team who helped students develop digital storytelling skills, inspired brainstorming sessions for topics to explore in the maps, and facilitated peer-review of StoryMap content prior to publication. Upon completion, students presented their work at a community open house; all StoryMaps created have been peer-reviewed by all project partners and are now published. The StoryMaps reflect a deepened understanding of relationships between geology, mining, and current environmental justice issues within our community. In the context of the Keweenaw, the European copper mining boom is most prominently interpreted in our place; students also reflected on the long history of mining, ways of mining, changing narratives, and missing human stories seen and experienced when visiting our landscape. Of note, respect, gratitude, and deepened relationships with land and water featured in all StoryMaps. Students shared reflections on reciprocity and their responsibility to help steward their place. Figure 1: Left - students deepen their understanding of mining impacts to Buffalo Reef and our local communities; Right: students brainstorm story arcs for the foundation of their StoryMap References [1] Deloria, V. and Wildcat, D. R. (2001). Power and place: Indian education in America. Fulcrum Publishing: Golden, Colorado. [2] Morton, R. and Gawboy, C. (2003). Talking Rocks: Geology and 10,000 Years of Native American Tradition in the Lake Superior Region. University of Minnesota Press. [3] Ricci, J. and Riggs, E.M. (2019). Making a Connection to Field Geoscience for Native American Youth through Culture, Nature and Community. Journal of Geoscience Education, special theme issues on Diversity in the Geosciences, DOI:10.1080/10899995.2019.1616273. [4] Semken, S. (2005). Sense of place and place-based introductory geoscience teaching for American Indian and Alaska Native undergraduates. Journal of Geoscience Education, 53 (2), 149-157. [5] Semken, S., Geraghty Ward, E., Moosavi, S. and Chinn, P.W.U (2017). Place-Based Education in Geoscience: Theory, Research, Practice, and Assessment. Journal of Geoscience Education, 65:4, 542562, DOI: 10.5408/17-276.1. [6] Tormey, D. (2019). New approaches to communication and education through geoheritage. International Journal of Geoheritage and Parks, 7 (4), 192-198, ISSN 2577-4441, https://doi.org/10.1016/j.ijgeop.2020.01.001. 108 �R.W. Boyle’s History of Geochemistry and Cosmochemistry WILSON, Graham C.1, BUTT, Charles R.M.2, GARRETT, Robert G.3 and ROBINSON, Heather A.4 1 Turnstone Geological Services Limited. P.O. Box 1000, Campbellford, ON K0L 1L0 Canada, CSIRO Minerals Resources, Kensington, Western Australia; 3 Geological Survey of Canada, 601 Booth Street, Ottawa, ON K1A 0E8 Canada, 4 25 Chester Crescent, Ottawa. ON K2J 2J6 Canada 2 Robert W. Boyle (1920-2003) was a well-respected geochemist with a long career at the Geological Survey of Canada. He is perhaps best remembered for geological and geochemical studies of gold, silver and other commodities, and for his association with mining camps, such as Yellowknife, N.W.T. and Keno Hill, Yukon. Retiring in 1985, Boyle devoted much of his final two decades to trips to far-flung libraries, gathering information exotic and/or obscure, and penning a major 3-volume review of the evolution of human knowledge of the nature and use of metals and other materials, and of diverse fields within geochemistry, cosmochemistry and biogeochemistry. The result, with copious help in compilation, was almost 2,000 pages of discussion (90% of it typed, fortunately!), backed by a formidable bibliography of almost 3,000 references. It was essentially complete at the time of his death, but altogether lacking in illustration. In time, his G.S.C. colleague Bob Garrett made some editorial notes, and then, in 2011, Charles Butt assessed the manuscript and made a detailed scientific and editorial review, heavy in marginalia. However, the work evidently arrived too late for the glory days of Survey and Society printing budgets, and it sat upon the shelf. In 2015, Ryan Noble, of the Association of Applied Geochemists, broadcast the existence of the manuscript, which attracted Wilson, who undertook to advance the work of the earlier editors, with encouragement from Boyle’s daughter, biochemist Heather Robinson. Volume 1 of the trilogy is set for publication in 2024 (Boyle, 2024). It covers the vast span of human time from the inception of mining and agriculture to the fall of Rome in the West (476 A.D.), and so ventures onto ground traditionally left to aspects of the Classics, Ancient History and Archaeology. Despite the western time frame, it is a worldwide review, covering, besides Europe and the Middle East, India and China and the Americas, every part of the globe where Boyle found relevant knowledge to impart. The intended volumes 2 and 3 explore, respectively, history through the critical 19th century, and then the 20th century (and so to the present). Volume 1 traverses the long development of early thought on the nature of matter. In addition to the various, often conflicting strains of philosophy, there is an equal treatment of the harnessing of materials (Stone, Bronze and Iron ages), and the early stages of the broad swathe of Earth sciences, mining and metallurgy. Early practical ideas on “Earth, air, fire and water” are discussed, e.g., geochemistry and mineralogy, cosmochemistry (meteorites), and early ideas on the hydrosphere and atmosphere. How was the raw manuscript processed? In brief, Wilson: a) ported Boyle’s references into a database, the easier to split up the long bibliography by chapter, rendering each section and volume a stand-alone story; b) utilised his MINLIB bibliography to update Boyle’s references, which for Volume 1 had ended in 1987; c) split up the seminal chapter 3, which provides reviews for some 29 metals and commodity groups (e.g., Au, Ag, Cu; Fe; Sn, Pb; industrial minerals, gemstones and organics); d) added a third layer of editing and consistency checks; and e) 109 �ultimately added 132 individual or composite illustrations in 93 numbered figures, including two original versions of the periodic table. Some of the additions (mostly to post-1987 research) are inserted in the text, others are collected in endnotes to each section. Some of the additions may be skating on thin ice (in which case, it is Wilson who falls through), a problem that one suspects would not have unduly worried the author of the original text. Boyle himself travelled widely across Canada, and the world. In terms of the Lake Superior region, Volume 1 has multiple references to the native copper and native silver of the Mesoproterozoic Midcontinent Rift (Fig. 1; see, e.g., Bornhorst and Barron, 2013; Wilson, 2023). One of two additional text boxes is devoted to native copper, while the other concerns the wider literature on the chemical elements, including some of the most accessible, popular titles. An explicit reference is made to native elements, of which Boyle was fond (e.g., Zn, Pb), including the obvious starting point of Au, Ag and Cu, and listing some 30 elements (many of them very rare in their unalloyed forms). Figure 1. Samples from famous occurrences of native metals in the Lake Superior region. Left: A spectacular, 4,264-kilogram mass of native copper, the exterior coloured by secondary Cu salts (Calumet, Keweenaw peninsula, Michigan). Right: native silver revealed in sawn and polished faces of calcite-veined fractured diabase from the Silver Islet mine in northwestern Ontario, a rich but short-lived venture on the east side of the Sibley peninsula, east of Thunder Bay, in Lake Superior. References Bornhorst, T.J. and Barron, R.J. (2013) Geologic overview of the Keweenaw peninsula, Michigan. Institute on Lake Superior Geology, v. 59, part 2: 1-42, Houghton, MI. Boyle, R.W. (2024) A History of Geochemistry and Cosmochemistry. Prehistory to the end of the Classical Period. Cambridge Scholars Publishing, Newcastle upon Tyne, England (Wilson, G.C., Butt, C.R.M., Garrett, R.G. and Robinson, H.A., editors), circa 600pp., in press. Wilson, W.E. (editor) (2023) Michigan Copper Country II. Mineralogical Record, v. 54 no.1: 196pp. 110 �Cooling of an Archean metamorphic terrane: garnet diffusion study of the Quetico Subprovince, Canada XU, Yiruo1 and HOLDER, Robert1 1 Department of Earth and Environmental Sciences, University of Michigan, 1100 North University Avenue, Ann Arbor, MI 48109 United States Archean metamorphic terranes are traditionally suggested to have cooled significantly slower than their Phanerozoic counterparts. Many have argued that the contrast in metamorphic timescale reflects changes in Earth’s tectonic regime. However, diffusion chronometry-based cooling rate data on Precambrian rocks are very limited. We present a case study of metamorphic timescales on the Neoarchean Quetico metasedimentary belt of the Superior Province, which has been hypothesized to represent a fore-arc accretionary prism. We combine conventional thermobarometry and phase-equilibrium modeling to constrain the peak temperature and pressure and estimate metamorphic cooling rates from major element diffusion in garnet. We then compare cooling rates across the Quetico Subprovince and with those from Phanerozoic metamorphic terranes of similar conditions. The results will contribute to the diffusion chronometry data available on Precambrian orogens for assessing any fundamental change in global tectonics. 111 � https://digitalcollections.lakeheadu.ca/files/original/462b1934e715e86830529c08f5039d63.pdf d45ee759251a8ad2fecc3255f6c3c6ab PDF Text Text 70th Annual Meeting Institute on Lake Superior Geology Houghton, Michigan May 15-18, 2024 Proceedings Volume 70 Part 2 - Field Trip Guidebook �70th Annual Meeting Institute on Lake Superior Geology Houghton, Michigan May 15-18, 2024 Sponsored by: A. E. Seaman Mineral Museum Great Lakes Research Center Department of Geological and Mining Engineering and Sciences Michigan Technological University Meeting Co-Chairs Theodore J. Bornhorst, Erika Vye, Patrice Cobin, and James DeGraff Proceedings Volume 70 Part 2: Field Trip Guidebook Compiled by Patrice F. Cobin and Theodore J. Bornhorst Cover Photo: The only known color photograph of in situ colorless calcite crystals with inclusions of native copper. Vug is about 15 cm across and 30 cm deep; located at the top of the Knowlton basalt lava flow at the 4th level, 850 ft stope, of the Caledonia Mine, Michigan. Photo taken in 1994 soon after the vug was blasted open. Native copper in the calcite crystals has not been visibly altered despite being about 1 billion years old. Photograph by Theodore J. Bornhorst i ��70th Institute on Lake Superior Geology Volume 70 consists of: Part 1: Program and Abstracts Part 2: Field Trip Guidebook Trip 1: Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan Trip 2: Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan Trip 3: Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty Trip 4: Keweenaw Fault System Geometry and Kinematics: Clues to Its Nature and Origin Trip 5: Geology and History of a Native Copper Mine: Adventure Mine, Ontonagon County, Michigan Trip 6: Southern Complex Granitoids, Gneisses, and Migmatites: New Data, Discoveries, and Perspectives Trip 7: Landslides on the Ontonagon River at Military Hill Reference to material in Part 2 should follow the example below: Authors, 2024, field trip title, 70th Institute on Lake Superior Geology, Abstracts and Proceedings, v. 70, Part 2, Field Trip Guidebook, p. xx-xx. Proceedings Volume 70, Part 1: Program and Abstracts and Part 2: Field Trip Guidebook are published by the 70th 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 ��Part 2: Field Trip Guidebook Table of Contents Trip 1: Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan………………………...1 Trip 2: Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan……………………………………………………….55 Trip 3: Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty…………………………………………………………………….79 Trip 4: Keweenaw Fault System Geometry and Kinematics: Clues to Its Nature and Origin………………………………………………………………………………..97 Trip 5: Geology and History of a Native Copper Mine: Adventure Mine, Ontonagon County, Michigan.…………………………………………………………137 Trip 6: Southern Complex Granitoids, Gneisses, and Migmatites: New Data, Discoveries, and Perspectives…………………………………..………….157 Trip 7: Landslides on the Ontonagon River at Military Hill…………………………………..173 iii ��Field Trip 1 Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan Theodore J. Bornhorst A.E. Seaman Mineral Museum, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Introduction The geology of the far western Upper Peninsula of Michigan consists of three temporally distinct episodes. During the Mesoproterozoic, about 1.1 Ga, up to 30 km of Keweenaw Supergroup volcanics and clastic sediments filled an intracratonic rift, the Midcontinent Rift (MCR) (Figures 1 and 2) (Cannon et al., 1989). After about 500 million years of erosion, the MCR rocks were buried by Phanerozoic sedimentary rocks from about 500 Ma to 175 Ma (Catacosinos et al., 2001). There are no exposures of rocks in the interval between 175 Ma to about 2.5 Ma (Velbel, 2009). Pleistocene continental glaciations, beginning about 2 million years ago, removed the Phanerozoic rocks from the Keweenaw Peninsula leaving only a few Phanerozoic outliers. About 10,000 years ago glaciers retreated from the Lake Superior basin and left behind a variety of unconsolidated clastic sediments. The geologic evolution of the far western Upper Peninsula is illustrated in Figure 3. Figure 1: Generalized geologic map of the Midcontinent Rift showing Grenville tectonic zone with interpretative cross-section in Figure 2. Modified from Bornhorst and Barron (2013). 1 �Figure 2: Generalized bedrock map showing the exposed rocks of the Midcontinent Rift around Lake Superior, the native copper occurrences, and the bedrock of the Upper Peninsula of Michigan modified from Bornhorst and Barron (2013). Interpretative cross section from Cannon et al. (1989). 2 �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). In the strictest sense, the geographic area of the Keweenaw Peninsula proper extends from L’Anse northwest to Lake Superior perpendicular to strike of the strata, however, the term Keweenaw Peninsula has also been applied to the area containing MCR rocks farther to the south from L’Anse to the White Pine area (Figure 4 and 5). The geologic descriptions in this field trip guide are mostly restricted to the Keweenaw Peninsula proper. The descriptions provided here were modified from a combination of Bornhorst and Barron (2011, 2013), Bornhorst and Lankton (2009), Bornhorst and Rose (1994) and Bornhorst et al. (1983). These sources are mostly used here without specific citation or quotation. 3 �Figure 4: Bedrock geologic may of the western Upper Peninsula of Michigan showing area of the Keweenaw Peninsula native copper district (from Bornhorst and Barron, 2013). Figure 5: Stratigraphic column the Keweenaw Peninsula, Michigan. 4 �Midcontinent Rift Strata The Keweenaw Peninsula is located on the southern margin of the Lake Superior segment of the MCR (Figures 1and 2). The rock units that are associated with the MCR have been termed the Keweenawan Supergroup (Figure 5). These rocks were deposited about 1.1 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 (Figures 2 and 3) (Hinze et al., 1990; Cannon et al., 1989). The MCR geology of the Keweenaw Peninsula can be divided into northwest-dipping, rift-filling volcanic and clastic sedimentary rocks under the central highlands and northwest flank of the Keweenaw Peninsula (Figure 4) and flat to low-dipping, rift-flanking clastic sedimentary rocks located on the southeast side. The Keweenaw Fault separates the rift-filling and rift-flanking strata. The rift-filling strata are subdivided into volcanic-dominated and clastic sedimentary rockdominated lithologies. (Figure 5). Portage Lake Volcanics The Portage Lake Volcanics of the Keweenaw Peninsula (Figures 4 and 5) are 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 which are located stratigraphically near the base of the exposed formation. Interflow reddish-colored conglomerate and sandstone layers are less than 5 % by volume and are stratigraphically scattered throughout the Portage Lake Volcanics although 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 near 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 et al., 1997 and reference therein). Much of the Portage Lake Volcanics erupted over 2 to 3 million years from 1,096.2+/-1.8 (Copper City flow, Figure 6) to 1,094.0+/-1.5 (Greenstone flow, Figure 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. Magmatic differentiation after eruption is especially significant in 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 of the Portage Lake Volcanics is cyclical with minor and major cycles superimposed on an overall trend. The basalt magmas were derived by partial melting of sub-continental upper mantle with an overall compositional trend towards younger more primitive basalt compositions as a result of less crustal contamination (Paces, 1988; Paces and Bell, 1989). The repeated magmatism at the rift axis and progressive crustal thinning provided pathways for magma with less extended contact with crustal rocks. 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 5 �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 enriched plume-related mantle (Nicholson et al., 1997; Nicholson and Shirey, 1990; Paces and Bell, 1989). 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. There is one thin hyaloclastic unit in the upper part of the formation (Johnson, 1985). 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 (Figure 6). The Greenstone flow has been correlated down dip across the Lake Superior syncline to Isle Royale (Longo, 1982; Huber, 1975). Table 1: Average representative geochemical data for least altered lavas of the Portage Lake Volcanics (from Paces, 1988). Tholeiites were grouped by Ni content. Primitive olivine tholeiite Intermediate olivine tholeiite Iron-rich olivine and quartz tholeiites Olivine tholeiite Olivine tholeiite 400-300 300250 250200 200-100 100-15 n=5 n=9 n=14 n=8 SiO₂ 47.82 47.34 48.03 Al₂O₃ 15.89 15.27 FeOt 9.77 MgO Andesite Dacite Rhyolite n=6 n=1 n=1 n=1 48.55 49.94 56.39 68.44 77.89 15.32 15.12 13.28 13.78 15.17 12.77 11.82 12.32 12.86 14.91 9.87 4.46 1.11 12.44 11.69 9.85 9.06 7.78 5.52 1.14 0.17 CaO 10.58 10.24 10.16 9.65 6.64 5.10 1.40 0.04 Na₂O 2.04 2.10 2.25 2.31 2.91 3.94 4.74 3.67 K₂O 0.19 0.22 0.33 0.42 1.43 2.27 3.86 4.28 TiO₂ 0.98 1.13 1.35 1.60 2.34 1.83 0.51 0.08 P₂O₅ 0.16 0.19 0.22 0.25 0.36 1.00 0.19 0.01 MnO 0.14 0.16 0.16 0.18 0.24 0.30 0.08 0.01 Ni 326 279 231 172 54 10 7 5 Cu 37 51 73 86 126 5 13 61 Zr 78 85 101 126 212 430 573 145 Ni (ppm) Wt.% PPM FeOt=total Fe as FeO 6 �The uppermost 5 to 20% of the tops of most individual lava flows are vesicular with between 5 and 50% vesicles (White, 1968). 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 are largely filled with secondary minerals, except for the stratigraphically uppermost lava flows; the filled vesicles are amygdules. Thus, local terminology is to call lava flows with vesicle-only tops, amygdaloids and those with brecciated tops fragmental amygdaloids. There are 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 into the exposed Portage Lake Volcanics. The rhyolitic volcanic setting is analogous to the shield-type central volcanoes of Iceland (Nicholson, 1991). Interflow clastic sedimentary rocks layers of the Portage Lake Volcanics are recognized as informal members since they are important stratigraphic markers in an otherwise monotonous succession of basalt lava flows. Many of them are given informal names (Figure 6). A few of them can be traced along strike for large distances, up to 90 km. These interflow sedimentary rock layers consist of redcolored conglomerates with lesser amounts of interbedded sandstone and occasional significant amounts of siltstone and shale. These informal members range in thickness from a few cm up to about 40 m (Merk and Jirsa, 1982; White, 1968; Butler and Burbank, 1927). The typical conglomerate is characterized by sub-rounded to angular pebbles in a sandy matrix. Clast size varies from granules 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, as 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 margins of the MCR toward the center of the rift basin (now under Lake Superior) (White, 1968). Copper Harbor Conglomerate The Copper Harbor Conglomerate is the oldest formation of rift-filling clastic sedimentary rocks and conformably overlies and interfingers with the top of the Portage Lake Volcanics (Figures 4 and 5). It consists of red-brown clastic sedimentary rocks with a maximum exposed thickness 2,000 m. Conglomerates and sandstones are the dominant lithologies in the Copper Harbor Conglomerate. 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 coalescing alluvial fans (Elmore, 1984). The diamictites are debris flow in origin. Sandstone interbeds are more common in the upper 2/3 of the formation. Sandstones are predominantly red-brown, subangular-to- 7 �angular lithic graywackes with volcanic lithic fragments. The sandstones exhibit current-ripples, trough-cross beds, current and parting lineations, and reduction spots. Abundant calcite cement in select conglomerate and coarse sandstone layers was probably deposited as vadose carbonate or caliche (Kalliokoski, 1986). Thin red-colored siltstone and shale interbeds have desiccation cracks and are interpreted as periodic drying of the surface. In the Copper Harbor area, there are also laminated cryptoalgal carbonate beds and ooid lenses. 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 stromatolites (genus Colleria). Figure 6: Generalized stratigraphy of the Portage Lake Volcanics in a strike parallel (longitudinal) section. Modified from Stoiber and Davidson (1959). Figure 4 shows location of Greenland-Mass subdistrict (Michigan, Caledonia, Mass, Adventure Mines). For decades The Copper Harbor Conglomerate CHC and overlying Nonesuch Formation (Figure 5) have been interpreted by many geologists as non-marine. Elmore (1984) interpreted the environment as a prograding coalescing non-marine 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). However, a number of sedimentological features could be interpreted as either non-marine or marine and thus, non-marine interpretations often relied on other evidence (Jones et al., 2020). In the stromatolite interval, Jones et al. (2020) cite bimodal (herring-bone) transport 8 �directions indicated by ripple marks that are mud draped and reactivated as evidence of a shallow marine environment. Hummocky cross stratification suggests waves on a marine shelf generated by storms (Jones et al., 2020). Periodic to rhythmic sedimentological features are indicative of “cyclical periodicity” of tidal deposition on a marine shoreline and are among evidence cited by Jones et al. (2020). Jones et al. (2020) conclude that the Copper Harbor Conglomerate and overlying Nonesuch Formation were deposited in a “braided fluvial-evaporitic shoreline-marine embayment” rather than fluvial-non-marine lacustrine setting. Geochemical evidence provided by Stüeken et al. (2020) also supports a marine estuary. The climate was probably arid with flashy seasonal streams. The highlands to the southeast from which the Copper Harbor Conglomerate was derived are now buried under the Jacobsville Sandstone. The Copper Harbor Conglomerate in the Keweenaw Peninsula includes a succession of subaerially deposited lava flows. Lane (1911) used the name, Lake Shore Traps, for this informal member (Figure 5). This member is well exposed near the tip of the Keweenaw Peninsula where the unit is composed of 31 lava flows and one interflow conglomerate with a maximum thickness of about 600 m (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 passive subsidence with rift-filling clastic sedimentation and little to no magmatism except for the Lake Shore Traps. These subaerial lava flows range from Fe-rich olivine tholeiitic basalt at the base to Fe-rich olivine-bearing tholeiitic basaltic andesites and tholeiitic andesites and are likely a shield volcano. Geochemical data are best 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 Conglomerate (Figures 4 and 5). It consists of dominantly black-to-gray-to-green to redgray siltstone and shale with a maximum thickness 240 m. Bornhorst and Williams (2013) provide a stratigraphic column of the entire Nonesuch Formation just south of the Porcupine Mountains State Park from exploration drilling. Exposures of the Nonesuch Formation in the Keweenaw Peninsula proper are limited with the best exposure at the Hancock campground and boat launch on M-203 (Stop 14). There are excellent exposures of the Nonesuch Formation along the Big Iron and Presque Isle rivers in the White Pine area (Woodruff et al., 2013). In areas with thicker stratigraphic section, 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, 2013). Well-laminated to massive black to dark-gray siltstone and shale are the dominant lithologies near the base of the Nonesuch Formation. The base of the Nonesuch Formation hosted 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, 2013; Williams and Bornhorst, 2023). A thin carbonate laminate yielded a Pb-Pb isochron age of 1,081 ± 9 Ma (Ohr,1993). The environmental setting of the Nonesuch is described above under the Copper Harbor Conglomerate. 9 �Freda Sandstone In the Keweenaw Peninsula the Freda Sandstone is the youngest rift-filling clastic sedimentary rock formations (Figures 4 and 5). The contact between the lower most Freda Sandstone and Nonesuch Formation is gradational (Bornhorst and Williams, 2013). The exposed thickness is greater than 3,700 m, with the top of the formation submerged beneath Lake Superior. The Freda Sandstone is generally poorly exposed except along the Lake Superior shoreline. This field guide provides an optional stop 13 at the McLain State Park where exposures of the Freda are visible during times of low levels of Lake Superior. Angular and tabular specimens are obtainable at the beach from outcrops just offshore. The last MCR magmatism was Bear Lake, an intrusive-extrusive dome of alkaline trachyandesite was emplaced near the middle of the exposed Freda Sandstone (Kulakov et al., 2018). Red-brown fine to very-fine sandstone, siltstone, and mudstone are the dominant lithologies in the Freda Sandstone. Fining-upward sequences occur on the scale of a few meters. The Freda Sandstone was deposited in an environment characterized by shallow meandering streams (Daniels, 1982). Based on regional correlations the Freda was likely deposited between 1,080 to 1,060 Ma. Jacobsville Sandstone The Jacobsville Sandstone was deposited in a rift-flanking basin (Figure 3D) and is outside the scope of this field guide. Its stratigraphic relationship with other formations is not determined. It occurs in a contiguous geographic region bound on the northwest side by the Keweenaw Fault and on the southeast by an unconformable contact with Paleoproterozoic and Archean basement rocks (Figure 4). The Jacobsville Sandstone is estimated to be more than 2,900 m thick and the top is not exposed (Kalliokoski, 1982). 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). The Jacobsville Sandstone was deposited in an environment characterized by shallow meandering streams (Kalliokoski, 1988). Faults, Folds, Fractures The last episode of the Midcontinent Rift was characterized by post-rift compressional inversion that facilitated hydrothermal formation of native copper deposits (Woodruff et al., 2020; Bornhorst, 1997). This compression transformed original normal faults into reverse faults, reactivated other extensional rift-related faults/fractures, and produced new compression-only faults/fractures and folds. Rather than being an inverted rift-related normal fault, the Keweenaw fault was likely a detached thrust (DeGraff and Carter, 2023). 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 (Figure 1) beginning as early as 1.08 Ga (Cannon, 1994; Cannon and Hinze, 1992; Hoffman, 1989). Final inversion of the MCR by compression during Grenville orogeny occurred between 1,010-980 Ma (Hodgin et al., 2022). 10 �The Mesoproterozoic Midcontinent Rift-filling strata of the Keweenaw Peninsula dip moderately toward the center of the rift with the angle of dip increasing toward Keweenaw fault where the stratigraphic base is truncated (Figure 7). The dip of the strata is interpreted as a combination of syn-depositional downwarpage and structural tilting in response to reverse faulting caused by regional continental compression (Woodruff et al., 2020; Cannon, 1994). There are many faults/fractures in the Mesoproterozoic rocks of the Keweenaw Peninsula. Some of these were exclusively formed during extension of the Midcontinent Rift when grabenbounding normal faulting was prominent along the margin (Figure 3). However, most faults/fractures were likely either reactivated by or directly produced by the regional compressional event. The Keweenaw Fault strikes and dips more or less parallel to the bedding of the truncated Portage Lake Volcanics (Figure 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). It is a detached thrust fault related to regional compression. 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. However, the rocks within and adjacent to the fault are altered by late-stage hydrothermal fluids. 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). 11 �Table 2: Production from 1845 to 1968 of refined copper from native copper deposits (after Weege and Pollock, 1971). Million lbs Produced Refined Copper Location Number see 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 Greenland-Mass Subdistrict 72 5 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 Several faults occur oblique to the strike of bedding. In the Eagle River area, fault-controlled native copper veins are common in association with high-angle faults whose displacement is from 0 to 200 m, (Figure 7; see also Figure 19; Butler and Burbank, 1929). The Allouez Gap fault (Figure 7) bisects the largest lava flow top hosted native copper deposit in the district (see Figure 16) and was likely a significant conduit for native copper mineralizing hydrothermal fluids (Bornhorst, 1997). Correspondence between the thickness of the Kearsarge lava flow and the Allouez Gap fault suggests this fault was active during deposition of the Portage Lake Volcanics. It is interpreted as having been reactivated during regional compression. 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 as well as those in the Greenland-Mass subdistrict (Broderick, 1931). Faulting occurred before, during, and after deposition of native copper and its associated alteration minerals based on fault brecciated and recemented alteration minerals. There is a close relationship between faulting/fracturing produced by or reactivated by compression and native copper deposits. The compressional structures acted as pathways for mineralizing hydrothermal fluids (Bornhorst 1997). 12 �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 late regional compression (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 ~20 billion lbs of copper (Bornhorst and Barron, 2011). Small quantities of native silver are temporally and spatially associated with the native copper. The major ore producing horizons are located in a 45 km-long belt in the Keweenaw Peninsula (Figures 4 and 7) and in a subdistrict to the southwest. Native copper and silver were the only economic metallic minerals and were co-precipitated with a suite of nonmetallic alteration minerals (Figure 8). Sulfide minerals, such as chalcocite, are uncommon in native copper deposits and when present only occur in trace amounts. Sulfide minerals occur in latestage veins (Figure 8). 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 (Woodruff et al., 2020; 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 hosted by the Portage Lake Volcanics where there is 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 which 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/ 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 (Table 2). 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 barren massive basalt of the same flow as the mineralized flow top and hanging wall interior of 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 amounts of copper. As brecciated/fragmental amygdaloidal transitions 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 brecciated/fragmental amygdaloid within the top of a 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 elongated, but 13 �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 on the incline below the surface (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 production of refined copper were hosted by them. These deposits were 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 overlying lava flow. The porosity of underlying brecciated/fragmental amygdaloid lava flow top is often greatly decreased by silt and sand filling the primary open space between fragments of the flow top hence, the originally porous flow top underlying a conglomerate bed acts more like an aquiclude in the paleohydrologic hydrothermal system. Native copper tends to be concentrated along specific stratigraphic bands within the conglomerate that are 0.5 to 5 m thick (Weege et al., 1972). The Calumet and Hecla Conglomerate was by far the largest single native copper deposit in the district producing 4.2 billion lbs. as compared to the next largest deposit, the Kearsarge flow top which produced 2.3 billion lbs. from a fragmental amygdaloid (Figure 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; 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 in the clastic sedimentary host rock 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. However, overall, the vein deposits are of slight economic importance in the district. 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 anastomosing filled open spaces. While brecciation within vein deposits is common, gouge is not present (Butler and Burbank, 1929). The lava flow tops and conglomerates adjacent to the vein 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 occurs as both finely disseminated and as masses weighing many tons. The grade of native copper in the veins has the nugget-effect making determination of grade difficult. Several small vein deposits are localized just beneath the thickest basalt flow in the district, the Greenstone flow. For these veins the hydrothermal fluids moved up along the cross fractures until blocked by the very thick impermeable massive interior of the Greenstone Flow. 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 14 �and Isle Royale faults (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 Figure 16). 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 also likely important for upward transport of ore fluids. 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 (Bornhorst, 1997). 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 area of major native copper deposits rocks were less altered at lower temperature. The intensity and degree of alteration also varies as a function of position within lava flows; the massive interiors of lava flows are 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 some flows, but secondary minerals exist in the massive interiors of all flows regardless of their thickness. While the thicker massive interiors of lava flows contain secondary minerals, their original igneous geochemical composition is often only slightly or essentially not modified by secondary hydrothermal processes. There are more than 50 different secondary alteration minerals in the Keweenaw Peninsula; most of them are related to hydrothermal processes and some are related to supergene processes. Only about 20 alteration minerals are major to less common minerals (Figure 8). 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 of 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). There is a district wide temporal (paragenetic) and spatial variation in the assemblage of alteration minerals which was first well described by Butler and Burbank (1929) and later summarized by White (1968). Recently Bodden et al. (2022) have refined the paragenetic and spatial variation of the hydrothermal minerals (excluding igneous and supergene related minerals) (Figure 8). The hydrothermal alteration minerals can be subdivided into main-stage which paragenetically overlap with the precipitation of native copper (Figure 8). While district-wide there is a well-defined mineral paragenesis, individual deposits may not exactly follow the district-wide timing of precipitation (compare Figure 8 to Stop 5). The main-stage is interpreted by Bodden et al. (2022) as formed during a continuous hydrothermal event. 15 �Figure 8: Paragenesis and relative abundance of secondary hydrothermal alteration minerals in the Keweenaw Peninsula native copper district. After Bodden et al. (2022). The late-stage minerals are widespread but volumetrically minor. They commonly occur in small veins/fractures which cross-cut the main stage minerals or as coatings on main-stage vug filling minerals. Late-stage alteration minerals are notably more abundant near the Keweenaw fault. The suite of late-stage minerals are distinguishable by the occurrence of sulfur-bearing minerals, sulfides and sulfates, and by an assemblage of lower temperature minerals in areas where they overprint an assemblage of main-stage minerals formed at higher temperatures. The timing of the late-stage hydrothermal event is uncertain. There could have been no time break or a major time break between the main-stage and late-stage hydrothermal events. Bodden et al. (2022) suggested that the main-stage and late-stage hydrothermal events are practically continuous with each other. Main-stage alteration minerals are spatially zoned perpendicular to stratigraphic strike as demonstrated for the Calumet area of the district (Figure 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 (Stop 5). The alteration mineral zones of the Portage Lake Volcanics are similar to the North Shore Volcanic Group of Minnesota (Schmidt and Robinson, 1997). Bodden et al. (2022) mapped the occurrence of 16 �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 modified from Bornhorst and Rose (1994). Figure 10: Main-stage hybrid metamorphogenic hydrothermal mineral zones of the Keweenaw Peninsula (modified from Bodden et al., 2022) 17 �alteration minerals of the Keweenaw Peninsula into those zones used for the North Shore Volcanic Group (Figure 10; Schmidt and Robinson, 1997). These zones can be equated to the temperatures of mineral formation (Bodden et al., 2022). The spatial zoning of alteration minerals is consistent with a thermal high associated with the major native copper deposits (compare Figure 4 and 10). The mineral zones dip more gently towards Lake Superior than the strata, implying that the strata were at least somewhat tilted prior to main-stage hydrothermal alteration (Livnat, 1983; Broderick, 1929). Native copper mineralization is younger than the Copper Harbor Conglomerate, which hosts rare veins of calcite and native copper (see Stop 7). White (1968) interpreted the age of native copper mineralization as after the deposition of parts or all of the Freda Sandstone. The Bear Lake igneous body within the Freda Sandstone is native copper mineralized. Minor amounts of native copper occur within the lower beds of the Jacobsville Sandstone near Rice Lake (Calumet and Hecla unpublished drill core log). 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 1050 Ma (+/- ~ 20 Ma) (Bornhorst et al., 1988). This age is consistent with the approximate age of 1060+/-20 Ma for regional continental compression that caused thrust faulting along the Keweenaw Fault (Cannon et al., 1993). Thus, the age of main-stage hydrothermal alteration is about 1060 to 1050 Ma contemporaneous with regional continental compression and some 30 million years after eruption of the Portage Lake Volcanics. Genesis of the Main-Stage Native Copper Deposits This section is summarized from Bornhorst and Mathur (2017) and Bodden et al. (2022) and illustrated in Figure 11. Native copper occurs throughout the MCR in Wisconsin, Minnesota, and Ontario (Figure 2). It formed during a regional hydrothermal event from about 1060 to 1050 Ma (Bornhorst et al., 1988). The regional Cu-bearing hydrothermal fluids are best explained as generated during burial metamorphism of rift-filling basalts with temperatures reaching a thermal maximum approximately 30 million years after the end (~1085 Ma) of widespread rift magmatism. The suite of main-stage hydrothermal minerals precipitated, except native copper, (Figure 7 and 8) is similar to those found elsewhere rocks have undergone very low to low grade burial metamorphism at less than about 300OC. The coincidence of regional continental compression with a burial thermal maximum (Woodruff et al., 1995) provided an integrated paleohydrologic system through reactivated and new faults and fractures. This allowed the upward movement of hydrothermal fluids to focus in sites of future copper deposits at the very time of greatest fluid availability (Bornhorst 1997). During generation of the regional burial metamorphogenic hydrothermal ore fluids, copper was leached at depth from the rift-filling basalt strata (Bornhorst and Mathur, 2017, 2018). More than sufficient amount of copper was available to have been leached from the buried rift-filling basalts. 18 �Figure 11: Cartoon cross sections showing conceptual genetic model of the native copper deposits of the Keweenaw Peninsula formed at about 1070 to 1040 million years ago. Modified from Bodden et al. (2022). A. Marine incursions and seawater penetration during deposition of volcanic and sedimentary rocks in MCR. B. Area prior to burial metamorphism with sulfur depleted evolved seawater providing salinity for ore-forming fluids. C. Burial metamorphism with generation of burial metamorphic hydrothermal fluids. Mixing the burial metamorphic fluids with evolved seawater produces copperbearing hybrid metamorphogenic ore-forming hydrothermal fluids. D. Precipitation of main-stage minerals as a result of mixing of the ore-forming fluids with meteoric water, decreasing temperature, and water-rock reactions. 19 �The rift-filling volcanic rocks were low in sulfur when they erupted and the little available sulfur was degassed prior to solidification (Bodden et al., 2022; Bornhorst and Mathur, 2017). These lowsulfur rift-filling volcanic rocks were buried into the source zone where burial metamorphogenic hydrothermal fluids were generated (Figure 3 and 11C). Since the very low sulfur rift-filling basalts in the source zone were the same as those in the fluid pathways to the zone of precipitation, the fluids remained sulfur poor. Since the native copper ore host rocks were again the same rift-filling very low sulfur volcanic rocks, the burial metamorphogenic hydrothermal fluids remained depleted in sulfur. The lack of sulfur resulted in the precipitation of native copper rather than copper sulfides. While the metamorphogenic hydrothermal fluids lacked sulfur, several studies (Kelly, 2022; Kelly, 2020; Püschner, 2001; Brown, 2006; Livnat, 1983; Jolly, 1974) have suggested that the main-stage hydrothermal fluids had at least moderate degree of salinity. The possible sources of salinity were evaluated by Bornhorst and Mathur (2017), Bornhorst (2021), and Bodden et al. (2022). Viable sources of salinity for the hydrothermal fluids need to also satisfy the constraint of very low sulfur. Bornhorst (2021) hypothesized that evolved formation water derived from sulfur depleted seawater could have been the source of salinity. As seawater penetrates midocean ridge basalts it is heated, reacts with the host basalt, and as a result of precipitation of minerals it becomes depleted in sulfur. If the sulfur content of the Mesoproterozoic seawater was low in sulfur, as proposed by Blattlet et al. (2022), then less depletion of sulfur would have been needed. During deposition of the youngest Portage Lake Volcanics and overlying Copper Harbor and Nonesuch formations there were probable incursions of an arm of the sea into the rift for a significant amount of time. This could have resulted in seawater deeply penetrating into the underlying rift-filling volcanic rocks (Figure 11A; Bornhorst, 2021). During burial, the riftfilling volcanic and clastic sedimentary rocks and contained seawater was progressively heated and thereby evolved to be depleted in sulfur (Figure 11B). Continued heating during burial resulted in burial metamorphic hydrothermal fluids which then thoroughly mixed with the evolved seawater to form a hybrid metamorphogenic-dominated ore-forming fluid (Figure 11C; Bodden et al., 2022). These main-stage ore-forming hydrothermal fluids moved upwards from the source zone through the same very sulfur poor strata as in the source rocks (Figure 11D). As they moved upwards they cooled, interacted with host rocks, and in the relatively shallow zone of precipitation they variably mixed with sulfur-poor, low salinity, reduced meteoric water (Figure 11D; Bodden et. al, 2022). These processes resulted in precipitation of native copper and main-stage hydrothermal minerals. Higher temperature main-stage mineral assemblages are spatially associated with the area of native copper deposits where the thermal anomaly was greatest because of focused hydrothermal fluids (Figure 10). The possible depth of the zone of precipitation is poorly estimated with a best guess at this time of 10 to 15 km (Kelly et al., 2022; Kelly, 2020). Within the native copper district, the suite of main-stage minerals, including native copper, is followed by late-stage minerals precipitated at lower temperatures than the main-stage hydrothermal fluids coincident with the native copper district. 20 �Late-Stage Hydrothermal Minerals The suite of late-stage minerals is widespread and similar throughout the Keweenaw Peninsula. The late-stage suite is readily distinguished in the main area of the native copper district since late-stage minerals are lower temperature (100 to 150OC) than the main-stage minerals in district itself. However, outside of the native copper district where main-stage minerals are expected to be formed at lower temperature, the main-stage and late stage are indistinguishable. Bodden et al. (2022) suggested that late-stage fluids are a variable mixture of hybrid metamorphogenic hydrothermal fluids, meteoric water, and shallow seawater, the latter being a source of sulfur in the late-stage fluids. Phanerozoic The Keweenaw Peninsula was subjected to a 500-million-year period of erosion, from about 1 Ga to 0.5 Ga (500 Ma) and multiple kilometers of rock were eroded exposing the native copper deposits at the surface (Figure 3). Downward percolating groundwaters supergene altered native copper and produced a suite of including cuprite, tenorite, malachite, and chrysocolla (Bornhorst and Robinson, 2004). The rocks of the Keweenaw Peninsula were subsequently buried by Paleozoic sedimentary rocks associated with the Michigan basin beginning about 500 Ma (Figure 3) and ending Precambrian supergene alteration. 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 Paleozoic outliers slightly south of the Keweenaw Peninsula (Figure 3). The last glacial episode exposed the native copper deposits at roughly the same erosional level as at 500 Ma or the end of the Precambrian (Bornhorst and Robinson, 2004). 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 submerged under 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. 21 �Figure 12: Geologic map of the far western part of the Upper Peninsula of Michigan showing field trip stops. Objectives of Field Trip This field trip is designed to provide an overview of the Mesoproterozoic Midcontinent Rift-filling strata and native copper deposits of the Keweenaw Peninsula (Figure 12). There are four rift-filling formations: Portage Lake Volcanics, Copper Harbor Conglomerate, Nonesuch Formation, and Freda Sandstone. The Nonesuch and Freda formations are poorly exposed in the Keweenaw Peninsula thus, only two optional stops are included in this field guide. The Jacobsville Sandstone is a rift-flanking formation and is outside the scope of this field trip. The rift-filling strata are overlain by unconsolidated Pleistocene glacial sediments. There is one glacial related stop. IMPORTANT NOTE TO ALL READERS: Many of the field trip stop descriptions and significant parts of the introductory geologic overview have been previously published especially in other Institute on Lake Superior guidebooks e.g., Bornhorst and Barron (2013) and the extensive guides by Bornhorst and Rose (1994) and Bornhorst et al. (1883). The stop descriptions in this field guide, as compared to previously published guides range from exact wording to significantly modified wording without specific citation. Two of the stops in this field guide, Stops 2 and 8, were not visited by previous field guides involving Bornhorst. These stop descriptions are new. 22 �Stop 1: Subaerial basalt lava flow cross-section at South Range Quarry Latitude: 47.07750N; Longitude: -88.64240W Directions: Drive west through downtown Houghton on US-41 to south M26. Drive about 4.5 miles to unmarked road on west side of M-26 just before church on south side of unmarked road. Proceed on road to tree line. Walk NE uphill to quarry. THIS STOP IS ON PRIVATE PROPERTY. PLEASE GET PERMISSION TO ENTER PROPERTY. Volcanic textures and structures typical of moderate-to-thick subaerial lava flows within the Portage Lake Volcanics are well exposed in this old quarry (Figure 13). As one traverses up the hill to the quarry along the rubbly path there is a low-profile exposure of a 4 m thick interflow conglomerate bed which is also exposed laterally along the SE slope face of this knob. The conglomerate layer is stratigraphically the National Sandstone member and is approximately near the middle of the exposed Portage Lake Volcanics stratigraphic section (Figure 6). The conglomerate is overlain by an 18 m thick lava flow (A). Figure 13: Geologic cross section of the South Range Quarry (modified from White, 1971). Laterally continuous interflow sedimentary beds provide critical stratigraphic markers within the Portage Lake Volcanics, an otherwise monotonous volcanic pile with many laterally discontinuous lava flows. The sedimentary unit exposed below the quarry has been correlated with the National Sandstone, a marker bed in the Mass-Rockland area (Figure 6). At South Range Quarry, the National Sandstone is a massively bedded, pebble-cobble framework conglomerate, composed of silicic with subordinate mafic volcanic clasts that are subangular to subrounded, within a matrix of poorly-sorted medium-to-coarse sand of similar composition. The Portage Lake Volcanics basalts in this part of the stratigraphic section are mainly olivine tholeiites and erupted as subaerial lava sheets. The principal lava flow exposed in the quarry walls illustrates many of the volcanological features of Portage Lake Volcanics lava flows. The top and bottom of the South Range Quarry lava flow (B) are exposed. Flow B was deposited directly on top of Flow A and consists of aphanitic chilled basalt. The base of Flow B occurs where amygdules disappear abruptly in the top of the underlying flow. 23 �The upper surface of the main flow (B) was brecciated slightly by movement of lava after the formation of an upper crust. The flow top breccia (locally termed fragmental amygdaloid) is laterally discontinuous. The fragmental amygdaloid rapidly grades downward to an unbrecciated, highly amygdaloidal (vesicular) flow top. Note the variation in vesicle size and shape downward in the flow. There are numerous layers of flattened amygdules (vesicles) in the flow top with their orientation parallel to the top and bottom of the flow (B). The orientation of these layers may represent laminar flow planes within the flow top. The section was tilted after emplacement. Slow cooling of the lava flow caused solidification toward the flow interior at a rate which allowed development of subophitic to ophitic textures (ophitic texture denoted by 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 (B). Before final solidification, small amounts of volatile-rich, differentiated residual liquid were likely injected into thin discontinuous layers and lenses (tabular openings produced during cooling). Most of these pegmatoid layers are subparallel to the bottom and top surfaces of the flow. A typical pegmatoid zone consists of a 5 to 10 cm border zone at the top and bottom which is composed of a medium-to-coarse grained aggregate of Ab-rich plagioclase, prisms of Fe-rich clinopyroxene and abundant Fe-Ti oxides, as well as accessory minerals such as apatite and zircon (Cornwall, 1951c). The cores of the pegmatoid zones are 5 cm to 1.2 m thick consisting of a green vesicular basaltic rock. Zircons extracted from pegmatoids within thick Portage Lake Volcanics basalt flows have yielded high-precision U-Pb dates (e.g., Davis and Paces, 1990). There are thin tabular layers and flattened amygdules (vesicles) in the top of the flow and in the massive interior of the flow (B) that are composed of a brown and sometimes green fine-grained material that have been described interpreted by White (1971) as detrital material. Alternatively, this green-to-red cherty rock, could be simply alteration minerals (quartz, prehnite, epidote, and pumpellyite) filling fractures and nearly complete pseudomorphic replacement of basalt. Numerous pegmatoid layers are exposed in the quarry walls, as well as in the glacially-polished surfaces above and to the north of the quarry. The effects of regional hydrothermal alteration can be observed within the vesicular flow top and pegmatoid zones. Vesicles are filled with a variety of secondary minerals including quartz, epidote (olive green), prehnite (waxy light green), calcite, pumpellyite (pale bluish green), chlorite (dark green to black) and traces of native copper (pinkish color). Pseudomorphic replacement of basalt by fine-grained secondary minerals is most intense where permeability was highest. The massive interior of the flow is only a little visibly altered, however plagioclase is altered to albite and pyroxene is altered to chlorite. In the vicinity of selected fractures there can be intense epidote or prehnite alteration. The massive interior was a relatively impermeable horizon in the paleohydrothermal system. Fracturing during late compression integrated the system and provided limited pathways for upward movement of ore fluids. 24 �Stop 2: Subaerial basalt lava flow, eastbound US-41 Houghton Latitude: 47.12144N; Longitude: -88.56474W Directions: Drive west through downtown Houghton and loop around (Yooper Loop) to head east for 0.8 miles along US-41 or Montezuma Avenue (one-way two-lane highway). Stay left (north side) as though going back through downtown and turn into unmarked gravel lot just before large outcrops on left (north). Use sidewalk to walk to outcrop. BE CAREFUL OF TRAFFIC. The rock cut at the southeast end of downtown Houghton provides an excellent example of the characteristics of Portage Lake Volcanics subaerial basalt lava flows (Figure 14). There is a sidewalk providing access to the Stop 2 south-facing road cut. While there is also a sidewalk on the other side of the road crossing the road is discouraged. These other outcrops can be accessed by parking uphill from them on the other side of the road. There are also north-facing exposures of this same stratigraphic interval on west bound US-41 (Shelden Avenue) east of the Houghton U.S. Post Office. The exposures at Stop 2 are located stratigraphically above the Scales Creek flow and below the Kearsarge flow, slightly closer to the Scales Creek flow (Figure 6). The lava flows of the exposure and vicinity strike approximately N30oE and dip toward the center of the rift (Lake Superior) at about 55o northwest (White, 1956). Figure 14: Geologic cross section of road cut eastbound US-41, southeast end of downtown Houghton. There are parts of two lava flows exposed at Stop 2 (Figure 14). The base of the stratigraphically lower of these two lava flows (A) is not exposed on the east end of the Stop 2 exposures. It is also not exposed on the other side of the road. The top of Flow A is well exposed and is overlain by massive basalt of Flow B. The contact itself is denoted by an abrupt change from underlying highly altered greenish, slightly brecciated amygdaloidal basalt of Flow A (about 0.5 m thick) that is overlain by massive dark grey to black basalt of Flow B. This greenish zone of the top of Flow A is underlain by about 3.5 m of amygdaloidal basalt with slightly fragmental (brecciated) basalt, also Flow A. The slightly fragmental basalt is gradational downward (east) towards the center of Flow A where the abundance of amygdules (filled vesicles) is sufficient to call the rock amygdaloidal basalt. Amygdaloidal basalt lacking fragments is about 4 m thick. There is an arbitrary boundary where the 25 �abundance of amygdules is too low to call the rock amygdaloidal, although it contains some amygdules. The abundance of amygdules progressively decreases towards the center of Flow A, the porous and permeable top of Flow A is about 8 m thick. Flow A is greater than 14 m thick as its base is not exposed. Notably, flattened amygdules (vesicles) occur along planar layers in the flow A top with their orientation roughly parallel to the top of the flow. Larger flattened amygdules are about 2 by 2 by 1 cm and between them, appearing to connect larger amygdules, are much smaller amygdules 1 to 2 mm thick. The orientation of these layers may represent laminar flow planes within the flow top. Layers of flattened amygdules are also numerous at Stop 1. Flow A is overlain by Flow B and the top of Flow B is not exposed at Stop 2. However, its top is poorly exposed on the other (south) side of the road. Flow B is about 30 m thick. There is a pegmatoid layer in the massive basalt interior of Flow B (Figure 14). Flow B is thicker than the average flow. The pegmatoid is distinguished as notably amygdaloidal. Pegmatoids are discussed further at Stop 1. The effects of regional hydrothermal alteration can be observed within the top of Flow A and the pegmatoid layer in Flow B. The massive basalt interior of Flow A and B are much less altered than the flow top. However, the primary magmatic plagioclase in the massive basalt has been replaced by albite and the primary mafic minerals are replaced by chlorite, pumpellyite, and iron oxides. The intensely altered basalt at the very top of Flow A is largely replaced by hydrothermal alteration minerals including epidote, prehnite, pumpellyite, quartz, chlorite, calcite, and trace native copper. Amygdules are frequently filled with colorless to white quartz. a mixture of prehnite pumpellyite and quartz, a mixture of pumpellyite, a mixture of quartz with lesser calcite, only quartz, and only pink inclusions of native copper in milky or colorless quartz. There is visible native copper in some amygdules. The massive interior of the flow is much less altered than the flow top and represents a relatively impermeable horizon in the paleohydrothermal system. In contrast, the flow top was a pathway for movement of hydrothermal fluids. 26 �Stop 3: Overview at Bumbletown Hill Latitude: 47.290100N; Longitude: -88.417250W Directions: From Portage Lift Bridge follow US-41 towards Copper Harbor and proceed through Calumet towards Allouez for about 15.5 miles to Bumbletown Road. Turn left (west) and proceed one mile up to the top of Bumbletown Hill via Cedar Street. Walk around outside of communication tower fence to get excellent views as described below. Figure 15: Geologic sketch map of Bumbletown Hill modified from White (1971). From the overlook on a clear day, Isle Royale may be seen 80 km to the northwest and the Huron Mountains may be seen beyond Keweenaw Bay, 60 km to the southeast. From the top of Bumbletown Hill the land slopes very gradually to the northwest toward Lake Superior. This slope is similar throughout much of the northwestern side of the Keweenaw Peninsula. The area is underlain mainly by conglomerates and sandstones of the Copper Harbor Conglomerate dipping at about 20 to 30 degrees NW. 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. Next to the Keweenaw fault beds of the Jacobsville Sandstone can be steeply dipping. Looking northeast along the strike of the Portage Lake Volcanics, one can see the cuesta form of the ridge underlain by the Greenstone flow. At Bumbletown Hill, the Greenstone flow is only 75 m thick (Figure 15), but the flow thickens abruptly to more than 400 m near this end of the cuesta ridge. The Greenstone flow dips northward at about 25o toward the center of the Lake Superior. It can be traced along much of the Keweenaw Peninsula (Figure 6) and has been stratigraphically and 27 �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 to 1,500 km3 (Longo, 1983; White, 1960). The Greenstone Flow is an enormous lava flow. It is possible that rather than having been a lava flow the Greenstone Flow was a lava lake. Regardless, the Greenstone Flow perhaps is the greatest single continuous outpouring of lava on Earth. Very slow solidification of this great mass of magma allowed extensive in-situ magmatic differentiation (Cornwall (1951a, 1951b). Magmatic differentiation resulted in a massive, ophitic (lath-shaped plagioclase surrounded by large irregular masses of clinopyroxene) base of the flow; an overlying zone of intercalated subophitic and pegmatoidal layers; an upper ophitic zone; and a finegrained, vesicular flow top. The lower ophitic zone experienced rates of undercooling low enough to allow growth of clinopyroxene oikocrysts up to 5 cm in diameter. The geochemical composition of the Greenstone Flow magma is more evolved than typical olivine tholeiites; which constitute the greatest volume of the Portage Lake Volcanics. Primitive olivine tholeiite and quartz tholeiite occur between the Greenstone Flow and the top of the Portage Lake Volcanics. Generally, magmas of the Portage Lake Volcanics become more primitive and less crustal contamination with time during the development of the Midcontinent Rift (Paces, 1988). At Bumbletown Hill, the Greenstone Flow is only 75 m thick and is composed of a thick amygdaloidal flow top with some exposures of fine-grained columnar basalt. To the left of the cuesta ridge the rocks consist of the top of the Portage Lake Volcanics and the bottom of the Copper Harbor Conglomerate. To the right of the ridge, the more distant hills are formed by lava flows near the base of the Portage Lake Volcanics. Bumbletown Hill is located on the southwest side of Allouez Gap, a NW- to SE-trending valley (see Figure 7). 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 (Figure 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 (Figure 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 Figure 16), which was the second largest copper producer in the native copper district. There was a readily visible line of poor rock piles, a little more than 1,500 m southeast of Bumbletown Hill, from the many mines which were producing native copper from the Kearsarge deposit. Many of these piles are now gone as they have been crushed for aggregate. The line is still visible in the fall when leaves are not on the trees. About 1,200 m N65oE of the hilltop, the Houghton conglomerate and the stratigraphically lower Iroquois flow produced 33 million pounds of copper. East of Bumbletown Hill but no longer visible is the Kingston Mine, one of the most recent native copper mines to open and last to close. It was discovered by exploration near the Allouez Gap fault. It only produced 20 million pounds of copper from 1963 to 1968. 28 �Stop 4: Interflow Conglomerate at Bumbletown Hill Latitude: 47.287136N; Longitude: -88.415365W Directions: From top of Bumbletown Hill turn around and head downhill 0.5 miles to pull over on left just before dirt road. Specimens of Allouez Conglomerate are scattered about the southeast flank of base of Bumbletown Hill. The remnants of Allouez Conglomerate poor rock piles are private property (Figure 15). Near this pullover there is the opportunity to collect specimens of the Allouez Conglomerate. The Allouez Conglomerate is one of a small number of interflow clastic sedimentary horizons within the Portage Lake Volcanics and is visible on the lower slopes southeast of Bumbletown Hill. Conglomerate layers within the Portage Lake Volcanics are important stratigraphic marker horizons (Figure 6). Correlation of basaltic lava flows along strike would be difficult without clastic sedimentary marker beds deposited during periodic waning of volcanism. The Allouez Conglomerate can be traced more than 120 km along strike from Mass to Delaware (Figure 6). The Allouez conglomerate is just below the Greenstone flow. At Bumbletown Hill the Allouez Conglomerate was mined for native copper (Figure 15) albeit it only yielded about 75 million lbs of refined copper (Table 2). Elsewhere the Allouez Conglomerate has yielded additional native copper. The Allouez Conglomerate consists of mostly red-colored conglomerate with lesser amounts of sandstone and siltstone. The largest clasts 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 by White (1971) gave the following results: 16% basalt, mostly amygdaloidal; 36% quartz porphyritic rhyolite; 11%, feldspar porphyritic rhyolite; and 37% felsic granophyre. The Houghton Conglomerate is almost entirely clasts of quartz porphyry as is the Kingston Conglomerate. There are conglomerates within the lower Portage Lake Volcanics near the tip of the Keweenaw Peninsula whose clasts are clearly sourced from an extrusive dome of rhyolite. This could be the explanation for the uniformity of clasts in the Houghton and Kingston Conglomerates. In contrast, the heterogeneity of the Allouez Conglomerate clasts suggests a less restricted source area (White, 1971). Evidence of native copper mineralization can be seen in some rocks of the Allouez Conglomerate at nearby 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.” Chalcocite is a product of late- stage hydrothermal fluids. Supergene alteration resulting from the downward percolation of groundwater is not common at depth in most the native copper deposits. Supergene alteration products are likely common when native copper ore bodies are at or very near the surface. At this stop, supergene alteration minerals are common including chrysocolla, malachite, and cuprite. The depth of occurrence is unknown although given the widespread distribution in rocks of the poor rock piles it seems likely at least some supergene alteration occurred at depth as documented elsewhere. The occurrence of supergene alteration minerals at depth in the native copper mines was used by Bornhorst and Robinson (2004) to hypothesize that at least some supergene alteration was Precambrian in age. 29 �Stop 5: Seneca Mine Rock Pile Latitude: 47.311915N; Longitude: -88.365818W Directions: At the intersection of Bumbletown Road and US-41, turn left (northeast). Drive about 2.9 miles almost through Mohawk to 1st Street when you will turn left (northwest). Proceed on 1st Street/Seneca Location Road 0.3 miles to gated road. Walk to rock piles. THIS STOP IS ON PRIVATE PROPERTY. PLEASE GET PERMISSION TO ENTER PROPERTY. The Kearsarge lode was worked by the Seneca Mine, one of multiple mines which produced native copper from the top of the Kearsarge lava flow over a strike length of more than 12 km and downdip as much as 2,500 m (Figure 16). 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 Calumet & Hecla 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 (Figures 6 and 16). It lies directly above the Wolverine Sandstone (Figure 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 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 are variable, but they can make up a large percentage and phenocrysts are up to 2.5 cm in length. This zone is probably the result in situ floating of plagioclase during surface crystallization of the flow. The phenocrysts likely formed in a shallow magma chamber. Specimens with abundant plagioclase phenocrysts are common on this rock pile. The basalt of 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 minerals are replaced by alteration minerals to varying degrees. The amygdule and interfragment spaces are filled with (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 (Figure 17). A zonal stratabound arrangement of amygdule minerals in the Kearsarge deposit is seen in the Ahmeek Shaft No. 3 (Figure 18). The zoning may be explained by deposition of secondary minerals from a hydrothermal solution moving along a permeable channel. Chlorite and microcline would have been deposited first, along the outer limits of the solution channel; 30 �followed by quartz and epidote in the center of the channel; and finally, deposition of calcite in the remaining openings. This is consistent with the paragenetic relationships seen in individual samples from the rock pile. 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 (Figure 16), 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 mineral zones. Figure 16: Thickness of the Kearsarge lava flow showing the productive area to be the thickest in the top diagram modified from Butler and Burbank (1929). The Kearsarge flow top ore body is bisected by the Allouez Gap fault. Bottom diagram shows strike parallel down-dip projection 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 downdip side of the line (lower) and K-feldspar is absent on the down-dip side lower line shown. The Kearsarge flow dips about 35 to 40o NW. Mine names and shaft numbers are noted. 31 �Figure 17: Paragenesis of secondary hydrothermal alteration minerals in the Kearsarge deposit at the Wolverine No. 2 Mine. Figure 18: 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 laumontite-quartz-calcite zone not shown here. 32 �The Allouez Gap Fault bisects the thickest segment of the Kearsarge Flow along its 55 km strike length (Figure 16). 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). The latter may be late-stage. 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 with 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% (microcline); pink feldspar (adularia) 15%; epidote, 17%; prehnite, trace; pumpellyite, trace and quartz, trace. Many specimens contain multiple minerals and illustrate paragenetic relationships. Stop 6: Eagle River Falls Latitude: 47°24'44.9N; Longitude: - 88°17'47.3W Directions: Return to US-41 from Seneca Mine and turn left (northeast) continuing 7.2 miles to junction of US-41 and M-26. Turn left (north) on M-26 towards Eagle River and proceed 2.3 miles across bridge and immediately right after the bridge into the parking lot. The waterfalls of Eagle River are near the contact between the top of the Portage Lake Volcanics and the base of the Copper Harbor Conglomerate (Figure 5). 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 of the peninsula. The tholeiitic basalt subaerial lava flows just below the contact are pahoehoe type with a ropy upper surface. The orientation of the ropes indicates that the flow 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 Harbor Conglomerate. The Copper Harbor Conglomerate consists of red-brown rhyolite-pebble conglomerate but includes many sandstones 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 Conglomerate. The environmental setting of the Copper Harbor Conglomerate is discussed further at Stop 8. This contact marks an abrupt change in the geologic evolution of the Midcontinent rift. Below this contact the dominant strata is a very thick succession of subaerial basalt lava flows erupted from fissure vents under Lake Superior that filled the progressively extended and down dropped rift basin 33 �during active Midcontinent rifting. Below this contact the Portage Lake Volcanics consists of more than 200 individual lava flows with a cumulative exposed thickness of about 5,000 m; the base is fault truncated (Figure 4 and 5). Abruptly above the contact lava flows are strikingly absent and clastic sedimentary rocks are the dominant strata deposited in a sagging basin after active extension ended. The clastic sedimentary strata above this contact consists mostly of conglomerate and sandstone with a cumulative exposed thickness of more than 5,700 m; the top is not exposed (Figure 4 and 5). While generally absent, a thin package of mafic to intermediate volcanic rocks (Lake Shore Traps) deposited as a shield volcano interfingers with clastic sedimentary rocks of the Copper Harbor Conglomerate. The very last magmatic activity in the MCR is the lone Bear Lake alkaline igneous body near the middle of the clastic sedimentary strata. Stop 7: Great Sand Bay Latitude: 47. 446140N; Longitude: -88. 216411W Directions: Continue driving northeast (right from parking area) on M-26 for 4.5 miles until the Great Sand Bay paved pullover with overview and beach access. The Great Sand Bay overlook provides a beautiful view of Lake Superior (Figure 19). 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. 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 Conglomerate. In the Keweenaw Peninsula the Lake Shore Traps are interbedded near the middle of the Copper Harbor Conglomerate (see Stop 9). The massive interiors of these lava flows are more resistant to erosion than the underlying and overlying conglomerates and sandstones of the Copper Harbor Conglomerate. As a result, harbors such as 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 crosscut the Eagle River shoals located about 0.5 to 1 km offshore. To date, there have been a total of 36 underwater copper veins discovered from the eastern tip of Great Sand Bay (visible at this stop) to Eagle River, about 3.2 km west. Some of these submerged veins likely connect with veins recognized from on land exposures (Figure 19). The native copper that is naturally on the bottom lands of Lake Superior are grouped as “lake copper.” The largest lake copper specimen ever recovered underwater was a massive 19-ton unattached copper slab in July of 2001. It was recovered from one of these vein deposits north of Jacobs Creek in about 9 m of water. This large underwater native copper vein is on display at the A.E. Seaman Mineral Museum in the outside copper pavilion. 34 �Many of the submerged veins 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 the 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 clay pockets associated with well-defined copper crystal specimens. Bornhorst and Barron (2017) provide additional information about the Guiness World record tabular 19-ton native copper mass. Figure 19. Geologic and location map of the 19-ton submerged native copper vein recovered from Great Sand Bay, Keweenaw Peninsula, Michigan (from Bornhorst and Barron, 2017). 35 �Stop 8: Copper Harbor Conglomerate, J. & M. Lizzadro Lakeshore Preserve By Daniel J. Lizzadro-McPherson Latitude: 47. 479128N; Longitude: -87.979576W Directions: Continue east-northeast on M-26 from Great Sand Bay to Eagle Harbor. Continue on M-26 towards Copper Harbor for 9.2 miles (14.8km) until the sign for the preserve appears on the lakeward side of the road. Park at the pull-off on the north-side of M-26 for beach access. PLEASE DO NOT DAMAGE THE OUTCROP. NO HAMMERS ALLOWED. COLLECTION OF BEACH ROCK ONLY. Figure 20: Overview of the J. & M. Lizzadro Lakeshore Preserve, which encompasses the entire shoreline between private properties (shaded, cross-hatched areas) along M26. The Joseph & Mary Lizzadro Lakeshore Preserve is located near the northern-most point on the Keweenaw Peninsula and is host to prominent outcrops of Copper Harbor Conglomerate, colorful cobble-pebble beaches, and iconic views of Lake Superior. This site is notable due to the diverse facies of CHC, rare occurrences of stromatolites (genus Colleria) and raindrop impressions found among the wave-washed bedrock exposures. Established as a preserve in 2003 by the Houghton Keweenaw Conservation District and the Keweenaw Land Trust, the Lizzadro Lakeshore Preserve protects over 640-feet of undeveloped shoreline and several small islands (Figure 20). The preserve showcases one of only three A-ranked Michigan occurrences of the globally rare imperilied plant community-type (G2) bedrock beaches. Thanks to a generous donation by Gina Nicholas, the preserve was named in honor of my grandparents, Joseph & Mary Lizzadro, with the hope that this significant Geoheritage site remains protected for many future generations to enjoy. 36 �The bedrock geology of the preserve consists of red-colored clastic sedimentary rocks grouped under the stratigraphic formation of the Copper Harbor Conglomerate (Figure 5). The Copper Harbor Conglomerate. ranges from 490 m thick near the Wisconsin border, to about 1,310 m thick on the Keweenaw Peninsula, reaching a maximum exposed thickness of 2,000 m on the shores of Isle Royale. Overall, the Copper Harbor Conglomerate is generally a medium reddishbrown colored wedge of fluvial siliciclastic conglomerates and sandstones that rapidly filled-in the rift basin as volcanism waned and subsequently terminated (Daniels, 1982; Cannon and Nicholson, 2000; Woodruff et al., 2020). The base of the Copper Harbor Conglomerate. locally interfingers with the uppermost subaerial basaltic lava flows of the Portage Lake Volcanics (Figure 5). Near the stratigraphic lower-third of the Copper Harbor Conglomerate there is a succession of basaltic to intermediate subaerial lava flows, informally named the Lakeshore Traps (Figure 5). Outcrops of Copper Harbor Conglomerate. exposed along the Lake Superior shoreline at the Lizzadro Lakeshore Preserve (Figure 21) are stratigraphically above Lakeshore Traps (Cornwall, 1954). Figure 11: Overview of geologic features present at the J. & M. Lizzadro Lakeshore Preserve. 37 �Nearby at well-studied Dan’s Point, outcrops of Copper Harbor Conglomerate exhibit lithologic facies that are characteristic of the upper two-thirds of the formation. The conglomeratic facies (Figures 21 and 22: Points A, C, and F) are primarily clast-supported and comprised of rounded to well-rounded poorly sorted clasts, consisting of a 2:1 silicic-to-mafic volcanic rock fragment ratio with minor pyroclastic, plutonic, and metamorphic rocks (Elmore, 1984). The conglomerate matrix is comprised of coarse sand-sized subangular grains cemented with carbonate and iron oxides. The sandstone facies (Figures 21 and 22: Points B and E) are predominantly subangular to angular, lithic graywackes which exhibit a number of sedimentary structures including: current-ripples, cross beds, and parting lineations. Many of the coarser sandstones and conglomerates in the upper section have calcite-rich cement in the matrix, consistent with a vadose or semiarid, caliche-style environment (Kalliokoski, 1986). Two noteworthy outcrop features include: 1) a thin continuous zone of laminated cryptoalgal carbonate where laterallylinked stromatolite are draped over cobbles and over contorted sandy siltstone layers (Figures 21 and 22: Point D); and 2) raindrop imprints preserved in fine- to medium-grained sandstone lenses (Figure 21: Point G). At the time of deposition, the region was nearly equatorial in geographic position and the climate was likely arid with seasonal rainfall patterns conducive to flashfloods and the development of vadose carbonate (Elmore and Vander Voo, 1982; Kalliokoski, 1986). The raindrop imprints in sandstone (Figures 21 and 22: Point G) support this paleoclimate assumption. Details surrounding the origin and depositional setting of the Copper Harbor Conglomerate and overlying Nonesuch Formation (Figure 5) have recently been reevaluated. The long-held inference of the origin for these two formations has been interpreted as a non-marine, fluvial-tolacustrine couplet (White and Wright, 1960; Elmore, 1983, 1984; Ojakangas et al., 2001). The Copper Harbor Conglomerate closely resembles modern-day examples of coalescing fluvial and prograding alluvial fan deposits with varying facies that exhibit proximal-to-distal braided streams, sheet flooding and sand flat features (Elmore, 1984). Isolated cryptoalgal carbonate and ooid lenses are interpreted by Elmore (1983) to have formed in shallow, medial fan lakes and possibly in abandoned or low-water stream channels with limited sediment input (Elmore, 1983). However, marine sedimentological features often resemble and can be easily mistaken for nonmarine features unless contextualized with additional evidence, such as isotopic geochemistry. New research presents evidence for a shallow-marine estuarine origin for at least the upper-third of Copper Harbor Conglomerate and overlying Nonesuch Formation based on new sedimentological observations (Jones et al., 2023) and geochemical analyses (Stüeken et al., 2020; Jones et al., 2020). Sedimentological observations include periodic to rhythmic flaserwavy-linsen-pinstripe bedding, superimposed sets of ripple cross-laminations with bimodal (herring-bone) sediment transport directions, desiccation cracks and hummocky crossstratification (Jones et al., 2023). Periodic to rhythmic textures are indicative of tidal-influenced marine depositional settings. Geochemical evidence indicates that gypsum evaporite fabrics have a marine sulfur isotopic composition (Stueken et al., 2020) and that pseudomorphs after gypsum formed in a saline-to-brackish waterbody (Jones et al., 2020). Both studies conclude that the upper section of the Copper Harbor Conglomerate and overlying Nonesuch formation were deposited in a braided fluvial-evaporitic, sabkha-like, tidally-influenced shallow marine embayment rather than fluvial-lacustrine non-marine depositional setting. 38 �Figure 22: Lithologic column of measured section at the Lizzadro Lakeshore Preserve (data collected by Lizzadro-McPherson, Bornhorst, and Vye (2023). Bedding about E-W strike and 35 to 40o N dip. 39 �Stop 9: Copper Harbor Conglomerate and interbedded Lakeshore Traps at Hunter’s Point Park Latitude: 47.474355N; Longitude: -87.899199W Directions: Continue driving east on M-26 for 3.7 miles to North Coast Road. Turn left (northwest) on North Coast Road and proceed 0.3 miles to Harbor Coast Lane. Turn right and drive 0.3 miles to the parking area for Hunter’s Point Park at end of the road. Figure 23: Geologic map of the Copper Harbor area taken directly from Cornwall (1955) showing the location of Hunter’s Point (Stop 9) and Brockway Nose (part of Stop 10). Geology from Cornwall (1954). 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 (Figure 23). Prior to becoming an official park the point was a popular hiking destination for visitors. The landowners subdivided the area for residential housing which would have restricted public access without its conversion into a park. The origin of the name 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 Conglomerate is overall composed of volcanogenic clastic sedimentary rocks, dominantly conglomerates with lesser sandstone, siltstone, and shale such as observed at Stop 8. 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-tointermediate+felsic composition of about 2:1 (Daniels, 1982). Towards the tip of the Keweenaw Peninsula, the Copper Harbor Conglomerate is informally subdivided into an inner (land side) “member” and an outer (lake side) “member.” Between these two “members” there is a thin 40 �succession of interbedded lava flows collectively known as the Lake Shore Traps (Figure 23). The Lake Shore Traps consist of Fe-rich olivine tholeiite, basaltic andesite, and andesite lava that were erupted during a time of waning volcanism within the MCR at 1087.2 +/- 1.6 Ma (Davis and Paces, 1990). 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 conglomerates of the Copper Harbor Conglomerate (Figure 23). The strike of bedding is about E-W and dip is about 36o 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. 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 flow top overlain by 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 (hydrothermal and weathering). 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. Secondary minerals filling amygdules include agate, chalcedony, quartz, laumontite, analcite, calcite, and smectite. 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. Highly visible red hematitic bands form circular patterns within the massive interior; this banding is interpreted to be the result of alteration related to weathering rather than hydrothermal fluids. 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 Conglomerate. 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 8 as these clasts are likely derived from erosion of strat equivalent to the Lake Shore Traps updip towards the highlands of the Keweenaw Peninsula 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 Conglomerate was deposited in an alluvial fan derived from highlands to the southeast in the vicinity of Keweenaw Bay. 41 �Additional outcrops of the Copper Harbor Conglomerate can be seen on the far western end of the pebble to cobble beach. These outcrops consist of interbedded conglomerates and sandstone that are typical of the formation as a whole. These conglomerates are similar to those described at Stop 8. There are several prominent, 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 Conglomerate and some of them contain native copper such as those described at Stop 7, Great Sand Bay. Stop 10: Overview at Brockway Nose Latitude: 47.467061N; Longitude: -87.898581W Directions: Return from Hunter’s Point to M-26. Turn left on M-26 towards Copper Harbor and drive 0.3 miles (0.5km) to Brockway Mountain Drive. Turn right, uphill, on Brockway and proceed 0.6 miles (1.0km) to Brockway Nose pullover on the left. Brockway Nose provides an excellent view of Copper Harbor and Lake Fanny Hooe (Figure 23). Copper Harbor and several other harbors between here and Eagle River have the Lake Shore Traps at the harbor entrance as the dipping massive interiors of these basaltic to andesitic lava flows are relatively more resistant to erosion. From Brockway Nose viewpoint, the town of Copper Harbor is a prominent visible feature. The town of Copper Harbor began as a boom town in 1843, following the nearby discovery of native copper. Porter's Island, at the mouth of Copper Harbor on the west side of the harbor's Lake Superior entrance (left) was the site of the first government land office. Hunter's Point is west of Porter's Island (Figure 23). On the east side of the mouth of Copper Harbor, the Copper Harbor Lighthouse, built in 1866, is visible. 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. The vein contained native copper and secondary copper alteration minerals. This location and others in the Keweenaw Peninsula became the foundation of the geological investigations of Douglass Houghton. Houghton's report to the Michigan legislature that sparked the first major mining rush in North America to the Keweenaw Peninsula. Lake Fanny Hooe is located southeast 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. Nearby, the Estivant Pines is a 0.8 mi2 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. 42 �Stop 11: Overview at Brockway Mountain Latitude: 47.464260N; Longitude: -87.969506W Directions: Continue uphill and towards Brockway Mountain for 3.4 miles (5.5km). The top of Brockway Mountain is accessed by continuing upwards from Brockway Nose. Brockway Mountain is a conglomerate ridge that reaches an elevation of over 400 m, with excellent views of the ridge and valley topography of the northern shore of the Keweenaw Peninsula. At Brockway Mountain, the Lake Superior shoreline is oriented about east-west 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, at the southwest end of the rock wall, one can view the dipping conglomerates of the Copper Harbor Conglomerates and see the lava flow near the base of the ridge. To the west, the Lake Shore Traps form island chains and a prominent ridge in the vicinity of Agate Harbor and Esrey Park. The ridges of the Lake Shore Traps and Copper Harbor Conglomerate along the Keweenaw Peninsula’s north shore are also the site of numerous shipwrecks. Lake Bailey (with the small island) and Lake Upson occupy a topographically low valley underlain by a finer-grained clastic horizon (sandstone and siltstone) within the Copper Harbor Conglomerate which was easier to erode by the glaciers than 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 Conglomerate and the uppermost basalt lava flows of the Portage Lake Volcanics. This contact was viewed at Stop 6. 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 as also seen at Stop 3. 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 not easily viewed from Brockway Mountain (better viewed from Brockway Nose). 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 roughly 50 miles (80km) 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 43 �Rift proper extends from the Keweenaw Fault, near 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 on the other side of Lake Superior, just northwest of Isle Royale. Glacial erosion exposed Keweenawan and pre-Keweenawan relatively hard and competent bedrock on the edges of the MCR. 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 1.1 billion year old MCR. 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 Duluth Glacial Lake was the largest of these glacial lakes and only elevations above roughly 400 m (1,300 ft) were emergent such as here at Brockway Mountain and the visible Mt. Bohemia. Stop 12: Float Copper at US-41 Calumet Latitude: 47.241989N; Longitude: -88.448427W Directions: Follow US-41 from Copper Harbor to Calumet. Across the street from the headquarters of the Keweenaw National Historical Park. Google Maps shows the float copper as “Float copper memorial.” Figure 24: Float copper exhibit along US-41 near headquarters of the Keweenaw National Historical Park. 44 �A glacially transported native copper mass is on exhibit at this stop (Figure 24). It weighs 4,263 kg (9,392 lbs) and was found about 7 miles SW of Calumet in less than three feet of surficial sediment/soil. Native copper deposits of the Keweenaw Peninsula were exposed at the bedrock surface at the time of Pleistocene glaciations. The glacial ice entrained masses laying at the surface from previous erosion and plucked masses of malleable native copper from the tabular lodes and veins/fissures. These originally irregular masses were largely cleaned of other minerals and were smoothed and flattened by abrasion from other rocks carried by the glacial ice. The native copper masses were "floating" in the glacial ice, hence locally called “float” copper. When the glaciers retreated about 10,000 years ago, unconsolidated rock debris was left behind by the melting ice as deposits of gravel, sand, and clay. Masses of native copper were scattered among the other sediments carried by the glacier. 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 strata exposed in the Keweenaw Peninsula. Most of the large float copper masses did not move far from their bedrock source in the Keweenaw Peninsula, but smaller masses have been transported quite far and have been found southward in Lower Michigan, Indiana, and Illinois (Bornhorst, 2017). The largest known float copper was discovered in the early 2000s and weighed about 35 tons (~70,000 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 world’s largest existing float copper weighs 26.6 tons (53,200 lbs) was discovered on Quincy Mine claims near Hancock. It is exhibited at a museum in China. The famous Ontonagon boulder was a 1,700 kg (3,708 lbs) float copper mass much smaller than the float copper on exhibit at this stop. The Ontonagon boulder was visited by numerous early European explorers. After Michigan became a territory, Henry Rowe Schoolcraft led an expedition in 1820 with a special goal of seeing the Ontonagon boulder. In 1831, Douglass Houghton accompanied Schoolcraft and visited the boulder too. Pieces of native copper were hacked off of the boulder by Houghton and one of these pieces is part of the University of Michigan mineral collection held by the A. E. Seaman Mineral Museum under the Michigan Mineral Alliance. The Ontonagon boulder was removed from the Keweenaw Peninsula to the nation’s capital in 1843 and is now part of the National Museum of Natural History, Smithsonian Institution’s collection. Float copper masses were altered by oxygenated groundwater and precipitation since the glaciers retreated. Many masses likely had smoothed fresh copper surfaces as abrasion in the glacial ice cleaned off and smoothed the surfaces. The alteration of these surfaces would have occurred in the last 10,000 years. This supergene alteration produced a surface coating on the native copper consisting of varying amounts of 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) (Figure 24). When small 10s of cm sized masses of float copper are cut, the typical surface alteration is often less than several mm thick. While native copper is not stable in contact with water at typical oxidizing surface conditions, the coating of copper oxides, in particular cuprite, inhibits surface oxidation and thereby protects the native copper from extensive alteration. An open access article provides more about float copper (Bornhorst, 2017). The basalt mine rock buildings are part of the Keweenaw National Historical Park. They were once all part of the Calumet and Hecla Mining Company (Bornhorst and Molloy, 2017). The Calumet and Hecla Mining Company was incorporated in 1871 as a consolidation of the Calumet 45 �(formed in 1865), Hecla (formed in 1866), Portland, and Scott Mining Companies. The buildings are built, as are many of the buildings, of local materials, including rock from the Calumet & Hecla Mining Company mines. The national 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 past 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 immigrated to the United States. Older mining districts, such as the Keweenaw Peninsula, 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 this stop whereas now, fewer people live in all of Houghton County. Behind the float copper stands the statue of Alexander Agassiz. Alexander was the son of famous Harvard biologist Louis Agassiz. Alexander was the president of the Calumet & Hecla Mining Company for over 40 years. The statue was moved here from its previous location near Agassiz Park near downtown Calumet in the 1960s. It now stands in front of the Keweenaw History Center, the location of the archives of the Keweenaw National Historical Park. This building was the Calumet & Hecla Library. It is said that at one time this library had more volumes in its collection than the Michigan State Library. Built in 1898, it served as an employee library and bathhouse. The baths were in the basement, until a new bathhouse was constructed in 1911 allowing the basement to be remodeled into additional library space. Stop 13: Freda Sandstone at McLain State Park Latitude: 47.238371N; Longitude: -88.613116W Directions: Follow US-41 to M-203 to McLain State Park. There is an entrance fee. From the pay station turn right towards the campground. Park near the large open area and walk towards the covered shelter and gazebo. Walk down a sandy slope from the gazebo to the lake shore and look around for blocks and slabs of red-colored sandstone. The Freda Sandstone is generally poorly exposed in the Keweenaw Peninsula except for numerous cliff exposures along the shore of Lake Superior southwest of this stop. At this stop, depending on the level of Lake Superior, slabs and blocks of Fred Sandstone can be found along the beach. If the lake level is low enough, then at the shoreline the bedrock of the Freda Sandstone is partially exposed in shallow water and just off shore. The Freda consists of red-colored fine sandstone and siltstone. The red is interrupted by whiteish reduced zones and spots. There are occasional outcrops of Freda Sandstone landward of the Lake Superior shoreline in the area north and south of Portage Channel which are shown on U.S. Geological Survey geologic quadrangle maps and generally located along creeks (Cornwall and Wright, 1956). There are good exposures of Freda Sandstone on the Lake shore between McLain State Park and Porcupine Mountains State Park. The bedding of the Freda at McLain State Park dips about 5OW as compared to the underlying Nonesuch at Hancock Campground and boat launch (Stop 14) where it dips 25OW. This shallowing of dip up-section is typical of the rift-filling strata, and is mostly due to syn-depositional down warping of the rift-filling strata. The Freda Sandstone is generally fine sandstone which is interpreted to have been deposited in a shallow fluvial environment. 46 �There are multiple excellent well-described stops in the vicinity of White Pine, Michigan to examine the Freda Sandstone and these stops are well described by Woodruff et al. (2013). Stop 14: Nonesuch Formation at Hancock Campground and Boat Launch Latitude: 47.133755N; Longitude: -88.620581W Directions: Follow US-41 to M-203 to Hancock Boat Launch and Campground. Drive towards the boat launch and park. At the shoreline you will find a small outcrop of Nonesuch Formation. Walk towards the tree area approximately perpendicular to the Portage Canal shore line and boat launch. About 150 ft from the pavement, you will find the long ago abandoned rock quarry. The Nonesuch Formation is generally poorly exposed in the Keweenaw Peninsula. At this stop the Nonesuch Formation crops out around the margin of a historic rock quarry which is northeast of the Hancock boat launch. Here the Nonesuch Formation is a fine- to-medium grained, gray-to reddish brown sandstone with subordinate interbedded reddish-brown laminated siltstone and shale. The attitude of bedding is about N30OE and 25OW (Cornwall and Wright, 1956). Overall, the Nonesuch Formation consists primarily of siltstone and shale with subordinate amounts of sandstone. At Hancock campground area the formation is coarser grained since this locality is on the northern fringe of the depositional basin centered some 60 km southwest of this stop near White Pine, Michigan. The Nonesuch can be distinguished from the formations below and above by its generally grayish color. Most Nonesuch is a ripple, laminated siltstone with reddish-gray partings. Siltstones and sandstones of the Nonesuch are composed of around 30 to 40 % rock fragments and 60 to 70 % mineral grains. The rock fragments are mostly volcanic with a 2:1 ratio of mafic-tosilicic + intermediate composition (Daniels, 1982). There are multiple excellent well-described stops in the vicinity of White Pine, Michigan to examine the Nonesuch Formation and these stops are well described by Woodruff et al. (2013). Acknowledgments I thank Allan Blaske for his review of this field guide. His comments made this a better guide. References Cited Blattler, C.L., Bergmann, K.D., Kah, L.C., Gomez-Perez, I., and Higgins, J.A., 2020 Constraints on Meso-Neoproterozoic ancient evaporite deposits: Earth and Planetary Science Letters, 532:115951. 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Seaman Mineral Museum, Web Publication 2, 8 p. Bornhorst, T.J. and Barron, R.J., 2013, Geologic overview of the Keweenaw Peninsula, Michigan: 59th Institute on Lake Superior Geology Proceedings, Part 2, Field Trip Guidebook, v. 59, p. 1-42. 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. and Mathur, R., 2018, Copper isotope constraints on the Genesis of the Keweenaw Peninsula Native Copper District, Michigan, USA: Reply. Minerals, v. 8, 508. https://doi.org:10.3390/min8110508. 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Williams, W.C. and Bornhorst, 2023, Controls on the stratiform copper mineralization in the western syncline, Upper Peninsula, Michigan: Minerals, v. 13, p. 927-950. https://doi.org/10.3390/min13070927 53 �Woodruff, L.G., Cannon, W.F., Nicholson, S.W., and Schulz, K.J., 2013, Geology of Keweenawan Supergroup, Porcupine Mountains, Ontonagon and Gogebic Counties, Michigan: 59th Institute on Lake Superior Geology Proceedings, v. 59, Part 2, Field Trip Guidebook, p. 69-96. Woodruff, L.G., Daines, M.J., Cannon, W.F., and Nicholson, S.W., 1995, The thermal history of the Midcontinent Rift in the Lake Superior region: implications for mineralization and partial melting: in International Geological Correlation Program, Field Conference and Symposium on the Petrology and metallogeny of volcanic and intrusive rocks of the Midcontinent rift system, Duluth, Minnesota, v. 336, p. 213-214. Woodruff, L.G., Schulz, K.J., Nicholson, S.W., and Dicken, C.L., 2020, Mineral deposits of the Mesoproterozoic Midcontinent Rift system in the Lake Superior region — A space and time classification: Ore Geology Reviews, v. 126, p. 1–21, https://doi.org /10.1016/j.oregeorev .2020 .103716. 54 �Field Trip 2 Mining History and Geology of the Quincy Mine, Keweenaw Peninsula Native Copper District, Michigan Theodore J. Bornhorst1, James M. DeGraff2, Tom Wright3, and Katherine Langfield2 1 Department of Geological and Mining Engineering and Sciences and A.E. Seaman Mineral Museum, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 2 Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 3 Quincy Mine Hoist Association, 49750 US-41, Hancock, MI 49930 Objectives of Field Trip The historic Quincy Mine was the fourth largest copper mine in the Keweenaw Peninsula native copper district. From 1851 to 1967, mining and processing of native copper ore produced ca. 488 million kilograms of refined copper via several shafts sunk along the top of the Pewabic basaltic lava flows. The Quincy Mine is interpreted and made accessible to the public by the Quincy Mine Hoist Association. The field excursion include tours of: the Quincy No. 2 shaft-rockhouse, the largest steam-driven mine hoist in the world, geology exposed along 675 m of an adit that intersects the historic Quincy Mine workings at the No. 5 shaft on the 7th level (107 ft below surface), the Quincy smelter works on the shore of Portage Lake, and two sites on Torch Lake for extraction of native copper from mined ore. Dress appropriately for the underground portion of the tour where the temperature averages 7o C (45 o F). Walk-in is on a wet surface and will access areas not normally visited by the public. Hard hats and lights will be provided. Brief history of the Quincy Mine Douglass Houghton's copper report in 1841 sparked the beginning of the first mining rush of North America to the Keweenaw Peninsula. This human event has its true beginning in the origin of the rocks and the native copper they host. The human story begins with native people's exploitation of native copper and their introduction of native copper to European explorers. Houghton's copper report led to human migration to the Keweenaw Peninsula, the discovery of the first native copper mine in 1845, the Cliff Mine, and the beginning of modern mining. The mining district quickly developed into the most significant copper district in the United States with peak copper production between 1890 and 1915. The district began to decline from 1920 to the end of mining in 1968 and the decline was accompanied by human exodus; the local population continues to decline today. The Keweenaw Peninsula exhibits the classic boom and bust cycle that is often associated with mining. In 1992, a national park was created to preserve and interpret the rich human and geologic history of the now dormant native copper district. 55 �The Quincy Mining Company was established in 1846 and was termed “Old Reliable” for its reliable payment of dividends to its shareholders. From 1866 to 1890 the Quincy Mine area developed into a robust mine. The Quincy Mine was one of the first mines in the Keweenaw Peninsula to evaluate their operations to maximize efficiency and hence, maximize profits. In 1866 they were the third mine to install a man-engine. They experimented with Burleigh power drills in 1868, but the large and cumbersome drills proved a failure. They built a rock house complete with rock crushers in 1873 which eliminated the time and labor involved in the old kiln houses. In 1879, black powder was replaced by dynamite and most hand drilling was eliminated when they adopted Rand Company's "Little Giant" power drills. While these innovations were adapted to increase efficiency, other changes were made to solve problems. The Quincy Mine could not find enough water in the abandoned shafts for its steam plant near the mine up on the top of Quincy hill which led to installation of a pumping station to pump water uphill from Portage Lake. In 1887, they closed their stamp mill on Portage Lake as the tailings they deposited into the lake began to hinder navigation. Quincy proceeded to build a new mill on Torch Lake. By building a new mill they were able to adopt steam-powered stamps rather than the old Cornish drop stamp technology (Lankton and Hyde, 1982). Quincy was able to continue to adapt because they had sufficient copper ore to mine. The Quincy Mining Company began to rebuild itself during the boom years for the whole Keweenaw Peninsula native copper district from 1890 to 1915. It acquired more powerful engines and hoists and began to raise two skips at a time in balance with each other. In 1891, they built the photogenic and functional No. 6 combined shaft and rockhouse. The man-engine reached its maximum depth in 1892 and as a result Quincy began using man-cars to raise and lower workers using the same hoist they used for ore skips. They built their own smelter on Portage Lake on the old Pewabic Mill site in 1898. In 1905 they began to replace their 2-cylinder engines with 4-cylinder engines to get better use of the steam and reduce coal consumption. Underground mining was changing too. Miners switched from candles to paraffin-based lamps in 1896-97 and to the less-smoky calcium carbide lamps in 1912-14. Miners began using machines to cut up masses in 1906 rather than drills. Quincy became the first mine in the district to modernize tramming and hauling when they installed a haulage system with a GE batterystorage locomotive in 1901. Quincy designed and patented their own side-dumping tramcars. By 1903 Quincy had 15 electric locomotives. They experimented without success with power shovels underground. In 1905, they reduced the wait time to dump cars underground by digging 500-ton underground storage bins into the hanging wall above the inclined shafts and thereby were able to hoist 25 % more ore to the surface. In 1910, Quincy operated 160 two-man drills, each of which weighed 245 pounds. The lighter one-man drills significantly decreased the cost of mining but this new technology left a miner alone underground which became one issue of the 1913 strike (Lankton, 1991). The Keweenaw Peninsula native copper mining district never fully recovered after the 1913 strike. The Quincy Mining Company continued to move forward after the strike, and they installed the world's largest steam hoist. Quincy constructed the No. 2 hoist house in 1918 and began operating the world’s largest steam hoist in 1920 as the shaft reached an incline depth of 7750 (2360 m) feet. In 1920 Quincy increased efficiency of processing and smelting by adding 56 �Wilfley tables to their stamp mill and adding a casting wheel at the smelter to eliminate hand pouring of ingots. As the underground workings went deeper, mining became much more difficult. In 1926, they had to install additional pumps to reduce the time used waiting for hoisting water-bailing skips rather than rock skips. In 1927, there was a major underground fire that impacted production. As the mine went deeper, they had to install fans to reduce the temperature at the bottom where miners still sometimes had to work at temperatures of 98oF (Lankton and Hyde, 1982). The Great Depression closed most Keweenaw mines. The Quincy Mine closed in 1930. The Calumet & Hecla Mining Company closed most of its mines except for a lucky few that operated on reduced shifts. The Quincy Mining Company reopened the No. 6 and No. 8 shafts on a limited basis in 1937. It wasn't until World War II and the advent of price controls that a larger production schedule was resumed. District-wide production resumed on a broader scale when price controls were lifted on August 31, 1945 and Quincy permanently ended its underground mining. The reclamation mill continued to operate and make a profit. In 1948, on its 100th anniversary, Quincy paid a dividend of 25 cents per share due only to the copper produced by the reclamation from tailings in Torch Lake plant (Lankton and Hyde, 1992). The Quincy smelter reopened in the mid-1950s after being closed for about 15 years as Quincy could no longer send reclamation concentrates to the Calumet & Hecla Mining Company smelter. The Quincy smelter remained open until the reclamation plant closed in 1967. The mines were allowed to begin flooding in 1970 (Thurner, 1994). The Keweenaw Peninsula native copper mining district has remained dormant in the 54 years to follow except for several episodes of exploration. Highland Copper Company, from 2011 to 2015, has been the latest explorer to attempt to reopen native copper mines in the Keweenaw Peninsula. The economies of mining districts are typically sustained by one principal industry, mining. When the mines are profitable and expanding the local communities also do well and when the mines suffer decline the local communities also suffer decline. This creates the boom to bust mining cycle. The local population follows this boom-and-bust trend. During the boom of mining Calumet was a vibrant city and since the mines have closed it has contracted. The last value obtained from the mining industry is selling the useful equipment to other mines followed by the dismantling of buildings to sell for scrap. Evidence of mining such as shaft-rockhouses and industrial buildings disappeared too. The surface rock piles left from mining once dotted the landscape of the Keweenaw Peninsula, but over time they too are disappearing as they are an inexpensive source of crushed stone for roads and other purposes. In the late 1980s Calumet community leaders envisioned that the past might be the key to the future of Calumet and they sought development of historical tourism. This led to creation of the Keweenaw National Historical Park in 1992. The park consists of limited park owned lands/structures and multiple Keweenaw Heritage Sites which are public, private, and non-profit. Together the park and its cooperative sites preserve and interpret the mining history of the Keweenaw Peninsula and supports historical tourism. The brief history of the Quincy Mine was slightly modified from text written by Larry Molloy, published by Bornhorst and Molloy (2016), and republished by Bornhorst (2022). 57 �Geologic Overview of the Quincy Mine Readers are referred to Bornhorst (this volume) and Bornhorst and Lankton (2009) for the geologic setting of the Quincy Mine and its native copper deposits hosted by the Portage Lake Volcanics. This overview of the geology of the Quincy Mine is from Bornhorst and McDowell (1992), Bornhorst et al. (1986), and Butler and Burbank (1929). The Portage Lake Volcanics comprise several hundred subaerial lava flows erupted within the Midcontinent Rift of North America about 1.1 billion years ago. There are occasional interbedded sedimentary rock layers dominated by conglomerate (Bornhorst, this volume). The native copper ores at the Quincy Mine occur in tabular bodies hosted by the originally porous and permeable tops of subaerial basalt lava flows. Ore-forming hydrothermal fluids precipitated native copper in the open spaces some 30 million years after the lava flows were erupted. About 50 lava flows of the Portage Lake Volcanics are exposed along the adit, twelve of them beneath the interflow sedimentary layer termed the Allouez conglomerate. The Allouez conglomerate and overlying Greenstone flow can be traced from the Houghton-Hancock area to the tip of the Keweenaw Peninsula (Bornhorst, this volume). A clay gouge along a bedding plane fault with undetermined slip occurs at the top of the Allouez conglomerate (Bornhorst and McDowell, 1992). Similar bedding plane faults are common at the tops and bases of conglomerate layers throughout the district. Native copper ore at the Quincy Mine is hosted by a group of relatively thin porous tops of lava flows that are difficult to correlate laterally without mapping the flows in detail (Butler and Burbank, 1929). The productive tops of the lava flows at Quincy, termed the Pewabic amygdaloids, are not brecciated as is common in the larger flow-top mines in the district. Cavernous zones within the Pewabic flows began as open spaces and subsequently were filled with hydrothermal minerals described above. There is much variability within the tabular lode from well to poorly banded and from high-grade of copper ore to practically barren poor rock (Butler and Burbank, 1929). Stratigraphically, about 12 lava flows with a cumulative thickness of about 100 m occur between the Allouez conglomerate and the overlying Pewabic flows which are the host rocks for the native copper deposits at the Quincy Mine. However, at this location along the adit, the Pewabic flows are not mineralized as they are on the northwest side of the Hancock fault. There are about 14 additional lava flows until the Hancock fault is reached. The Hancock fault is marked by a distinctive clay gouge, almost pure corrensite, and a green corrensite-rich brecciated mineralized zone adjacent to the gouge (Bornhorst and McDowell, 1992). The Hancock Mine produced native copper from a mineralized segment of the Hancock fault, suggesting that it may have been a feeder of ore-forming hydrothermal fluid into void spaces in the tops of the Pewabic flows (Bornhorst and McDowell, 1992). The amygdaloidal flow tops exposed by the adit are filled with a number of hydrothermal alteration minerals. Butler and Burbank (1929) describe the main-stage hydrothermal minerals at the Quincy Mine: quartz and calcite are abundant throughout the lode, commonly as euhedral crystals in open cavities; pumpellyite is less abundant but is present throughout the lode; epidote is less abundant than 58 �pumpellyite but is a common hydrothermal mineral; chlorite is particularly abundant in amygdules near the bases of Pewabic lava flows and is locally replaced by quartz and calcite; prehnite is present but not common; and datolite is present only in upper levels of the mine. The Pewabic lode of the Quincy Mine is notable for spectacular euhedral calcite with visibly unaltered pink to rose colored inclusions of native copper, which are highly sought by mineral collectors. Native silver is closely associated with native copper, although abundance is low. Laumontite is sparse and associated with small fissures. Thus, it may be a late-stage hydrothermal mineral rather than main-stage (Bornhorst, this volume). Butler and Burbank described early pumpellyite and epidote followed by quartz, calcite, and native copper. Bumgarner (1980) described amygdules indicating that prehnite and chlorite were early hydrothermal minerals followed by quartz, then chlorite, and lastly calcite. Overview from native copper ore to copper products At the Quincy Mine and elsewhere in the Keweenaw Peninsula native copper was extracted from tabular ore bodies by underground mining methods. Quincy Mine yielded about 42,870,000 tons of ore (1 ton = 2,000 lbs) with recovered refined copper totaling 1,077,000,000 lbs at an average grade of 1.26 % copper per ton of ore. The amount of silver in the native copper ore can be estimated from incomplete production statistics in Butler and Burbank (1929) to be roughly 0.2 oz of silver per ton of ore. The ore is blasted underground into small enough size to be able to be hoisted to the surface. Rock lacking sufficient native copper was sent to the poor rock pile or into abandoned underground mine openings. Prior to being sent to the mineral processing plant to recover native copper, the broken ore is sized by a slatted grating, “grizzly”. The broken ore passing through the grating (most of the ore) is sent to the mineral processing plant and those fragments too large are broken further at the surface near the shaft rock-house with a steam driven hammer. One reason a fragment could be too large is because it is mostly native copper. Native copper is malleable and large masses are not readily fragmented by underground blasting. Once most of the rock was removed from the larger fragments, the copper-dominant fragments were put into barrels (barrel copper) or, if they were too big, they were put onto a flat rail car and shipped directly to the smelter instead of to the mineral processing plant. The ore from the Quincy Mine was processed to separate native copper from the barren host rock and barren minerals in order to produce a product sufficiently enriched in copper to be smelted. The ruins of the Quincy processing plant will be visited at Stop 4. The first essential step in processing was crushing the ore into sand and smaller sized fragments using stamp mills. The crushing aims to produce fragments which are mostly native copper or mostly rock and thereby liberates the native copper from the rock. Because native copper is much denser than the host rock (about 3 times denser than the barren host rock), fragments of native copper can be separated from the barren host rock using water and gravity methods such as jigging. Hence, the mineral processing plant usually was constructed near a body of water. The mineral processing plant produced a “concentrate” which was composed of sand-sized and finer native copper with some fragments or partial fragments of barren rock because no separation method is able to completely separate every particle. The ore 59 �mined at Quincy averaged about 25 lbs of copper in a ton of ore. The copper concentrate was more than 50 % copper, hence most of the sand and smaller size fragments from the crusher ended up being waste, which still contained some copper. This waste, called tailings, was transported by water slurry and dumped into lakes, especially Torch Lake. The amount of native copper in a sand-sized fragment may have been so low that it was correctly separated into the tailings. In other fragments there was enough copper in them but they were incorrectly separated into the tailings. For example, the copper could have been finer grain size than the fragment and it was diluted by barren host rock and minerals. By crushing such fragments to a finer grain size, more copper could have been liberated from the barren host material. Quincy Mine kept track of the lost copper as it dumped tailings into Torch Lake. They later went back and recovered much of the lost copper by reprocessing the tailings using newer more efficient technology. The mineral processing plant and recovery of lost copper are discussed at Stop 4. The copper concentrate from the processing mill was shipped to the Quincy Smelter Works which is discussed further at Stop 2. The smelting and refining process resulted in solid Quincy copper ingots of up to 99.8 % pure copper. The copper ingots transported to markets by ships. The copper mined and processed by Quincy was sufficiently pure to be fabricated into a variety of usable copper products such as electrical wire. Copper ore to copper products text from previously published field trip guide by Bornhorst (2022) with modification. Field Trip Stops Figure 1: Map showing approximate stop locations of Field Trip 2 to the Quincy Mine. 60 �Stop 1: Quincy Mine, Keweenaw Heritage Site of the Keweenaw National Historical Park Latitude: 47.137137N; Longitude: -88.574875W Directions: From Michigan Tech drive west on US-41 (left from parking lot by the MUB) and continue through downtown Houghton across the Portage Lake lift bridge through downtown Hancock and uphill to the Quincy Mine No. 2 shaft rock-house turning right into the parking lot just past the shaft rock-house. The Quincy Mining Company was incorporated in 1846 and operated until 1967. Quincy mined underground from nine shafts on the Pewabic flow top and there was an industrial complex associated with mining (Figure 2). Throughout its history, Quincy Mining Company paid dividends on such a regular basis it was nicknamed "Old Reliable". Quincy Mining Company produced a total of 1.08 billion pounds of refined copper and approximately 100 million oz of silver from approximately 43 million tons of ore at an average grade of 25.1 lbs. of copper per 2000 lb. ton of ore (including copper reclaimed from Quincy tailings). The Quincy Mine ranks as the fourth largest mine in the native copper district. In 1921 the No. 2 shaft was the world's deepest. The Nordberg steam hoist is the world’s largest. Figure 2: The Quincy Mining Company complex ca. 1900. The No. 2 shaft-rockhouse is near the center of the drawing. From Molloy (2011) with permission. 61 �The No. 2 shaft of the Quincy Mine opened in 1858. At the beginning of mining a simple house was built over the shaft. By 1892, Quincy introduced the concept of hoisting the ore and doing initial crushing and sorting of the ore in the same building, a shaft-rock house. The current Quincy No. 2 shaft rock house was built in 1908 (Figure 3). The Quincy No. 2 shaft rock house is 147 feet (45 m) tall and the angle on the side of the building facing US-41 is at the dip angle of the native copper deposit. Behind the shaft rock house, there are two of the original eight pulley stands and stanchions that were used to support a steel cable extending to the No. 2 hoist house built in 1919. Mining at the No. 2 shaft ended in 1931. Figure 3: Quincy Mine No. 2 shaft-rockhouse. This drawing is based on Historic American Engineering Record drawing, MI- 2,19/34, Durward W. Potter, Jr., 1978 and Richard K. Anderson, Jr., 1979. From Molloy (2007). By 1917, the No. 2 shaft had reached such great depths that the hoist engine housed in the 1895 hoist house was no longer adequate. The Quincy Mine needed a large and faster hoist to continue its production. In 1918, the No. 2 hoist house was constructed but World War I delayed 62 �delivery of a new hoist until 1919. The Nordberg hoist began operating in 1920 as the shaft reached an incline depth of 7750 (2360 m) feet. The Nordberg hoist consists of four cross-compound steam engines that work as one (Figure 4a and 4b). The new hoist could move an ore skip carrying 10 tons of rock (13 tons total weight) up at 3200 feet per minute (36 miles per hour) and was more energy efficient than the hoist it replaced. The Nordberg hoist, the world's largest, operated 24 hours per day for 11 years until mining ended in 1931; to a depth of over 9000 feet (2743 m) on the incline (Molloy, 2007). Stairs - down Entrance from the 1895 Hoist House Down Hoist Rope Slots DisplaysModel of #6, Mine Cross Sections Up Overhead Crane Low Pressure Cylinder High Pressure Receiver Oil For Hydraulics Stored Under Here Low Pressure Receiver High Pressure Cylinder Condenser Under Here Oiler's Gallery Hoisting Drum 30' Maximum, 16' Minimum Diameter D i s p l a y s Vacuum Pump Low Pressure receiver Miniatures High Pressure Cylinder Displays High Pressure Receiver Operator's Platform Overhead Door Stairs to Platform Top Of Water Circulating Pump Lily Hoist Controler Displays Oil Pump Drive Low Pressure Cylinders Up Corliss Steam Engine Flywheel Display Of Large Tools Down Figure 4a: Quincy Mining Company No. 2 shaft Nordberg hoist diagram. This drawing is based on HAER drawing, MI-2,14/34, Durward W. Potter, Jr., 1978. From Molloy (2007) A steam engine functions by using alternating intake and exhaust valves to allow steam to enter a chamber, expand inside of the chamber, and use the force of the expansion to push a piston as the steam expands. The double-acting pistons used here can be pushed both up and down in the cylinder by the expanding steam. The entire engine occupies 60 feet by 54 feet of floor space and is 60 feet tall. It is a cross-compound steam engine, an engine where steam is used twice. This common technique accounts for the “choo-choo-choo-choo” sound one hears near steam powered trains. Today the hoist remains an engineering marvel and is still the world’s largest steam mine hoist. 63 �Miniatures H ois ting D rum Operator's Platf orm Brak e Main C rank Oiler's Gallery Low Pres s ure Throttle H igh Pres s ure Throttle Steam Supply Line H igh Pres s ure Low Ex haus t to Pres s ure Equalizing Line Pres s ure C ondens er C y linder C y linder H igh Press ure Low Pres s ure Steam R ec eiv er Steam R ec eiv er Figure 4b: This diagram illustrates the major features of the hoist and traces the flow of steam through the hoist. It is based on HAER drawing, MI-2,13/34, Jon R. Carter, 1978. From Molloy (2007). At the Quincy Mine and elsewhere in the Keweenaw Peninsula native copper was extracted from tabular ore bodies by underground mining methods. The ore is blasted underground into small enough size to be able to be hoisted to the surface. Prior to being sent to the mineral processing plant to recover native copper, the broken ore is sized by a slatted grating, “grizzly.” Those fragments too large are broken further by steam hammers. The native copper-dominant fragments were put into barrels (barrel copper) or, if they were too big, they were put onto flat rail cars. The barrel copper and larger masses were shipped directly to the smelter located downhill from the shaft rock-house. Next to the Quincy Mine No. 2 shaft rock house is a specimen of mass copper weighing hundreds of lbs. that would have been shipped directly to the smelter. Text from previously published field trip guide by Bornhorst (2022) with limited modification. 64 �Stop 2: Quincy Smelter Works Latitude: 47.126688N; Longitude: -88.565290W Directions: From Stop 1 turn left from Quincy Mine returning downhill to Hancock, just before the Portage Lake lift bridge follow M-26 through underpass and from bridge continue 0.4 miles (0.65 km) to the Quincy Smelter Works. The Quincy Smelter was constructed in 1898 and initially consisted of four or five reverberatory furnaces until switching to two larger furnaces in 1920. The purpose of the smelter was to produce bars of copper with as few as possible impurities. Copper with low enough impurities was ready to be made into copper products. The copper concentrate from the mill, barrel copper, and larger masses are melted, turned from solid to liquid in a furnace. Smelting is melting at temperatures higher than the melting point. By being higher than the melting point the liquid metal separates from the liquid silicate rock. Since the density of metallic copper is much higher than other components in the liquid rock, the liquid copper sinks to the bottom of the furnace while the other components, molten slag (waste liquid) float to the top. Men skimmed off the molten slag and it was dumped on a pile where it was left to cool into glass and minerals. Across from the parking lot opposite the agent’s house/office is a pile of solidified slag. The molten copper was tapped from the bottom of the furnace into a refining furnace where it was rabbled and poled. Rabbling (stirring) agitates the molten copper and allows the introduction of small amounts of air which oxidizes the impurities (copper has less tendency to oxidize than impurities). The oxidized impurities are lighter than liquid copper and thus, rise to the surface and are skimmed off the molten copper. After the impurities are skimmed off of the molten copper, green sapling poles are inserted into the molten copper (poling). Poling introduces carbon (wood) which reduces the amount of copper oxides in the molten copper. The refined copper was ladled and poured into molds and sprayed with water to cool it. The solid copper shapes were then dumped into a tank of water to complete cooling. Quincy cast copper ingots of several shapes and sizes (bars, wedges, and cakes). Refined copper from Quincy was up to 99.8 % copper which was sufficient for fabrication of copper products in the early 1900s. Some copper ingots had to undergo further refining. Today, the copper would undergo electrolytic refining to make it more than 99.99 % pure copper. Quincy constructed the smelter to be able to refine and ship its own copper. It also accepted custom work from neighboring mining operations. While underground mining ended in 1945, the Quincy smelter remained open until 1971. After closure of the underground mine, Quincy was actively recovering copper from reprocessing of tailings from Torch Lake. Fortunately, the smelter site remained intact after closing and is now open for guided tours. In 2014, the Keweenaw National Historical Park Advisory Commission acquired the smelter from Franklin Township (Figure 5). Text from previously published field trip guide by Bornhorst (2022) with limited modification. 65 �Figure 5: The historic Quincy Smelter Works. View July 2022 towards the north from Houghton waterfront. Stop 3: Quincy Mine Adit to 7th Level Underground Workings Latitude: 47.130383°N; Longitude: 88.573824°W Directions: Turn left onto M-26 heading west and drive 0.5 mi (0.8 km) to its junction with US41 north of the Portage Lake lift bridge. At the stop sign, merge onto US-41 and continue west for two blocks (~0.1 mi, ~0.15 km) to Dunstan Street on the right before a BP gas station. Turn right onto Dunstan and drive uphill for three blocks on the right (~0.15 mi, ~0.25 km) to Mason Avenue. Turn right onto Mason and drive east about 0.2 mi (0.3 km) to near a sharp right-hand curve. Just before the curve, turn left to enter an unpaved driveway that leads to the adit. The Quincy adit was first opened in 1892 by the Quincy Mining Company and subsequently widened in the 1970s by Michigan Tech mining engineering faculty and students for an underground educational and research facility. The adit enters the south side of Quincy hill at a point located ~690 m SSE of the iconic Quincy No. 2 shaft house along US-41 at the top of the hill. From the portal, the adit follows an azimuth of 331° for 675 m, where it intersects the 7th level workings at the Quincy No. 5 shaft (Bumgarner, 1980). The average strike of PLV layers here is 213° (right-hand rule), which means that the adit cuts across them in a direction 30° clockwise from the dip direction. Approximately fifty lava flows and one conglomerate layer, dipping on average 50° NW, are exposed along the adit (Bumgarner, 1980; Bornhorst et al, 1986). Geologic sites of interest in the adit and side passages are described as sites 3A to 3E in the following text and shown in Figure 6. Underground workings of the Quincy Mine were developed along a series of parallel, relatively thin, lava flow tops, referred to collectively as the Pewabic flow top within the Portage Lake Volcanics. The mine was developed to a vertical depth of 1897 m (6225 ft) comprising 92 levels. The ore body decreases in dip from 54o at the surface to 32o at the bottom levels. Pewabic flows are characterized by large cavernous zones up to 1.5 m thick, interpreted by Butler and Burbank (1929) as coalescing vesicles and large gas cavities. Alternatively, the large cavities could have 66 �been caves or lava tubes formed by lava drainage beneath a solidified flow surface. Openings were connected up to 100s m along formation strike and, where especially well developed, they may host 2 to 10 high-grade copper zones. An interconnected series of such openings would have provided continuous flow paths for ore-forming hydrothermal solutions. Several prominent and steeply dipping veins extend throughout the mine. They probably helped to integrate the hydrothermal system as their mineralogy is similar to that of the hydrothermal mineral assemblage in the flow tops. Figure 6: Map of Quincy adit and 7th level mine workings (modified from Bumgarner, 1980). Six geological sites of interest are labeled 3A through 3F. Numbers 1 – 19 along adit are distances in hundreds of feet. 67 �Site 3A: Contact between two basaltic lava flows. [~75 m from portal] The contact between the two lava flows at this site shows many characteristics that are common in the Portage Lake Volcanics of the Keweenaw Peninsula. Contacts between successive subaerial lava flows are recognized by textural and color differences between the top of the older layer and base of the younger one, as well as by the geometry of their boundary. Massive flow interiors grade into margins with finer grain size and abundant vesicles often arranged in bands parallel to flow contacts. Flow tops have the highest abundance of vesicles. Upper flow surfaces were exposed to the atmosphere and escaping gases during and after emplacement, and later to hydrothermal fluids permeating along flow tops that filled many voids (amygdules) with an assemblage of minerals (Butler and Burbank, 1929; Stoiber and Davidson, 1959; White, 1968; Bornhorst, 1997), thus modifying their color as seen here. Looking at the adit’s northeast wall (right side walking in), the top of the older flow toward the southeast is highly vesicular and has a lighter greenish tone relative to the younger to the northwest. In this case, the distinctive lighter tone and greenish tinge of the flow-top rock results from secondary epidote, pumpellyite, and various copper minerals, which are not present in the base of the overlying flow. The adit cuts nearly perpendicular across the contact between the flows, such that their boundary can be traced from one side of the adit across its ceiling to the other side. This exposure is nearly ideal for measuring strike and dip of the contact because it is exposed in 3-D over a sufficient distance to allow visual averaging of irregularities that are typical of flow contacts. Site 3B: Allouez conglomerate and the Greenstone flow. [~290 m from portal] The Allouez conglomerate and overlying Greenstone flow are well correlated stratigraphic units along the strike of the historic copper mining district (Butler and Burbank, 1929; Stoiber and Davidson, 1959). They can be traced from near the tip of the Keweenaw Peninsula to southwest of the Quincy Mine, a strike length of 80 kilometers. In the Quincy adit, the Allouez conglomerate is a 2-m-thick layer and the overlying Greenstone flow is only 10-15 m thick, far less than its maximum thickness elsewhere. The conglomerate here is typical of most conglomerates interbedded with mafic lava flows of the Portage Lake Volcanics. It is largely clast-supported and contains rounded to subrounded cobbles and pebbles of felsic volcanic rocks and subordinate mafic rocks. At this site, the bedding surface between conglomerate and basalt is marked by a clay seam, or “fluccan” following Cornish mining terminology (Hubbard, 1898). Hubbard (1898) systematically documented such clay seams in mines accessible at the time and he noted that they typically follow contacts between interflow sedimentary layers and lava flows, but also occur at contacts between flows. The clay seams are commonly associated with polished surfaces, fault gouge, brecciation and alteration of adjacent units, and secondary mineralization in fractured zones generally less than a couple of meters thick. As noted by Hubbard (1898), these phenomena indicate that many layer boundaries in the Portage Lake Volcanics have slipped during one or more deformation events. Such surfaces are essentially layer-parallel faults that slipped because of their weaker mechanical strength relative to the layers themselves. 68 �Site 3C: Hancock fault at the adit and in drift to the northeast. [~570 m from portal] The Hancock fault is a major splay of the Keweenaw fault system that intersects the main Keweenaw fault about 14.5 km northeast of here and downdip at a depth of around 3 km (Cannon and Nicholson, 2001; DeGraff and Carter, 2023). Over its 18-km length, the Hancock fault strikes 234° (right-hand rule) and cuts obliquely across the exposed Portage Lake Volcanics in a direction that is ~25° clockwise from formation strike and dips northwest at a steeper angle than layering. It cuts the Pewabic flows of the Quincy and Hancock mines. The Hancock fault had a strong influence on the distribution of copper in these two mines as follows: 1. A portion of the fault in the Hancock Mine had copper ore that was exploited; 2. Copper is also abundant along the Hancock fault in the Quincy Mine; 3. Copper mineralization in Pewabic beds of the Quincy Mine is restricted to the hanging wall northwest of the fault. The Hancock fault has been proposed as an important pathway for ore-forming fluids to access porous zones at the Quincy and Hancock Mines (Bornhorst et al., 1986). Figure 7: Cross-sectional view of the Hancock fault looking northeast at the wall of the adit. Height of the view is about two meters. The slip surface is marked by a clay seam. 69 �The Hancock fault is best exposed and investigated in the Quincy Mine workings, where it is cut by the adit and drifts leading northeast and southwest from the adit (Figure 7). At this site, the Hancock fault is observed crossing the adit and following the northeast drift for about 50 meters, so that measuring its orientation is easy. At the adit, the fault strikes 239° and dips 54° NW, in contrast to stratigraphic layering that strikes 213° and dips a bit less than layering. The fault zone here has a thin (~3 cm) medial clay seam, marking the main slip surface, that is contained within a breccia envelope of disaggregated fragments. Outside of the breccia envelope, country rock is highly fractured near the fault but the rock is still intact. As the fault is followed northeastward along the drifts orientation varies and the thicknesses of its clay seam and surrounding breccia increase and decrease and are not always symmetrically arranged. Site 3D: Ropy pahoehoe at the top of a Pewabic flow in the first northeast drift. [~570 m from portal; ~75 m along drift] Farther along the same drift at Site 3D is a good exposure of the surface of one of the Pewabic flows (Figure 8). Like subaerial mafic lava flows elsewhere, the tops of Portage Lake Volcanics flows show characteristics of two fundamental types – a smooth or ropy geometry (pahoehoe, locally termed amygdaloid) and a rough blocky geometry (aa, locally termed fragmental amygdaloid). Most Portage Lake Volcanics lava flows have smooth flow tops and yet ropy pahoehoe is rarely observed, probably because plan views of flow surfaces are uncommon and because of degradation of flow surfaces between successive extrusive events. The ropy pahoehoe observed on the northwest side of the drift is well preserved. The view here is upward at the base of an overlying flow, and so the pahoehoe feature is really a mold of the upper surface of the underlying flow top that has been removed along the drift. The curved geometry of pahoehoe ropes results from local movement of partly solidified lava crust that forms arcuate patterns that are convex in the local direction of flow (Fink and Fletcher, 1978; Self et al., 1998). The local flow direction indicated by the convexity of the ropes seen here is updip and toward the southwest, i.e. from the center towards the edge of the rift. Most flow-top native copper ore bodies are hosted by brecciated flow tops because their relatively large volume of open space was well connected and provided relatively easy movement of ore-forming fluids (Butler and Burbank, 1929; White, 1968; Bornhorst, 1997). Flows with smooth to ropy pahoehoe tops are usually not favorable because vesicles have limited volume of not well connected open space which hindered movement of ore-forming fluids. However, Pewabic flows with pahoehoe tops had economic grades of copper mineralization because they had large connected openings up to 1.5 m wide and extending up to 100s m along strike. Butler and Burbank (1929) proposed that the exceptional Pewabic lode was the result of an extremely gaseous flow that produced a great abundance of vesicles that coalesced to produce large connected voids. An alternative hypothesis is that the large openings resulted from lava draining from pools and tubes beneath a solidified flow surface. 70 �Figure 8: Ropy pahoehoe in the hanging wall of the first drift northeast from the adit. View is to the northwest. The two highlighted patterns indicate flow updip and toward the southwest. Height of the view is about six meters. Site 3E: Hancock fault in drift to the southwest. [~675 m from portal; ~55 m along drift] This site provides another opportunity to examine the Hancock fault where it follows much of the drift leading southwest from the No. 5 shaft to the No. 7 shaft (Figure 6). From Site 3C at the adit, the fault angles across the corner of unmined rock between the adit and southwest drift and is next exposed on the southeast wall of the drift at Site 3E. From here, the fault trace rises up the southeast wall, gradually angles across the ceiling, and descends the northwest wall where it enters the drift’s hanging wall before reaching the No. 7 shaft. Native copper is frequently found in the fault’s hanging wall up to the breccia zone but has not been found in the footwall. Here, the “medial” clay seam is generally thicker than at Site 3C but its thickness varies considerably along the fault trace, as does the thickness of the disaggregated breccia envelope. It is difficult to estimate the thickness of the still-intact zone of fractured rock 71 �that encloses the breccia zone because, as seen at other faults examined in detail (Caine et al., 1996; Caine and Forster, 1999), its intensity gradually decreases away from the fault to a background value for the system. Study of exposures of the Hancock fault and nearby satellite faults in the Quincy Mine is needed to better understand the slip characteristics and wall-rock modification of the Hancock Fault. This could also help understand the influence of the Hancock fault on copper mineralization. Fault slip is being investigated by the measurement and analysis of slip indicators – slickenlines and steps – on the surfaces of the many small faults associated with the Hancock fault. Observations of how rocks along the Hancock fault were modified physically and chemically have been mostly qualitative so far, though with some exceptions (Bornhorst and McDowell, 1992; Langfield et al., 2023). Site 3F: Hanging wall of Pewabic flow in drift northeast of No. 5 shaft. [~675 m from portal; ~30 m along drift] Along the drift leading northeast from the No. 5 shaft to the No. 4 shaft (Figure 6), the hanging wall of the drift and attached stope exhibit slickenlines directed parallel to dip of the lava flows. One may need to search a little and use oblique lighting to see these fault surface features, which manifest layer-parallel dip slip on the boundary between two lava flows. Another surface with similarly oriented slickenlines occurs in a parallel drift a few tens of meters northwest beyond the hanging wall of this drift. As discussed at Site 3B, Hubbard (1898) documented 12 layer boundaries with layer-parallel slip, but this surface and the other northwest of here were not among them. He mentioned polishing and slickensides on such surfaces but did not specify their orientation. It is likely that such layer-parallel slip is far more common than has been observed because, like the ropy pahoehoe texture, it can only be observed where an ideal exposure permits viewing. The documented layer-parallel slip within the PLV section is strong evidence for a detached style of thrusting, which was used recently to model the cross-sectional geometry of the Keweenaw and Hancock faults (DeGraff and Carter, 2023). 72 �Stop 4: Quincy Mining Company Processing Plant and Dredge Latitude: 47.146295N; Longitude: -88.460474W, Directions: From Quincy Smelter Works turn right (east) and continue 5.7 miles (9.2 km) to ruins of the Quincy processing plant on left carefully pulling into open lot on west side of ruins. The Quincy Mine dredge No. 2 is located on the edge of Torch Lake on the other side of the road. Figure 9: The Quincy Mining Company mills, 1890-1928. Star shows the location of the Quincy dredge No. 2. From Molloy (2011) with permission. The Quincy Mining Company had to move their mill from its Portage Lake site below Quincy Hill because tailings deposited in the lake were beginning to hinder navigation. Construction began on the Quincy mill at this site in 1888 (Figure 9). Mill No. 1 began with three rock crushing stamps and two additional were added in 1892. The building closest to the dredge on the north side of the road contained the 1890 mill which was modified over time. The square building adjacent to it was a turbine building that was built in 1921. As production increased, Stamp Mill No. 2 was built to the north of the No. 1 mill in 1900 and had three stamps. Underground mining activities at the Quincy Mine focused on mining ore (rock with sufficient recoverable copper to make a profit). Inevitably the mine also produced waste (uneconomic) rock along with the ore rock. Some of this waste rock can be seen today throughout the Keweenaw Peninsula as poor rock piles. Some of the waste rock was used underground as fill for already mined out stopes. The broken ore from underground blasting that is small enough to pass through a coarse grating was sent by train to the processing mill. 73 �The purpose of the mill was to remove as much rock as possible from the native copper producing a product where the percent of contained copper is up to 20 times higher than the percent of copper in the ore. The copper-rich product is sand to smaller sized fragments of native copper mixed with similar sized rock and mineral fragments (termed “copper concentrate”). The inefficiencies in separating native copper from rock and minerals fragment results in some amount of waste rock and minerals in the copper concentrate. At the end of processing in the mill, most of the rock and minerals end up as fine sand to sand of waste containing only a small amount of copper (termed tailings). The Quincy Mill used several techniques to separate the copper from the rock and minerals. The essential first step in milling begins with crushing the ore rock to a small enough size to liberate the unwanted rock and minerals from the native copper. The ore rock and minerals containing disseminated copper of various sizes was crushed by steam stamps. Much of the rock was liberated from the native copper by crushing the broken ore fragments derived from mining to fragments 0.0165 to 0.188 inches in diameter (fine to very fine sand size). Larger masses of native copper were removed prior to crushing at the mine and sent directly to the smelter. Since the density of native copper is much higher than the liberated particles of rock and mineral, gravity and water methods could be used to separate the waste rock and minerals (tailings) from the copper. Early mines used jigging to concentrate the native copper from the tailings. Jigging was a well-developed mineral separation technology prior to the first mining of native copper in the Keweenaw Peninsula beginning in 1845. Jigging is accomplished by placing the sand sized particles on a screen where pulsating water allows the heavier copper particles to settle while the lighter particles rise to the top and overflow the screen or are skimmed off the top as waste tailings. Jigging resulted in a “concentrate” of sand sized native copper particles with some rock and minerals that were not copper-bearing (tailings), about 50 % copper and the rest waste rock and mineral. The concentrate was sent to the smelter to be turned into nearly pure copper. The waste tailings were dumped into nearby Torch Lake. However, using jigging up to 25 % of the native copper was incorrectly classified as waste (tailings) and dumped the copper was lost as the tailings were dumped into nearby Torch Lake. Over time new technologies were introduced into the mills by Quincy to increase efficiency and loose less copper into the tailings. Quincy used froth flotation in its processing circuit as early as 1920s. Violent agitation of copper and rock/mineral particles in an oily water along with frothing agents and other chemicals cause bubbles to rise to the surface with copper particles attached to their surface. The bubbles and copper particles were captured from the top of the flotation tank while the waster rock/mineral (tailings) sank to the bottom. and was slurried to Torch Lake. The heavier copper particles were carried upward by the floatation bubbles rather than sinking due to gravity by jigging. The mineral processing circuit could begin with jigging and with the tailings after jigging sent to floatation cells to recover more copper. Later Wilfley tables were added to the mineral processing circuit to help minimize loss of copper from finer grain sizes. A Wilfley table utilizes gravity and water to separate denser particles. A shaking ribbed table with a film of water flowing along the long axis results in the higher density copper particles concentrating in beds behind the riffles. 74 �The Quincy Mining Company knew that the tailings contained a lot of copper that they were unable to recover with technology available at that time. The company had considerable foresight and kept detailed maps each year of the copper content of the tailings that were deposited into the lake. As technology improved, they were able to reclaim copper from the tailings at a profit. Quincy built a special reclamation processing mill near here in 1942-43 for $1.2 million. The main building, 124'x255', had six Harding ball mills to grind the tailings even finer than the stamp mills in order to release more fine particles of copper from the rock/mineral and facilitate addition of Wilfley tables and flotation cells to mineral processing circuit. Across the road is Quincy Mining Company dredge No. 2 on the shore of Torch Lake. Torch Lake was filled with several 100 millions of tons of tailings since not only did Quincy Mine operate a mill on its margin so did many other companies, especially Calumet and Hecla Mining Company. The tailings were sucked up via the dredge and sent to the reclamation mill to recover more copper from the tailings. The reprocessed tailings were redeposited back into Torch Lake. This dredge was built in 1913 by Calumet & Hecla Mining Company. In 1951, the Quincy Mining Company purchased the dredge and it became known as Quincy Dredge No. 2. It could process over 10,000 tons of tailings per day and it had 141 ft suction pipe that could work 115 ft below the surface of the lake (Figure 10). Tailings were conveyed to the mill via a tube held up with pontoons. Quincy recovered 100 million lbs. of copper from tailings in Torch Lake from 1943 until it closed in 1967 or about 10 % of total production. In the 1800s and early 1900s, depositing 100s of millions of tons of tailings into Torch Lake was acceptable practice and the environmental consequences were not considered. Today, the environmental impact of tailings must be carefully considered to obtain a permit, "social license" to operate a mine. The mining companies did not just put tailings into Torch Lake, they also used it to dispose of other waste such as that from electrical systems and barrels filled with chemicals. The early processing plants used only water to separate the copper from the rock. Later floatation cells were used in processing ore from the mine and in processing reclaimed copper from the tailings. The floatation cells used chemicals in water. These chemicals along with the tailings were put into Torch Lake. Discovery of fish with tumors in Torch Lake led to its being designated an U.S. Environmental Protection Agency Superfund site. Some of these chemicals were biodegradable and are no longer present in the Torch Lake water but would have been present decades ago, thereby could have readily caused tumors in older fish. There are far too many tons of tailings in Torch Lake to remove them and thus, the mitigation strategy is to simply cover them with soil. Fortunately, the copper ores of the Keweenaw Peninsula lack pyrite that is known under certain environmental settings to produce acid drainage and in turn can result in significant environmental impact. The lack of acid generating potential and the otherwise mostly inert minerals in the waste rock has greatly lessened the environmental impact of the tailings themselves. Today crushed poor rock from the mines is used directly as aggregate and is also crushed for use as aggregate. Text from previously published field trip guide by Bornhorst (2022) with limited modification. 75 �Figure 10: Quincy Mining Company dredge No. 2 working on Torch Lake, ca. 1920s. From Molloy (2011). Stop 5: Ahmeek Mining Company Stamp Mill Latitude: 47.168633N; Longitude: -88.435810W Directions: From Quincy Mill Ruins and Dredge carefully pull out of gravel lot and continue east on M-26 for 1.9 miles (3 km) to stamp mills of the Ahmeek Mining Company. This is the only remaining steam-powered stamp in the Keweenaw Peninsula. The Ahmeek Mining Company had four of these Nordberg compound steam-powered stamps installed in 1910, and four additional stamps were added in 1914. Rail cars brought rock to the mill using the trestle, located above the trees across the street from the stamp. The compound-expansion nature of the stamps represented a major improvement in processing of copper ore. The stamp could strike approximately 104 24-inch blows per minute. The mill could process approximately 7,000 tons of ore in a 24-hour period. This is not part of the Quincy Mine but we stop here to see a steam-powered stamp used for crushing the native copper ore. Text from previously published field trip guide by Bornhorst (2022) with limited modification. 76 �References Cited Bornhorst, T.J., this volume, 2024, Mesoproterozoic Midcontinent Rift-filling Strata and Native Copper Deposits of the Keweenaw Peninsula, Michigan Field Trip 1: Institute on Lake Superior Geology, Field Trip Guidebook, 70th Annual Meeting, Houghton, MI v. 70, part 2, p. this volume. 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., 2022, Field guide to the geology and native copper mining history of the Keweenaw Peninsula: Field Trip Guidebook for American Institute of Professional Geologists 2022 National Conference held in Marquette, Michigan August 6-9, 68p. Bornhorst, T.J., Kalliokoski, J.O., and Paces, J.B., 1986, The Keweenaw native-copper district: in Brown, A.C. and Kirkham, R.V. (eds.), Proterozoic Sediment-hosted Stratiform Copper Deposits of Upper Michigan and Belt Supergroup of Idaho and Montana: Geological Survey of Canada, Contribution Series 13386, p. 21-36. 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., and McDowell, S.D., 1992, Michigan Tech Earth Science Laboratory and Experimental Mine connecting with the Quincy native copper mine, Michigan: Society of Economic Geologists Field Conference Guidebook Series, v. 13, p. 100-104. Bornhorst, T.J., and Molloy, L.J., 2016, Geological and Historical Field Trip to the Keweenaw Peninsula: A Tribute to Douglass Houghton "Michigan's Pioneer Geologist". Michigan Basin Geological Society, 65 p. Bornhorst, T. J., and Molloy L.J., 2016, Geological and historical field trip to the Keweenaw Peninsula, A tribute to Douglass Houghton: “Michigan’s Pioneer Geologist: Michigan Basin Geological Society Geological and Historical Excursion September 10th-12th, 89 p. Bumgarner, E.L., 1980, The Geology of the Portage Lake Volcanics in the M.T.U. Mining Laboratory, Hancock, Michigan: Michigan Technological University, M.S. thesis, 138 p. Butler, B. S., and Burbank, W. S., 1929, The copper deposits of Michigan: U.S. Geological Survey Professional Paper 144, 238 p. Caine, J.S., Evans, J.P., and Forster, C.B., 1996, Fault zone architecture and permeability structure: Geology, v. 24, no. 11, p. 1025-1028. Caine, J.S. and Forster, C.B., 1999, Fault zone architecture and fluid flow: Insights from field data and numerical modeling: in Faults and Subsurface Fluid Flow in the Shallow Crust, American Geophysical Union, Geophysical Monograph 113, p. 101-128. 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. 77 �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. 135, no. 1/2, p. 449–466. https://doi.org/10.1130/B36186.1 Fink, J.H. and Fletcher, R.C., 1978, Ropy pahoehoe: surface folding of a viscous fluid: Journal of Volcanology and Geothermal Research, v. 4, p. 151–170. 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. Langfield, K.M., DeGraff, J.M., and Gamet, N.G., 2023, Slip kinematics of the Keweenaw and Hancock faults within the Midcontinent Rift System, Upper Peninsula of Michigan: Institute on Lake Superior Geology, 69th Annual Meeting, Eau Claire, Wisconsin, Part 1 – Program and Abstracts, v. 69, p. 5051. Lankton, L.D., 1991, Cradle to Grave: Life, Work, and Death at the Lake Superior Copper Mines, Oxford University Press, 319 p. Lankton, L.D. and Hyde, C. K., 1982, Old Reliable: An Illustrated History of the Quincy Mining Company, Quincy Mine Hoist Association, 159 p. Molloy, L. J., 2011, A guide to Michigan's historic Keweenaw copper district: published by Great Lakes Geoscience LLC, 122 p. Molloy, L. J., 2007, A Visitor's Guide to the Historic Quincy Mine: published by Great Lakes Geoscience LLC, 61 p. Self, S., Keszthelyi, L., and Thordarson, T., 1998, The importance of pahoehoe: Annual Review of Earth and Planetary Sciences, v. 26, p. 81-110. 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. Thurner, A. W., 1994, Strangers and Sojourners: A History of Michigan's Keweenaw Peninsula, Wayne State University Press, 404 p. 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. 78 �Field Trip 3 Geoheritage of Buffalo Reef: Industrial Impact on Land, Culture, and Fish Sovereignty Erika Vye Great Lakes Research Center, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Charlie Kerfoot Great Lakes Research Center, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Stephanie Swart Michigan Department of Environment, Great Lakes, and Energy, 525 West Allegan Street, Lansing, MI, 48909 Dione Price Keweenaw Bay Indian Community, Natural Resources Department, 14359 Pequaming Road, L’Anse, MI 49946 Evelyn Ravindran Keweenaw Bay Indian Community, Natural Resources Department, 14359 Pequaming Road, L’Anse, MI 49946 INTRODUCTION Buffalo Reef is a geoheritage site with scientific, educational, cultural, and aesthetic value. This field trip explores the relationship between geology, mining waste, and culture of Buffalo Reef a 2,200-acre natural cobble feature of Lake Superior’s lakebed southeast of the Keweenaw Peninsula and about 20 miles northeast of Houghton. Finely crushed waste rock - stamp sand from copper ore milling operations at the community of Gay have been moved by currents along the shoreline of the Keweenaw Peninsula to Big Traverse Bay, thus covering this highly productive spawning ground for lake trout and whitefish. This has negative implications for commercial fisheries, local economies, subsistence uses, and the spiritual, physical, and cultural well-being of tribal nations that identify as fishing people. Further, this has adverse impacts on tribes’ ability to exercise their treaty rights to fish in this area (Buffalo Reef Task Force, 2024). Geoheritage is a nascent yet 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). More 79 �distinctively, geoheritage emphasizes the importance of the varied personal values people have for geologic features and explores the wide-ranging relationships people have with landscape. The rich geodiversity of the Keweenaw has fostered relationships for millennia and has imbued our place with significant sites that provide opportunities to broaden both Earth science and cultural literacy (Rose & Vye with Martin, 2017; Vye, 2016). The rich geosites of the Keweenaw provide an accessible platform for people to learn about deep time, Earth's dynamic processes, and importantly, the diverse relationships and reciprocity people have with our geologic underpinnings. The Keweenaw is renowned for Earth's largest native copper deposits that became the first great copper mining district in the United States. From 1845 to 1968, over 11 billion pounds of refined copper were produced in Keweenaw mines, making it a cornerstone of the American economy in the 19th and 20th centuries (Bornhorst & Lankton, 2009; Bornhorst & Barron, 2011). Interpretations and public educational programming depicting the relationship between people and geology have largely focused on stories and heritage associated with the European diaspora that fueled the Copper Boom of 1845-1968 and how copper’s prolific use for transatlantic cables, telegraphs, electricity, and the auto industry ultimately modernized the country (Bornhorst & Lankton, 2009). Less interpreted, yet central to our history, is how geology has shaped, and continues to shape, the ancestral, traditional, and contemporary lands, waters, and livelihoods of the Anishinaabeg- the Three Fires Confederacy of Ojibwe, Odawa, and Potawatomi peoples and their many more-than-human relatives. As such, this field trip visits three stops (Figure 1) to explore mining impacts on culture, subsistence uses, and the wellness of all beings. Figure 1: Map of field trip site locations, courtesy of D.J. Lizzadro-McPherson, MTU 80 �Stop 1 - Mohawk, MI: The trip begins at the poor rock pile where the Mohawk No. 4 mine shaft once stood - a native copper host rock from which some of the Gay stamp sands were generated. Stop 2 - Gay, MI: Participants will then travel to the town of Gay to walk the stamp sands and learn how Tribal, State, Federal, and academic partnerships are collaborating to mitigate environmental damage and ultimately restore Buffalo Reef to the ecological resource that has sustained both tribal and non-tribal communities for generations. Stop 3 - Big Traverse Bay & Buffalo Reef, MI: From Gay, we will travel to Big Traverse Bay to learn more about why shoreline and habitat restoration efforts are necessary and how stakeholders and rights holders are working together to address this environmental justice issue in our community. Participants will benefit from visiting Buffalo Reef aboard Michigan Tech’s Research Vessel Agassiz. This part of the trip enables participants to compare healthy parts of the reef with areas that have been inundated with stamp sands. STOP 1 - MOHAWK, MI At this site, we will learn more about the source rock for the Gay stamp sands after a brief overview of the geological history of the Keweenaw region. Participants will have time to explore the poor rock pile before departing for Gay. Directions: Leave downtown Houghton and head north on U.S. 41 towards Calumet. Stay on U.S. 41 until you reach Mohawk; turn right on 6th street, just past E Phoenix Street there is a place to park on the left-hand side of the road. This space provides access to the Mohawk No. 4 poor rock pile (Figure 2). Coordinates: 47.301203, 88.363369 Figure 2: Parking location to visit Mohawk No.4 Geological Origin Story Michigan’s Keweenaw Peninsula has a rich and globally significant geodiversity in tandem with a fascinating cultural story. This is the site of the largest native copper dominated deposits known on Earth, the oldest metal workings in the Western Hemisphere, and a recent diaspora of 81 �European cultures that flocked to the region for copper mining in the late 1800s – this has shaped an entangled mosaic of cultural, mining, and industrial heritage. Importantly, the Keweenaw Peninsula offers an important window to Earth’s past, exposing the heart of the Mesoproterozoic Mid-Continent Rift. Located on Lake Superior in Michigan’s Upper Peninsula (Figure 3); the abundant geodiversity sites are the result of a flood basalt sequence comprised of hundreds of voluminous lava flows, interbedded and covered by a redbed clastic sedimentary rocks. An upwelling of heat and magma from a hot spot initiated great lava flows erupting from the rifting of supercontinent Rodinia. The rift created a ~3000 km U-shaped feature in the center of North America that extends from Kansas, through what is now Lake Superior and terminated in what is now Michigan (Cannon, 1994, Cannon and Nicholson, 2001, Stein et al., 2015). Figure 3: Extent of rifting associated with the Mid-Continent Rift (K. Schulz, USGS) Some of the largest lava flows on Earth erupted and ponded in the Mid-Continent rift over many centuries (Huber, 1983). During quiet times, red-brown conglomerates and sandstone were deposited between flows in high-energy alluvial fans (Elmore, 1984). The interbedded lava flows and sedimentary layers were normally faulted resulting in a syncline feature that extends from Isle Royale to the Keweenaw Peninsula, now the basin for Lake Superior (Figure 4). Copper was brought to the surface by hydrothermal systems mineralizing permeable layers such as the amygdaloids of lava flow tops and the cracks and crevices of conglomerate rocks (Bornhorst & Barron, 2011, Nicholson et al, 1997, Cannon, 1984). Figure 4: Cartoon depicting interbedded lava flows and minor sedimentary rocks (Huber, 1983. 82 �The Keweenaw is noted for mining nearly 9000 years ago (Martin, 1995, 1999). The valued red metal has since been traded up and down the Mississippi by the Ojibwa caretakers of this landscape and used in ceremony today. From 1845 to 1968, ~11 billion pounds of refined copper were produced from Keweenaw Peninsula mines making it one of the cornerstones of the American economy and the first great metal mining district in the United States (Bornhorst and Barron, 2011, Bornhorst and Lankton, 2009). The region has been extensively mapped and researched since the mid-1800s because of the discovery of copper and the subsequent mining boom. This field trip explores the impacts related to one of the many mining sites in operation from the late 1800s to mid-1900s in the Keweenaw region - Mohawk Mine, a native copper host rock from which some of the Gay stamp sands were derived. Mohawk No. 4 This field stop explores the site of the Mohawk mine, specifically the poor rock pile associated with the Mohawk No. 4 shaft (Figure 5). Rock from this site is the main source for the Gay stamp sands. Following the discovery of copper on the property in 1896 by lumberman Ernest Koch, the Mohawk Mining Company was founded. It was incorporated in 1898 and lasted until 1932 (John Stanton as president, later replaced by Joseph Gay). In 1923 the Mohawk Mining Company took over the neighboring Wolverine Figure 5: A group of Keweenaw youth explore the poor rock Copper Mining Company and the pile at Mohawk No. 4 (photo credit Erika Vye) Michigan Copper Mining Company. In 1934 the company was purchased by the Copper Range Company (Molloy, 2008). The Mohawk Mine had six shafts numbered 1 through 6 running from north to south along what is now US 41. The was a profitable mine paying out over $15 million in shareholder dividends between 1906 and 1932 (Clark, 1978). The No. 4 shaft of the Mohawk Mine was constructed at this site in 1901 and stayed open until the mine closed in 1932 (Figure 6). In the early days of operation, the shaft reached a depth of 501 feet. Technological advances in 1904 led to equipping the No. 1, 2, and 4 shafts with Nordberg Conical Drum Hoists (Figure 7) enabling depths of up to 6000 feet. With this technology, the shaft reached a depth of 900 feet by 1906, 1,175 feet by 1908, and by 1922 a final depth of 2,832 feet. In 1914, the No. 4 shaft was producing between 450 and 500 tons of ore per day. Mining continued from the No. 4 shaft until 1924, with a brief lull in operations until it was reopened in 1926, eventually closing in 1932 (Clark, 1978). 83 �Figure 6: No 4 shaft in Mohawk, MI (photo courtesy of the MTU Archives) Figure 7: Image of Nordberg hoist; the world's largest Nordberg hoist can be visited at the Quincy Mine in Hancock, MI, pictured here (photo courtesy of the MTU Archives) 84 �Ore Deposit and Significant Minerals The Kearsarge lode at Mohawk is a native copper-dominated deposit with minor native silver hosted by basalt lava flows of the Portage Lake Volcanics (about 1.1 billion years old). The strike of the Kearsarge orebody is N45E, dips at 35W, and is 2.9 meters thick, 1,490 meters wide, with a total length of 3,250 meters spanning an area of 250 hectares. The productive part of the fragmental amygdaloidal lode was richer at greater depths than nearer to the surface where the lode consisted of more massive basalt. The Kearsarge lode was the second largest in the district and averaged 6 to 13 feet thick (Butler & Burbank, 1929; White et al, 1953; and USGS, 2005). The fragmental amygdaloid ore of the Mohawk mines is associated with arsenic-rich minerals, more common than in most of the other Keweenaw deposits. The Mohawk mine is well known among mineral collectors for its large occurrence of mohawkite - a mixture of algodonite, domeykite, and copper (Moore, 1971). Mohawkite was first found on the property in 1901, on the first level north of the No. 1 shaft. In addition to copper, the mine also produced a small amount of native silver (Figure 8). Figure 8: Left -Polished mohawkite, a rare mixture of copper and copper arsenides, is named after the Mohawk-Ahmeek area of the Keweenaw (Photo by Robert M. Lavinsky); Right – native copper in amygdaloidal basalt. Geologic features of note include the occurrence of the Mohawkite Fissure that crossed the Kearsarge lode. From 1900-1901, 105 metric tons of mohawkite ore was produced from the Mohawk Mine at the No. 1 shaft. From 1902 through 1925 the No. 1 shaft produced about 117,000 metric tons of refined copper. Additional veins were discovered south of the No. 2 shaft in 1901 resulting in 230,000 pounds of mohawkite (Butler and Burbank, 1929). Mohawkite proved to be challenging in the Keweenaw as smelters were not able to process it on account of the high arsenic levels and the deadly fumes produced. A unique smelter was built in Hackensack Meadows, New Jersey for the specific purpose of processing the ore; it became 85 �operational in 1901. The mohawkite ore contained mostly copper and arsenic, it also contained small amounts of nickel and cobalt, as well as about 20 ounces of silver per ton of ore. (Stevens, 1902) Copper-bearing rocks were transported 11 miles by rail to the town of Gay for milling near the mouth of the Tobacco River on Traverse Bay. This is the next stop on the field trip. STOP 2 – GAY, MI This part of the field trip begins at the Gay Sands sign just outside of the town of Gay. This stop is intended to engage participants in a discussion of how the stamp sands were generated and to learn how Tribal, State, Federal, and academic partnerships are collaborating to mitigate environmental damage and ultimately restore Buffalo Reef to the ecological resource that has sustained both tribal and non-tribal communities for generations. Participants will have an opportunity to walk the stamp sands (accessed just south of the Gay Sands sign): note that at times large trucks are moving through this area; always exercise caution and stay together. Directions: From Mohawk, travel to the town of Gay; follow the Mohawk-Gay Road. When you arrive in Gay, turn right on Main Street, then left on 2nd Street (which becomes the Gay Lac la Belle Rd), park on the right-hand side of the road beside the smokestack and local signage that shares information about the Gay stamp sands (Figure 9). Coordinates: 47.226766, -88.161933 Figure 9: Left - Gay Sands signage; Right – Gay mill smokestack (photo credit Erika Vye) How were the stamp sands created? Liberating the copper from the host rock required a large water source - in this instance, Lake Superior. In 1900, the Wolverine mine and mill opened in Gay. This piqued the interests of New York financiers of the Central, Atlantic, and Baltic mining companies and shortly thereafter the Mohawk mines and mills were opened. The two mills were built in Gay by the Mohawk and Wolverine mining companies enabling the processing of copper ore. A railroad was built to 86 �transport the copper-bearing rocks from Mohawk to Gay; poor rock was left at the mine sites in Mohawk remaining in piles today. A dock was built nearby to import coal to power operations and to load copper on ships to transport to other places (Keweenaw County Historical Society). Note: A QR is included in the resources section of this guide leading to the recently developed Gay Mills Exhibit created by the Keweenaw County Historical Society. An unusual landscape & vastly different soundscape When visiting this site take note of what you hear; try to imagine what the soundscape of this place might have been like in the early 1900s. At that time, the stamps at both mills were running constantly. Imagine the ground shaking with over 70 stamps per minute crushing pieces of rock brought up from underground to free the copper from the matrix. The next step required washing the material to separate the copper and produce a copper-rich mineral concentrate to then be sent to the smelter and formed into copper ingots (Lankton, 2005). The material was brought to the top of the mill with a railway line. To process the copper and create the concentrate, water was pumped from Lake Superior, and the tailings and waste slurry were dumped back into the lake with conveyors. The length continued to grow as the tailings built up along the shoreline (Keweenaw County Historical Society). Figure 10 shows an image of this process on the left, with what remains of this structure today. Figure 11 illustrates the milling process. Figure 10: Left - Conveyor sending tailings slurry back to the lake (Photo credit: MTU Archives) during milling operations in Gay; Right - wood beams left from conveyor structure as sands are continually swept away by wind and currents over time (Photo by Erika Vye) What is happening to the sands? From 1900-1930 the mills flushed over 22.7 million metric tons of tailings to Lake Superior leaving a large bank of black sand, the consistency of kitty litter, along the shore. This finely crushed waste rock has moved, and continues to move, along the southeast shoreline of the Keweenaw Peninsula to Grand Traverse Bay Harbor threatening the nearby Buffalo Reef. Since the 1930s Lake Superior’s fierce storms and strong currents have eroded the stamp sand bank, pushing the tailings into the lake and gradually moving them southward along the shoreline to 87 �the Traverse River. These stamp sands contain high amounts of copper and cover 1,426 acres of shoreline and lakebed. Currently, 30% of the reef has been impacted by stamp sands; modeling predicts that by 2025 stamp sands will impact 60% of the reef. The next stop focuses on the importance of understanding the impacts of these predictions and what is happening for restoration. Before traveling to the last stop we will have lunch at the Gay Park. This site also hosts the Gay Historic School and Museum open from 1 PM to 4 PM on Wednesday and Saturday. 88 �Figure 11: This image describes how copper was processed in the mills - first by gravity and large amounts of water through steam stamps, the rock was funneled into a trommel that sorted and classified the pieces by size. A jig then separated the copper from the mine rock. Next, a series of vibrating Wilfley tables separated the tailings from the small-sized copper. This image was recently modified for a Keweenaw County Historical Society exhibit on the Gay Mills (a link for the exhibit is found in the Resources section of this guide). Original delineation by the Historic American Engineering Record, Heritage Conservation & Recreation Service, Eric M. Hansen, 1978. Modification by David A. Vago, MTU, for the Houghton County Historical Society, 2005; updated 2021. STOP 3 - BIG TRAVERSE BAY & BUFFALO REEF. MI At this site, we will explore how the Gay sands are impacting life and culture in the region, learn about remediation efforts, and how we can share in educating ourselves and others about this environmental justice issue. Directions: From the Gay Park by the Historic School and Museum travel left on Lake Street, then right on 1st Street (becomes the Lake Linden-Gay Rd). Turn left onto Rice Lake Rd, which becomes the Big Traverse Rd. Turn left on the Traverse River Rd, please drive slowly along this road as there are families and small children frequently playing here. This takes approximately 15 min from Gay to Big Traverse. Note that the dock is on the other side of the harbor; to get there simply head back to Big Traverse Rd, then turn left. The dock is about ¼ mile ahead. Coordinates: Big Traverse Bay 47.189857, -88.236644, Dock to board Agassiz 47.190769, 88.237154 What is Buffalo Reef? Buffalo Reef is a natural cobble 1 feature in Lake Superior, located just off the eastern edge of the Keweenaw Peninsula in the U.P. of Michigan (Figure 12). The reef has historically maintained an invaluable spawning habitat for fish species such as lake trout and lake whitefish. Buffalo Reef is estimated to supply 33% of the tribal harvest of lake trout and lake whitefish from Michigan waters of Lake Superior (i.e. 136,375 pounds of whitefish and 61,830 pounds of lake trout average annual yield 2001-2016). Figure 11: Location of Buffalo Reef, also not the locations of Additionally, if the stamp sands both Field Trip Stop 2 and 3, marked by stars. migrate south of Grand Traverse Harbor, they will threaten the undisturbed native sand that serves as habitat for juvenile whitefish (Kerfoot et al, 2012 & 2021). Juvenile whitefish produced on Buffalo Reef migrate south of 89 �Grand Traverse Harbor to sandy, shallow-water habitat where they feed before migrating to the deeper water they inhabit as adults. Without this habitat, whitefish recruitment in the vicinity of Buffalo Reef would be greatly diminished. Whitefish provide the bulk of the commercial, cultural, and spiritual value for the tribal communities that use this resource (BRTF, 2024). How are the stamp sands impacting Buffalo Reef? Most of the cobbles are glacial rocks, scattered around the Jacobsville sandstone bedrock highs of the reef. During spawning, fish drop eggs into the crevices between rocks. Cobbles are coated with a natural organic film, the basis of a food web for feeding fish. Stamp sands move into the northern cobble field of Buffalo Reef burying cobbles and killing living communities on rocks along the leading edge of the sands. (Figure 13). These impacts to fish habitat create a risk for a decline in commercially and culturally important fish keynote species. Importantly, a decline in species impacts the ability to exercise treaty rights. The sands also have high enough copper concentration (about two-eighths of a percent) to be toxic to fish and other organisms that live in Grand Traverse Bay. There are additional impacts to human health through fish consumption. Other concerns include the impacts of the copper-toxic sands on coastal wetlands medicine chests to the KBIC - as the material migrates up streams thereby affecting plant and amphibian habitat (Kerfoot et al, 2012, 2021, in prep). The sands also impact the aesthetics and sense of place inspired by the landscape; the aerial photos in Figure 14 illustrate the movement of the stamp sands inundating white sand Figure 12: Healthy reef habitat progressively beaches along the south shore of the inundated with stamp sands until covered completely Keweenaw Peninsula. The stamp sands also (Photos by Charlie Kerfoot) create challenges for people coming and going from the Big Traverse Bay Harbor; the sands have repeatedly blocked the harbor stopping boat traffic. For example, in 2015, Big Traverse Bay Harbor was dredged removing 4,500 cubic yards of stamp sand. Dredging happens repeatedly here, without this continued maintenance there is a great risk of harming the region’s fishing industry. As such, local stakeholders are working with the Army Corps of Engineers on a long-term solution. 90 �Figure 13 Figure 14: Aerial views of the entry of Traverse River into Lake Superior. The regular movement of stamp sands over the top of the breakwater requires annual or biennial dredging to keep the small Traverse River Harbor, home of a commercial fishing fleet, open to entry. Note the black sands on the left to the north of the channel compared to the white (natural) sands to the right. Left photo credit: Neil Harri; Right photo credit: Charlie Kerfoot Impacts on Fish Sovereignty Buffalo Reef lies within the ceded-territory homelands established by the Treaty of 1842 where 11 Lake Superior Bands of Ojibwa retain rights and responsibilities to harvest fish from the Michigan waters of Lake Superior. Buffalo Reef has always been considered a culturally significant harvesting ground for local communities. Fishing is the strand of the cultural core that ties history to the present day and the future; it is a vital part of the foundation for cultural beliefs and values, traditional lifeways, and even individual identity (Gagnon, 2018; KBIC, 2017). The KBIC identify as fishing people and have always had a strong focus on cultivating and protecting relationships within ecosystems that support healthy food sovereignty initiatives within the community (Figure 15). The Anishinaabeg teachings and ways of life emphasize that landscapes and waters are an intricate system of diverse relationships, and interconnected rights and responsibilities rooted in an acknowledgment and understanding that humans and nature are relatives. These sentient elements have been honored in ceremony since time immemorial and these important teachings Figure 14: Fresh catch on Keweenaw Bay 91 �and traditions continue today. These relationships are founded in the long-standing nation-tonation agreements between the Anishinaabe Ojibwa and all orders of creation from rock, water, fire, and wind; the physical world of sun, stars, moon and earth; plant beings; animal beings; and human beings - all rooted in the First Treaty with Gichi Manidoo (the Creator), also known as Sacred Law, Original Instructions, and Natural Law (Johnston, 1976; Keweenaw Bay Indian Community, 2021). Today, Tribal, State, Federal, and Academic partnerships are combining efforts to mitigate damages and ultimately restore Buffalo Reef as the ecological resource that has sustained both tribal and non-tribal communities for generations. Buffalo Reef Remediation Efforts This is a complex environmental justice issue requiring local, regional, and state stakeholders and rights holders to work together on research and restoration efforts for Buffalo Reef. It was tribal fishermen who first alerted our communities of this issue, underscoring the vital need to bridge knowledges for a holistic understanding of what is happening within our landscape and to recognize the gifts that different knowledges share. Research over the past years has included: a) modeling wave action, currents, and how the sands are migrating across the coastal shelf and along the shoreline; b) use of LiDAR and sonar assessments to understand the extent of the sands underneath water; c) submerging remotely operated underwater vehicles to take pictures of the boulder and cobble fields to assess the quality of the remaining fish habitat; d) researching environmental impacts on the benthic and fish communities; e) exploring considerations for building a revetment to hold back stamp sands from further entering the bay; f) conducting archaeology to understand industrial heritage, mining legacies and Native American uses of the area prior to colonization; and f) and meeting with community members to understand social and economic impacts on the communities involved. After years of observation, scientific study, and dredging at Big Traverse, the EPA funded a feasibility study for a long-term solution. In 2017 the USEPA endorsed the formation of a Buffalo Reef Task Force (BRTF) comprised of multiple state, federal, and tribal agencies. In addition, several academic institutions and private entities have joined the team, recognizing that this issue is larger than any single entity can accomplish on its own. The Buffalo Reef Final Alternatives Analysis Report (Report) was recently on public notice from January 30 to March 1, 2024. This report provides an overview of efforts made by the BRTF “in 2019, the BRTF identified 13 potential alternatives to remediate the stamp sands and restore the habitat. The alternatives were screened based on constructability, operation and maintenance requirements, environmental impact, ecological sustainability, initial costs and legacy costs, regulatory requirements, public input and opinion, time needed for implementation, impact to local populations, and potential for beneficial use of the stamp sands”. Three alternatives were selected for further consideration that included the dredging and disposal of stamp sands in the following locations: 1) White Pine Mine, a closed tailings basin; 2) a Lakeside Placement site; and 3) an Upland Placement site. In 2022, a public meeting was held to discuss the one alternative that would be feasible - the removal of the stamp sands to a regulated and lined landfill to be constructed near Gay. These remediation efforts require land ownership, 92 �both at the shore for a proposed jetty to impede migration of the stamp sands and an upland placement area to move the mining waste. The Report indicates that the BRTF’s preferred remedial alternative and “Potential Plan” identifies the “Upland Placement Alternative” as the BRTF’s preferred remedial alternative (Figure 16). The study compares three scales of implementation for the Upland Alternative are compared in the report. “The scale of implementation differs based on the volume of stamp sands to be dredged and disposed of during the implementation phase. Additionally, the study provides an assessment of the impact the volume of stamp sands removed during implementation has on the cost and duration of the operation, maintenance, repair, replacement, and rehabilitation phase.” Figure 16: Upland Placement Alternative: Small, Medium, All Dredging Scales represented (Left to Right). S Shoreline – South Shoreline; M Shoreline – Middle Shoreline; and N Shoreline – North Shoreline. The Upland Placement footprint required to dispose of the volume of stamp sands removed during project implementation varies per dredging scale. In the placement site measure, the solid white line represents the estimated placement site footprint to contain the stamp sands dredged during the project implementation phase. The dashed white line represents the estimated placement site footprint to contain stamp sands dredged during the OMRR&R project phase (sourced from page ES-5 of the Buffalo Reef Final Alternatives Analysis Report). RETURN TO HOUGHTON Billion-year-old geologic processes brought copper to the surface where people could find it - a gift from the deep Earth. Relationships with the red metal have varied over millennia resulting in different values and actions within our shared lands and waters. Further efforts are required by all stakeholders and rights holders to communicate and elevate the importance of this environmental issue with decision-makers, government officials, and environmental advocacy groups. This field trip is meant to deepen our understanding of these relationships and to reflect on the following questions: What are the relationships between people, Earth, and Buffalo Reef? Why does this matter? How have people impacted Buffalo Reef in the past, and how will they in the future? What can you do to help with the remediation efforts of this place? 93 �Directions: From Big Traverse Bay Harbor, head north on S Big Traverse Rd and continue along Rice Lake Rd. Take M-26 S to US-41 S in Houghton. EDUCATIONAL RESOURCES Access the following websites and videos by following the QR codes. Saving Buffalo Reef Website (DNR) Saving Buffalo Reef Video (GLIFWC) Buffalo Reef Restoration (Great Lakes Now, Ep. 1006 Segment 3) Gay Milling at Gay A Lake Superior Story - REFERENCES CITED Bornhorst, T. J. and Barron, R. J. (2011). Copper deposits of the western Upper Peninsula of Michigan. Geologic Society of America, Field Guide, 24, 83-99. Bornhorst, T. J. and Lankton, T. J. (2009). Copper Mining: A Billion Years of Geologic and Human History, in Schaetzl, R., Darden, J. and Brandt, D. (eds.) Michigan Geography and Geology. United States of America: Pearson Custom Publishing, 150-173. 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. Buffalo Reef Task Force (2024). DRAFT Buffalo Reef – Final Alternatives Analysis. Retrieved from: https://www.michigan.gov/dnr/-/media/Project/Websites/dnr/Documents/Fisheries/BuffaloReef/00DRAFT-Buffalo-Reef-Main-ReportJAN2024.pdf?rev=1ebf68cee9ae428881d8d6991b4d7471&hash=48BD2B6EA77CD5C430FD63757863 2B28 94 �Butler, B.S. & Burbank, W.S. (1929). The Copper Deposits of Michigan. USGS Professional Paper No. 144 Cannon, W. F. (1994). Closing of the Midcontinent Rift: a far-field effect of Grenvillian compression. Geology 22, pp. 155-158. Cannon, W. F. and Nicholson, S. W. (2001). Geologic Map of the Keweenaw and Adjacent Area Michigan 1:100,000. USGS Map I-2696 Clarke, D. (1978). Copper mines of Keweenaw; no. 12: Mohawk Mining Company. ASIN B0066RONSQ. 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, pp. 610-617. Gagnon (2018). A Tribute to Our Fisherman -Minaadowenjigaaziwaat Gidoo giigoonkeninii-minaanik. Retrieved from: https://nrd.kbicnsn.gov/sites/default/files/A%20Tribute%20to%20our%20Fishermen%20-%20handout.pdf Geological Society of America (2017). GSA Position Statement: Geoheritage. Retrieved from: https://www.geosociety.org/documents/gsa/positions/pos20_Geoheritage.pdf. Huber, N. K. (1983). The geologic story of Isle Royale National Park. United States Geologic Survey Bulletin (1309), pp. 66. Johnston, B. (1976). Ojibway Heritage. Toronto: McClelland and Stewart. Kerfoot C., Yousef F., Green A., Regis R., Shuchman R., Brooks N., Sayers M., Sabol B., and Graves M. (2012). LiDAR (Light Detection and Ranging) and multispectral studies of disturbed Lake Superior coastal environments. Limnol. Oceanogr. 57: 749–771. https://doi.org/10.4319/lo.2012.57.3.0749 Kerfoot C., Hobmeier M., Swain G., Regis R., Raman V., Brooks C., Grimm A., Cook C., Shuchman R., and Reif M. (2021). Coastal Remote Sensing: Merging Physical, Chemical, and Biological Data as Tailings Drift onto Buffalo Reef, Lake Superior. Remote Sensing. 2021, 13 (13), 2434. https://doi.org/10.3390/rs13132434 Kerfoot. C., Swain, G., Regis, R., Raman, V.K., Brooks, C., Cook, C and Reif, M. (in prep). Coastal Copper Tailings Dispersal: 3D Mapping and Shoreline Impacts, Particle Migration, Leaching, And Toxicity. Remote Sensing, 2024, 16. Keweenaw Bay Indian Community (2017). Testimony of Warren C. Swartz Jr. Before the Senate Committee on Commerce, Science & Transportation. Retrieved from: https://www.michigan.gov/dnr//media/Project/Websites/dnr/Documents/Fisheries/BuffaloReef/TribalTestimonyADA1.pdf?rev=2c9d195 864e44557bc3bc315ed93e4a6 Keweenaw Bay Indian Community (2021). Who We Are. Keweenaw Bay Indian Community Natural Resources Department. Retrieved from http://nrd.kbic-nsn.gov/about-us. Keweenaw County Historical Society (2022). Copper Milling at Gay - A Lake Superior Story. Retrieved from: https://www.keweenawhistory.org/Copper-Milling-At-Gay 95 �Lankton, L. (2005). Keweenaw National Historical Park Historic Resource Study. Prepared for the National Park Service, United States Department of the Interior. Retrieved from: http://npshistory.com/publications/kewe/hrs.pdf Martin, S. R. (1995). Michigan prehistory facts: The state of our knowledge about ancient copper mining in Michigan. The Michigan Archaeologist, 41(2-3), pp. 119-138. Martin, S. R. (1999) Wonderful Power: The Story of Ancient Copper Working in the Lake Superior Basin Wayne State University Press, Detroit, MI, p. 284 Molloy, Lawrence J. (2008). A Guide to Michigan's Historic Keweenaw Copper District: Photographs, Maps, and Tours of the Keweenaw, Past and Present. Hubbell, Michigan: Great Lakes GeoScience. p. 66. ISBN 978-0-979-1772-1-7. Moore, P. (1971). Copper-Nickel Arsenides of the Mohawk No. 2 Mine, Mohawk, Keweenaw Co., Michigan. American Mineralogist, Volume 56, pages 1319-1331. National Park Service and American Geosciences Institute (2015). America’s Geologic Heritage: An Invitation to Leadership. NPS 999/129325. National Park Service, Denver, Colorado. 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. Can. J. Earth Sci., v. 34, pp. 504-520. Reynard, E. and Brilha, J. (2017). Geoheritage: Assessment, protection, and management. Elsevier, ISBN: 9780128095317. Rose, W.I. and Vye, E. with Martin, V. (2017). How the Rock Connects Us: A Geoheritage Guide to Michigan’s Keweenaw Peninsula and Isle Royale. Isle Royale and Keweenaw Parks Association, ISBN 9780935289213. Stein, C. A., Kley, J., Stein, S., Hindle, D. and Keller, G. R. (2015). North America’s Midcontinent Rift: When Rift met LIP. Geosphere, 11(5), pp. 1607-1616. Stevens, Horace J. (1902). The Copper Handbook: A Manual of the Copper Industry of the United States and Foreign Countries. Vol. II. Houghton, Michigan: Mines Publications. US Geological Survey (USGS) (2005). Mineral Resources Data System (MRDS). Vye, E. (2016). Geoheritage of the Keweenaw Peninsula (Doctoral dissertation). Michigan Technological University. White, W.S., Cornwall, H.R., & Swanson, R.W. (1953). Bedrock Geology of The Ahmeek Quadrangle. USGS Map GQ-27, Scale 1:24000 96 �Field Trip 4 Keweenaw Fault System Geometry and Kinematics: Clues to Its Nature and Origin James M. DeGraff, Katherine M. Langfield Department of Geological and Mining Engineering and Sciences Daniel J. Lizzadro-McPherson Geospatial Research Facility, Great Lakes Research Center Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 Objectives This field trip is designed to demonstrate four fundamental attributes of the Keweenaw fault system along the Keweenaw Peninsula: (1) geometry of the fault network and individual faults; (2) style of deformation of hanging-wall and footwall strata; (3) zonation of deformed rocks in fault zones; and (3) kinematic slip indicators on individual faults. Stops in this guide were chosen and grouped into four areas to illustrate the segmented geometry of the Keweenaw fault system and to point out fault blocks defined by intersections between faults in three main directions. Kinematic slip indicators observed at some stops manifest a preponderance of oblique slip with components of dextral strike slip and northwest-side-up reverse slip. Based on this information and fault network geometry, the overall nature of the fault system is interpreted to be transpressional and consistent with forcing by the Grenville orogeny. Introduction The Keweenaw fault extends southwest from the tip of the Keweenaw Peninsula in Michigan’s Upper Peninsula to near Ashland, Wisconsin (Fig. 1), a distance of about 250 kilometers. It is one of two faults along the south shore of western Lake Superior with reverse movement that juxtaposes Mesoproterozoic volcanic layers against Meso-Neoproterozoic clastic sedimentary strata, the other being the Douglas fault in Wisconsin and Minnesota. Irving and Chamberlin (1885) first established the existence of the Keweenaw fault by combining detailed observations in outcrops and trenches with astute reasoning to settle a long-running debate about the nature of the contact between volcanic layers to the northwest and sandstone strata to the southeast. Prior to their work, Wadsworth (1884) had argued that the sandstone was older and dipped northwest beneath the volcanic layers based on their similar dip at some locations along their contact. Afterward, this idea was resurrected intermittently until the mid-1970s, when deep drilling for oil and gas established the younger age of sandstone in the Keweenaw fault’s footwall once and for all. 97 �Figure 1: Mesoproterozoic Keweenawan Supergroup of the Midcontinent Rift System (inset map modified from Stein) in the Lake Superior region. For sources of geologic units and faults, see DeGraff and Carter (2023). KF‒Keweenaw fault; LOF‒Lake Owen fault, DF‒Douglas fault; IRF‒ Isle Royale fault, MIF‒Michipicoten Island fault, TF‒Thiel fault; MF‒Munising fault (provisionally named). Red “U” on upthrown sides of faults. Tectonic Evolution The Keweenaw fault and related major faults along the southern edge of the western Lake Superior basin cut rocks of the Midcontinent Rift System (Fig. 1) and have slipped several kilometers in a reverse sense (Cannon and Nicholson, 2000, 2001; DeGraff and Carter, 2023). Ideas to explain the origin and evolution of the Keweenaw fault differ in relation to the roles of rifting and orogenesis on fault initiation and movement over time. The Midcontinent Rift System (MRS), formed by extension of Laurentian crust in Mesoproterozoic time, was accompanied by voluminous mafic volcanism followed by a period of crustal sag with siliciclastic sediment filling the resulting basin (Cannon et al., 1989; Cannon, 1992; Hinze et al., 1990; Stein et al., 2015). Woodruff et al. (2020) defined stages of magmatism and tectonism in MRS evolution as follows: (1) Plateau Stage with widespread volcanism and distributed minor extension (c.1112 to c.1105 Ma); (2) Rift Stage when volcanism and extension reached maximum intensity along a central subsiding rift basin (c.1102 to c.1090 Ma); (3) Late-Rift Stage with declining volcanism transitioning to sedimentary infill of the rift basin (c.1090 to c.1083 Ma); and (4) Post-Rift Stage with sedimentary infill of a sag basin (c.1083 to c. 1060 Ma). Prior to recognition of the MRS, the Keweenaw fault was interpreted as a thrust that formed during an unspecified compressional event following eruption of lava flows and deposition of siliciclastic strata in the Lake Superior basin (Irving and Chamberlin, 1885; Butler and Burbank, 1929, White, 1968). After recognition of the rift, a new model for formation and evolution of the Keweenaw fault and others around the Lake Superior basin proposed that they formed as normal faults during stages 1 and 2, and then were inverted as reverse faults by post-rift Grenville compression (Cannon et al., 1989; Hinze et al., 1990; Cannon, 1994), i.e. the fifth 98 �Compressional Stage of Woodruff et al. (2020). A recent model, discussed below in relation to the inversion model, is that the Keweenaw fault formed by thrusting during post-rift Grenville compression and did not exist as a normal fault during MRS extension or sag (DeGraff and Carter, 2023). In some ways, the latter model reverts to earlier ideas about the Keweenaw fault. Ideas about the timing of reverse slip on the Keweenaw fault have evolved from the original ideas of Irving and Chamberlin (1885), who inferred two episodes of movement – one prior to deposition of the footwall sandstone and a second afterward. Early thrusting elevated the region northwest of the fault, creating a highland of mafic volcanic layers that shed debris into the tectonic basin southeast of the fault. A second episode of thrusting occurred after most of the sandstone deposition based on deformed footwall strata with steep to overturned attitudes toward the southeast. Later workers have suggested quasi-continuous fault slip that began before Jacobsville deposition and continued during most or all of its deposition (Cannon and Nicholson, 2000), and others have argued for a late reactivation of the fault system during one or more phases of the Appalachian orogeny (Craddock et al., 1997). Stratigraphy Extension during MRS evolution followed by compression during the Grenville Orogeny produced a broad syncline with strata dipping toward an axis beneath Lake Superior (Fig. 1). The related rocks in the vicinity of the Keweenaw Peninsula are well described by Cannon and Nicholson (2000, 2001), whose stratigraphic column and nomenclature are adopted here (Fig. 2). Here, the main geologic units of the Keweenawan Supergroup are from oldest to youngest: (1) Siemens Creek Volcanics and Kallander Creek Volcanics (Powder Mill Group) with basaltic to andesitic lava flows; (2) Portage Lake Volcanics (Bergland Group) mostly composed of basaltic to andesitic flows with lesser conglomerate and sandstone layers; (3) Oronto Group strata beginning with the Copper Harbor Conglomerate and locally interbedded basalt-andesite flows, continuing with the Nonesuch Formation composed of siltstone and shale, and ending with the Freda Sandstone composed of lithic sandstone and siltstone; and (4) Jacobsville Sandstone mostly composed of quartzose to subarkosic sandstone and generally correlated with the Bayfield Group in Wisconsin, USA. Archean granite and gneiss of the Superior craton, with a Paleoproterozoic cover of graywacke and slate (Michigamme Formation), lie beneath the Keweenawan rocks south of Keweenaw Bay. The following descriptions focus on the two geologic units that are juxtaposed along the Keweenaw fault system, namely the Portage Lake Volcanics in the hanging wall and the younger Jacobsville Sandstone in the footwall (Fig. 3). Both units are stratified and have compositional and textural variations between layers that likely influenced their mechanical behavior, which in turn influenced fault initiation and propagation. A more complete description of other Precambrian units along the Keweenaw Peninsula is provided elsewhere in this field trip volume (see Field Trip 1 in this volume). 99 �Figure 2: Left. Regional stratigraphy following Cannon and Nicholson (2001). For sources of radiometric ages in red italics, see DeGraff and Carter (2023). Right. Stratigraphy of the Portage Lake Volcanics from USGS bedrock geology maps of the Keweenaw Peninsula. Intense green units are major lava flows: pcc‒Copper City; pgf‒Gratiot; psc‒Scales Creek; pk‒Kearsarge; pg‒ Greenstone. Letter codes on left side are sedimentary layers, except for one: pu‒unnamed conglomerate (cgl), pbc-pl‒Baltic-Lac La Belle cgl; ps‒St. Louis cgl; pb‒Bohemia cgl; poc‒Old Colony sandstone; pw‒Wolverine sandstone; pkc‒Kingston cgl; pc‒Calumet and Hecla cgl; ph‒ Houghton cgl; pa‒Allouez cgl; pp‒Pewabic West cgl; paf‒ashbed flow top; phc‒Hancock cgl. Bold red bars mark units with layer-parallel slip observed in mines and trenches. Letters on right side indicate where slip occurred (T‒top; B‒bottom; A‒top and bottom). The Portage Lake Volcanics (PLV) exposed along the Keweenaw Peninsula is truncated on the southeast by the Keweenaw fault system, leaving an undetermined thickness of PLV in the footwall beneath Jacobsville Sandstone. The PLV section in the fault’s hanging wall is estimated to be 3000 to 5000 m thick and to contain about 300 flows (Cannon and Nicholson, 2000). It mostly consists of subaerial basaltic flows with less abundant andesitic flows, rhyolitic to dacitic extrusive domes, and associated pyroclastic layers (Butler and Burbank, 1929; White, 1968; Cannon and Nicholson, 2000, 2001). Six named basalt flows are sufficiently traceable along strike due to their 100 �great thickness and distinctive texture to serve as marker units (Fig. 2). The thickest of these, the Greenstone flow, in the upper part of the section is traceable for 80 kilometers in outcrop and drill holes and has a precise radiometric age date of 1091.6 ± 1.3 Ma (Swanson-Hysell et al., 2019; cf. 1094.0 ± 1.5 Ma, Davis and Paces, 1990). The next thickest flow, the Copper City flow, in the lower part of the section is traceable for 25 kilometers and is precisely dated at 1093.4 ± 1.4 Ma (Swanson-Hysell et al., 2019; cf. 1096.2 ± 1.8 Ma of Davis and Paces, 1990). About 3% of the PLV section consists of conglomerate and sandstone deposited between some of the lava flows (Butler and Burbank, 1929; White, 1952; Merk and Jirsa, 1982). Eleven named interflow units (Fig. 2) are traceable along strike for considerable distances up to 90 km (Bornhorst and Barron, 2011), and thus are useful for correlation. Interflow sedimentary layers mostly consist of clast-supported, pebble-to-cobble conglomerate with a gravelly to sandy matrix, which is well indurated and tends to form prominent strike ridges. Conglomerate layers are coarsely bedded and poorly stratified, except where sandy lenses or conglomeratic sandstone occur. Less common are sandstone and siltstone that may occur as interbeds within a conglomerate layer or locally may make up an entire interflow layer. Interflow sedimentary layers may be up to 40 m thick and locally may pinch to near zero thickness (Butler and Burbank, 1929; Merk and Jirsa, 1982; White, 1968). Clasts and matrix grains in the interflow layers were derived from volcanic rocks of the rift basin, with felsic material generally being far more abundant than expected from the small proportion of felsic volcanic rocks that make up the PLV section (< 1% according to Nicholson, 1992). These so-called felsic conglomerates may occur by themselves or in association with mafic conglomerates, in which case the felsic layer tends to lie atop the mafic layer lying atop the underlying lava flow (Butler and Burbank, 1929). Stops of this field trip are all within the lower part of the PLV section exposed in the fault’s hanging wall. For the first half of the trip, the thick Copper City flow is an important stratigraphic reference that be traced for 25 km along strike. Another important stratigraphic reference with a greater strike extent is the St. Louis conglomerate (#6 on USGS bedrock geology maps), which lies at the base of the Copper City flow in many places but elsewhere is separated from it by a thinner basalt flow. For the second half of the trip, the key stratigraphic reference is the Lac La Belle conglomerate (#3 on USGS maps), which lies 580 m to 1050 m stratigraphically below the St. Louis conglomerate. Other lava flows and interflow sedimentary layers in the lower part of the PLV have not yet been correlated well enough to be useful as regional stratigraphic references. A few intrusions cut the lowermost PLV layers along the Keweenaw Peninsula, such as the Mt. Bohemia syenodiorite stock (Cornwall, 1954a) and scattered dikes largely of intermediate composition (Robertson, 1975). Jacobsville Sandstone (JS) in the Keweenaw fault’s footwall fills a sub-basin that extends southeast beneath Keweenaw Bay to an onlap with Paleoproterozoic metasedimentary rocks, northeast to near Stannard Rock, and southwest to Lake Gogebic (Fig. 1). The unit may reach to 23 kilometers in depth based on gravity modeling and early seismic reflection data (Bacon, 1966; Aho, 1969; Kalliokoski, 1982). An exploratory drill hole (Mayflower #41) through the fault near Calumet penetrated JS to a total depth of 803 m (White, 1985), and a deeper drill hole 8.6 km southeast of the fault (Rice Lake #1) reached a total depth in sandstone at 1106 m (Keweenaw NHP, 2016). Regionally correlated units are not well defined for the unit because of lateral facies variability and lack of marker beds. However, Hamblin (1958) recognized the following 101 �succession: (1) basal “conglomerate facies” with locally-sourced clasts deposited adjacent to topographic highs; (2) “lenticular sandstone” facies dominated by planar and trough cross-bedding; (3) “massive sandstone” facies with laterally persistent layers and occasional planar cross-bedding; and (4) “red siltstone” facies with thinly bedded, fine-grained, siliciclastic layers and local trough cross-bedding. Based on this facies succession, Hamblin (1958) interpreted Jacobsville depositional environments as transitioning from (1) alluvial-fluvial at the base, (2) to dominantly fluvial in the lower cross-bedded part of the unit, (3) to dominantly lacustrine in the upper massive part of the unit, and (4) to variably lacustrine and fluvial in the uppermost thinly bedded part of the unit. He inferred that the general sequence of facies and depositional environments was time transgressive, representing both a vertical time sequence at any given point as well as a basinward change of depositional environment at any given time. The JS section along the Keweenaw Peninsula consists of siliciclastic strata that generally conform to the facies of Hamblin (1958) but with some variations. Proximal to the fault system, the lower parts of exposed sections commonly have brownish conglomerate layers consisting of subangular volcanic clasts chaotically dispersed in a poorly indurated muddy matrix, and interbedded with soft, fissile, reddish-brown siltstone and shale (Irving and Chamberlin, 1885; Hamblin, 1958; DeGraff, 1976; Brojanigo, 1984). In such sections with muddy conglomerate and shaly strata, interbeds of whitish to orangish to pinkish, fine- to medium-grained sandstone occur as isolated thin beds near the base and become more abundant and thicker upward in the section until muddy strata are rare. The quartzose to subarkosic sandstones are locally conglomeratic and often crossbedded, which is typical of most JS strata exposed away from the fault system to the southeast. Sand-prone JS strata also occur near the fault system where the muddy conglomerate and shaly facies is absent. Brojanigo (1984) interpreted the muddy conglomerate layers as debris flows derived from an upland of PLV rocks to the northwest, and the more mature sandy strata as fluvial deposits derived from older quartzo-feldspathic rocks to the south. We interpret this bimodal alluvial-fluvial assemblage to be basal Jacobsville, whose PLV-derived detritus is evidence of prior reverse slip on the Keweenaw fault system that elevated the northwest hanging wall and allowed volcanics there to be weathered, eroded, and transported southeastward into the JS sub-basin. The time span of JS deposition has been difficult to determine due to a lack of fossils and igneous rocks interbedded with or crosscutting Jacobsville strata (Kalliokoski, 1982). Deposition probably did not occur prior to reverse slip on the Keweenaw fault system (KFS) because conglomerate low in the JS section near the KFS contains PLV detritus sourced from an elevated region to the northwest (Brojanigo, 1984; Cannon and Nicholson, 2000). The earliest slip on the fault system is generally taken to be ~1060 Ma based on Rb-Sr ages of gangue minerals in fractures associated with native copper mineralization of the Keweenaw Peninsula (Bornhorst et al., 1988) and reset biotite ages in Archean granite gneiss uplifted along related faults west of Lake Gogebic (Cannon et al., 1993). Later reverse slip on the Keweenaw fault that deformed adjacent Jacobsville strata is relatively well constrained at 985 ± 30 Ma by a U-Pb date for late-kinematic vein calcite in the fault zone, in combination with a maximum depositional age of ~993 Ma for deformed strata high in the Jacobsville section (Hodgin et al., 2022). Most Jacobsville deposition probably occurred in the period 1060-985 Ma, although somewhat younger layers could have been deposited in the region (Craddock et al., 2013; Malone et al., 2016). 102 �Keweenaw Fault System – Key Observations and Concepts The Keweenaw fault has been described as a strike fault because of the general parallelism between its map trace and PLV layers in its hanging wall (Fig. 3). Butler and Burbank (1929) also noted a remarkable similarity in dip of hanging-wall layers and the fault, such that PLV layers change dip along strike by about the same amount and in the same sense as changes in the underlying fault surface. Where differences in dip exist, the fault cuts 5° to 17° more steeply and upward across PLV layers in a southeasterly direction, i.e. in the thrusting direction (DeGraff and Carter, 2023). Related to these surface observations, Hubbard (1898) had systematically catalogued in mine workings many layer-parallel faults along contacts between volcanic and interflow sedimentary layers (Fig. 2). These bedding-plane faults are manifested by fault gouge, polished surfaces, breccia with alteration, and secondary mineralization in fractured zones up to 8 m thick (65% ≤ 1.5 m). Hubbard’s observations indicate that PLV layer boundaries were relatively weak and became detached during one or more deformation episodes. Furthermore, his observations along the Keweenaw Peninsula mining district northeast of Portage Lake document layer-parallel slip throughout the PLV section. The surface and subsurface evidence of layer-parallel slip within the PLV section is consistent with knowledge about PLV stratigraphy. The lava flows have variations in texture and mineralogy across their thickness that likely influence their mechanical properties. Massive holocrystalline flow interiors grade into margins with finer grain size and abundant vesicles and amygdules, often arranged in bands parallel to the contacts. Flow tops have the highest abundance of amygdules, are commonly brecciated, and often have been modified by weathering between eruptions or by mineralizing fluids that later migrated along these once-permeable zones (Butler and Burbank, 1929; Stoiber and Davidson, 1959; White, 1968; Bornhorst, 1997). Where one flow lies directly atop another, a significant mechanical contrast exists across their comparatively weak contact. Contrasts in layer characteristics and properties are even more significant at contacts between interflow sedimentary layers and lava flows. Altogether, the evidence just summarized implies a detached style of formation for the Keweenaw fault and related splays (DeGraff and Carter, 2023). Although the Keweenaw fault has been described as a high-angle reverse fault, its northwesterly dip at the surface varies from as low as 20° up to 70° (Butler and Burbank, 1929). Cross-section construction along a transect with abundant constraining data from mining operations and surface mapping implies that the fault is nearly horizontal northwest of its surface trace at Hungarian Falls (DeGraff and Carter, 2023, their Fig. 6), one of the stops on this trip (see Stop 2-1, Fig. 8). Several cross-sections on USGS bedrock geology map sheets from the 1950s show the fault dipping ≤ 25° or show nearly horizontal PLV layers near surface that imply a similarly-dipping fault below (e.g., Laurium, Ahmeek, Mohawk, and Bruneau Creek quadrangles). Thus, the fault and overlying PLV strata dip between nearly horizontal up to 25° NW at several locations and, there, the hanging-wall PLV block has overridden footwall JS strata by as much as 2.5 kilometers. Between 2017 and 2022, three mapping projects funded by the USGS EDMAP program have refined fault and stratal geometries along the Keweenaw fault, and the results indicate that what was largely considered to be a single fault is better characterized as a fault system (Tyrrell, 2019; Mueller, 2021; Lizzadro-McPherson, 2023; Gamet, 2023, Langfield, 2024). The Keweenaw fault 103 �system (KFS), as used here, refers to an ensemble of fault segments whose collective motion has a significant component of dextral strike slip in addition to the long-recognized component of reverse slip with northwest side up. The fault system consists of segments that define three main directional sets: (1) a dominant set that defines the trend of the fault system and locally separates more steeply dipping PLV layers to the northwest from less steeply dipping PLV layers to the southeast; (2) splay faults that are angled 15-30° clockwise from set 1; and (3) connector faults angled 35-75° counterclockwise from set 1 that join footwall splays to the main fault trend (Fig. 3). The three directional fault sets maintain these angular relationships among each other as the curved KFS changes direction from a 35° azimuth near Houghton to a 95° azimuth near the tip of the peninsula. The interconnected nature of the three fault sets defines fault blocks with long dimensions parallel to the local trend of the KFS. Fault-slip indicators measured on ~400 fault surfaces show that the collective ratio of reverse slip to dextral slip on the KFS varies from 1:1 near Houghton to 1:2 or more near the tip of the peninsula. In other words, strike slip is twice as large as dip slip on the KFS near the tip of the peninsula where fault azimuth is 60° clockwise from its azimuth near Houghton (Tyrrell, 2019; Mueller, 2021; Lizzadro-McPherson, 2023; Gamet, 2023; Langfield, 2024). This change in the ratio of strike slip to dip slip is expected as the fault’s azimuth approaches the estimated 105° azimuth of maximum compressive stress that is attributed to the Grenville orogeny. The leftstepping arrangement of footwall splay faults and the fault blocks they help to define is consistent with a transpressional fault system as indicated by the fault-slip data. The relatively short, north- to northeast-trending, connector faults (set 3) are associated with similar-trending fold axes at relatively high angles to the estimated maximum compression direction and shortening direction. The connector faults are inferred to have mostly reverse slip with west side up and to accommodate northeast to east transport of footwall fault blocks associated with dextral transpressional slip of the entire KFS. Why concern ourselves with the geometric details and timing of slip on the KFS? From a science perspective, the arrangement of fault segments in the system along with fault-slip indicators like slickenlines provide clues to the mechanics of fault initiation and kinematics of fault slip. Improved understanding of these aspects of the KFS should apply to similar faults around Lake Superior and help to understand their causative tectonic events and stress regimes. From a practical perspective, faulting along the Keweenaw Peninsula and on strike to the southwest provided pathways for upward migration of mineralizing fluids (Bornhorst, 1997), leading to copper deposits that once supported a thriving mining industry and are still prospective today. Native copper is commonly found along major and minor faults in the mining district. The Hancock fault cutting the Quincy and Hancock mines is an example of a major fault with copper mineralization concentrated in its hanging wall. Another example is the Allouez Gap fault that bisects the Kearsarge flow-top copper deposit, the largest of this type in the district (Bornhorst, 1997). Faults likely provided pathways for upward-moving ore fluids into vesicular flow tops at the Baltic and Isle Royale deposits and others in the Greenland-Mass subdistrict (Broderick, 1931). The earliest native copper deposits exploited in the district east of Eagle River were along subvertical fracture zones and minor transverse faults with displacements less than 200 m (Butler and Burbank, 1929). Therefore, a better understanding of fault geometry and timing may provide insights about the distribution of known deposits and about the possible presence of undiscovered deposits. 104 �Figure 3: Geologic map of the Keweenaw Peninsula showing the Keweenaw fault system north of Portage Lake and field trip stops within the four focus areas. Faults are indicated by curved black lines. Field Trip Stops The field trip stops described below are grouped into four areas to illustrate themes related to fault system geometry, stratal relationships and deformation, fault blocks, deformation fabric in fault zones, and slip kinematics. Each area has an introductory overview of the key themes and nature of the stops, followed by individual stop descriptions. Reported strike and dip values follow a right-hand-rule convention. The dip value is followed by letters indicating the cardinal direction of dip, which is redundant but makes for clarity. About half of the stops in this guide are on public land. Those on private land are specified in the stop descriptions as requiring permission from the owners to access. 105 �Figure 4: Geologic map of Area 1 with the Keweenaw fault system (KFS) and field trip stops 1-1 to 1-3. Geologic unit codes: pbc – Baltic conglomerate, ps – St. Louis conglomerate, pb – Bohemia conglomerate, psc – Scales Creek flow, pk – Kearsarge flow with Wolverine sandstone at its base. Teeth along reverse faults are on upthrust sides. 106 �Area 1: Hungarian Fault Block ‒ Southwest End STOPS IN THIS AREA ARE ON PRIVATE LAND AND REQUIRE PERMISSION TO ACCESS. The three stops in Area 1 are within the vehicle testing grounds of Michigan Tech’s Keweenaw Research Center (KRC), under lease from the Houghton County Memorial Airport (Fig. 4). From west to east, the stops are on the main Keweenaw fault zone (1-1), in a basalt quarry within a faultbounded block (1-2), and on a footwall splay fault (1-3) that passes through Hungarian Falls to the northeast in Area 2. Within the fault-bounded block, an unconformity between the PLV and JS is interpreted to dip shallowly southwest. These three stops illustrate an en echelon, overlapping, fault geometry that define a fault block whose long dimension is parallel to the trend of the Keweenaw fault system (KFS). Recent mapping shows that such fault patterns and fault-bounded blocks are repeated along the KFS north of Portage Lake (Fig. 3). A similar configuration of faults with the enclosed block having a west-dipping PLV-JS unconformity occurs in Area 4 (Figs. 3 and 13). Stop 1-1: Gooseneck Creek fault exposure Directions: From the Portage Lake lift bridge between Houghton and Hancock, follow US-41 north for 6.7 mi (10.8 km) to Airpark Boulevard on the right. Turn right toward Houghton County Memorial Airport and drive 0.5 mi (0.8 km) to the Keweenaw Research Center (KRC) on the right. Enter the KRC parking area and wait in vehicles. We may be escorted by KRC staff for 1.5 mi (2.4 km) to the south edge of their vehicle testing grounds, where we will park and walk. [Lat: 47° 9.259'N | Lon: 88° 29.895'W] The fault exposure on Gooseneck Creek is part of the main Keweenaw fault zone that is traced by means of outcrops, drill holes, and water wells along a line bearing 36° from the Michigan Tech campus to the intersection of Airport Park and Forsman roads, about one kilometer southwest of this stop (Fig. 4). Northeast from that intersection, the fault’s map trace curves eastward so that here it trends 72°, as do the hanging-wall PLV strata. This stop is the most northeasterly site where this segment of the KFS can be observed because the flat upland to the northeast, where the airport is located, has little to no bedrock exposure. Hubbard (1898) seems to be the first geologist to describe this site. While his report puts the outcrop about 80 m south of its actual location, his geologic description is very similar to what was observed and measured during the 2021-2022 EdMap project (Langfield, 2024). “A conglomerate here, underlain by trap, strikes N. 72° E., and dips northerly 44°, the trap being in contact on the south with the sandstone, which is much broken and disturbed but appears to dip rather flat to the N. E.” (Hubbard, 1898) The fault’s main slip surface is not exposed but its position is determined to within a couple of meters by the proximity of hanging-wall PLV outcrops to footwall JS outcrops. PLV strata in the hanging wall, striking 252° and dipping 40-44° NW, change stratigraphically upward from amygdaloidal basalt at the fault at creek level to a felsic pebble conglomerate, and then to a series of basaltic flows for as far as outcrop exists to the north and northwest. The stratigraphic 107 �position of the conglomerate layer suggests that it is the Baltic (#3) conglomerate shown on the USGS Hancock and Chassell bedrock geology maps (Cornwall, 1956a; White, 1956). Abundant fractures within the basalt flows are not obviously systematic, although detailed work might reveal dominant sets. Jacobsville strata in the footwall commonly appear massive due to being thickly bedded to locally cross-bedded, which makes their attitude difficult to measure. In general, JS strata dip gently both toward and away from the fault, suggesting an open anticlinal structure with an axis that trends approximately east-west (Fig. 4). Figure 5: Gooseneck Creek cross-section in progress using outcrop and drill hole data (Langfield, 2024). Red lines below topography are New Arcadian drill hole trajectories. Colors and letter codes of geologic units are explained in Figure 2. KFS-P = Keweenaw fault system - Portage segment; KFS-H = Keweenaw fault system - Hungarian segment. Dashed orange line above KFSP marks what may be the Baltic (#3) conglomerate. Diamond drill holes (DDH) of the Calumet and Hecla Consolidated Copper Company provide data that are critical for the interpretation of this site, situated on a NW-trending line of four New Arcadian holes drilled in 1911-1912 (Fig. 4). Southeast of here, three drill holes penetrated the following beneath glacial overburden (hole depth converted to vertical depth): DDH #17 about 140 m SE cut 46 m of JS sandstone followed by 109 m of PLV basalt; DDH #15 about 205 m SE cut 69 m of JS sandstone and conglomerate followed by 14 m of PLV basalt; DDH #13 about 355 m SE cut 37 m of JS sandstone (Fig 5). Southeast of the fault, therefore, drilling reveals a veneer of JS strata less than 70 m thick overlying PLV basaltic lava flows. About 135 m northwest of the fault, DDH #11 inclines 52° toward this site and reaches a total depth of 457 m. It cuts a single felsic conglomerate layer between depths of 93 m and 109 m (16 m apparent thickness) that we correlate to the 6-m-thick conglomerate layer seen here in outcrop, which yields an average stratal dip of 53° NW. Below the conglomerate, DDH #11 penetrated ~2 108 �meters of sludge that we interpret as fault gouge, and then continued through basaltic flows to its bottom at a vertical depth of 363 m below the DDH #17 surface location to the southeast. The position of the fault below the conglomerate in DDH #11 and not far below the conglomerate at the surface indicates that the fault dips nearly parallel to hanging-wall PLV strata. While at the surface the fault juxtaposes PLV strata against JS strata in the footwall, at depth in DDH #11 the fault’s hanging wall and footwall both have PLV strata, which is consistent with the thin veneer of JS strata penetrated by DDH #17 and DDH #15 southeast of here. Stop 1-2: Basalt quarry of the Keweenaw Research Center Directions: Return to vehicles and drive north 0.15 mi (0.25 km) to KRC perimeter road. Turn right and drive southeast for 0.75 mi (1.2 km) to a basalt quarry on the right. Pull off road into a gravel terrace on the right and park. [Lat: 47° 9.144'N | Lon: 88° 29.023'W] The KRC basalt quarry was opened in 2018-2019 to provide materials for expansion of their vehicle testing facilities. The quarry lies between the fault segment just visited at Gooseneck Creek and a parallel segment to be visited at the next stop (Fig. 4). The quarry exposes portions of at least two basalt lava flows that dip shallowly southwest. The shallow dip is roughly manifested by subhorizontal benches of the quarry and can be measured on the first dry bench west of the quarry pond, where the upper surface of a lava flow has been exposed. Inspection of the irregular subhorizontal surface reveals isolated patches of sediment whose stratification yields an average strike of 125° and dip of 19° SW. The shallow southwesterly dip of PLV strata is important to the understanding of structural geometry. Early geologists noted that older PLV strata near the Keweenaw fault and its splays commonly have anomalous orientations relative to younger PLV strata away from the fault zone, which have a well-defined regional trend (Hubbard, 1898; Butler and Burbank, 1929). In most of the PLV section, strata generally strike northeast to east and dip 35° – 55° northwest to north. At this stop and for ~3 kilometers north and northeast, PLV strata dip less than 25° in various directions, including counter-regional to the southeast. Such anomalies are clues to the structural configuration of the area along the Keweenaw fault system. Specific to this stop, the shallow southwesterly dip of PLV stratal is consistent with data from DDHs #17 and #15, where a thin veneer of JS strata overlies PLV basaltic rocks. The nature of the PLV-JS contact is not described in the core logs, but the normal stratigraphic sequence indicates that it is an unconformity. The unconformity dips shallowly southwest to south based on the two cited DDHs and DDH #20 located ~510 m southwest, which penetrates the unconformity ~35 m lower relative to a common datum. Projecting the unconformity updip brings it to the surface southwest of this stop. Therefore, we interpret that basalt at the quarry roughly correlates with basalt below JS strata in the three DDHs by passing beneath an erosional wedge of JS strata on a SW-dipping unconformity (Langfield, 2024). This new interpretation differs from one involving a transverse fault shown on USGS bedrock geology maps for the Hancock and Laurium quadrangles (Cornwall and Wright, 1956a, 1956b). 109 �Stop 1-3: East branch of Quincy Creek fault exposure Directions: Return to vehicles and drive east on the perimeter road for 0.25 mi (0.4 km) to an intersection with a gravel road on the left. Turn left and drive ~0.35 mi (~0.6 km) to a clearing on the right. Turn into the clearing and park. [Lat: 47° 9.282'N | Lon: 88° 28.548'W] The east and west branches of Quincy Creek cross a segment of the Keweenaw fault system that is about 750 meters southeast of the fault segment at Gooseneck Creek and subparallel to it. Both branches of Quincy Creek constrain the position of the fault to within 8-10 m by the proximity of hanging-wall PLV outcrops to footwall JS outcrops, but the fault zone itself is not exposed. Outcrops along the west branch of Quincy Creek are the most southwesterly constraint on the position of this fault segment because of glacial deposits that cover bedrock southwest of here. However, we interpret the fault to continue southwesterly to an acute intersection with the fault segment seen on Gooseneck Creek at a point about halfway to Portage Lake (Fig. 4). Northeast from this stop, the fault segment is easily traced by means of outcrops and water wells along an azimuth of 42° for a distance of 3 km to the upper Hungarian Falls in Area 2. We will focus on the east branch of Quincy Creek because it provides better exposures near the fault. Again, Hubbard (1898) appears to have been the first geologist to describe outcrops at this site as well as other outcrops in stream valleys to the northeast that cross the fault line. Although his report puts the outcrop 150-175 meters north-northeast of its actual location, his geologic description of part of the outcrop is very similar to what was observed and measured during the 2021-2022 EdMap project (Langfield, 2024). “In Sec. 22 . . . occur outcrops of sandstone and of a conglomerate with a very sandy matrix. The pebbles in the conglomerate are subangular and some of them are of quartz porphyry. The dip is about 50°-54° N. W., strike about N. 45°-50° E.” (Hubbard, 1898) The outcrop just described begins 60 meters downstream from the gravel access road and extends another 30 m downstream. The sedimentary layers here are well indurated and well stratified, having an average strike of 225° and dip of 35° NW. A reddish-brown sublithic sandstone is the dominant rock type along the semi-continuously exposed section in the creek bed, with subordinate conglomeratic sandstone and pebble-to-granule conglomerate that ranges from matrix-supported to clast-supported. Clasts are mostly subrounded to subangular and have a variety of compositions but are dominantly felsic. Whereas Hubbard (1898) thought that these were JS strata, we interpret them as PLV strata because of their induration, lithic nature, and attitude that are similar to other PLV sedimentary units in the area and differ from JS sandstone strata to be seen downstream. The south end of the 30-m extent of indurated sandstone and lesser conglomerate coincides with a sharp deflection in the creek by ~20 meters east before the creek resumes its southerly course at a sharp right bend. Downstream from this point for ~30 meters, intermittently exposed sandstone and minor conglomerate differ in many aspects from the sedimentary strata upstream of the creek’s deflection. The sandstone is lighter toned and locally streaked off-white to beige, is less 110 �indurated and commonly friable, and has a more quartzose composition. The pebble-to-granule conglomerate layer in the first outcrop downstream from the creek’s deflection has a muddy to silty matrix, is poorly indurated, and has subrounded to rounded clasts with a variety of compositions. These characteristics of the sandstone and one conglomerate layer are typical of Jacobsville strata observed elsewhere in fault juxtaposition with hanging-wall PLV strata. The sharp deflection in the creek is interpreted, therefore, as marking a fault contact between PLV sedimentary strata upstream and JS strata downstream. This interpretation is supported by structural observations. Strata in the first outcrop south of the creek’s deflection strike 220° and dip 65° NW (overturned), but become less steeply dipping over a short distance as the creek is followed downstream to the south. These stratal attitudes and their significant change going downstream from the creek’s deflection contrast with stratal orientation and its consistency over a similar distance upstream of the deflection. About 15 meters south of the creek’s deflection, a small fault striking 225° and dipping 80° N cuts thickly bedded JS strata in a larger outcrop along the west bank of the creek. South of the fault, deformation bands in the sandstone are expressed as quasi-linear to broadly sinuous ridges on the outcrop surface. Deformation bands are cataclastic shear zones that develop during compression of partly indurated, porous, clastic material (Fossen et al., 2007). They are generally more cemented than the host material, accounting for their raised relief on erosional surfaces, and have not been observed in PLV sedimentary strata. Farther downstream, JS strata are abundantly displayed for a distance of 1.4 km, nearly down to the derelict Quincy Mining Company stamp mill along state highway M-26. Over that distance, JS strata generally dip less than 15° NW and display a few broad open folds with NNE-trending axes. Area 2: Hungarian Fault Block ‒ Northeast End The first two stops in Area 2 are on the same fault segment seen at Quincy Creek (2-1) and on its curved portion (2-2) that connects back to the main Keweenaw fault zone to the north (Fig. 6). A PLV sedimentary layer, usually conglomerate but locally sandstone, is traceable in the hanging wall from Quincy Creek through the east branch of Dover Creek near Hungarian Falls, where it begins a smooth northward curve from 42° to 345° azimuth, a change of over 55°. From this curved stratal geometry in map view, we infer a single smoothly curved fault along the southeast and east edges of the Hungarian fault block. The long straight part of the fault is the block’s southeast edge that parallels the overall KFS, whereas the short curved part is the block’s east edge and connects the fault segment back to the main Keweenaw fault zone. This geometry and a northwesterly decrease in stratal dip based on drill hole data define a single scoop-shaped fault rather than an intersection of two distinct faults. Stop 2-2 illustrates a common dynamic of the KFS – northeast and east edges of fault-bounded blocks were thrust eastward along west-dipping reverse faults. The third stop in Area 2 is on the main Keweenaw fault zone (2-3), which aligns with its counterpart in Area 1 (Fig. 3). The relationship between the main fault zone and the connector fault may involve the connector fault terminating against a main fault zone that continues to the southwest (Fig. 6 at “?”). Another option is that the hanging-wall sedimentary layer is continuous from Stop 2-2 to Stop 2-3 and is draped over a lateral ramp in the KFS that steps up in stratigraphy to the northeast. The stratigraphic relationships across the Hancock fault are discussed in the Stop 2-3 description. 111 �Figure 6: Geologic map of Area 2 with the Keweenaw fault system (KFS) and field trip stops 2-1 to2-3. Geologic unit codes: ps – St. Louis conglomerate, pcc – Copper City flow; pb – Bohemia conglomerate, psc – Scales Creek flow, poc – Old Colony sandstone; pk – Kearsarge flow with Wolverine sandstone at its base. Teeth along reverse faults are on upthrust sides. See Figures 4 or 13 for symbology. 112 �Stop 2-1: Dover Creek fault exposure at Upper Hungarian Falls Directions: Return back to the KRC perimeter road, then turn right and follow it back to the KRC entrance on Airpark Boulevard. Turn left and drive 0.5 mi (0.8 km) to US-41. Turn right and drive 1.6 mi (2.5 km) to Oneco Road. Turn right onto Oneco and drive 3.1 mi (5.0 km) to Amygdaloid Street. Turn left and follow Amygdaloid and Tamarack Hill Road for 0.3 mi (0.5 km) to M-26. Turn left and drive 0.35 mi (0.56 km) to 6th Street in Hubbell. Turn left and then, after the second house on the left, veer left onto Golf Course Road going uphill. Drive 0.5 mi (0.8 km) to a rutted gravel road on the left. Depending on conditions, we may turn left and take the gravel road. Otherwise, park along Golf Course Road and walk to the stop at Upper Hungarian Falls. [Lat: 47° 10.440'N | Lon: 88° 27.086'W] The fault segment crossed by Dover Creek is near the northeast end of the segment seen at Quincy Creek, and is where the fault’s surface trace begins to curve northward (Fig. 6). Dover Creek has three sets of falls that are worthwhile visiting. The two downstream falls have larger drops entirely over Jacobsville Sandstone. We will visit the upper set of falls at the fault contact between PLV strata upstream in the hanging wall and JS strata in the footwall. This site has been visited by geologists since the mid-1800s and is considered a classic exposure of the Keweenaw fault (original sense). The most insightful description by early geologists is in the work of Irving and Chamberlin (1885), whose diligent field work, keen observations, and logical reasoning convinced the geological community of the time that the contact between PLV strata and JS strata was a large fault. Prior to their work, some geologists argued that JS strata lay stratigraphically below PLV strata based in part on their similar dip here and at Houghton-Douglass Falls on Hammell Creek (Stop 2-3). Of historical interest, Roland Duer Irving is this year’s nominee for recognition by the ILSG as a Pioneer of Lake Superior Geology, and Thomas Chrowder Chamberlin is another wellknown geologist of his time, famous for his classic 1890 work “The method of multiple working hypotheses” published in Science. Both were contemporaries of John Wesley Powell, a U.S. Army officer during the American Civil War, famed explorer of the American west, and second director of the U.S. Geological Survey from 1881–1894. The faulted relationship between PLV and JS strata at this stop was revealed by excavations made by a “force of miners” hired to trench across the PLV-JS contact at three locations on the southwest valley wall (Fig. 7). The trenches exposed a fault zone dipping 30-35° NW along the contact, which has the following internal zonation from hanging wall to footwall, as summarized from Irving and Chamberlin (1885) and converted to true thickness. 1. Trap [basalt]: highly fractured but in place. (2.57 m) 2. Trap debris: disintegrated basaltic fragments in a lumpy crudely laminated clay, having a transitional boundary with zone 1. (0.39 m) 3. Clay: red and “shaly” with light grayish-green spots, some sandy seams, and occasional lumps of disintegrated trap. (0.13 m) 4. Trap debris: similar to zone 2 but with more clayey material, whose dark color contrasts with adjacent red clay. (0.13 m) 113 �5. Trap debris mixed with red clay: trap debris is similar to zone 4 but red clay and minor sand gives zone 5 a reddish tint. (0.17 m) 6. Sandstone: light reddish and quartzose. (0.51 m) Footwall JS strata near the fault contact dip 10° or less, but at the fault contact they are bent downward to be nearly parallel with the fault surface. The trench observations indicate that the fault propagated upward across subhorizontal JS strata and the overriding PLV strata crushed and abraded the truncated edges of JS strata. Figure 7: Trenches at upper Hungarian Falls that exposed relationships across the fault, marked by the red lines (Irving and Chamberlin, 1885). Downstream from the trenches, JS strata generally dip 10-20° NW toward the fault but exhibit a few broad open folds similar to what is observed downstream along Quincy Creek. Upstream from the trenches, PLV strata in the hanging wall begin with a fractured basalt flow at the minor falls below the main falls, and then progress stratigraphically upward to a felsic cobble-pebble conglomerate at the main falls, followed by a series of basaltic flows intermittently exposed for over two kilometers upstream. Stratal dip decreases upstream from 25-30° NW at the main falls to flat-lying and then to 15° SE to define an open syncline, which is succeeded upstream by an open anticline before reaching the Hancock fault (Fig. 8). A well-laminated, 1-m-thick sandstone layer at the base of the 6-m-thick conglomerate layer provides a confident formation strike of 205° and dip of 30° NW, which is essentially parallel to the underlying fault surface. This geometry indicates that the hanging-wall PLV section at this location became detached along 114 �layering somewhere downdip to the northwest and was thrust southeast and upward to its current position along a ramp that cuts upward across younger layers. The fault segment exposed at this stop is penetrated by the vertical Oneco #9 DDH located 1.7 km west-northwest, at a depth of 556 m, which gives an average fault dip over this distance of 19° NW after accounting for topography (Fig. 8; DeGraff and Carter, 2023). Because the fault dips 30-35° NW at the surface, it must dip less than 19° NW over some portion of its trajectory between here and the Oneco #9 DDH. That is, its dip must shallow going in a northwesterly direction similar to the shallowing of PLV stratal dips observed upstream along Dover Creek, which is characteristic of a detached style of faulting with layer-parallel slip. Figure 8: Cross-section along Dover Creek based on outcrop and drill hole data (DeGraff and Carter, 2023). Long bar inclination of L-shaped symbols at land surface shows apparent dip. Thin black lines below topography are drill hole trajectories. Colors and letter codes of geologic units are explained in Figure 2. KF‒Keweenaw fault; HF‒Hancock fault; B‒Bacon (1966) seismic experiment. Stop 2-2: Beaudoin Creek fault exposure Directions: Return to Golf Course Road and turn left to go uphill. Drive 0.5 mi (0.8 km) on Golf Course Road to Beaudoin Creek. We will park along the west shoulder of the road north of the creek culvert. This will require turning vehicles around at the first driveway north of the creek. [Lat: 47° 10.806'N | Lon: 88° 27.049'W] THIS STOP IS ON PRIVATE PROPERTY. PERMISSION IS REQUIRED TO OBTAIN ACCESS. 115 �The small creek crossed by Golf Course Road is unofficially called Beaudoin Creek after the property owner, who also owns the Wood’n Spoon specialty food and gift shop in Mohawk, MI. We will walk upstream to the west for 300 meters to outcrops first described by Hubbard (1898), who traced the PLV-JS fault contact throughout this area and noted remarkable changes in its strike. Between Quincy Creek and the west branch of Dover creek (not visited), the Hungarian fault segment has an azimuth of 42°, but it begins to curve northward at the east branch of Dover Creek, where it has an azimuth of 25° (Fig. 6). The gradual change in fault direction continues from the last stop to this one such that here the fault trace and hanging-wall PLV strata have an azimuth of 345°, which completes a total change in fault strike of 55-60° along a smooth arc. The NNW-trending portion of the fault segment extends more than a kilometer north to a point along another smooth curve of the fault toward the northeast, which re-aligns the fault with the trend of the Portage fault segment. It is unlikely that the position of the second broad curve in the fault north of here is a coincidence, and more likely that it manifests in some way a continuation of the KFS-Portage segment seen at Stop 1-1. Along Beaudoin creek east of the fault, JS strata are intermittently exposed over a distance of 170 meters and mostly consist of yellowish quartzose sandstone with occasional layers of reddish siltstone and conglomerate. Over most of this distance, JS strata dip less than 15° W toward the fault, but within 20 meters of the fault they dip more steeply at 70° E to vertical. The fault contact between PLV basaltic rock to the west and JS sandstone to the east is located to within a meter by the exposures, but the fault zone is not well exposed due to the degraded basaltic rock in the hanging wall. PLV stratigraphy here is very similar to that observed at Hungarian Falls, beginning with a basaltic lava flow of low relief that extends upstream to a small pond below a 6-m-tall waterfall. The waterfall is over a NNW-trending ridge of felsic cobble-pebble conglomerate ~10 m thick that has a meter-thick basal layer of siltstone to fine-grained sandstone, as seen at the previous stop. West of the conglomerate ridge, a series of basaltic flows crops out along the creek bed for another 200 meters. A reliable measurement of PLV strata orientation from the basal unit of the conglomerate layer gives a strike of 170° and dip of 45° W, which probably is also the attitude of the fault surface by analogy with Hungarian Falls and based on the interpretation of a detached fault system. The smooth curve of the fault segment and subparallel PLV strata that are traceable from southwest of Stop 2-1 to north of Stop 2-2, combined with the decrease in dip of the fault and PLV strata toward the Oneco #9 DDH, imply a curved fault surface in three dimensions. A number of geometries involving smaller faults and folds are possible, but the overall geometry implies a larger scoop-shaped fault surface that plunges between southwest and west. Further work to integrate surface mapping with subsurface DDH data are needed to fully define the fault and stratal geometries in this area. For now, we interpret the curved fault segment as part of the Hungarian segment of the KFS, which defines the long southeast edge and shorter east edge of the Hungarian fault block (Fig. 6) with relatively shallow dipping PLV strata. 116 �Stop 2-3: Hammell Creek fault exposure at Houghton-Douglass Falls Directions: Drive southeast on Golf Course Road 1.0 mi (1.6 km) to M-26. Turn left and drive 2.6 mi (4.2 km) to 10th Street in Lake Linden, where M-26 makes a 90° left turn. Turn left and drive 2.0 mi (3.2 km) to a parking area on the right. Turn in and park. [Lat: 47° 12.413'N | Lon: 88° 25.611'W] Houghton-Douglass Falls (a.k.a. Douglass Houghton Falls) was purchased by the State of Michigan in 2018 and, as of this writing, is being converted into a day park with walking trails and signage. The work is not yet complete and access is still somewhat limited but allowed. We will view the site from overlooks at the top of the steep canyon walls and will not descend to the fault contact near the base of the falls due to time constraints. The total vertical drop of 34 m (110 feet) makes Houghton-Douglass Falls the tallest in Michigan. This stop is another classic exposure of the Keweenaw fault (original sense) that has been investigated by geologists since the mid-1800s. The trail to the overlook area crosses the Hancock fault, not yet found in outcrop, and PLV strata in the Keweenaw fault’s hanging wall. The Hancock fault has been traced 16 km on an azimuth of 55° by means of drill holes from the Hancock and Quincy mines to an intersection with the Keweenaw fault one kilometer northeast of the falls (Fig. 3). The Hancock fault seems to have had an important role in the occurrence and distribution of native copper at those two mines (Bornhorst et al., 1986; Field Trip 2 in this volume). Copper mineralization at the base of Houghton-Douglass Falls next to the Keweenaw fault was investigated in an adit opened by the Douglass Houghton Mining Company, organized in 1845 (Stephens, 1902), and recently observed in veins (Gamet, 2023). The geometry of two PLV layers near the intersection of the Hancock and Keweenaw faults is critical to understanding geologic relationships at Houghton-Douglass Falls (Fig. 6). The Laurium bedrock geology map shows the St. Louis (#6) conglomerate and overlying Copper City flow in the hanging walls of the Hancock and Keweenaw faults north and west of their intersection but not to the southwest in the acute fault wedge containing Houghton-Douglass Falls (Cornwall and Wright, 1956b). The two units are easily recognized and correlated in drill holes and outcrops from many kilometers northeast of Calumet-Laurium in a southwesterly direction to the Hancock fault. The St. Louis conglomerate is a felsic, pebble-cobble, clast-supported conglomerate with locally significant lithic and conglomeratic sandstone that is associated with rhyolitic rocks along strike (Hubbard, 1898; White et al., 1953; Nicholson, 1992; Gamet, 2023). The Copper City flow is recognized for its anomalous thickness of 180 m at a point 7 kilometers to the northeast, coarse grain size, and pegmatitic segregations that have been precisely dated (Fig. 2). Recent mapping aided by drill hole data has allowed these two units to be correlated across the Hancock fault into the acute fault wedge, where they are identified in outcrop at Houghton-Douglass Falls (Fig. 6). The main drop at Houghton-Douglass Falls is over the Copper City flow, which extends from upstream of the waterfall down to a ledge ~10 meters above its base. From the north side of the gorge, a planar fabric with meter-scale spacing over much of the flow thickness dips gently 117 �upstream. Observations at the top of the waterfall suggest that this fabric results from amygdule layers that dip 20-25° NW. Below the main waterfall on the ledge above the base, a 3-m-thick sedimentary layer, first noted by Hubbard (1898), consists of a felsic pebble-granule conglomerate and coarse lithic sandstone whose layering provides a reliable strike of 225° and dip of 20° NW, i.e. parallel to amygdule layers at the top of the waterfall. This is interpreted to be the St. Louis (#6) conglomerate below the anomalously thick Copper City flow. Both layers project northeast along strike to intersect the Hancock fault at a point where their map offsets across the fault match the offsets of PLV layers previously correlated across the fault (Fig. 6). Below the St. Louis conglomerate is a basalt lava flow that becomes increasingly fractured at the foot of the falls, at which point sheared basaltic rock is observed along the south face of the gorge. Figure 9: Trenches below Houghton-Douglass Falls that exposed relationships across the fault, marked by the bold red line (Irving and Chamberlin, 1885). Irving and Chamberlin (1885) focused their investigation of Houghton-Douglass Falls along the bottom and walls of the gorge below the falls, again combining careful field observations with trenching across the fault surface. Two trenches running up and down the south valley wall and a third trench along the fault surface (Fig. 9) revealed similar relationships to those observed in the trenches at Hungarian Falls but with some important differences. Basaltic rock in the hanging wall has a clay seam at the fault surface that transitions upward to a clayey breccia and then to fractured but intact basalt as seen in the Hungarian trenches. However, footwall JS strata here are generally poorly indurated, red, shaly conglomerate with subordinate whitish quartzose sandstone layers, which is the opposite relationship of dominant to subordinate lithologies at Hungarian Falls. The downward flexing of subhorizontal JS strata to dips of 20-30° close to the fault surface is again 118 �observed, but the flexing is spread over a larger horizontal distance than at Hungarian Falls. The wider zone of flexure here may result from the poor induration and ductility of the red shaly conglomerate here relative to the moderately indurated sandstone strata at Hungarian Falls. Irving and Chamberlin (1885) estimated a fault dip of 25-30° NW based on their excavations, which is similar to their estimate of PLV stratal dip of 25° NW in the hanging wall. Recent field work yielded a fault dip of 20° NW based on a three-point method using surveyed points across the gorge and by visually siting upward along the fault surface (Gamet, 2023). This result for fault dip matches the PLV stratal dip measured on the St. Louis conglomerate and on amygdule layers in the Copper City flow. While earlier and recent dip values are slightly different, both studies agree that hanging-wall PLV strata are essentially parallel to the underlying fault surface, which is indicative of a detached thrust system. A recent cross-section through Houghton-Douglass Falls uses concepts related to detached thrusting, conservation of volume, and ductile portions of the JS section to model fault geometry and deformation of strata in the hanging wall and footwall of the fault system (Fig. 10). Figure 10: Cross-section along Hammell Creek at Houghton-Douglass Falls based on outcrop data and drill hole correlations (Gamet, 2023). Geologic unit codes: pcc – Copper City flow; psc – Scales Creek flow, pk – Kearsarge flow; pg – Greenstone flow. KFS-M = Keweenaw fault system Mayflower segment; HF = Hancock fault. Area 3: Snake Creek Fault Block – Keweenaw Fault Zone at Lake Gratiot The three stops in Area 3 (Fig. 11) are on the east boundary fault of the Snake Creek block (3-1), on the main Keweenaw fault zone north of Lake Gratiot (3-2), and at the intersection of these two fault trends (3-3). The Keweenaw fault zone roughly parallels the north edge of Lake Gratiot and, relative to Areas 1 and 2, it trends more easterly with an azimuth of 72° and dips more steeply northwest. Other fault segments in Area 3 are similarly oriented clockwise relative to their counterparts in Areas 1 and 2. South of the main Keweenaw fault zone, two fault-bounded blocks have long dimensions oriented parallel to the KFS (Fig.3). West of Lake Gratiot, the Snake Creek block is bounded on the southeast by a NE-trending connector fault, along which PLV layers are 119 �thrust southeast over JS strata (Fig. 11). Deformed hanging-wall PLV strata and footwall JS strata have NE-trending fold axes parallel to their fault contact. Fold axes in the hanging wall plunge 23° with an azimuth of 44° (Mueller, 2021). The southern boundary fault of the Snake Creek block is an ESE-trending footwall splay of the main Keweenaw fault zone (Fig. 3) that juxtaposes PLV strata on the north against JS strata to the south. Similar to the northeast end of the Hungarian block (Fig. 6), this footwall fault splay appears to curve north and merge seamlessly into the connector fault along the southeast edge of the Snake Creek block. Figure 11: Geologic map of Area 3 with the Keweenaw fault system (KFS) and field trip stops 3-1 to 33. Geologic unit codes: pb – Bohemia conglomerate, pgf – Gratiot flow. Teeth along reverse faults are on upthrust sides. See Figures 4 or 13 for symbology. The main Keweenaw fault zone in this area is well exposed along several creeks that empty into Lake Gratiot and, based on outcrop relationships and cross-section models, it consists of two parallel branches (Fig. 11). The northern branch has a well-developed gouge and breccia zone that is at least 20 m wide in places and perhaps as much as 45 m wide. This branch juxtaposes PLV basaltic flows on the north against presumably younger basaltic flows on the south, whereas the southern branch juxtaposes basaltic flows on its north against JS strata on the south. The double fault zone north of Lake Gratiot continues in a west-southwest direction past the termination of the 120 �connector fault intersecting from the southwest and forms the northern boundary of the Snake Creek block. This main fault zone trajectory differs from the 1950s USGS bedrock geology maps and was first proposed by Cannon and Nicholson (2001) in their map compilation. We have adjusted the position of their proposed fault based on a combination of outcrop relationships, drill hole data, and topographic features, and we consider evidence for its existence to be compelling. Stop 3-1: Unnamed creek fault exposure along Iron Gate Road Directions: Continue driving north on M-26 for 2.0 mi (3.2 km) to Hecla Street. Turn right and drive through historic downtown Laurium for four blocks (0.5 mi, 0.8 km) to 1st Street, a.k.a. School Street. Turn left and drive 0.35 mi (0.56 km) to US-41. Turn right and drive 17.6 mi (28.3 km) past the road to Eagle Harbor to historic Central Location. Turn right onto Gratiot Lake Road and drive ~4.8 mi (~7.7 km) to unpaved Iron Gate Road on the right. Turn right and drive 0.5 mi (0.8 km) to an unimproved dirt road on the right. Turn into the side road and park. [Lat: 47° 21.222'N | Lon: 88° 9.340'W] This stop will be a quick show-and-tell to explain fault and stratal geometries at the east edge of the Snake Creek fault block that extends 5 kilometers to the west (Fig. 11). A short walk northwest along the two-track road leads to a 3-m-tall outcrop of JS in the creek northeast of the road. This outcrop is much larger than other JS outcrops sometimes exposed for 35 meters upstream in the creek bed, depending on spring run-off. About 80 meters upstream is the first PLV basalt outcrop where the creek emerges from the upland to the northwest. The boundary between the upland with PLV bedrock and the lower flatter area to the southeast with JS bedrock has a trend of 42° and it extends ~3 kilometers from the southeastern rounded corner of the Snake Creek fault block to the main Keweenaw fault zone at Nine Thirty Two Creek (Stop 3-3). Several small creeks flowing southeast from the upland toward Lake Gratiot cross this boundary and expose bedrock, constraining the position of the geologic contact but not exposing it. The sandstone strata here strike 36° and dip 63° SE in the inferred direction of stratigraphic up, indicating that they have been rotated down to the southeast (Lizzadro-McPherson, 2023). Similar orientations of JS strata are noted at the mouths of other creek valleys where they emerge from the upland. Scattered JS outcrops to the southeast have nearly flat-lying strata, indicating that tilting of JS strata is negligible beyond 100 to 150 meters from the PLV-JS contact. In the upland northwest of the contact, fractured PLV strata with veins of secondary minerals generally do not present good opportunities to determine strike and dip. Where possible to measure, consistent northeasterly strikes with dips both to the northwest and southeast define an anticlinesyncline pair with NE-trending fold axes that parallel the PLV-JS contact. The relationships along the east edge of the Snake Creek block are evidence of a NE-trending fault along which PLV strata to the northwest were thrust over JS strata to the southeast. Similar to the curved fault at the northeast end of the Hungarian block, the fault at the east end of the Snake Creek block also has a curved geometry where it wraps around the block’s southeast corner and gradually changes direction by 55° to a west-northwesterly trend. Along the fault’s 121 �NE-trending section, its position and nature in the latest map (Fig. 11) do not differ much from what is shown on the Bruneau Creek bedrock geology map (Wright and Cornwall, 1954). However, the earlier map labels this fault the “Keweenaw fault,” whereas we interpret it to be part of a curved splay in the footwall of the main Keweenaw fault that lies to the north, which we will visit at the next stop. Stop 3-2: Main Keweenaw fault zone at Eister Creek and Falls Directions: Return 0.5 mi (0.8 km) on Iron Gate Road back to Gratiot Lake Road. Turn left and drive 0.2 mi (0.3 km) to East Gratiot Lake Road. Turn right and drive 1.0 mi (1.6 km) to a parking area along the road. [Lat: 47° 21.975'N | Lon: 88° 7.967'W] Eister Creek near the falls provides excellent exposures across the northern branch of the main Keweenaw fault zone where PLV strata in the hanging wall are juxtaposed against presumably younger PLV strata in the footwall (Figs. 11 and 12). Jacobsville strata that commonly occur in the footwall of the main fault zone are not present here, though they may constitute bedrock south of the southern branch of the main fault zone. Elsewhere nearby in the footwall, the JS unit has thicknesses ranging up to 130 m confirmed in outcrop and greater than 135 m in a water well northwest of Lake Gratiot. In general, the thickness of the JS unit in this area appears to be much less than to the southwest near Houghton, where the unit is known to be at least 785 m thick at the fault system and at least 1,100 m thick away from it to the southeast, but could be 2,000 to 3,000 m thick based on geophysical data. Figure 12: Cross-section along Fault Creek east of Eister Creek (Lizzadro-McPherson, 2023). Geologic unit codes: pb – Bohemia conglomerate; pgf – Gratiot flow; psc – Scales Creek flow. KFS = Keweenaw fault system. Walking north up Eister Creek from the parking area, scattered outcrops of PLV basaltic lava are first encountered at the mouth of the incised creek valley. About 80 meters into the narrow gorge 122 �where it becomes deeper, the basaltic rocks become more fractured and are locally sheared, brecciated, and gougy. Over the next 40-50 m upstream to the base of Eister Falls, fracture intensity increases and cataclastic shearing becomes more prominent until the rock is essentially a breccia cut through by shear zones. The area around the base of the falls is at the core of the main Keweenaw fault zone that is up to 45 meters wide along the creek, depending on how its northern and southern edges are defined. A rock wall perhaps 4 meters wide that projects from the eastern wall of the gorge consists of fault breccia and gouge that is better indurated that the surrounding material, which is also brecciated and gougy. This ridge is inclined steeply upstream, and is taken to define the orientation of the fault zone as striking 255° and dipping 76° N. Upstream from the projecting wall of breccia, the long gradual rise of Eister Falls exposes highly fractured basaltic rocks with some shear zones, but the rock mass is mostly intact and unlike the completely brecciated fault core. The fault-core relationships seen at Eister Creek are even better displayed along another creek located ~800 m east-northeast of here at a locality informally named “Fault creek” (Fig. 12). Stop 3-3: Nine Thirty Two Creek Directions: Return 1.0 mi (1.6 km) on East Gratiot Lake Road to Gratiot Lake Road. Turn right and drive 0.4 mi (0.6 km) to a narrow driveway on the right. Either turn into the driveway and park where possible or park along the right shoulder of the paved road. [Lat: 47° 21.711'N | Lon: 88° 8.730'W] THIS STOP IS ON PRIVATE PROPERTY. PERMISSION IS REQUIRED TO OBTAIN ACCESS. A walk of about 300 meters down the overgrown old road to Lake Gratiot leads to a hairpin turn in Nine Thirty Two Creek that marks the intersection of two fault trends (Fig. 11). The south branch of the main Keweenaw fault zone generally follows the creek valley upstream from the hairpin turn toward the west and, in the opposite direction, it follows an ENE-trending path to the north shore of Lake Gratiot. The thrust that defines the eastern edge of the Snake Creek block follows the western side of the creek downstream from the hairpin turn and terminates northward against the main Keweenaw fault zone. The Keweenaw fault zone upstream of the hairpin turn is manifested by basaltic rocks that are fractured, brecciated, and altered for a few hundred meters to the west. Its east-northeast path is marked by a linear depression that connects to the neighboring creek valley where altered basalt occurs in a cutbank. PLV strata north of the Keweenaw fault zone generally strike east-west and dip moderately north based on two nearby creek traverses. These hanging-wall strata are juxtaposed against presumably younger footwall PLV strata west of the creek’s hairpin turn and against footwall JS strata east of it (Fig. 11). This change in juxtaposition of geologic units along the Keweenaw fault zone occurs because the thrust fault intersecting it on the footwall side raises PLV strata on the west over JS strata to the east. South of the creek’s hairpin turn, JS strata crop out for 235 meters along the creek bed and valley walls and consist of yellowish to reddish, medium-grained, quartzose sandstone that is poorly 123 �laminated along with subordinate reddish, muddy-to-silty, clast-supported conglomerate that is poorly indurated. The exposed JS strata are generally subhorizontal but their dip increases to 69° SE with a strike of 58° within 30 meters of the hairpin turn. The rotation of JS strata downward to the southeast is consistent with components of reverse slip on the intersecting faults that contain the sandstone within their obtuse angle (Fig. 11). Above the sandstone outcrops to the west, the upland caprock is a layer of trachyandesite overlying a basaltic layer, whose contact dips about 10° NE. This stratal orientation may represent a plunging fold axis that is parallel to fold axes near the southeast corner of the Snake Creek block. Some important questions arise from the geometric relationships observed at the intersection of the two faults. For example, what happens to the main Keweenaw fault zone west and south of this fault intersection? Wright and Cornwall (1954) show the Keweenaw fault coming from Eister Creek as bending southwest to follow the PLV-JS fault contact discussed at Stop 3-1. They also show a splay of the Keweenaw fault continuing along Nine Thirty Two Creek as far as Gratiot Lake Road. Cannon and Nicholson (2001) continued this splay fault in a broad arc parallel to regional strike to a reconnection with the Keweenaw fault of Wright and Cornwall (1954) about 6 kilometers east of Mohawk. Based on new mapping and integration of DDH data, we have modified the trajectory of the fault extension proposed by Cannon and Nicholson (2001) and we propose that this is actually the main Keweenaw fault zone. The corollary to this interpretation is that the thrust fault along the east edge of the Snake Creek block is a footwall splay of the main fault zone. In other words, we think that the main Keweenaw fault zone is not always the one that juxtaposes PLV strata in the hanging wall against JS strata in the footwall. Area 4: Keweenaw Fault Zone with Footwall Splays, Lac La Belle to Bête Grise The three stops in Area 4 are along part of the KFS that changes direction from a 72° azimuth to nearly east-west (Fig. 3). The first stop is in the deformed hanging wall of the main Keweenaw fault zone (4-1) that runs along the northwest edge of Lac La Belle, past the southern base of Mt. Bohemia, and along most of the paved road to Bête Grise Bay east of here (Fig. 13). In a westerly direction, the main fault zone forms the northern boundary of the Deer Lake block, whose long dimension is again parallel to the KFS. The Deer Lake block is limited on the south by a footwall splay that diverges from the main fault zone north of Lake Gratiot and juxtaposes PLV basaltic flows on the north against nearly vertical JS strata at the Little Gratiot River (Lizzadro-McPherson, 2023; DeGraff, 1976). Based on diamond drill hole data, the eastern edge of this fault-bounded block is a connector fault where PLV strata are thrust eastward over JS strata, whereas the block’s western edge has a thin cover of Jacobsville Sandstone unconformably overlying weathered basaltic rock. Field relationships and magnetic data suggest that the footwall fault splay curves north and merges into the connector fault along the east edge of the Deer Lake block. The other two stops in Area 4 are at historic localities on the shore of Bête Grise Bay that Irving and Chamberlin (1885) investigated as part of their USGS Bulletin 23 titled “Observations on the junction between the eastern sandstone and the Keweenaw series on Keweenaw Point, Lake Superior”. The Bête Grise block is defined by the main Keweenaw fault zone on the north, by a footwall splay that diverges from it onshore and passes offshore while remaining visible in shallow 124 �Figure 13: Geologic map of Area 4 with the Keweenaw fault system and field trip stops 4-1 to 4-3. Teeth along reverse faults are on upthrust sides. 125 �water (4-2), and by an inferred thrust along the east edge of the block (4-3). The thrust is inferred from a NE-trending belt of intense fracturing and veining in altered basaltic rock that follows the shoreline parallel to the strike of the PLV-JS unconformity exposed to the east along the shore. Thrusting of hanging-wall PLV strata toward the southeast has tilted footwall PLV strata and JS strata above the unconformity 50° SE. Stop 4-1: Haven Falls Directions: Get back onto Gratiot Lake Road and drive 4.3 mi (6.9 km) back to US-41. Turn right and drive 5.7 mi (9.2 km) to Lac La Belle Road. Turn right and drive 4.2 mi (6.8 km) to a Yintersection. Veer right and then through the sharp right curve for 0.5 mi (0.8 km) to Haven Falls Park on the right. Enter the park and park. [Lat: 47° 22.913'N | 88° 1.716'W] The small but beautiful park at Haven Falls spans a terrace formed by a previous higher stand of Lake Superior that was about 10 meters above the current lake level. The cliff of felsic conglomerate over which Haven Creek flows was probably a shoreline cliff at the time of that higher lake level. This stop is in the proximal hanging wall of the main Keweenaw fault zone (Fig. 13) and its stratigraphy is exposed almost continuously from south of Lac La Belle Road, which is privately owned, to well upstream of the falls. About 650 meters east of here and south of the main fault zone, an east-directed thrust fault, cored by the Deer Lake #2 DDH, defines the east edge of the Deer Lake block (Fig. 13A). Slip on this thrust fault has pushed PLV basaltic lavas on the west up and over JS strata to the east (Lizzadro-McPherson, 2023) in a manner similar to what occurs at the east edge of the Snake Creek block. The Haven Falls stop is, therefore, analogous to the Eister Creek stop except that the fault core here is not as well exposed south of the paved road (cf. Figs. 12 and 14). The stratigraphic sequence in the hanging wall begins with a highly fractured, veined, and locally brecciated and sheared lava flow that crops out on both sides of the paved road, but please stay on the north side to respect the private property on the south side. This strongly deformed lava flow lies along a topographic step up from the lowland at the lakeshore to the old lake terrace, and it probably marks the northern edge of the core of the fault zone. Upstream from this first lava flow, a 12-m-wide belt of conglomerate is followed by a 30-m-wide belt of ophitic basalt that locally is highly fractured, veined, and sheared. The felsic cobble-pebble conglomerate at the main falls is largely clast-supported and has a massive appearance. It extends for a horizontal distance of 16 meters along the creek to well above the top of the falls, where it is brecciated along a fault zone that cuts slightly up section toward the east. A reliable formation strike of 245° and dip of 65° NW comes from a sandstone layer above the falls near the northern edge of the second conglomerate layer. The two conglomerate layers here are collectively known as the Lac La Belle conglomerate (Cornwall, 1954a) and have been tentatively correlated with the Baltic (#3) conglomerate near Houghton. Although there is considerable uncertainty about this correlation due to the distance involved, the Lac La Belle conglomerate lies well below the St. Louis (#6) conglomerate seen at many of the earlier stops of this field trip. This means that the Keweenaw fault system here cuts the PLV section at a significantly deeper level than near Calumet, Laurium, and Lake Linden. 126 �The stratigraphic units at the falls have been traced by detailed mapping along strike in both directions. About 350 meters east-northeast, the two conglomerate layers merge into one layer where the intervening basalt flow pinches out, but otherwise the layers can be traced continuously along strike. The only structural complications observed along strike are a few faults that cut at acute angles upward across layers toward the east and have small reverse offsets, similar to the fault at the top of Haven Falls. Detailed mapping in 2019-2020 did not find evidence of four transverse faults shown on the Delaware bedrock geology map as offsetting the Lac La Belle conglomerate, neither in terms of offsets nor enhanced fracturing (LizzadroMcPherson, 2023). In fact, the kinematics of slip along the main Keweenaw fault zone would argue for a system of smaller subparallel faults rather than a series of faults normal to the main fault zone. Figure 14: Cross-section along Haven Creek at the falls (Lizzadro-McPherson, 2023). Geologic unit codes: pb – Bohemia conglomerate; psc – Scales Creek flow. KFS = Keweenaw fault system. THE NEXT TWO STOPS ARE ALONG APACHE LANE, WHICH IS PRIVATELY OWNED AS ARE ALL PROPERTIES ALONG IT. PERMISSION IS REQUIRED TO OBTAIN ACCESS. Stop 4-2: Bête Grise Shoreline, Irving & Chamberlin Historic Site Directions: Exit Haven Falls Park and drive east back to the stop sign at the Y-intersection. Turn sharply right onto Bête Grise Road and drive 3.2 mi (5.2 km) to Apache Lane on the left. [Geology Note: at 1.9 mi / 3.1 km along BG Road, an outcrop to the north is the location of the dated calcite vein in the fault zone.] Turn left, drive 0.5 mi (0.8 km), and then park along the right side of the road. [Lat: 47° 23.345'N | Lon: 87° 56.905'W] We will walk down a moderately steep slope to the shore, where Irving and Chamberlin (1885) used a “force of miners” to strip the shoreline bare in order to better expose the contact between PLV layers on the north and JS strata to the south (Figs. 13B and 15). This is probably not a field 127 �practice we could get away with today even if the site was not privately owned. Depending on the lake level, we may be able to see the contact between highly fractured, veined, and partly altered basaltic rock at the base of the shoreline scarp and JS strata that strike 105° and dip 55° S. An aerial image of the shoreline at this stop shows a sharply defined line with an azimuth of 105° that separates uniformly dark-toned PLV basaltic rocks on the lake bottom and along the shore to the east from variably lighter-toned layers of JS strata to the south and west (Fig. 16). This is the fault line that may be observed onshore or can be closely constrained by nearby outcrops. The aerial image shows many small faults as darker lines and narrow zones that cut JS strata at angles approaching 90° and offset strata by less than a meter or two. Near the fault line, splay faults break the JS unit into blocks up to 30 meters long parallel to the fault line and 4 meters wide. Figure 15: Cross-sectional view of a segment of the KFS exposed by excavation along the Bête Grise Bay shoreline. Fault strike = 100°, dip = 55° S (Irving and Chamberlin, 1885). Circle with black dot indicates movement toward viewer; circle with cross indicates movement away. Onshore north of the fault line, PLV strata are highly fractured, veined, and sheared for at least 55 meters along the shoreline to the east, which is about 25 meters perpendicular to the fault. Basaltic outcrops along Apache Lane and to the north do not exhibit such deformation. Following the shore to the west and south, JS strata change in terms of both facies and structural orientation. The basal part of the section consists of reddish, thin-bedded, siltstone and mudstone with minor fine-grained sandstone (6-7 m true thickness), followed upward by a lighter-toned package of silty to fine-grained sandstone with minor silty pebble conglomerate that forms a resistant ridge (~2 m), and then an interval of reddish siltstone and muddy conglomerate (~4 m). 128 �From this point southward and up section, JS strata tend to be sand-prone but alternate between: (1) pinkish to orange, fine to medium grained, quartzose sandstone; (2) red, thinly bedded, siltstone to mudstone; and (3) muddy to silty, poorly indurated, conglomerate layers. In summary, the JS unit tends to clean upward from a silt- and mud-dominated basal section with conglomeratic units near the fault to a quartzose sand-prone section away from the fault. Along with these stratigraphic changes, the dip of JS strata decrease from 55° S near the fault to about 20° S at a perpendicular distance of 60 meters from the fault, and presumably becomes subhorizontal not much further to the south. Figure 16: Aerial image of Stop 4-2 showing the south boundary fault of the Bête Grise block. The fault is interpreted to have dominant strike slip. Stratigraphic up in the JS is to the south. It is clear that movement along the fault has juxtaposed older PLV strata to the north against younger JS strata to the south and that the north side has a component of upward movement, as noted elsewhere. However, what is the nature of this fault? The work by Irving and Chamberlin (1885) and their team exposed a fault surface that dips about 55° S, essentially parallel to adjacent JS strata (Fig. 15). Based on textbook definitions, this would be a normal fault with younger JS strata in the hanging wall to the south above older PLV strata in the footwall to the north. Something about this interpretation seems paradoxical, however, because at previous stops 129 �the main fault zone dipped north or northwest and also because the tectonic setting of the Keweenaw fault system was compressional. Part of the paradox results from textbook definitions of reverse and normal faults that are based on idealized planar fault surfaces that have mostly dip slip. In the case of curved or corrugated fault surfaces with mostly strike slip, the textbook nomenclature for dip-slip faults may lead to confusion. Depending on the portion of a corrugated fault surface that is exposed, a mostly strike-slip fault with a smaller dip-slip component may exhibit an apparent normal or reverse component of dip-slip even though the actual dip-slip component is the same everywhere along the fault. The Keweenaw fault system in this area trends nearly east-west and is dominated by right-lateral strike slip with lesser north-side-up dip slip (2:1 ratio of strike-to-dip slip), and we infer that motion on this fault is dominantly right-lateral strike slip. The fault forms the southern edge of a fault-bounded block (Fig. 13B) whose east side will be visited next. Stop 4-3: Bête Grise Shore, PLV-JS Unconformity and Fault Directions: Continue driving east on Apache Lane for 0.4 mi (0.6 km) to the end of the road. Park where space allows. [Lat: 47° 23.437'N | Lon: 87° 56.350'W] We will walk 100 meters east and 50 meters south to access the shoreline, where geologists traveling along the coast by boat in the mid-1800s described impressive layers of sandstone on the rocky bottom of Lake Superior (Figs. 13B and 17). This area was examined later by none other than Irving and Chamberlin (1885) and then by Cornwall (1954b). The aerial image of the shoreline and offshore region reveals what the early explorers reported, a set of well-defined parallel layers that curve sharply from NE-trending layers on the western side to EW-trending layers along the shore to the east. The shoreline stop is at the western edge of the JS strata where an unconformity between PLV and JS strata is tilted about 50° SE. Depending on water level and shoreline erosion, we may be able to see the unconformity between saprolitized PLV basaltic lava to the northwest and JS strata to the southeast. If the saprolite is exposed, please do not disturb it by digging, picking at it, or walking on it. The basaltic protolith has been completely converted to clay minerals and still retains its original textures, including whitish veins that are approximately normal to the tilted unconformity. Northwest of the unconformity, i.e. deeper below the paleosurface, PLV basaltic rocks do not exhibit such alteration. Saprolitic basaltic rock that retains original textures, such as ophitic texture, has been observed elsewhere in the area where the PLV-JS unconformity is inferred, such as the east end of the Deer Lake block (Lizzadro-McPherson, 2023; DeGraff, 1976). East of the unconformity along the shore, a recessive basal part of the JS section consists of thinly bedded, reddish siltstone and mudstone with minor interbedded fine-grained sandstone. The recessive strata here strike northeast and dip moderately southeast over a distance of 25 meters to the first small resistant sandstone layer (1-2 m thick). Next in the section is another recessive interval (6-7 m wide) of reddish siltstone with interbeds of poorly indurated muddy conglomerate, followed by a larger ridge of resistant sandstone that begins a sequence of 130 �alternating resistant sandstone beds, recessive reddish siltstone, and reddish poorly indurated conglomerate. Thus, the basal JS section here is very similar to the section at Stop 4-2. The main difference is in their structural attitude, which differs in strike by 55-60°. At Stop 4-2, the dip direction of JS strata is 195° and away from the inferred strike-slip fault on the south side of the Bête Grise block, whereas the dip azimuth of JS strata on the east side of the fault block is 138° and away from the tilted PLV-JS unconformity (Fig. 13B). We infer that southeast tilting of the PLV-JS unconformity resulted from southeast thrusting along the east edge of the Bête Grise fault block, based in part on intense fracturing observed along the shoreline northwest of the unconformity at stop 4-3. Figure 17: Aerial image of Stop 4-3 showing the north and east boundary faults of the Bête Grise block. The east boundary fault is interpreted to have dominant dip slip with west side thrust eastward, whereas the north boundary fault is interpreted to have dominant strike slip. Stratigraphic up in the JS is to the southeast and south. If time permits, we will ascend the shoreline scarp to the flat bench above and walk another 100 meters to reach the eastern edge of JS outcrop along the shore. Here, JS strata strike nearly eastwest and are vertical to slightly overturned to the south, forming narrow ridges and eroded furrows and clefts due to differential erosion of the resistant and recessive layers (Fig. 17). Near the eastern edge of JS outcrop, vertical JS strata are flanked on the north by PLV strata that begin 131 �with a felsic conglomerate and continue upslope with PLV basaltic lavas. The contact between PLV strata on the north and vertical JS strata to the south is a major fault that intersects the shoreline, where a fault breccia is exposed several meters east of the last onshore JS outcrop. From this point, the fault turns eastward and runs along the shoreline for 200 meters before continuing offshore and splitting into two branches. This is the last stop and we hope that you had a good experience that will help you to understand other fault systems. Thank you for your participation! Acknowledgements We thank the following M.S. graduates and their assistants, whose field mapping was funded by U.S. Geological Survey EDMAP projects G17AC00115, G19AC00140, and G21AC10681: Colin Tyrrell (M.S.), Sophie Mueller (M.S.), Nolan Gamet (M.S.), Graham Hubbard, Ian Gannon, Ginny Hemmila, Gabe Ahrendt, Jack Hawes, Braxton Murphy, Breeanne Heusdens, and Dillon Breen. We also thank many who have expressed interest in this work and have provided helpful comments. References Cited Aho, G.D., 1969, A Reflection Seismic Investigation of Thickness and Structure of the Jacobsville Sandstone, Keweenaw Peninsula, Michigan: Michigan Technological University, MS thesis, 104 p. Bacon, L.O., 1966, Geological structure east and south of the Keweenaw fault on the basis of geophysical evidence: in Steinhart, J.S. and Smith, T.J. (eds), The Earth Beneath the Continents: Am. Geophys. Union, Geophysical Monograph 10, p. 42-55. 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, 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., 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. Broderick, T.M., 1931, Fissure vein and lode relations in Michigan copper deposits: Econ. 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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. 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., 2000, Geologic Map of the Keweenaw Peninsula and Adjacent Area, Michigan to accompany Map I-2696: U.S.G.S Pamphlet, 7 p. 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., 1954a, Bedrock Geology of the Delaware Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-51, scale 1:24,000. Cornwall, H.R. and Wright, J.C., 1956a, Geologic Map of the Hancock Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Mineral Investigations Field Studies Map MF-46, scale 1:24,000. Cornwall, H.R. and Wright, J.C., 1956b, Geologic Map of the Laurium Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Mineral Investigations Field Studies Map MF-47, scale 1:24,000. Craddock, J.P., Pearson, A., McGovern, M., Kropf, E., Moshoian, A., and Donnelly, K., 1997, Postextension shortening strains preserved in calcites of the Midcontinent Rift: Geological Society of America, Special Paper 312, p. 115-126. https://doi.org/10.1130/0-8137-2312-4.115 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 Mesoproterozoic Midcontinent Rift, Lake Superior region, USA: Journal of Geology, v. 121, no. 1, p. 57-73. 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. DeGraff, J.M., 1976, Structural and Age Relationships of Rocks Associated with the Lac La Belle Magnetic Anomaly, Keweenaw County, Michigan: Michigan Technological University, M.S. thesis, 130 p. 133 �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, p. 449–466. https://doi.org/10.1130/B36186.1 Fossen, H., Schultz, R.A., Shipton, Z.K., and Mair, K., 2007, Deformation bands in sandstone: a review: Journal of the Geological Society, v. 164, p. 755-769. Gamet, N.G., 2023, Structural Analysis and Interpretation of Deformation along the Keweenaw Fault System from Lake Linden to Mohawk, Michigan: Michigan Technological University, M.S. thesis, 122 p. Hamblin, W.K., 1958, The Cambrian Sandstones of Northern Michigan: Michigan Department of Conservation, Geological Survey Division, Lansing, MI, Publication 51, 55 p. 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, no. 2, p. 303-310. Hodgin, E.B., Swanson-Hysell, N.L., DeGraff, J.M., Kylander-Clark, A.R.C., Schmitz, M.D., Turner, A.C., Zhang, Y., and Stolper, D.A., 2022, Final inversion of the Midcontinent Rift during the Rigolet Phase of the Grenvillian orogeny: Geology, v. 50, no. 5, p. 547-551. Hubbard, L.L., 1898, Keweenaw Point with particular reference to the felsites and their associated rocks: Geological Survey of Michigan, v. 6, part 2, 155 p. Irving, R.D. and Chamberlin, T.C., 1885, Observations on the junction between the eastern sandstone and the Keweenaw series on Keweenaw Point, Lake Superior: U.S. Government Printing Office, Washington, D.C., U.S. Geological Survey, Bulletin No. 23, 58 p. Kalliokoski, J., 1982, Jacobsville Sandstone: in Wold, R.J. and Hinze, W.J. (eds.), The Geology and Tectonics of the Lake Superior Basin, Geological Society of America Memoir, No. 156, p. 147-155. Keweenaw National Historical Park, 2016, Calumet & Hecla Records – 00019/004.02.01.03-007 Microfiche Drill Core Log Library: Calumet, Michigan, National Park Service, U.S. Department of the Interior, on microfiche. Langfield, K.M., 2024, Slip Kinematics and Structural Analysis of the Keweenaw Fault System from Lake Linden to Hancock, Michigan: Michigan Technological University, M.S. thesis, in preparation. Lizzadro-McPherson, D.J., 2023, Structural Analysis and Slip Kinematics of the Keweenaw Fault System between Bête Grise Bay and Gratiot Lake, Keweenaw County, Michigan: Michigan Technological University, M.S. thesis, 140 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, no. 4, p. 1271-1282. 134 �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. Mueller, S.A., 2021, Structural Analysis and Interpretation of Deformation Along the Keweenaw Fault System West of Lake Gratiot, Keweenaw County, Michigan: Michigan Technological University, M.S. thesis, 69 p. Nicholson, S.W., 1992, Geochemistry, Petrography, and Volcanology of Rhyolites of the Portage Lake Volcanics, Keweenaw Peninsula, Michigan: U.S.G.S. Bulletin 1970, Chapter B, p. B1-B57. Robertson, J. M, 1975, Geology and mineralogy of some copper sulfide deposits near Mount Bohemia, Keweenaw County, Michigan: Economic Geology 70 (7) 1202-1224. 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. Stoiber, R.E. and Davidson, E.S., 1959, Amygdule mineral zoning in the Portage Lake Lava Series, Michigan copper district, parts I and II: Economic Geology, v. 54, p. 1250–1277 and 1444-1460. doi:10.2113/gsecongeo.54.7.1250/ 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, v. 131, nos. 5-6, p. 913-940. Tyrrell, C.W., 2019, Keweenaw Fault Geometry and Slip Kinematics – Bête Grise Bay, Keweenaw Peninsula, Michigan: Michigan Technological University, M.S. thesis, 30 p. Wadsworth, M.E., 1884, On the relation of the Keweenawan series to the eastern sandstone in the vicinity of Torch Lake, Michigan: Boston Soc. Natural Hist. Proc., v. 23, p. 172-180. White, W.S., 1952, Imbrication and initial dip in a Keweenawan conglomerate bed, J. Sediment. Petrol., v. 22, pp. 189-199. White, W.S., 1956, Geologic Map of the Chassell Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Mineral Investigations Field Studies Map MF-43, scale 1:24,000. White, W.S., 1968, The native copper deposits of northern Michigan: in Ridge, J.D., ed., Ore Deposits of the United States, 1933-1967 (Graton-Sales Vol.): New York, Am. Inst. Mining Metall. Petroleum Engineers, v. 1, p. 301-325. White, W.S., 1985, “Unpublished diamond drillhole core logs”: U.S. Geological Survey, Field Records Collection, Boxes 282, 287-290. Woodruff, L.G., Schulz, K.J., Nicholson, S.W., and Dicken, C.L., 2020, Mineral deposits of the Mesoproterozoic Midcontinent Rift system in the Lake Superior region – A space and time classification: Ore Geology Reviews, v. 126, p. 1-21. 135 �136 �Field Trip 5 Geology and History of a Native Copper Mine: Adventure Mine, Ontonagon County, Michigan Theodore J. Bornhorst Department of Geological and Mining Engineering and Sciences and A.E. Seaman Mineral Museum, Michigan Technological University, 1404 E. Sharon Avenue, Houghton, MI 49931 Matt Portfleet Department of Geological and Mining Engineering and Sciences, Michigan Technological University, 1400 Townsend Drive, Houghton, MI 49931 [Latitude: 46.777224; Longitude: -89.081906] 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). Stay on M-38 towards Ontonagon. Follow signs to Adventure Mine. The Adventure Mine is privately owned and operated as a publicly available mine tour Introduction The historic Adventure Mine is part of the Greenland-Mass subdistrict of the Keweenaw Peninsula native copper district of the western Upper Peninsula of Michigan (Figures 1 and 2). The Adventure Mining Company began mining native copper in 1850. It permanently ceased mining in 1917 at the time when most small mines of the subdistrict ceased mining operations. From the 1980s to today, the mining activities at the Adventure Mine have yielded specimens of massive native copper and copper crystals for purchase by tourists and mineral collectors. The Adventure Mine is a Keweenaw Heritage Site of the Keweenaw National Historical Park. It is located in the Greenland-Mass subdistrict about 40 km southwest of the Baltic Mine in Painesdale, the southernmost major native copper mine in the district (Figure 2). The subdistrict yielded about 85 million lbs (39 million kg) of refined copper at grades ranging from 0.5 to 1.25 percent (Butler and Burbank, 1929; Weege and Pollock, 1971). The Adventure Mine was the second largest producer in the subdistrict, yielding about 11 million lbs. (5 million kg) of refined copper from the tops of five different basaltic lava flows (Butler and Burbank, 1929). Despite the relatively low production of copper from the Greenland-Mass subdistrict (~0.8 % of the total district production of ~ 5 billion kg (11 billion lbs), the geologic characteristics of the Greenland-Mass subdistrict deposits are typical of the native copper deposits elsewhere in the Keweenaw Peninsula (e.g., Stoiber and Davidson, 1959; Butler and Burbank, 1929; etc.). Since 137 �mining in the Keweenaw Peninsula native copper district ceased in 1968, underground access to observe or study the native copper deposits has been limited. Currently there is access to only two mines in the main area of the district (Quincy and Delaware) and two mines in the Greenland-Mass subdistrict (Adventure and Caledonia). These mines are operated as underground experiences for tourists, except for Caledonia. This field trip guide relies on existing publications by Bornhorst and Whiteman (1995); Bornhorst et al. (2013); Bornhorst and Barron (2011); and Butler and Burbank (1929). The overview of the geologic and part of the human history is summarized from Bornhorst and Lankton (2009), Bornhorst and Mathur (2016), and Bodden et al. (2022). Brandon Erickson prepared a brief history of the Adventure Mine which is included in this guide. Figure 1: Simplified bedrock geology of the Mesoproterozoic Midcontinent Rift around Lake Superior. Modified after Bornhorst et al., 2013) Overview of Geologic History The largest known accumulation of native copper in the world is the Keweenaw Peninsula native copper district (Figure 2). In comparison to other copper mining districts where major copper ore 138 �minerals are sulfides, nearly all of the copper in the district occurs as native copper. About 5 billion kg (11 billion lbs) of refined copper were extracted from about 380 million tons of ore between 1845 and 1968 via underground mines (Weege and Pollock, 1971). Small quantities of native silver occur with the native copper. An estimate using incomplete records suggests the amount of silver was between .05 to .5 oz per ton of ore. Native copper and silver were coprecipitated. The native copper deposits are hosted by the Mesoproterozoic Midcontinent Rift (MCR) (Figures 1 and 2). More than 25 km of volcanic rocks and 8 km of clastic sedimentary rocks fill the center of the MCR (Cannon et al., 1989 and 1993). These rift-filling rocks were emplaced between about 1.15 to ~1 Ga (Cannon et al., 1989; Davis and Paces, 1990; Kulakov et al., 2018; Heaman et al., 2007). Figure 2: Simplified geologic map of the Keweenaw Peninsula and vicinity modified from Bornhorst et al. (2013). At the White Pine mine “Keweenawan” native copper cuts across diagenetic the shale-hosted chalcocite deposit. Early eruptions of MCR basaltic lava flows were scattered on a broad land area above a developing mantle plume. These early eruptions were followed by many eruptions from fissure vents concentrated in the center part of the MCR (now buried under the center of Lake Superior). These eruptions were dominated by fissure volcanoes along linear faults. The Portage Lake Volcanics and Porcupine Volcanics are rift-filling basaltic volcanic rocks erupted about 1.1 billion years ago 139 �during the active rifting of the MCR (Figure 3). The MCR was bounded on the edges by down dropped normal faults resulting in the MCR being a faulted basin. The basin progressively dropped down by stretching and by magma erupted at the surface during active rifting. Eruptions of basalts of the Portage Lake Volcanics were on land surface (subaerial). Subaerially erupted basalt lava flows have either a vesicular or brecciated and vesicular top (pahoehoe or aa lava flow). Mineral filled vesicles are termed amygdules and flow tops dominated by vesicles are termed amygdaloids and those dominated by breccia clasts of amygdaloidal basalt are termed fragmental amygdaloids. The top of a subaerial lava flow is underlain by a massive (relatively vesicle-free) basalt. Massive basalt flow interior in thinner flows is fine grained and in thicker flows it is coarse-grained (ophitic). Thin flows can be vesicular throughout with vesicles more abundant at the top. The typical flow is 10 to 20 m thick. Eruptions of basaltic lavas were cyclical and during eruptive hiatuses minor gravel and sand were deposited on top and infiltrated in the tops of occasional lava flows. These clastic sedimentary layers are overlain by basalt lava flows. A desert environment 1.1 billion years ago resulted in red coloration of these clastic sedimentary rocks. Figure 3: Stratigraphic column for the Adventure Mine region with approximate ages. Active rifting and basaltic volcanic activity ended over a short period of time, but the rift basin continued to passively sag and was progressively filled with clastic sedimentary rocks from the Copper Harbor Conglomerate to the Freda Sandstone (Figure 3). In the Adventure Mine region, the 140 �rift-filling volcanic rocks (Porcupine Volcanics and Portage Lake Volcanics, Figure 3) were first covered by red-colored gravels and sands (Copper Harbor Conglomerate. Overlying the Copper Harbor Formation are black- to gray-colored muds and silts (Nonesuch Formation). Lastly the rift was filled with a thick section of red-colored fine sandstones (Freda Sandstone). Today, the rocks of the Keweenaw Peninsula and Adventure Mine region span the edge of the MCR, and consist of a thick section of rift-filling subaerial basaltic lava flows overlain by a thick section of rift-filling clastic sedimentary rocks (Figures 1, 2, and 3). 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 resulted in reverse and thrust faults as well as folding and fracturing of rift-filling volcanic and clastic sedimentary rocks. Native copper and related minerals were emplaced during this regional compressional event (Bornhorst, 1997). From the beginning of regional compression at about 1.06 to 1 billion years ago the Jacobsville Sandstone was deposited in a rift-flanking basin until about 1 billion years ago (Figure 2). There were no recorded geologic events by rocks of Michigan’s Upper Peninsula from about 1.0 billion years ago to 500 million years ago. Erosion likely exposed the native copper deposits at the surface and downward percolating oxidizing groundwaters had access to alter the native copper (Bornhorst and Robinson, 2004). The MCR rocks were buried by Phanerozoic rocks deposited from 500 to 175 million years ago (Catacosinos and others 2001). The native copper deposits of the Keweenaw Peninsula were again at the surface after erosion of overlying rocks by Pleistocene glaciers over the last 2.5 million years. Native Copper Deposits of the Keweenaw Peninsula The cumulative pre-mining geologic copper resource of the Keweenaw Peninsula native copper district totaled about 9 billion kg of copper (20 billion lbs; Bornhorst and Barron, 2011). About ½ of the geologic resource was recovered. Speculative concentrations of copper in rocks are even greater. These concentrations have not been mined for various reasons such as too low of grade or too deep below the surface. Permeable and porous primary geologic settings that have sufficient open spaces that were sufficiently connected with each other facilitated the movement of ore-forming hot waters (hydrothermal fluids) from which native copper and other minerals were precipitated. Compression of the MCR strata integrated the primary permeability and porosity of the hydrothermal plumbing system with compression generated faults/fractures (Bornhorst, 1997). The absolute age of the main-stage of hydrothermal activity coincides with the age of the regional compressional event at about 1.06 to 1.04 billion years ago (Bornhorst et al., 1988). The permeable and porous tops of amygdaloidals and fragmental amygdaloids hosted ~58.5% of produced native copper. Horizons of conglomerate and sandstone between lava flows, which are also permeable and porous, hosted ~39.5% of produced native copper. Native copper ore bodies are "sandwiched" on the bottom side by the massive basalt interior of the flow whose top hosts the native copper ore body. On the top side the ore bodies are sandwiched by massive basalt of the overlying lava flow. The sandwiched ore-bodies are geometrically approximately tabular 141 �(called lode) with a thickness between 3 and 5 m and the same orientation as surrounding host lava flows. The typical lode extends down-dip 1.5 to 2.6 km and has a lateral extent of 1.5 to 11 km (Butler and Burbank, 1929; White, 1968). Open spaces in amygdaloidal lava flow tops (vesicles) and sandstones/conglomerates are typically up to a cm across. They are typically filled dominantly by gangue minerals and lesser native copper. Less frequently the entire open space was filled with masses of native copper. Open spaces between breccia fragments in the top of a lava flow (fragmental amygdaloid) or between clasts in conglomerate will tend to have larger masses of native copper and can weigh up to several lbs, to tens of lbs to hundreds of lbs. and rarely weighing tons. A minor amount of the total produced native copper, ~ 2%, was from sub vertical tabular open spaces (veins when filled with minerals) that follow faults and fissures that perpendicularly cut across the volcanic-dominated strata. Ore-forming hydrothermal fluids readily moved along faults and fissures since they have a relatively large amount of interconnected open-space; these ore-bodies are also tabular lodes. Since the size of open space is large the corresponding size of masses of native copper can also be large weighing multiple tons with the largest masses being several hundred tons. Native copper is closely associated with about 22 common and many more uncommon minerals (Bodden et al., 2022; Butler and Burbank, 1929; White, 1968). These minerals fill the same open spaces along with and instead of native copper. The suite of minerals is similar to those found where rocks have undergone very low to low grade burial metamorphism at less than about < 300OC (Bodden et al., 2022). Thermal modeling suggests that peak burial metamorphic conditions at depth were between 400 to 500oC (Woodruff, 1995). Burial metamorphic processes resulted in ore-forming hydrothermal fluids carrying copper leached from the tops of buried riftfilling basalt lava flows. Batches of metamorphogenic-dominated hydrothermal fluid generated over time were similar to one another. The rift-filling volcanic rocks were very low in sulfur when they erupted, and the little contained sulfur degassed into the atmosphere. During subsequent deposition of rift-filling clastic sedimentary rocks there was an incursion of seawater into the rift for a significant amount of time. This resulted in seawater deeply penetrating into the underlying rift-filling volcanic rocks (Figure 4A). During burial, the rift-filling volcanic and clastic sedimentary rocks were progressively heated and during initial heating the seawater evolved to be depleted in sulfur similar to expelled modern sea floor hydrothermal fluids (Figure 4B). Continued heating during burial resulted in burial metamorphic-dominated hydrothermal fluids with copper leached from the rift-filling volcanic rocks. These fluids were well mixed with the evolved seawater resulting in hybrid metamorphic-dominated ore-forming fluids (Figure 4C). These main-stage hybrid metamorphic-dominated ore-forming hydrothermal fluids moved upwards from the source zone through the same very sulfur poor strata as in the source rocks (Figure 4D). As they moved upwards they cooled, interacted with host rocks, and in the relatively shallow zone of precipitation they variably mixed with sulfur-poor reduced meteoric water (Figure 4D; Bodden et. al, 2022). These processes resulted in precipitation of native copper and main-stage hydrothermal minerals. Higher temperature main-stage mineral assemblages are spatially associated with the area of native copper deposits where the thermal anomaly was greatest 142 �because of focused hydrothermal fluids. Within the native copper district, the suite of main-stage minerals that is associated with native copper and is followed by late-stage minerals precipitated at lower temperature than the main-stage. Figure 4: Cartoon cross sections showing conceptual genetic model of the native copper deposits of the Keweenaw Peninsula formed at about 1060 to 1040 million years ago. Modified from Bodden et al. (2022). A. Marine incursions and seawater penetration during deposition of volcanic and sedimentary rocks in MCR. B. Area prior to burial metamorphism with sulfur depleted evolved seawater providing salinity for ore-forming fluids. C. Burial metamorphic fluids mixing with evolved seawater produce hybrid ore-forming fluids. D. Precipitation of main-stage minerals, including native copper, as a result of mixing of ore-forming fluids with meteoric water, decreasing temperature, and water-rock reactions. 143 �Figure 5: Bedrock geologic map and cross sections of the Greenland-Mass subdistrict of the Keweenaw Peninsula Native Copper District. Modified from Whitlow (1974). 144 �The Evergreen Succession Butler and Burbank (1929) recognized the Evergreen lava flow and a succeeding number of lava flows of the Portage Lake Volcanics as having distinctive lithologies and hosting the native copper deposits in the Greenland-Mass subdistrict (Figure 3 and 5). These are informally termed the Evergreen Succession (Figure 3). The Evergreen Succession is stratigraphically about 150 m (500 ft) above the Bohemia (No. 8) conglomerate (Butler and Burbank, 1929). The Evergreen Succession basaltic lava flows are slightly more intermediate in composition than other lava flows within the PLV (Butler and Burbank, 1929). They are characterized by porphyritic or glomerporphyritic texture although thicker flows are ophitic. It is difficult to correlate individual lava flow with one another except in developed areas for mining of native copper where individual lava flow can be traced along strike from mine to mine. In the Greenland-Mass subdistrict most of the native copper was produced from the tops of lava flows of the Evergreen Succession. Those flow tops hosting native copper are generally fragmental amygdaloidal lodes with the best areas for native copper being where the flow top is thicker. Thin amygdaloidal only flow tops or those with areas of massive basalt mixed in the flow top are typically lower grade. The Evergreen Succession is at a similar stratigraphic position as those lava flows developed at the Isle Royale Mine to the north of the subdistrict in the Houghton area of the main district (Butler and Burbank, 1929). The Evergreen flows were also developed for native copper in the Winona area in the middle between the main district and the subdistrict. The individual copper-rich lava flows within the Evergreen succession were each informally named (Figure 3). The Evergreen flow is a 3 to 15 m thick plagioclase porphyritic lava flow. The Ogima flow is a 30 to 43 m thick slightly plagioclase glomerophyritic basalt lava flow. The Butler flow is a 15 to 27 m thick plagioclase glomerporphyritic basalt lava flow. The Mass and Merchant flows are up to about 25 m thick. The South Knowlton flow is up to 15 m and is a plagioclase glomeroporphyritic basalt. At the top of the Evergreen succession is the Knowlton flow which is a 9 to 21 m thick plagioclase glomeroporphyritic basalt. Between the Butler and Knowlton flows there are a number of thin flows of plagioclase glomeroporphryitic basalt with total thickness of 75 to 90 m thick (Calumet and Hecla, 1958). The tops of the Evergreen lava flows were productive over a strike length of about 5 km. Native copper mined from the Evergreen Succession was extracted from many different mines with some of them connecting with others. The Butler flow top yielded the most copper followed by the Evergreen and Knowlton flow tops which also yielded significant amounts of copper. Vesicle- and inter-fragment void-fillings consist of quartz, calcite, K-feldspar, epidote, prehnite, pumpellyite, and chlorite (Table 1). Less abundant main-stage minerals are native copper, native silver, and datolite. Laumontite and adularia are common late-stage minerals. 145 �Table 1: Percent amygdule-filling minerals estimated from rock piles adjacent to mines/shafts of the Greenland-mass subdistrict. Unpublished data by Stoiber and Davidson 1959). % Amygdule-Filling Mineral Quartz Calcite Mine/Shaft Adventure #1 Adventure #2 Adventure #3 Adventure #4 National #2 Old Mass Mass C Mass 1 &2 Mass B Mass A Michigan Michigan Flintsteel #1 Flintsteel #2 Butler Knowlton Red KFeldspar Epidote Prehnite Pumpellyite Chlorite 52 5 6 26 2 9 tr 30 20 5 40 0 1 4 11 13 30 37 5 1 3 36 22 17 22 55 19 45 63 21 18 23 20 30 27 17 21 10 23 22 16 27 30 52 45 45 38 0 3 40 45 0 40 trace trace 35 18 18 20 15 2 5 15 19 22 18 35 5 3 5 5 5 6 24 53 4 0 0 0 trace 5 10 0 0 0 8 10 0 2 3 trace trace 3 0 0 trace 0 10 0 1 0 1 1 trace 1 1 0 0 7 9 0 8 The Evergreen Succession in the Greenland-Mass subdistrict dips about 45o NW and forms a local broad open anticline (Fig. 6). The largest mine, the Mass Mine, occurs near the maximum bend in this anticline. Most faults have displacement of < 1 m while those faults with significant vertical displacement are uncommon. There are multiple veins in tension fractures in the area of maximum bend that cut perpendicular across the lava flows (Butler and Burbank, 1929). There are some veins that are parallel to the 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-forming fluid into the top of the lava flow. The Adventure Mine The Adventure Mine was very small, producing only about 5 million kg (11 million lbs) of refined copper, in context of all mines in the district which produced about 5 billion kg (11 billion lbs. Most of the production of native copper from the Adventure Mine came from the top of the Knowlton basalt lava flow (Knowlton lode; Figure 6 and 7). There was also significant 146 �production from the Butler lode and minor production from the Evergreen, Ogima, and Merchant lodes (Figure 6 and 7). Figure 6: Historic 1902 sketch cross section showing the lodes of the Evergreen succession at the Adventure Mine. Adits perpendicular to strike of the lava flows are shown and drifts parallel to the strike are indicated by black squares. The field excursion utilizes the adit near the No. 1 shaft. Figure 7: Longitudinal sections (parallel to strike) of the Evergreen succession native copper lodes showing underground workings (openings) for the Adventure Mine shafts #1 to #4. Section from Butler and Burbank (1929). 147 �The Knowlton was the focus of native copper mining at the Adventure Mine. The Knowlton lava flow top is a fragmental amygdaloid. In the subdistrict, the Knowlton flow top was developed for about 3000 m along strike and to a maximum depth of about 375 m. At the nearby Mass Mine (Fig 5), the Bulter lava flow top was the principal focus of native copper mining. It was the second focus of mining at the Adventure Mine. In the subdistrict the Butler lava flow top was developed for about 2000 m along strike and to a maximum depth of 300 m down dip. The most abundant secondary minerals in the Butler are quartz and calcite with slightly lesser amounts of K-feldspar and epidote (Table 1). Prehnite and pumpellyite are usually much less abundant and chlorite is present in amounts < 1 %. The Butler contains a high number of veins. Usually, the veins 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). 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 strikeparallel fractures and veins which were likely feeders of hydrothermal fluids (Bornhorst et al, 2013). At the Adventure Mine, on average the most abundant main-stage minerals filling amygdules and spaces between fragments is quartz which is closely followed by epidote and then calcite and red K-feldspar (Table 1). There are lesser amounts of prehnite, pumpellyite, and chlorite. Native copper is present in small amounts with average grades of between 0.5 to 1.25 % copper with native copper associated epidote, quartz, and calcite. Native silver and datolite are present in much lesser amounts. Least abundant are the late-stage hydrothermal minerals precipitated after native copper that occur in open space fillings as coatings on earlier formed minerals; late-stage minerals include calcite, laumontite, and adularia and in cross cutting fractures and veins. Alteration of hydrothermal mineral is most obvious for native copper. 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 likely formed by downward-percolating groundwater when the native copper deposits were sufficiently near the surface (supergene alteration) as they are today. In addition to tenorite and cuprite, there are occasional copper carbonate minerals (such as malachite), brochantite (hydrated Cu sulfate) and atacamite (hydrated Cu chloride). These are likely to be supergene in origin. At least one mineral, gerhardtite (hydrated Cu nitrate) is the result of chemical reactions involving explosives. At the Adventure Mine a near horizontal cross-cut adit beginning at Shaft No. 2 connects to the near horizontal Butler drift (Figure 8). To the southeast the Butler drift daylights at the Overview Entrance/Exit. To the northwest the Butler drift connects with the Shaft No. 1 cross-cut adit and with the cross-cut adit to the Ogima lode where there is a large, in place mass, of native copper (Figure 8). The underground of the Adventure Mine is at a stable temperature of about 6oC and is relatively dry and regular field shoes are usually sufficient; hard hats and lights are required and provided by Adventure Mining Company (tour operator). The field trip involves an easy walk 148 �underground to observe the character of native copper mineralization in a horizontal adit that cross cuts the lava flows and a horizontal drift that parallels the strike of the top of the Butler lava flow which hosts a tabular native copper ore body (lode) (Figure 8). The character of native copper mineralization is readily observable in adits, drifts, and stopes (Figure 9). Figure 8: A. Longitudinal section of the Butler lode at the Adventure mine showing workings/openings projected to the vertical and locations for the field trip. B. Geologic map of the Adventure Mine showing workings/openings projected to the horizontal and locations for the field trip. Modified from Butler and Burbank (1929). 149 �Figure 9: Cross section sketch of the topography at the Adventure Mine showing top of the Butler lava flow and the Shaft No. 1 crosscut adit. Overview of the Human History As the land surface of the Keweenaw Peninsula emerged above the progressively retreating glacial lake levels by ca. 7,000 years ago native people took an interest in native copper since its malleability facilitated making tools. At first native peoples likely found boulders of native copper deposited from the glaciers (locally termed float copper) along with gravels and sands derived from erosion of local bedrock and bedrock north of Lake Superior in Canada. These boulders of native copper have a distinctive weathered surface crust of malachite, a green copper-bearing mineral. The green color would have made the relatively infrequent float copper boulders stand out among the other brown, gray, red, and white rocks. The float copper would have also been much heavier than other rocks of the same size. The native people shaped the native copper into tools and decorations. After they depleted the float copper boulders on the surface, they needed a new source of native copper to be able to continue making and trading these items. The native peoples likely found native copper in bedrock because of the green coloration as compared to black- or red-colored host rocks and then became prehistoric miners. There are many shallow mine pits throughout the Greenland-Mass subdistrict. Early European explorers were shown specimens of float copper by native inhabitants which created interest in the Keweenaw Peninsula. In 1841, Douglass Houghton’s report to the Michigan legislature (Michigan’s first state geologist) sparked the first major mining rush in North America. The first significant discovery of native copper was in 1845 at the Cliff Mine which in 1849 became the first profitable native copper mine in the Keweenaw Peninsula native copper district. There were only a few profitable native copper mines from 1845 to the early 1860s. The Minesota Mine, in the Greenland-Mass subdistrict (Figure 6) southwest of the Adventure Mine, became profitable a few years after the Cliff Mine. Many discoveries led to the opening of many mines in the early 1860s. But by 1880, the fate of most of these mines was the same as described below. Initial excitement of possible riches from mining copper was promoted by discoveries of mass copper that implied high grade ore (Figure 10C). Unfortunately, the existence of masses of copper did not necessarily indicate high grade ore. These masses represent the sampling problem termed the “nugget effect.” When the deposit formed there was a clustering of copper in distinct parts of the ore body into a “nugget”. If the mass of copper was missed during exploration the estimate of the grade of the ore body could be far too low and if a 150 �Figure 10: Historic photos from the Adventure Mine. mass is found the grade can be far too high. Mining decisions, such as putting in a shaft or constructing an oversized mill, made from discovery of masses of copper can be costly mistakes especially when funds are limited. Many of the mining projects were underfunded making it difficult to succeed and as a result of lack of funds the operations had to frequently close and in many cases the company merged with another company if they could convince shareholders of potential to discover a large ore body with high grade. There were also frequent shareholder assessments. Rather than paying a dividend to each shareholder from excess funds (profits) an 151 �assessment is the opposite and required each shareholder to pay the company a fee for each share. The fate of the Adventure Mine briefly described below follows this fate and was permanently closed and abandoned by 1917. Adventure Mine History (modified with permission from text provided by Brandon Erickson) In 1848 the Adventure Mining Company began exploration for native copper in the GreenlandMass subdistrict. Exploration activities were focused on a topographic bluff where several different lodes were exposed at the surface. By 1850, the Butler lode appeared to have best potential and the first production of native copper began in 1850. Despite the initial promise the mine struggled to turn a profit as the Butler lode was very rich in copper in some areas and in other areas it was barren. By 1855, the Adventure Mining Company itself ceased mining and to survive in 1855 the company introduced a tributing system. Under tributing, miners would receive a percentage of copper profits instead of a daily wage. In 1856, during financial troubles, a water-powered stamp mill was built along nearby Adventure Creek. By the start of the Civil War in 1861, the richest known ore shoots had been mined out and in 1864 the mine was sold to new investors. The new Adventure Copper Company explored the Butler lode on the eastern side of the bluff but also started an exploration crosscut on the northern slope. The purpose of this adit was to cut across all the lodes. Today’s mine tours enter the mine using this adit (Figure 8, Shaft No. 1 crosscut.) By 1869, the adit had reached the Butler lode and the miners commenced drifting along the lode. Several years later this zone proved rich enough to warrant the sinking of a shaft from the top of the bluff, which is seen on today’s tours as the “skylight stope.” (Figure 8). Mining was centered around this area until an economic downturn in 1877 forced the mine to cut costs and once again only support a small handful of tribute miners. In 1890, the tribute miners uncovered the Knowlton lode, which, unlike the others, cropped out at the base of the bluff. The deeping of shaft No. 1 started immediately and soon reached sufficient depth to begin drifting along the Knowlton lode. The Knowlton lode was very rich in stamp rock (fine sand sized copper disseminated throughout the ore body and lacking the nugget effect). The high-grade copper ore incentivized the miners to continue sinking the shaft to 200 feet deep. At these depths, work was severely hindered by the lack of more modern equipment, and without financial backing the Knowlton efforts ceased by 1893. Adventure Mine was revived again in 1898, as the Adventure Consolidated Mining Company and listed on the Boston stock exchange. This new company was backed by 2.5 million dollars perhaps prompted by ”riches” from mass copper (Figure 10C). The new investment at Adventure Mine was sufficient to build a company town, put in a railroad spur, purchase modern drills, construct and equip a state-of-the-art stamp mill, and install an electric tram line. The company first dewatered the No. 1 shaft and started sinking No. 2 shaft. The promising initial results prompted the company to start a new fully modern third No. 3 shaft (Figure 10A and 10B). This shaft is near the present-day parking lot. By 1903, the No. 1 shaft production had declined and was abandoned at a depth of 700 feet. In response, the company started a fourth shaft on the 152 �eastern limit of their lands (Figure 8). This shaft was a failure and was abandoned after a short time. Shipments of ore from the No. 3 shaft (Figure 10A and 10B) to the stamp mill declined in tonnage and grade. The mine was forced to use diamond drilling to explore for a new ore shoot. In 1909, they found a series of promising new lodes, but to reach them a vertical shaft would need to be sunk. This endeavor around 1910 was Adventure Mines’s “ last hope.” The Adventure Consolidated Mining Company diverted all resources to sink a shaft to 1,500 feet and explored several different lodes along the way. Unfortunately, this exploration was a failure as the ore bodies were not large enough or with high enough grade to make them profitable. The Adventure Mine was more or less abandoned in 1910 but one last attempt was made in 1916 when copper prices increased. The No. 3 Shaft (Figure 10A and 10B) was dewatered down to 700 feet and soon after they were shipping 300 tons of concentrate per day to the smelter. This effort was short-lived and on October 27, 1917 the company ceased all mining and processing activities. Keweenaw National Historical Park The historical significance of the Keweenaw Peninsula native copper district can readily be deduced from the fact that 80% of the new copper for the entire United States was produced from the district in 1880. This was the peak of significance of native copper mining in the district. By 1900, the Keweenaw Peninsula produced only 25 % of the United States new copper. However, absolute copper production from the district peaked in 1916 at an annual production of 121 million kg (267 million pounds). Mining ended in the Keweenaw Peninsula native copper district in 1968. The Adventure Mine peaked in production of copper between 1902 and 1907 at a total of about 3.9 million kg (8.5 million lbs.). 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. The national park visitors center in Calumet provides an excellent overview of the historical significance of the district. Keweenaw Heritage Sites are affiliated with and support activities of the national park. The A.E. Seaman Mineral Museum and Quincy Mine are heritage sites and support the activities of the national park. The Adventure mine is also a Keweenaw Heritage Site and welcomes tourists and visitors seeking an underground mining experience. ACKNOWLEDGMENTS We thank Brandon Erickson for his brief Adventure Mine history summary which we modified for this field guide. We thank Allan Blaske for his review of this field guide that provided significant improvements to this guide. REFERENCES CITED Bodden, T.J., Bornhorst, T.J., Bégué, F., and Deering, C., 2022, Sources of hydrothermal fluids inferred from oxygen and carbon isotope composition of calcite, Keweenaw Peninsula native copper district, Michigan, USA: Minerals, v. 12, 474. https://doi.org:10.3390/min12040474 153 �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., Barron, R.J., and Whiteman R.C., 2013, Caledonia Mine, Keweenaw Peninsula native copper district, Ontonagon County, Michigan: 59th Institute on Lake Superior Geology Proceedings, v. 59, part 2, p. 43-57. 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. and Mathur, R., 2017, Copper isotope constraints on the genesis of the Keweenaw Peninsula native copper district, Michigan USA: Minerals, v. 7, 185, https://doi.org:10.3390/min7100185 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 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: 41st Institute on Lake Superior Geology Proceedings, v. 41, part 1, p. 34. 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. 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. 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. 154 �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. 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. 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. Woodruff, L.G.; Daines, M.J.; Cannon, W.F.; Nicholson, S.W., 1995, The thermal history of the Midcontinent Rift in the Lake Superior region: implications for mineralization and partial melting: in International Geological Correlation Program, Field Conference and Symposium on the Petrology and metallogeny of volcanic and intrusive rocks of the Midcontinent rift system, Duluth, Minnesota, v. 336, p. 213-214. 155 �156 �Field Trip 6 Southern Complex Granitoids, Gneisses, and Migmatites: New Data, Discoveries, and Perspectives Chad D. Deering Michigan Technological University, 1400 Townsend Dr., Houghton, MI 49931 Introduction The Superior Province is part of the Archean Canadian Shield in North America and represents one of the oldest and most stable cratonic regions on Earth, encompassing parts of Canada and the United States, including Michigan, Wisconsin, and Minnesota. The Archean craton in Northern Michigan is divided into a Northern Complex and Southern Complex, which are separated by the Great Lakes Tectonic Zone (GLTZ) (Morey and Sims, 1976; Sims et al., 1980). The northern portion of the Southern Complex consists of classic ‘dome-and-keel’ structures characterized by domes of Archean basement surrounded by keels of Paleoproterozoic Marquette Range Supergroup lithologies (Annhaeusser et al., 1969). The Archean rocks include a complex assemblage of granitoids, gneisses, and migmatites intruded by numerous mafic dikes and/or sills. Our recent research on the area has revealed new aspects of the igneous and metamorphic evolution that have improved our understanding of the assembly of this large igneous-metamorphic complex through the Archean-Proterozoic transition, but numerous questions regarding the origin evolution of these rocks remain unresolved. This field excursion will include the exploration of a number of different terrains representative of the magmatic, metamorphic, and structural evolution of the Southern Complex and the overlying metasedimentary rocks; highlighting new discoveries while at the same time providing an opportunity to investigate still unresolved questions regarding the geologic evolution of the region. Figure 1. Regional map outlining the location of the Southern Complex Mineral District near Marquette, Michigan. 157 �Evolution of the Superior Craton Mesoarchean The Mesoarchean evolution of the Superior Craton in North America encompasses a critical period of crustal evolution between 3.2 to 2.8 billion years ago. During this time, significant tectonic and magmatic processes shaped the early Earth's crust and laid the foundation for the stable continental core we recognize today. In the early Mesoarchean, around 3.2 billion years ago, smaller continental fragments began to accrete and amalgamate due to tectonic processes related to subduction and associated magmatic activity (Percival et al., 2012; Thurston et al., 2008; Wyman, 2010). These proto-continents served as the building blocks for the Superior Craton (Percival et al., 2012). Intense magmatic activity occurred, leading to the generation and growth of the continental crust within the Superior Craton, as magma intruded into and solidified within the existing crust (King et al., 1998). Greenstone belts, characterized by volcanic and sedimentary rocks, also began to form during this period. These belts, such as the Abitibi and Wawa greenstone belts of Canada (Thurston, 2002) and the Ishpeming greenstone belt of the Upper Peninsula, Michigan, USA (Bornhorst and Johnson, 1993), are important features of the Superior Craton and provide insights into early Earth processes, including volcanic activity and the nature of oceanic environments. The rocks of the Superior Craton underwent significant metamorphism and deformation during the Mesoarchean and high temperatures and pressures caused by tectonic activity led to the development of foliations and other prominent structural features in the rocks. Neoarchean During the Neoarchean Eon, which lasted from approximately 2.8 to 2.5 billion years ago, the crust of the Superior Craton underwent further significant geological evolution. This involved the continued accretion of smaller continental blocks and terranes through tectonic processes such as subduction, collision, and magmatic activity (Mole et al., 2021). In particular, the Minnesotan orogeny occurred around 2.7 to 2.6 billion years ago. This was a period of intense tectonic activity characterized by the collision and amalgamation of various smaller continental blocks and island arcs, leading to the formation of a larger continental mass (Schmitz et al., 2018). The collisional process resulted in the growth of the Superior Province and the formation of the granite-greenstone terranes that comprise much of the region. The amalgamation of these crustal fragments contributed to the expansion and stabilization of the Superior Craton. Neoarchean rocks include extensive granitic intrusions, which formed through the partial melting of existing crustal rocks or through the emplacement of mantle-derived magmas (Mole et al., 2021). These granitic intrusions contributed to the growth of the continental crust and are often associated with mineralization and hydrothermal activity (Mole et al., 2021). Greenstone belts, characterized by volcanic and sedimentary rocks, continued to develop during the Neoarchean within the Superior Craton (Polat et al, 1998; Polat and Kerrich, 2000). These belts represent ancient oceanic crust and island arc environments, and they are interspersed with granitic intrusions. Metamorphic processes affected both the greenstone belts and the granitic intrusions within the craton. Hydrothermal activity continued to play a significant role in the formation of mineral deposits within the Superior Craton during the Neoarchean. Ore deposits such as gold, iron, and copper formed in association with granitic intrusions, greenstone belts, and hydrothermal alteration zones, contributing to the economic significance of the region (Mole et al., 2022). 158 �Southern Complex, Marquette District: Compeau Creek and Bell Creek batholith The southern complex is located south of the Marquette synclinorium and is dominated by the Archean Bell Creek batholith, which consists primarily of coarse-grained megacrystic, high-K igneous rocks with minor amounts of mafic gneiss and metasedimentary rock layers typically found concordant with the foliation of the gneiss. Bell Creek includes trondhjemite-tonalitegranodiorite (TTG), granites, gneisses, and migmatites. Migmatite comprises only a small portion of the complex, distributed at irregular intervals throughout the region and is assigned to a unit referred to as Compeau Creek. The age and origin of the Southern complex of the Marquette region has been debated for decades. It was originally thought to be genetically related to similar lithologies found in the nearby Northern Complex (Cannon and Simmons, 1973; Van Schmus and Woolsey, 1975). However, the Southern Complex is separated from the Northern complex by the Great Lakes Tectonic Zone (GLTZ), which is a continental scale suture/fault zone (Morey and Sims, 1976; Sims et al., 1991). The only age information available before our study of the Bell Creek batholith was obtained by Tinkham (1997) from a single zircon with a U-Pb age of ~2.61 Ga. This period marks the onset of the Archean-Proterozoic transition, which is associated with a shift from the production of dominantly mantle-derived magmas that differentiated to form new continental crust to the early stages of significant recycling of crustal material (Taylor & McLennan, 1995; Valley, 2005). This is a crucial period in Earth’s history due to a decrease in global heat flow, and potentially the onset of ‘modern style’ subduction (Brown et al., 2020) and an increased number of sedimentary environments (Taylor and McLennan, 1995). The origin of trondjemite-tonalite-granodiorite (TTG) and high-K granites during the Archean Eon, including those produced in the Bell Creek batholith, is closely linked to the evolution of continental crust. Several models have been proposed to explain the formation of TTGs during the Archean: 1) Partial melting of pre-existing continental crust. During the Archean, the Earth's crust was thicker and more mafic (rich in magnesium and iron) compared to modern crust (Tang et al., 2016). As a result, when portions of this crust were subjected to high temperatures and pressures, particularly in subduction zones or during collisional events, they could undergo partial melting to produce high potassium granites. 2) Partial melting of the mantle. In this scenario, mantle-derived melts ascend through the crust, assimilating and interacting with continental crust along the way. These interactions can lead to the enrichment of potassium and other incompatible elements in the melts, ultimately resulting in the formation of high potassium granites. 3) Derivation from hybrid sources. It is also possible that the formation of high potassium granites during the Archean involved a combination of crustal and mantle processes. This hybrid model suggests that both the crust and mantle contributed to the source materials for the granites, with melting and mixing occurring at various depths within the Earth's crust. New petrogenetic insights on the origin of Bell Creek batholith Our new bulk-rock major and trace element data and U-Pb zircon dates (including oxygen and LuHf isotopes) provide insight into several aspects of the origin of these rocks. First, extensive U-Pb zircon dating of Bell Creek and Compeau Creek rocks from several recent Michigan Tech geology graduate student studies (Table 1; Petryk, 2019 and Barth, 2023) indicates that the bulk of the magmas were emplaced during a single tectonic event between ~2.4 to 2.6 Ga and is attributed to the collision of the Paleoarchean Minnesota River Valley Terrane (MRVT). Second, bulk-rock 159 �major and trace element compositions of the Compeau Creek migmatite/gneiss and Bell Creek granitoids/gneisses are consistent with formation in a continental arc type tectonic environment (Figure 2). Table 1. Age summary for Bell Creek granitoids (data from Petryk, 2019) Bell Creek granitoid subgroups Sample U-Pb crystallization age (Ga) Inherited grains (Ga) Metamorphic grains (Ga) CLG-14B 2.42 ±0.042 2.8-3.6 2.3 CCG-12A BCG-4A BCG-1A BCG-8B CCG-9D BCG-7C CCG-9E 2.43 ±0.160 2.51 ±0.021 2.58 ±0.056 2.54 ±0.040 2.59 ±0.028 2.61 ±0.047 2.56 ±0.038 2.7 2.8-3.9 (4.2) 2.7-2.8 3.1 3.3-3.5 2.7-3.3 2.7-2.8 2.0 2.3, 2.4 2.3, 2.4 Fine-grained CCG-6A 2.53 ±0.070 - 2.3, 2.4 Clotted BCG-8A CCG-1A 2.55 ±0.017 2.55 ±0.033 2.7-2.9 2.7-3.2 2.2, 2.3 1.8, 2.3 Normal' Foliated 2.3 Figure 2. Left: Tectonic discrimination diagram showing a volcanic arc to syn-collision tectonic setting for Bell Creek granitoids and gneissic rocks (blue) and two younger alkaline granites consistent that formed within-plate following the emplacement of the Bell Creek batholith. Right: Arc rock discrimination diagram showing the high-K nature of most of the Bell Creek and Compeau Creek rocks. Note that the mafic rocks plotted here represent dikes/sills of varying and relatively unknown age. 160 �Third, the oxygen isotopic composition of zircon from the granitoids is dominated by mantle-like values (5.3±0.3) indicating additional of a significant amount of juvenile magma to the crust; however, the hafnium isotopic compositions of the same zircons are heterogeneous and have model ages that indicate the involvement of older Mesoarchean crustal lithologies at the source and during emplacement in the mid- to upper-crust. Our current interpretation is that the granitoids, gneisses, and migmatites were formed through a hybrid process, initially by partial melting of an isotopically heterogeneous Mesoarchean lower crust followed by assimilation of mid- to uppercrustal lithologies. Continental arc subduction produces TTGs when high heat flow facilitates partial melting of lower crust and potentially later as the magma reaches the level of final emplacement in the mid- to upper-crust. Interestingly, the high oxygen isotopes (δ18O > 6‰) and εHf isotopes, with a range between -2 and -24, together indicate a potentially greater role for older Mesoarchean crust in the formation of the Southern Complex magmas than what has been found in the Canadian portion of the Superior craton. Therefore, it appears as if the evolution of the southernmost region of the Superior Craton involved a much greater contribution from crustal lithologies than what has currently been found in the well-studied Canadian segment of the Superior Craton (Mole et al., 2019). Figure 3. Hf-O isotopic compositions of zircon from Bell Creek TTGs Mid- to upper-crustal level mixing and assimilation The Archean granitoids of the Bell Creek batholith also display field evidence of assimilation and mafic-felsic magma mixing at the level of emplacement. Small (up to a cm or two) xenoliths that appear as clots throughout the granitoids in the area consist of biotite, chlorite, garnet and quartz indicative of assimilation of pelitic crust. However, basement crustal lithologies of this type are apparently not well exposed. In our investigation we have identified several outcrops that have quartzite and schist in contact with, or within, gneissic or granitoid host rocks, but the ages have yet to be determined (in progress). These metasedimentary rocks outcrop within the igneous intrusions and are considered to be the best possible candidates for remnants of Archean supracrustal material that was assimilated into the felsic magmas. In addition, our new O and Hf isotope data from single, inherited zircons indicate incorporation of Mesoarchean age crustal lithologies (Table 1 and Figure 3). Major and trace element data reveal a slightly peraluminous 161 �character (Figure 4), high-K (Figure 2), and enrichment of incompatible trace elements (Figure 2) and are best explained as reflecting the assimilation of metasedimentary crustal lithologies. The mafic intrusions also show evidence of felsic ‘blebs’ and complex interactions with the host granitoids throughout the region and are interpreted to reflect magma mixing, which would indicate that at least some of them are sills rather than dikes. However, later generations of mafic intrusions clearly crosscut the dominant foliation, have sheared boundaries, and/or sharp contacts with the host rock and represent at least four to five distinct episodes of magmatism related to younger events unrelated to the granitoids. Figure 4. Shand's index for peraluminosity of Compeau Creek and Bell Creek granitoid, gneissic and migmatitic rocks. Paleoproterozoic During the Paleoproterozoic Eon (roughly 2.5 to 1.6 billion years ago), the Superior Craton underwent significant geological evolution, marked by a series of tectonic, magmatic, and metamorphic events. The early Paleoproterozoic was dominated by the development of rift basins during the breakup of the Superia supercraton during rifting that began ~2.1 Ga, which separated the Wyoming Province from the Superior Province (Drenth et al., 2021). These rift basins accumulated thick sequences of sedimentary rocks, including sandstones, shales, and iron formations that are today known in the Lake Superior region as the Marquette Range Supergroup and Huronian Supergroup (Ojakangas et al., 2001). The Penokean Orogeny, which occurred around 1.85 to 1.75 billion years ago, resulted from the collision of the Superior Craton with other continental blocks that contributed to the growth of Laurentia, leading to crustal thickening, mountain-building, and metamorphism (Schulz and Cannon, 2007). Hydrothermal activity during the Paleoproterozoic played a significant role in the formation of mineral deposits that include iron-rich sedimentary deposits, such as banded iron formations (BIFs), which were deposited during this time, along with important mineral deposits such as iron, copper, and gold (DeMatties, 2022). Following the main tectonic events of the Paleoproterozoic, the Superior Craton experienced episodes of post-orogenic magmatism, marked by the emplacement of large igneous provinces and granitic intrusions. 162 �Field Trip Objectives This field trip is designed to provide a geologic overview of the formation of the Neoarchean Bell Creek batholith including associated metasedimentary rocks and the Paleoproterozoic metasedimentary rocks that filled the deep basins. Figure 5. Generalized geologic map of the Southern Complex Mineral District. 163 �Stop 1: Paleoproterozoic Negaunee Iron Formation (FR/Xn) and mafic dike Directions: Leaving Michigan Technological University drive south along US-41 ~63 miles to the Michigamme Roadside Park. N 46° 32’ 20” W88° 5’ 31” The outcrop is prominently exposed along the northern side of US-41 across from the Michigamme Roadside Park, which overlooks Lake Michigamme to the south. Paleoproterozoic Negaunee Iron Formation that is weakly magnetic containing hematite-goethite with banded chert as fine laminations. The formation is part of a westward plunging syncline conformably overlying Siamo Slate or Ajibik quartzite with a gradational contact and lies unconformably over the Archean basement complex (Gair and Thaden, 1968). This unit includes sideritic slates, grunerite-magnetite-schists; ferruginous slates, ferruginous cherts and jaspilite (Van Hise and Bayley, 1895; Cannon and Gair, 1970). Here, the iron formation is dipping ~65° to the SW and is in sharp contact with the adjacent massive, medium-grained to porphyritic amphibolite dike. A younger, near vertical dike can be observed on the east side of the outcrop in sharp contact with the older amphibolite dike/sill. Stop 2: Paleoproterozoic Ajibik quartzite (Xa) Directions: Heading east along US-41 ~0.5 miles exposure of Ajibik quartzite outcrops along the northern edge of the road. N 46° 32’ 35” W88° 4’ 16” The Ajibik consists of basal conglomerates, slates, and graywackes that grade into the overlying quartzite. At this location, the small exposure of the Ajibik is massive, thick bedded white to buff orthoquartzite. Some relic cross-bedding may be present, but clear evidence of ripple marks or other features reflecting the original depositional environment are not present. A mafic dike crosscuts the quartzite roughly NW-SE and is highly sheared along a sharp contact. Bedding is difficult to identify, but it appears to be dipping ~60° to the SW. Stop 3: Paleoproterozoic Michigamme formation (Xms) Directions: Head east on US-41 S toward Orange Rd. 8.4 miles. Turn right onto M-95 S, and drive ~7.3 miles, crossing the Michigamme River basin to destination. N 46° 24’ 37” W87° 59’ 45” This exposure is of the lower slate member, which includes laminated iron-rich rock consisting of biotite-garnet-cummingtonite-quartz schist with thin beds of quartzite. Minor fold axes are prominent here as chevron folds plunging to the NW and reflect the regional fold axis orientation. Boudinage, which form when single, competent layers are stretched into separate pieces through plastic and/or brittle deformation mechanisms and reflect the presence of minor quartzite layers or lenses within the schist. 164 �Stop 4: Archean Bell Creek batholith granitoid (Wbcf) Directions: Head southwest on M-95 N ~0.5 miles to destination. N 46° 24’ 13” W87° 59’ 52” The most common form of Bell Creek batholith is exposed here as a medium to coarse-grained granitoid that is locally porphyritic with sparse megacrysts of alkali-feldspar (up to several cms). Minerals include alkali-feldspar, oligoclase, biotite, oxides with apatite and zircon as common accessory phases. The abundance of mafic xenoliths (clots) (Figure 6) within the granitoids is correlated with the degree of peraluminosity of the host rock. Some minor pegmatitic veins are also present. Small mafic injections are in sharp contact with the granitoid indicating emplacement that post-dates the main magmatic episode. There is some weak alignment of alkali-feldspar megacrysts that can be best observed on the top of the outcrop. Kfs Grt Iron-Chl 250 μm Figure 6. Left: Mafic clot containing Fe-chlorite or biotite, K-feldspar, and garnet (almandine). Garnet-biotite geothermometry yields an equilibration temperature range between 450°C and 550° C. Right: Example of mafic xenoliths clustered within granitoid. 165 �Stop 5: Undivided granitic rocks (Wgu) Directions: Head south on M-95 toward Co Rd 601 ~4.8 miles to destination. Outcrop is located just south of the city park along the Michigamme river. N 46° 20’ 13” W087° 58’ 22” A metasedimentary sequence is exposed here, which doesn’t appear on the regional geologic map. Blonde and gray/black feldspathic quartzite forms alternating beds with gradational contacts dipping to the north (Figure 7). The blonde quartzite is dominated by quartz, with minor amounts of feldspar and sparse garnet (1-2mm). The gray/black quartzite has similar proportions of quartz and feldspar to the blonde quartzite but includes muscovite, biotite and oxides nearing the abundance of feldspar. Complex interaction with mafic dikes/sills also appear in this outcrop with minor ptygmatic folding and mafic enclaves throughout. Vertical dike contacts are sharp, whereas an apparently older generation of mafic intrusions have more complex contact margins with the host rock. Near vertical mafic dikes on both ends of the outcrop have sharp contacts with the host rock penetrating locally as splays or fingers. Figure 7. Metasedimentary sequence dominated by blonde quartzite and a grey to black quartzite. Lunch Stop: Leif Erickson Roadside Park along the Michigamme River 166 �Stop 6: Archean Bell Creek batholith and Compeau Creek gneiss/migmatite (Wbcf/Wccg) Directions: Drive north along M-95 towards Welsh’s Rd. ~5.7 miles to destination. Outcrop exposed prominently on the east and west sides of the road. N 46° 24’ 54” W87° 59’ 31” The contact between Compeau Creek to the south and Bell Creek batholith to the north is clearly exposed at this outcrop. A zircon U-Pb date of 2426±160 Ma was obtained for the Bell Creek granitoid at this location. The granitoid is fine- to medium-grained, typically equigranular and has an apparently gradational contact with the adjacent migmatite/gneiss. Migmatite/gneiss with ptygmatic folding is best exposed in the outcrop on the west side of the road, but the highly deformed continuation of this rock type is prominently exposed with mafic intrusions on the east side of the road. Stop 7: Archean Bell Creek batholith (Wbcg) Directions: Drive north along M-95 ~2.1 miles to destination. N 46° 26’ 16” W87° 57’ 48” Exposure of Bell Creek megacrystic granitoid similar to Stop 4. Minerals include alkali-feldspar, oligoclase, biotite, oxides and apatite with apatite and zircon as common accessory phases. A strong foliation is not apparent here, but the mafic xenoliths (clots) are well represented as mm- to cm-sized dark spots typically clustered and randomly oriented. Pegmatitic veins are common and vary in width up to tens of centimeters. Stop 8: Archean Migmatite (Wbsc); Compeau Creek? Directions: Drive north along M-95 ~0.8 miles to destination. N 46° 26’ 53.7” W087° 57’ 20.0” This exposure is likely the Compeau Creek quartofeldspathic migmatite/gneiss dipping to the south. The highly deformed migmatite/gneiss includes numerous mafic inclusions and enclaves (Figure 8). However, it is unclear how much of the mafic-felsic segregations are representative of the initial magma mixing followed by deformation or melanosome-leucosome complementary rocks derived by melting. Ptygmatic folding is prominently displayed and likely represents felsic segregations that have buckled during deformation. A zircon U-Pb date of 2525±70 Ma was obtained for the gneiss, which overlaps in time with the dates obtained for other Bell Creek granitoids along the M-95 corridor. 167 �Figure 8. Near vertical mafic dike cross-cutting migmatite with chaotic mixture of mafic-felsic components. Stop 9: Archean Compeau Creek migmatite/gneiss (Wccg) Directions: Drive north along M-95 ~1.4 miles to destination. N46° 27’42” W87° 56’ 05” This outcrop is mapped as Compeau Creek migmatite/gneiss that is presumed to be older than the adjacent Bell Creek granitoid/gneiss. A zircon U-Pb date for the migmatite is slightly older (2633±46 Ma) than the pink, clotted granite, which has been dated to 2550±33 Ma. Inherited zircon grains range in age from 2717 to 3200 Ma. The Compeau Creek gneiss/migmatite on the northern end of this outcrop appears to include a sliver of what might be Archean quartzite (Figure 9); note that the Paleoproterozoic Goodrich quartzite is mapped directly to the north and would, therefore, be in direct contact with the migmatite/gneiss. There are numerous mafic dikes/sills that manifest as either single generation dikes that clearly cross-cut the existing foliation or more complex dismembered dikes (possibly sills) associated with mixing at the time of felsic magma emplacement (Figure 10). Inclusions of felsic material (either bulk rock or individual feldspar crystals) can be found in some of the composite dikes/sills indicative of magma mixing that may have occurred contemporaneously between the mafic magma and a liquid dominant felsic host magma. The foliation is steeply south-dipping (~70°). 168 �Figure 9. Band of quartzite concordant with the foliation of gneiss/migmatite fabric. Relict feldspar crystals ~ 4 cm Figure 10. Mafic intrusion with felsic xenocrysts/xenoliths interpreted to have likely formed through mixing with the host magma during emplacement. This particular intrusion has a Nd model age of ~2.57 Ga, which is within error of the age obtained for the Bell Creek gneiss at this location. 169 �References Barth, Elana G. Age and chemistry of Bell Creek batholith (MS report), Michigan Technological University (2023) Bornhorst, Theodore J., and Rodney C. Johnson. Geology of volcanic rocks in the south half of the Ishpeming greenstone belt, Michigan. No. 1904. US Government Printing Office, 1993. Brown, Michael, Tim Johnson, and Nicholas J. Gardiner. "Plate tectonics and the Archean Earth." Annual Review of Earth and Planetary Sciences 48 (2020): 291-320. Cannon, W. F., and J. E. Gair. "A revision of stratigraphic nomenclature for middle Precambrian rocks in northern Michigan." Geological Society of America Bulletin 81, no. 9 (1970): 2843-2846. Cannon, W. F., and George C. Simmons. "Geology of part of the southern complex, Marquette district, Michigan." J. Res. US Geol. Surv 1 (1973): 165-172. DeMatties, Theodore A. "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 (2022): 104489. Gair, Jacob Eugene, and Robert E. Thaden. Geology of the Marquette and Sands Quadrangles, Marquette County, Michigan. No. 397. 1968. King, Elizabeth M., John W. Valley, Don W. Davis, and Garth R. Edwards. "Oxygen isotope ratios of Archean plutonic zircons from granite–greenstone belts of the Superior Province: indicator of magmatic source." Precambrian Research 92, no. 4 (1998): 365-387. Mole, D. R., C. L. Kirkland, M. L. Fiorentini, S. J. Barnes, K. F. Cassidy, C. Isaac, E. A. Belousova, M. Hartnady, and N. Thebaud. "Time-space evolution of an Archean craton: A Hfisotope window into continent formation." Earth-Science Reviews 196 (2019): 102831. Mole, D. R., P. C. Thurston, J. H. Marsh, R. A. Stern, J. A. Ayer, L. A. J. Martin, and Y. J. Lu. "The formation of Neoarchean continental crust in the south-east Superior Craton by two distinct geodynamic processes." Precambrian Research 356 (2021): 106104. Mole, D. R., B. M. Frieman, P. C. Thurston, J. H. Marsh, T. R. C. Jørgensen, R. A. Stern, L. A. J. Martin, Y. J. Lu, and H. L. Gibson. "Crustal architecture of the south-east Superior Craton and controls on mineral systems." Ore Geology Reviews 148 (2022): 105017. Morey, G. B., and P. K. Sims. "Boundary between two Precambrian W terranes in Minnesota and its geologic significance." Geological Society of America Bulletin 87, no. 1 (1976): 141-152. Ojakangas, R. W., G. B. Morey, and D. L. Southwick. "Paleoproterozoic basin development and sedimentation in the Lake Superior region, North America." Sedimentary Geology 141 (2001): 319-341. 170 �Percival, John A., Tom Skulski, Mary Sanborn-Barrie, Greg M. Stott, Alain D. Leclair, M. Tim Corkery, and Michel Boily. "Geology and tectonic evolution of the Superior Province, Canada." In Tectonic styles in Canada: the LITHOPROBE perspective, vol. 49, pp. 321-378. Saint‐John's, Newfoundland: Geological Association of Canada, 2012. Petryk, Brandi. "The origin of an Archean batholith in Michigan’s Upper Peninsula.”. (MS thesis) Michigan Technological University (2019). Polat, Ali, and Robert Kerrich. "Archean greenstone belt magmatism and the continental growth– mantle evolution connection: constraints from Th–U–Nb–LREE systematics of the 2.7 Ga Wawa subprovince, Superior Province, Canada." Earth and Planetary Science Letters 175, no. 1-2 (2000): 41-54. Polat, A., R. Kerrich, and D. A. Wyman. "The late Archean Schreiber–Hemlo and White River– Dayohessarah greenstone belts, Superior Province: collages of oceanic plateaus, oceanic arcs, and subduction–accretion complexes." Tectonophysics 289, no. 4 (1998): 295-326. Schmitz, M. D., D. L. Southwick, M. E. Bickford, P. A. Mueller, and Scott Douglas Samson. "Neoarchean and Paleoproterozoic events in the Minnesota River Valley subprovince, with implications for southern Superior craton evolution and correlation." Precambrian Research 316 (2018): 206-226. Sims, Paul Kibler. Great Lakes tectonic zone in Marquette area, Michigan: implications for Archean tectonics in north-central United States. No. 1904. US Government Printing Office, 1991. Tinkham, D. K. Tectonic Evolution of the Southern Complex Regiong of the Penokean Orogenic Belt, Upper Peninsula Michigan: The Formation of Precambrian Dome-and-Keel Architecture (MS), University of Illinois. (1997). Tang, Ming, Kang Chen, and Roberta L. Rudnick. "Archean upper crust transition from mafic to felsic marks the onset of plate tectonics." Science 351, no. 6271 (2016): 372-375. Taylor, Stuart Ross, and Scott M. McLennan. "The geochemical evolution of the continental crust." Reviews of geophysics 33, no. 2 (1995): 241-265. Thurston, P. C. "Autochthonous development of Superior Province greenstone belts?." Precambrian Research 115, no. 1-4 (2002): 11-36. Van Hise, Charles Richard, and William Shirley Bayley. Preliminary report on the Marquette ironbearing district of Michigan. US Government Printing Office, 1895. Valley, J. W., J. S. Lackey, A. J. Cavosie, C. C. Clechenko, M. J. Spicuzza, Miguel Angelo Stipp Basei, I. N. Bindeman et al. "4.4 billion years of crustal maturation: oxygen isotope ratios of magmatic zircon." Contributions to Mineralogy and Petrology 150 (2005): 561-580. 171 �Wyman, Derek, and Robert Kerrich. "Mantle plume–volcanic arc interaction: consequences for magmatism, metallogeny, and cratonization in the Abitibi and Wawa subprovinces, Canada." Canadian Journal of Earth Sciences 47, no. 5 (2010): 565-589. 172 �Field Trip 7 Landslides on the Ontonagon River at Military Hill Stanley J. Vitton Civil, Environmental, and Geospatial Engineering, Professor Emeritus, Geological and Mining Engineering and Sciences, Adjunct Professor Emeritus, Michigan Technological University, Houghton, MI 49931 Mohammad Sadeghi Civil, Environmental, and Geospatial Engineering, Assistant Professor, Geological and Mining Engineering and Sciences, Affiliated Assistant Professor Michigan Technological University, Houghton, MI 49931 Figure 1: 2003 landslide east of US-45 on the East Branch of the Ontonagon River. 173 �Introduction An impressive sight, at least to geologists and engineers, are the landslides along US-45 as it descends into the Ontonagon River Valley at Military Hill. US-45 follows the trail used by Indigenous peoples such as the Menominee, Dakota and Anishinaabe (Ojibwe/Chippewa) tribes, and later fur traders making their way to and from to the mouth of the Ontonagon River at Lake Superior. In 1844, U.S. Secretary of War William Wilkins, proposed a road between Fort Howard in Green Bay, WI to Fort Wilkins just north of Copper Harbor, MI to create a supply route to the Army posts and to improve access to frontier communities such as Copper Harbor and other neighboring mine sites. Congress, however, did not fund the road. In 1862, with British troops in nearby Ontario and the prospect of Great Britian entering the war on the side of the confederacy, the road became a national security issue. According to a Michigan historian, Le Roy Barnett, “More than half the copper used in the United States came from mines along the proposed frontier passage. As a report of the U.S. Senate Committee on Military Affairs noted, "In case of hostilities with Great Britain, and a descent upon this important portion of our lake coast, there would be no means of affording succor without a road [to the region], and these valuable deposits of copper and other ores might be lost to us. Senator Jacob Howard pressed the point, explaining that if the Soo Canal were "seized and closed ... there would be no means of getting into the [Keweenaw] country or out of the country, either with troops or munitions of war or without them, except by means of some such road as this." (The Free Library, 2024) On March 3, 1863, Abraham Lincoln signed an Act of Congress authorizing construction of the military road. Road construction was started in 1863 but was not completed until 1872. Following construction, the road was used primarily for timber and other commercial interests until the railroad network in the western Upper Peninsula became operational a few years later. Again, according to LeRoy Barnett, “Despite the years of work put into the route, it never became the avenue of settlement and commerce its supporters anticipated. As legislators in the early 1860s had feared, railroads--built shortly after the road's completion--soon siphoned traffic. Furthermore, parts of the road were poorly maintained. "There was no surfacing on most of the road," wrote Upper Peninsula author/historian Knox Jamison of a trip in 1915, "and if it had rained recently, you had better not try it." Getting up and down the Military Hill--a massive body of slippery, wet clay through which the Ontonagon River had cut a 300foot gorge--took him at least two hours.” In 1923, the Michigan Highway Department added Military Road into the state’s transportation network as a section of State Highway M-26 but did not appreciably improve the road, leaving it with a gravel surface and in general following the route’s topography. In 1957, the Michigan State Highway Department added this section of M-26 to the federal highway system as US-45. 174 �Becoming a federal highway requires the road to meet both state and federal road design standards, such as the highway having a maximum grade of less than 8%. This required significant excavation of the glacial clay slopes and most likely the start of the Military Hill landslides. Construction was completed in 1959 with a concrete pavement and a grade meeting state and federal specifications. While the Military Hill clay soils are known for their difficult road construction issues, the highway department most likely did not fully appreciate the difficulty with the excavation and long-term behavior of the slopes. An investigation, either upstream or downstream of the US-45 river crossing, for example, would have shown many old and new landslides not to mention the red silt laden river itself that never changes throughout the seasons. The landslides are in the glacial lacustrine and till sediments in the Ontonagon Basin formed during the area’s deglaciation. One of the more recent landslides near the river crossing, which occurred in 2003, is shown in Figure 1, is just upstream of US-45 and one of the sites we will visit. Figure 2 shows the US-45-Ontonagon River crossing and the location of the landslide north of the confluence of the East and Middle Branches of the Ontonagon River. The steep slopes and the meandering of the river are evident. Figure 2 also shows the stops that we will be making at Military Hill. The clays in the Military Hill area were deposited in the final stages of the Late Wisconsin glacial period, estimated to be at its maximum at about 26,000 BP with deglaciation ending about 9,500 years BP when the Superior Lobe of the Laurentide Ice Sheet retreated into the Superior Basin for the last time (Attig et al., 1985). Fortunately, the northern Great Lakes region is known as the “type area” for the stratigraphic subdivisions of the late-glacial period of North America so the area has been investigated in some detail (Evenson et al., 1976). The first and the most detailed investigation of the glacial sediments in the Ontonagon area was conducted between 1905 and 1919 by Frank Leverett who identified the area’s main glacial features including Glacial Lake Ontonagon and Duluth (Leverett, 1928). Later, Hatch (1965) investigated the postglacial drainage evolution and stream geometry in the Ontonagon area while detailed mapping was conducted by Peterson (1985 and 1986). The area’s prominent feature is the Copper Range, a topographic ridge formed by the Portage Lake Volcanics. The Copper Range was influential in the formation of the Ontonagon River Basin shown in Figure 3 and Figure 4. As shown in these figures, the Ontonagon River System converges at a gap along the Copper Range. According to Leverett (1928), the Copper Range formed an ice-margin boundary during the last couple of glacial advances resulting in the formation of a series of proglacial lakes shown in Figure 5. The proglacial lakes Ontonagon, Ashland, Brule and Duluth formed along this ice-margin with drainage flowing westward to the St. Croix River. As isostatic rebound occurred and the Superior Lobe retreated into the Superior Basin, the drainage merged through a gap in the Copper Range. Leverett notes that it was probable that Lake Goebic was in a pre-glacial valley that extended northward to Lake Superior but was prevented from draining into Lake Superior by a moraine that filled the gap. Thus, the West and South Branch of the Ontonagon River flowed eastward along the Copper Range to where it drained into the present Ontonagon River. 175 �Another distinct feature of the area is a series of parallel rivers that formed between Copper Range and Lake Superior as shown in Figure 6 and 7. Hatch (1965) investigated the drainage system indicating that “it was strongly grooved by glacial flutings parallel to the direction of ice motion. In places the grooves are buried by lacustrine sediments of glacial Lake Duluth.” It is possible that the Ontonagon River, as it began flowing through the Copper Range gap at Military Hill, followed one of the parallel drainages. Due to a greater volume of water, however, it has significantly cut through the Lake Ontonagon and Duluth sediments as shown in Figure 6. Figure 2: Military Hill along US-45 showing Field Trip Stops (Google Maps, 2024). Figure 3: Ontonagon River Basin location (USACE, 2010). 176 �Figure 4: Ontonagon drainage basin showing river recreation, science, and wild designations (Ontonagon, 2023). Figure 5: Development of proglacial lakes along the Superior Lob (Farran and Drexler, 1985). Figure 6: Parallel drainage system on the Ontonagon Plain by glacial grooving (Hatch, 1965). 177 �Figure 7: Aerial photo of the parallel river systems on the Ontonagon till plain. Figure 8: Bedrock geology of the Military Hill area (Cannon et al., 1995). 178 �Background Information Bedrock Geology The landslides at Military Hill are located over the Jacobsville Sandstone, just south of the Keweenaw Fault, as shown in Figure 8. Ontonagon River Watershed Overview The Ontonagon River Watershed, at 1,348 sq. miles, is the second largest watershed entering Lake Superior behind the St. Louis watershed west of Duluth, MN, at 3,584 sq. miles. What makes the Ontonagon River watershed distinctive, however, is the large amount of red sediments that continually flows into Lake Superior. In 1987, the National Geography Magazine, ran an article on “The Great Lakes Trouble Waters” showing an aerial view of the sediment discharge from the Ontonagon River into Lake Superior with the caption, “Each year millions of tons of sediment – like this red clay silt spilling into Lake Superior from Michigan’s Ontonagon River – enters the lakes from tributary stream. Many contaminated with agricultural chemicals and industrial wastes (National Geographic, 1987).” The cover and photo of the Ontonagon River are shown in Figure 9. In the early 2000s, the US Corps of Engineers (USACE, 2010) conducted a 516e sediment study on the Ontonagon River to determine the source of this large quantity of sediments to estimate future dredging requirements. After its investigation, the USACE reported: “The Ontonagon River watershed is primarily undeveloped and consists of forested land uses with little urbanization or agriculture. Much of the watershed experiences significant erosion of the incised valley walls due to the highly erodible soils associated with the lacustrine geology of the lower reaches. Based on the comparison of historic and present-day river morphology it is concluded that the valley walls and river banks have been a large contributor of sediment long before logging or other anthropogenic disturbances were present in the watershed. Therefore, the primary contributor of sediment yield in the watershed is natural processes, rather than anthropogenic alterations. Based on this conclusion, a quantification of geologic time scale sediment yields was conducted.” The report went on the state: There is evidence that the deep valleys and steep walls that exist today are due to natural sediment transport processes and landscape scale geomorphic evolution that have carved out the valleys and deepened the Ontonagon River and its tributaries. Frequent examples of mass wasting of the valley walls exist throughout the downstream reaches of the Ontonagon River and a flat terrace with steep valley walls adjacent to a meandering channel with a narrow beltwidth are typical in the lacustrine areas of the watershed…..Moreover, an average sediment yield of 2.4 million tons per year is more than an order of magnitude greater than current sediment yields of similar watershed sizes in the Great Lakes.” The USACE report did not investigate which branch of the river produced the greatest amount of 179 �sediments. Based on a study by Weidner et al., (2019), however, it would appear the East and Middle Branches produce the most sediment based on the frequency of observable landslides. This would be constitent with Peterson’s 1985 mapping showing that the East and Middle Branches are mostly located in the Glacial Lake Ontonagon sediments as seen in Figure 10, which is a portion of Peterson’s glacial geology map. The length of the East and Middle Branches of the Ontonagon River can also be seen in the soils map from the USDA Natural Resources Conservation Service landform map shown in Figure 11 where the East and Middle Branch extend to the southern end of the lake sediments. Figure 9: July 1987 issue of the National Geographic Magazine (1987). Review of Landslide Activity on the Ontonagon River Travelers along US-45 can view a landslide just to the west of the US-45 – Ontonagon River crossing. A 1,000-foot walk upstream on the East branch of the Ontonagon River will provide the opportunity to inspect a large-scale landslide that occurred in 2003. If you continue to walk upstream from the 2003 landslide, about every bend in the river will show landslide activity, either current or past as shown in Figure 12. Weidner et al., (2019) studied the landslide activity in the Military Hill area (Figure 13a) for the development of a landslide susceptibility map based on riverbank erosion-triggering. Additional landslide studies were conducted by Koons (1965), Dyl (1979), and Smith (2012), all master theses or reports at Michigan Tech. The Weidner et al., study utilized aerial and satellite imagery from the United States Geological Survey (USGS) EarthExplorer web tool for the years 1992 through 2016 identifying 21 landslides. The landslide locations in the study area are shown in Figure 13(a). 180 �Figure 10: A portion of the USGS map of the glacial history of Iron River 1° x 2° Quadrangle (Peterson, 1985). Figure 11: Landforms map of the study area adapted from Jerome (2006), showing the main soil regimes. River hydraulic data were obtained from a USGS gauging station downstream. The USGS Scoops3D limit-equilibrium analysis software was used to develop a “factor of safety” map of the study area, which is shown in Figure 13(c). A significant portion of the Ontonagon River slopes have a factor of safety less than one, which is supported by the observable landslide activity shown in Figure 13(b) through the lacustrine sediments of Glacial Lake Ontonagon and Duluth. 181 �Figure 12: Landslide activity upstream of US-45 on the East Branch of the Ontonagon River (Google Map, 2024). Figure 13: Weidner et al., Military Hill landslide investigation, (a) study area, (b) observed landslides between 1992 and 2016, and (c) Scoops3D factor of safety assessment. 182 �Objectives of the field trip: The objective of the field trip is to investigate landslides in the sediments deposited in the former Glacial Lakes Ontonagon and Duluth. In general, the sediments consist of lacustrine sediments deposited in Glacial Lake Ontonagon and later in Lake Duluth which overlay older till deposits. The first stop, however, will be at Quincy Hill, north of Hancock, MI to observe glacial grooving and striations in the Portage Lake Volcanics indicating the direction of glacial lobe movement across the Keweenaw Peninsula. The second and third stops will be in the US-45 Military Hill area. A summary of the stops are as follows: • Stop 1: Top of Quincy Hill, north of Hancock, to view glacial grooving in the Portage Lake Volcanics indicating the direction of Keweenaw Bay Lobe movement westward into the Ontonagon Lobe that formed Lake Ontonagon. • Stop 2A: The north side of Military Hill to observe the effects of vegetation on limiting landslide development. • Stop 2B: US-45 Military Hills Roadside Park: 2003 Large landslide on the East Branch of the Ontonagon River • Stop 2C: Lower Military Hill Erosion with Slope Movement • Stop 2D: Middle Military Hill with Partial Vegetation and Some Slope Movement • Stop 2E: Middle Military Hill with More Vegetation and Limited Slope Movement • Stop 2F: Middle Military Hill with Some Vegetation and Varved Clay Slope Movement • Stop 2G: Upper Military Hill with Active Slope Movement on both Sides of Highway in non-varved clay • Stop 3: Slope movement two miles south of Military Hill, across from Primrose Acres, where a recently excavated slope on the east side of US-45 has been moving for about eight years. Stop 1: Glacial Grooves at the Quincy Hill Historic Park Lookout Directions: From Michigan Tech drive west through Houghton on US-41 and cross the Houghton-Hancock Bridge, staying on US-41 going into Hancock. In Hancock, go straight uphill, passing Hancock’s main street, onto East White Street. East White Street will take you to US-41 bypassing downtown Hancock. At the US-41 stop, take a right turn, going uphill to the Quincy Mine Hoist. Directly across from the entrance from the Quincy Mine Hoist take a left turn onto No. 2 Road. Follow No. 2 Road a short distance until you come to a two-track road on your left that takes you to the Keweenaw National Historic Park’s “Quincy Mine Dryhouse Ruins” parking lot. The glacial grooves are a short distance from the parking lot. From Michigan Tech to the parking lot is 3.6 miles. Lat: 47.135904°, Lon: -88.577986° 183 �Figure 14: View, looking west, of glacial grooves sculptured into the Portage Lake Volcanics at the Keweenaw National Historic Park. Glacial grooving and striations are common along the Keweenaw Peninsula as the glaciers moved over the Portage Lava Volcanics, which were resistant to glacial erosion. It is generally assumed glaciers came from northeastern Canada moving south to southwest but are surprised to see grooves heading due west as shown in Figure 14. The reason for this direction is the movement of the Keweenaw Bay Lobe filled Keweenaw Bay and then moved west, south, and east as shown in Figure 15. The Keweenaw Bay Lobe then was stopped by the Ontonagon Lobe and to the west by the Michigamme Lobe to the east. Figure 15: Location of glacial lobes in the western Lake Superior Basin (from Attig et al., 2013). Stop 2: Military Hill – Seven Stops – Safety instruction will be provided at the site Directions: From Michigan Tech drive west through Houghton on US-41 to M-26. At the Houghton-Hancock Bridge go straight through the intersection to access M-26. Stay on M-26. Going southwest 37.8 miles to the M-26/M-38 intersection. Take a left turn, staying on M-26 through Mass City, to the M-26/US-45 intersection, 5.5 miles. Take a left turn at the M-26/US-45 intersection headed south 1.8 miles to Stop 2A. Lat: 46.705726°, Lon: -89.159581° 184 �Stop 2A: North US-45 - Slope with Vegetation and Small Slope Movement In the Military Hill area, the north side slopes required less excavation than on the south side and therefore there is less slope movement. In addition, where the slopes were excavated, they tend to be more vegetated as seen in Figure 16. This is due (possibly) to the north facing slope losing the spring snow before the south facing slopes and thus having longer growing season. Figure 16: Stop 1 (a) descending Military Hill going south and (b) vegetated slopes with some slope movement. Stop 2B: US-45 Military Hill Roadside Park: 2003 Large landslide on the East Branch of the Ontonagon River Directions: From Stop 2A to Stop 2B go 0.4 miles south on US-45 to the Military Hills Roadside Park parking lot on the east side of US-45. Starting at the rest area facilities (outhouses) on the northside of the parking lot, walk northeast through the woods to the landslide. There is no path, but the woods are relatively easy to walk through. The 2003 large scale landslide is approximately 1,400 ft (411 m) northeast of the parking lot as shown in Figure 17. There is a smaller landslide north of the path, which has been active for many years, but is difficult to access. The slope debris, below the landslide, has a thick undergrowth, which is difficult to walk through. Military Hills Roadside Park: Lat: 46.699650°, Lon: -89.158627° Military Hill 2003 Landslide: Lat: 46.703234°, Lon: -89.154132° In 2003, a large-scale landslide occurred on the East Branch of the Ontonagon River as shown in Figure 1 and Figure 17. The landslide’s stratigraphy consisted of lacustrine varved clay over a clean alluvial sand that grades downward into a red silty sand, silt and then clay till as illustrated in Figure 18 and Figure 19. While the failure mechanism is unknown, it is speculated that the spring runoff and possibly high-water table caused the alluvial sand to liquefy causing the massive landslide. As the liquefied sand lost strength, it started to flow outward into the river channel as shown in Figure 20(b) and (c). During the slope’s collapse, a sliding plane formed in the varved clay as shown in Figure 20(a) allowing the varved clay to fail over the alluvial sand. Liquefaction boils are shown in Figure 20(b) and (d). 185 �It will take about 20 minutes to walk to this landslide through the woods from the roadside park. While there is no path, the woods are fairly open and easy to walk in. However, the landslide itself is more difficult due to the landslide debris and the vegetation that has developed over the years. Caution must be observed when walking over this site and it is highly recommended that the landslide slopes are not accessed. Figure 17: Stop 2B Military Hills Roadside Park and path to the 2003 landslide. Figure 18: Assumed stratigraphy of the 2003 landslide (Smith, 2012). 186 �Figure 19: 2003 landslide illustrating the varved clay directly over an alluvial sand. Figure 20: Landslide illustrating (a) a sliding plane on the varved clay, (b) liquefaction boils in the varved clay, (c) the flow debris into the river's channel, and (d) a large sand boil. 187 �Stop 2C: Lower Military Hill Erosion with Slope Movement Directions: From Stop 2B go 0.8 miles south on US-45 across the US-45 bridge to Stop 2C, which is the start of the lower landslides with significant erosion. Lat: 46.690087°, Lon: -89.165513° At the following five stops we will travel up Military Hill’s southside. In general, the soils traveling up the south side of Military Hill are like the sediments seen at Stop 2B except for the alluvial sands. The varved clays overlie a sandy brown silt, which have significant erosion (Figure 21). The base of the varved clay can be seen moving over the sandy silt. As we moved up US-45 at Stop 2C a short distance, MDOT has been attempting to stabilize a small landslide just below the base of the varved clay as shown in Figure 22. The clay content appears to be higher in this area. On the west side of US-45 you can observe the steep US-45 embankment that descends down to Sandstone Creek, exposing the Jacobsville Sandstone. Figure 21 Stop 2C showing the contact between the varved clay and a brown/red sandy silt, which is eroding. Figure 22 Stop 2C small landslide. 188 �Figure 23 Stop 2C showing the US-45 embankment that descends to Sandstone Creek on the Jacobsville Sandstone. Stop 2D: Middle Military Hill with Partial Vegetation and Some Slope Movement Directions: From Stop 2C go 0.23 miles south on US-45 Stop 2D, which is at the middle landslides. Lat: 46.686869°, Lon: -89.164320° Stop 2D is a short distance up US-45 and is in the same sediments as Stop 2C, which also are sloping uphill as seen in Figure 24. Movement of the varved clay base can be seen in the upper portion of Figure 24. Figure 24 Stop 2D varve clay movement over the lower sandy silt but without the erosion seen at Stop 2C. 189 �Stop 2E: Middle Military Hill with More Vegetation and Limited Slope Movement Directions: From Stop 2D go 0.30 miles south on US-45 Stop 2E, an excavated slope that is now partially vegetated. Lat: 46.682564°, Lon: -89.164857° At Stop 2E, the slope has much more vegetation with limited apparent slope movement. Figure 25 Stop 2E, slope with more vegetation and limited slope movement. Stop 2F: Middle Military Hill with Some Vegetation and Slope Movement Directions: From Stop 2E go 0.34 miles south on US-45 Stop 2F, a slope with vegetation that had was excavated for the construction of US-45. Lat: 46.677892°, Lon -89.166195° At Stop 2F, there is again less vegetation but more slope movement Figure 26. The clay soils at this elevation are not varved. Figure 26: Stop 2F showing partial vegetation and upper slope movement. 190 �Stop 2G: Upper Military Hill with Active Slope Movement on both Sides of Highway Directions: From Stop 2F go 0.23 miles south on US-45 Stop 2G, a slope with active slope movement. Lat: 46.674640°, Lon -89.167126° At Stop 2G, landslides occur on both sides of the highway as shown in Figure 27. Figure 27(a) shows slope movement on the west side of US-45, which has less movement than on the eastside of US-45 shown in Figure 27 (b), (c) and (d). These landslides have been moving for many years and appear to be in a non-varved clay. While limited soil investigation has been conducted in these sediments, it appears that at the higher elevations in the Military Hill area, a non-varved lacustrine soil overlies the varved clay soils indicating that it possible that this sequence might indicate when Glacial Lake Duluth and Lake Ontonagon combined. An interesting feature of Glacial Lake Ontonagon is that it is at a much higher elevation at 1,320 feet than Glacial Lake Duluth. Isostatic rebound would have had to occur for the lakes to combine. Figure 27 Stop 2G near the top of Military Hill showing slope movements on both sides of the highway with (a) west side, (b) top of east side, (c) south section of slope movement on east side, and (d) north portion of slope movement on the east side. 191 �Stop 3: US-45 Recent Excavation and Resulting Slope Movement Directions: From Stop 2G go 3.4 miles south on US-45 Stop 3, a slope with active slope movement. Lat: 46.627240°, Lon -89.178520° Stop 3 is at a recent location where MDOT excavated a natural slope to improve drainage along the eastside of US-45. The natural slope was excavated at a 2H:1V angle, the same slope as the highway embankment as can be seen in Figure 28. Soon after excavation, however, the slope started to move and has been moving ever since. Figure 29 was taken one year after the photo in Figure 26. The US-45 embankment was constructed with the local clay soil but was compacted, whereas the natural slope was not. The clay soils at this site and at Stop 2G have relatively high plasticity and are at the interface of the low (CL) and high (CH) plasticity soils defined by the Unified Soil Classification System (USCS) with moisture contents in the 20 to 25% range. As we saw at Stop 2F, the lacustrine clay soils are not stable at the angles at which they were excavated, unless compacted. Figure 28 Stop 3 showing a recent excavation into the clay soil at a 2(H):1(V) angle. Figure 29 Stop 3 - excavated slope one year after the photo in Figure 26 was taken. 192 �References Cited Attig, J.W., Clayton, L., and D.M. Mickelson, 1985. Correlation of late Wisconsin glacial phases in the western Great Lakes area, Geological Society of America Bulletin, v. 96, p. 1585-1593. Black, R.F., 1969, Valderan glaciation in Western Upper Peninsula, Proc. 12th Conf. Great Lakes Res., Internat. Assoc. Great Lakes Res., p. 116-123. Cannon, W.F, Nicholson, L.G., Woodruff, C.A, Hedgman, C.A. and K.J. Schulz, 1995. Geologic Map of the Ontonagon and Part of the Wakefield 30’ x 60’ Quadrangle, Michigan, USGS. Creech, C., Selegean, J., and T. Dahl, 2010. Historic and modern sediment yield from s forested watershed, 2nd Joint Federal Interagency Conference, Las Vegas, NV, June 27 - July 1. Dyl, Stanley, 1979. Engineering geologic factors affecting the stability of slopes in the Ontonagon Clay at the Military Hill Slide, U.S. Highway 45, Ontonagon County, Mi. In: M.S. Thesis. Michigan Technological University, pp. 92. Farran, W.R and C.W. Drexler. 1985. Late Wisconsinan and Holocene History of the Lake Superior basin. In Quaternary Evolution of the Great Lakes, eds. P.F. Karrow and P.E. Calkin, Geological Association of Canada Special Paper 30:17–32. Gunderman, B.J. and E. A. Baker, 2008. Ontonagon River Assessment, Michigan DNR, Fisheries Division, Special Report 46, Hack, J.T. Postglacial drainage evolution and stream geometry in the Ontonagon area, Michigan. US Geological Survey, 1965. Jerome, D.S., 2006. Landforms of the Upper Peninsula of Michigan. 1:750,000. USDA Natural Resources Conservation Service, pp. 17. Koons, G.J., 1969. Some geologic and engineering properties of the Pleistocene Ontonagon Clays at Victoria, Ontonagon County, Mi. In: M.S. Thesis. Michigan Technological University, pp. 110. Leverett, Frank, 1928, Moraines and shorelines of the Lake Superior region: U.S. G. S. Prof. Paper 154A, p. 1-72. Ontonagon River. (2023, Nov. 12). In Wikipedia. https://en.wikipedia.org/wiki/Ontonagon_River Peterson, W. L., 1985, Surficial geologic map of the Iron River 1° x 2° quadrangle, Michigan and Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map I-1360-C, scale 1:250,000. Peterson, W.L., 1986. Late Wisconsinan glacial history of northeastern Wisconsin and western upper Michigan. USGS Bulletin 1652, p. 1-14. Smith, J., 2012. Large scale landslide on the Ontonagon River, Michigan. In: M.S. Report. Michigan Technological University. http://digitalcommons.mtu.edu/etdrestricted/146. The Free Library. S.V. 2024. Through the wilderness, Retrieved Mar 19 2024 from https://www.thefreelibrary.com/Through+the+wilderness.-a0452051967. USACE, 2010. Ontonagon River Watershed 516e Sediment Study, USACE Great Lakes Hydraulics and Hydrology Office, Detroit District, p. 90. 193 �Weidner, L. DePrekel, K., Oommen, T. and Vitton, S., 2019, Investigating large landslides along a river valley using combined physical, statistical, and hydrologic modeling. Engineering Geology, v. 259, p. 1-12. http://doi.org/10.1016/j.enggeo.2019.1051169 194 � 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, 2024 Description An account of the resource Institute on Lake Superior Geology. Houghton, Michigan. May 15-18, 2024. 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 2024-05 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. 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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 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/c8c4c3103eec111ade66a79fd1788f93.pdf 21961203f4cc0eba4f8e4a0fe3ee69de https://digitalcollections.lakeheadu.ca/files/original/522bc7409122f20f5388a823c7517dab.pdf 665265caad9a54ea311fb8a943ae7153 PDF Text Text 68th ANNUAL MEETING Sudbury, Ontario — May 10-11, 2022 INSTITUTE ON LAKE SUPERIOR GEOLOGY Part 2 — Field Trip Guidebook �Thank you to our sponsors! INDIVIDUAL CONTRIBUTORS TO STUDENT TRAVEL SCHOLARSHIP: MARY KAY ARTHUR, AL MACTAVISH, MARK & LAURIE SEVERSON JIM DEGRAFF, MICHAEL & MONICA EASTON, DICK HEGLUND JIM DEGRAFF, BOB MAHIN, MIKE BEAUREGARD JOANNA HODGE, TERRY BOERBOOM, JIM MILLER BEN BERGER, DEAN PETERSON, GRAHAM WILSON �Proceedings of the 68th ILSG Annual Meeting – Part 2 68th ANNUAL MEETING INSTITUTE ON LAKE SUPERIOR GEOLOGY May 10-11, 2022 Sudbury, Ontario HOSTED BY Michael Easton and Wouter Bleeker Co-Chairs Ontario Geological Survey and Geological Survey of Canada Proceedings - Volume 68 Part 2 – Field Trip Guidebook Compiled and edited by Michael Easton Cover Photos. Upper Left — Signage at the Sudbury ore discovery site (Trips 1 and 4). Upper Right— Footwall Breccia in the Crean Hill Mine area (Trip 3). Lower Left — Arkosic sandstone overlain by matrix- to clast-supported conglomerate, Gowganda Formation. On Highway 108 north of Elliot Lake (Trip 5). Lower Right — Deformed, migmatitic gneiss in the Grenville Front tectonic zone showing garnet dispersed throughout the rock. Scale card is 9 cm long. On Highway 537 southeast of Sudbury (Trip 2). �Proceedings of the 68th ILSG Annual Meeting – Part 2 �Proceedings of the 68th ILSG Annual Meeting – Part 2 68th INSTITUTE ON LAKE SUPERIOR GEOLOGY VOLUME 68 CONSISTS OF: PART 1: PROGRAM AND ABSTRACTS PART 2: FIELD T RIP GUIDEBOOK Trip 1: A TRAVERSE ACROSS THE SUDBURY IMPACT STRUCTURE Trip 2: GEOLOGY OF THE GRENVILLE FRONT AND THE GRENVILLE FRONT TECTONIC ZONE IN THE SUDBURY AREA Trip 3: MAGMATISM AND BRECCIATION IN THE FOOTWALL ROCKS IN THE SOUTHWESTERN SUDBURY STRUCTURE Trip 4: AN OVERVIEW OF THE GEOLOGY OF THE SUDBURY STRUCTURE Trip 5: A CROSS-SECTION THROUGH THE HURONIAN SUPERGROUP AT ELLIOT LAKE, ONTARIO Reference to material in Part 2 should follow the example below: Gordon, C., Généraux, C-A. and Clarke, B. 2022. Magmatism and Brecciation in the Footwall rocks of the southwestern Sudbury Structure; in Easton, R.M. (Ed.), Institute on Lake Superior Geology Proceedings, 68th Annual Meeting, Sudbury, Ontario, Part 2 – Field trip guidebook. v.68, part 2, p.147-180. Published by the 68th 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 68th ILSG Annual Meeting – Part 2 Part 2: Field Trip Guidebook Table of Contents Introduction, safety considerations and acknowledgements 1 Field trip 1 — A traverse across the Sudbury Structure Earth’s largest preserved impact crater (2 days) 2 Field trip 2 — Geology of the Grenville Front and the Grenville Front tectonic zone in the Sudbury area 102 Field trip 3 — Magmatism and brecciation in the Footwall rocks of the southwestern Sudbury Structure 147 Field trip 4 — Overview of the Sudbury Structure 182 Field trip 5 — A cross-section through the Huronian Supergroup at Elliot Lake 200 Figure 1. Map showing the location of the five field trips offered in 2022. vi �Introduction, safety considerations and acknowledgements Michael Easton Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, Ontario P3E 6B5 and Wouter Bleeker Geological Survey of Canada, 601 Booth Street Ottawa, Ontario K1A 0E8 Sudbury is located near the boundary between 3 major geological provinces (the Archean Superior Province, the Paleoproterozoic Southern Province, and the Mesoproterozoic Grenville Province) and the largest preserved ancient impact crater on Earth. Consequently, it is an ideal setting for geological field trips. Despite the impact of Covid19 related health-measures on meeting planning and organization, 3 pre-meeting and 2 postmeeting field trips were available for delegates to the 68th Annual Institute on Lake Superior Geology (ILSG) meeting in 2022 (see Figure 1, opposite). In the case of Trips 1 and 3, some stops are on property owned by mining companies, who granted special permission to the ILSG trip leaders and participants to access these properties. It will not be possible for the average guidebook user to revisit these stops. 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 (Manuel Duguet, Peter MacDonald, Alinda Aubin, Matthew Eles, and Monica Easton). In addition, the efforts of Johanne Roux and Carlo Casrechino in producing the guidebooks in a timely fashion is greatly appreciated. This volume is intended to serve not only as a guide for 68th 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/or staked land, please be sure to obtain the land owners’ permission before entering their land. Contact the staff of the Resident Geologist Program of the Ontario Geological Survey in Sudbury for current ownership information. We also appreciate the assistance and cooperation of the exploration and mining companies in providing access and information concerning their properties, particularly Vale Canada, Limited Crean Hill Mine, (Whistle Mine), Lonmin PLC, Wallbridge Mining Company Limited (Parkin, Trill, Hess), SPC Nickel Corporation (Worthington), KGHM International (Podolosky), Sudbury Integrated Nickel Operations (a Glencore Company) and North American Nickel (distal Whistle). We also thank the City of Greater Sudbury for providing access to a trip stop on field trip 4. The field trips, for the most part, will be visiting stops along either major highways or municipal roads. Please take care when crossing or parking along these roads. 1 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Frontispiece: Classical “shatter cones” in the shocked footwall and target rocks of the 1850 Ma Sudbury impact crater, well developed in quartzites of the circa 2.4 Ga Mississagi Formation. With the newly recognized knowledge, in the mid- to late-1950s, that these conical, radiating fracture surfaces represent unique “trace fossils” for the high-velocity, extremely high-pressure shock waves associated with meteorite impacts (Dietz, 1959), Sudbury was quickly recognized, in 1962, as an astrobleme—the scar of an ancient impact crater (see Dietz, 1964; and Dietz and Butler, 1964). 2 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Field Trip 1 – A Traverse Across the Sudbury Structure—Earth’s Largest Preserved Impact Crater Wouter Bleeker Ontario Geological Survey of Canada, 601 Booth Street Ottawa, Ontario K1A 0E8 Sandra Kamo Jack Satterly Geochronology Laboratory, University of Toronto, 22 Ursula Franklin Street, Toronto, Ontario M5S 3B1 Henning Seibel and Michael Lesher Mineral Exploration Research Centre, Harquail School of Earth Sciences, Laurentian University, 935 Ramsey Lake Road, Sudbury, ON P3E 2C6 This two-day field trip involves a traverse across the deformed 1850 Ma Sudbury impact structure, from the older country rocks and brecciated target rocks to the south of the preserved melt sheet, across the entire folded impact structure, the melt sheet, and crater fill, and onto its northern rim. Day 1 will concentrate on the regional cross-section/traverse and will introduce many of the key aspects—and controversies— associated with the structure. Day 2 will finish the regional traverse and allow time to examine some of the ore environments in more detail, along the basal contact of the differentiated melt sheet and into the brecciated footwall. Note 1: Ore environments to be examined are dependent on company approval to access properties, as well as the evolving situation regarding COVID-19. Note 2: This trip is a more detailed examination of the Sudbury impact structure than that offered by post-meeting Trip #4. Trip #4 provides more of an overview of the structure. 3 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Day 1 Wouter Bleeker1 and Sandra Kamo2 1 2 Geological Survey of Canada, Ottawa Jack Satterly Geochronology Laboratory, University of Toronto 1. Introduction This two-day fieldtrip will involve a traverse across the deformed Sudbury structure, one of the world’s largest and oldest preserved meteorite impact structures: from deformed and brecciated older country rocks in the south, across the folded impact structure, the melt sheet and crater fill, onto the northern rim, and into the footwall rocks below (Figure 1). Day 1 will focus on the regional cross-section and traverse and will introduce many of the key aspects—and controversies—associated with this unique structure. Day 2 will finish the regional traverse and allow time to examine some of the ore environments in more detail, along the basal contact of the differentiated melt sheet and into the brecciated footwall. Due to the on-going but hopefully waning Covid-19 situation, the issue of access to company properties remains somewhat fluid and some last-minute changes may have to be made in terms of field trip stops and localities. Nevertheless, this two-day trip will allow most major points of interest regarding the Sudbury impact structure to be addressed and discussed. With well over a century of geological research in the area, the literature on the Sudbury structure is voluminous (e.g., Coleman, 1905; Thomson, 1956; Hawley, 1962; Dietz, 1964; Souch et al., 1969; Naldrett et al., 1970; Brocoum and Dalziel, 1974; Krogh et al., 1982; Pye et al., 1984 and all contributions therein; Faggart et al., 1985; Grieve et al., 1991; Dickin et al., 1992; Butler, 1994; Wu et al., 1995; Spray et al., 2004; Lightfoot and Zotov, 2005; Zieg and Marsh, 2005; Ames et al., 2002, 2008a,b; Bleeker et al., 2015; Papapavlou et al., 2018; and numerous other papers listed in the reference list). Yet many key questions remain. For instance, how big was the original impact structure? And where was “Ground Zero”, i.e. the centre of the impact? Is there a global Sudbury fall-out layer and, if so, where is it? How much of the melted footwall geological heterogeneity has been inherited in the differentiated melt sheet? Given the size of the structure and its transient crater, did it trigger any mantle melting? And, with respect to the complex spectrum of observations on the ores and the footwall rocks: where do impact processes and cratermodification processes stop, and where do regional deformation and hydrothermal-metamorphic processes take over? In the context of an Institute of Lake Superior Geology (ILSG) fieldtrip, one of the questions raised above is particularly pertinent: is the accretionary lapilli layer recognized at the disturbed top of the Gunflint Formation in the Lake Superior area indeed the Sudbury event and fall-out layer (Figure 2), as is permissible and as has been argued based on current evidence (e.g., Addison et al., 2005, 2010; Cannon et al., 2010), or does that horizon represent a different event layer? From an ECREE perspective ("extraordinary claims require extraordinary evidence", Carl Sagan), a unique link of this prominent event layer to the Sudbury area still remains to be documented, and other localities of an 1850 Ma global event layer need to be demonstrated (Figure 2). Among the largest preserved impact craters on Earth (Dietz, 1964; Dietz and Butler, 1964), Sudbury’s very thick (~3–5 km) and differentiated melt sheet (the “Sudbury Igneous Complex”, hereafter SIC; Pye et al., 1984 and contributions therein) is unique. Why? Perhaps this can be reasoned away by other comparable melt sheets not being preserved (e.g., Vredefort) or not yet fully explored (Chicxulub), or perhaps was Sudbury bigger than the current consensus (~200 km final crater diameter; see Grieve et al., 1991; Grieve, 1994; Butler, 1994; Spray et al., 2004), and is it in a class of its own? 4 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 1: Geological map of the Sudbury area centered on the erosional remnants of the deformed Sudbury impact structure, the Sudbury Igneous Complex (SIC); after Bleeker et al. (2015), and adapted from Dressler (1984) and Ames et al. (2005). Day 1 fieldtrip stops are indicated, as is the area we will visit on Day 2 (and a possible alternate area for Day 2). The arcuate red dashed line marks the outer limit of observed shatter cones in the footwall. The circa 1.0 Ga Grenville Front truncates the area in the southeast. 5 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Finally, Sudbury is fascinating from a science history perspective: how it was discovered and how its interpretation was slow to change from, originally, a largely mafic igneous complex or lopolith — “the Sudbury Irruptive” — in part or wholly derived from mantle melting (e.g., Wilson, 1956; prior to the work by R. Dietz in the early 1960s); to a hybrid igneous and impact-generated structure (see, for instance, the various contributions in Pye et al., 1984; see also Dietz, 1964 1); and, finally, to an impact-only crater, melt sheet, and crater fill complex with little to no mantle input and merely modified by deformation (e.g., Stöffler et al., 1989; Grieve et al., 1991; Grieve, 1994). With respect to the latter perspective, the pendulum only changed in the late 1980s and early 1990s when detailed isotopic data became available showing that most if not all of the melt sheet was derived from melting of the crust (e.g., Faggart et al., 1985; Stöffler et al., 1989; Walker et al., 1991; Dickin et al., 1992; Deutsch, 1994 and references therein). From this slowly evolving historical perspective, the brilliant early 1960s papers by Robert Dietz, who in 1962 quickly confirmed “shatter cones” in the footwall rocks around the Sudbury complex (see frontispiece of this guidebook), and from there confidently posited a meteorite impact origin, stand out as even more remarkable. 2 This field trip will allow participants to become familiar with this amazing structure and to discuss and debate all these first-order questions. The current fieldtrip guidebook should be viewed as a preliminary offering, as many details remain to be expanded on. 2. The Sudbury Structure: Geological Setting The deformed Sudbury impact structure, now preserved as a broadly doubly plunging, synclinal, erosional remnant ~60 km long by ~30 km wide, is situated in the southern Canadian Shield, approximately on the boundary between two main structural provinces, the Archean Superior Province to the north and the Paleoproterozoic Southern Province to the south (Figures 1 and 2). The former is represented by the circa 2.85-2.64 Ga granite-greenstone terrain of the southern Superior craton, whereas the latter is dominated by moderately to tightly folded and faulted Paleoproterozoic strata of the circa 2.50-2.30 Ga Huronian Supergroup (Young, 1973 and contributions therein; Bennett et al., 1991; Young et al., 2001; Rasmussen et al., 2013) overlying Superior craton basement (Figures 3 and 4). The Southern Province trends roughly E-W and, to the west, extends into the Lake Superior area where it is known as the Penokean fold belt, part of the larger 1.87-1.83 Ga Penokean orogen (e.g., Brocoum and Dalziel, 1974; Schulz and Cannon, 2007). Just to the south of the city of Sudbury, the Paleoproterozoic Southern Province fold belt is truncated by the relatively sharply defined deformation front of the circa 1.1-1.0 Ga Grenville Front sensu stricto (see front #4, Figure 2). This front trends from northeast to southwest and represents the northern boundary of the complex, multi-cyclic and very extensive Grenville orogen that rims Proterozoic Laurentia to the southeast and played major role in building the late Proterozoic supercontinent Rodinia. In more detail, the Grenville Front sensu stricto represents, in the Sudbury area, the relatively well-defined deformation front of the terminal collisional phase of the Grenville orogen, an orogen that also involves older Paleo- to Mesorproterozoic events, some of which are represented by rocks units right along the Grenville Front (e.g., see front #3 in Figure 2). Although quickly recognizing the fundamental impact origin of the structure in 1962, based on his identification of shatter cones in the footwall rocks, in his seminal 1964 paper Dietz maintains an intrusive origin for much of the Sudbury Igneous Complex. 1 2 A bibliography of relevant papers by Dietz is included at the beginning of the reference list and they make for an interesting read. After convincing himself, in the mid-1950s, that shatter cones were a key indicator for high-energy meteorite impacts, he quickly clarifies a large number of impact structures, such as Sudbury, which were previously seen as enigmatic and attributed to “crypto-volcanic” (i.e. endogenic) processes. In 1961 he coined the term “astroblemes” for these structures (i.e., “star wounds” or impact scars; Dietz, 1961) to highlight the fundamental role of impact processes here on Earth. 6 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 2: General location of the deformed Sudbury impact structure on the approximate boundary of the Archean Superior Province and the Paleoproterozoic Southern Province. The latter is broadly part of the “Penokean fold belt” and the associated orogen that rims the southern margin of the Superior craton (e.g., Card et al., 1972; Brocoum and Dalziel, 1974; Card et al., 1984; Holm et al., 2007; Schulz and Cannon, 2007). Structural fronts younger than the circa 1.87-1.84 Ga Penokean sensu stricto (front #1) also contributed to the deformation of the Southern Province (i.e. fronts #2 and 3). South of Sudbury, the Southern Province is truncated by the main circa 1.1-1.0 Ga Grenville Front (front #4; Davidson, 1997). Also shown is the approximate footprint of the circa 0.6–0.5 Ga Ottawa-Bonnechere Graben (in light grey), which is manifested in the Sudbury area by E-W-trending olivine-bearing diabase dykes. Red star symbols indicate the localities where the accretionary lapilli event layer at the top of the Gunflint Formation and correlative strata has been identified (e.g., Cannon et al., 2010). The figure also draws attention to other circa 1870-1840 Ma sequences preserved in the Canadian Shield, only at marginally larger distances than the Mesabi Range locality, where a comparable Sudbury fall-out layer remains to be identified. Map adapted from Wheeler et al. (1996), Young (1983), and Cannon et al. (2010). 7 �Proceedings of the 68th ILSG Annual Meeting - Part 2 A simplified NNW-SSE cross-section through the area illustrates the main pre-impact, litho-tectonic elements that define the wider Sudbury area (Figure 3). A summary stratigraphic column, highlighting the main features of the Paleoproterozoic Huronian Supergroup in more detail, is shown in Figure 4. The age range of the Huronian Supergroup is constrained by the onset of rifting and associated large-scale mafic magmatism at circa 2500-2480 Ma (Krogh et al., 1984), defining its base, and final emplacement of extensive mafic sill complexes at circa 2250 Ma (May Township sills, Bleeker et al., in prep.) and circa 2217 Ma (Nipissing sills; Corfu and Andrews, 1986; Noble and Lightfoot, 1992; Bleeker et al., 2015; Davey et al., 2019). The latter intrusive episodes provide a minimum age for the supergroup. Thin felsic ash layers in the fine-grained sedimentary rocks of the Gordon Lake Formation, in the upper part of the supergroup, have an approximate age of circa 2308 Ma, based on SHRIMP dating of zircons (Rasmussen et al., 2013) on a sample obtained from drill core through the formation. Figure 3: A simplified cross-section through the Sudbury area from NNW to SSE, showing the main lithotectonic elements that define the geology of the area, prior to Penokean and younger deformation and prior to the 1850 Ma impact. Note the lower Huronian rift structure and associated rift fill (the Elliot Lake Group, E), which is well developed and exposed in the Sudbury area, in part due to impact-induced uplift and exhumation. The estimated position of “Ground Zero” is indicated above the section, based on both ringlike structures in the foreland (Butler, 1994) and a statistical intersection of shatter cone axes (Bleeker, unpublished). A circa 2.0 Ga passive margin sequence, comparable to the Gunflint Formation in the Lake Superior area, is indicated for general comparison only but is not preserved in the Sudbury area. Postimpact, the area was covered by the depositional wedge of Penokean foreland basin deposits, the upper shale dominated Onwatin Formation and the turbiditic Chelmsford Formation, which are preserved only within the doubly plunging syncline of the “Sudbury Basin”. These turbiditic deposits are comparable and likely directly correlative to the Rove Formation of the Lake Superior area. The three stars on the right side of the section indicate glacially-influenced formations of diamictite, from bottom to top: the Ramsay Lake, Bruce, and Gowganda Formation diamictites. Based on global correlations and U-Pb age dating in the Transvaal Basin of South Africa, we know that the Ramsay Lake Formation glacial episode ended at circa 2426 Ma (Gumsley et al., 2017). 8 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4: Summary stratigraphic column of the 2.50-2.30 Ga Paleoproterozoic Huronian Supergroup, unconformably overlying Archean basement of the southern Superior craton. Compiled from numerous sources: Young (1973) and contributions therein; Bennett et al. (1991); Young et al. (2001); Gumsley et al. (2017). Age data from: Krogh et al. (1982) and Krogh et al. (1984); Corfu and Andrews (1986); Noble and Lightfoot (1992); Heaman (1997); Rasmussen et al. (2013); Bleeker et al. (2015); Davey et al. (2019). 9 �Proceedings of the 68th ILSG Annual Meeting - Part 2 3. More on the 2.50-2.30 Huronian Supergroup Age range, thickness, Archean-Proterozoic boundary The Huronian Supergroup (Figure 4) is of more than just local interest. With its extended age range and ~6–12 km total stratigraphic thickness, it represents one of the principal, and best-preserved Paleoproterozoic stratigraphic records in the world, perhaps only rivalled by the Transvaal Supergroup of southern Africa, and correlative units of the Hamersley Basin overlying the Pilbara craton. Its basal unconformity and low-grade metamorphic state, overlying Archean granite-greenstone terrain, played a major role in the debate and final definition of the Archean–Proterozoic boundary, one of the most fundamental boundaries of the terrestrial geological time scale. GOE: the Great Oxidation Event, pyritic placer deposits, continental red beds Equally important from a historical point of view, the Huronian Supergroup straddles the circa 2.4–2.3 Ga “Great Oxidation Event” (GOE, see Figure 4; e.g., Roscoe, 1973 3; Holland, 1978, 1984, 2002; Prasad and Roscoe, 1996; Bekker et al., 2004; Gumsley et al., 2017), which marks the initial rise of atmospheric oxygen across the critical threshold above which oxidized surface environments promoted the deposition of hematite-stained red sandstones (“red beds”). The world’s oldest genuine terrestrial red beds occur in the Cobalt Group of the upper Huronian, specifically the Gowganda and overlying Lorrain formations. The elevated oxygen levels no longer allowed prolonged preservation of detrital pyrite and associated heavy minerals, in terrestrial sedimentary environments, thus resulting in pyrite placer deposits largely disappearing from the geological record. Hence, no such deposits are known from the upper Huronian Supergroup, but, with associated gold and uraninite, they form minor but important components of the lower part of the supergroup that was deposited prior to the GOE transition. Pyrite-uraninite placers form important placer deposits in the basal Matinenda Formation (Figures 3 and 4) and were mined extensively for uranium in the Elliot Lake area west of Sudbury. Gold-bearing pyrite placers are currently the subject of active exploration northeast of Sudbury and are thought to be part of the conglomerates and cross-bedded sandstones at the base of the Mississagi Formation (Long et al., 2011; Whymark and Frimmel, 2018). The emergence of these important insights involving atmospheric evolution, increasing oxygen levels, detrital pyrite, and the first red beds were largely based on work in the Huronian Supergroup and correlated sequences, such as the Snowy Pass Supergroup overlying the Wyoming craton (e.g., Roscoe and Card, 1992, 1993; Prasad and Roscoe, 1996). GOE: disappearance of the mass-independent sulphur isotope fractionation signature Correlated with the observable transition “detrital pyrite out, red bed sandstones in”, more recent research has shown that there is also a fundamental shift in global S-isotopic signatures, particularly the disappearance of anomalous “mass-independent fractionation” (MIF) of the 33S isotope relative to 32S and 34 S isotopes (Farquhar et al., 2000; Farquhar and Wing, 2003). This MIF signature of 33S is understood to result from UV-induced processes in the upper atmosphere and this MIF signal can only survive and be transmitted to the sedimentary record at very low total oxygen atmospheric levels, estimated at <10-5 times present atmospheric level (PAL) of oxygen (Farquhar and Wing, 2003). These insights have revolutionized the understanding of both the shallow and deep sulphur cycles in the last two decades and shed new light 3 Although now widely referred to as the “Great Oxidation Event” (GOE), following Holland and others, Stew Roscoe through his work in the Elliot Lake area, on the pyrite-uraninite placer deposits there, was one of the first to draw attention to this important transition and called it the “oxyatmoversion”. In his 1973 paper he also draws attention to the obvious potential of this transition in terms of a first-order time scale boundary. 10 �Proceedings of the 68th ILSG Annual Meeting - Part 2 on a host of fundamental recycling processes. As shown in Figure 4, detrital pyrite and the MIF-S signature are absent above the Mississagi Formation. Glacial episodes, Snowball Earth events, cap carbonates Stratigraphic and sedimentological studies of the Huronian Supergroup also played a key role in the recognition of early Precambrian glacial deposits, and possibly global-scale glacial events. Unsorted diamictite deposits, almost certainly tillites or reworked tillites, are recognized at three stratigraphic levels within the Huronian Supergroup (Figure 3 and 4): in the Ramsay Lake Formation, the Bruce Formation, and in the post-GOE Gowganda Formation. Some of these diamictite deposits have other associated features that confirm a glacial origin, such as dropstones in overlying varve-like siltstones, and glacially polished clasts with striae. This is certainly the case for the Gowganda Formation (Young, 1983; Young and Nesbitt, 1985; Young et al., 2001). Paleomagnetic evidence places the Superior craton at low latitudes in the earliest Paleoproterozoic (e.g., Evans and Hall, 2010; Salminen et al., 2014). This, together with glacial deposits in what was at least in part a marine basin, thus places glacial deposits at sea level, and close to the equator and far from the poles. In turn, observations such as these argue for glacial events of global significance, and events that can be correlated to other cratons and similar sedimentary successions such as the Transvaal basin of southern Africa (e.g., Gumsley et al., 2017). This is particularly relevant for the second of the three glacial events represented by the Bruce Formation diamictites. This formation is overlain by the only carbonates in the Huronian succession, the Espanola Formation (see Figure 4). These carbonates may represent deposits that are thought to have formed in response to rapid deglaciation of global ice cover and, in this context, are referred to as “cap carbonates” (e.g., Kirschvink, 1992; Hoffman et al., 1998; Hoffman and Shrag, 2000; Kirschvink et al., 2000). If so, these carbonates would allow global correlations and constitute an ideal, globally synchronized time scale boundary, as they do in the Neoproterozoic. Epicontinental rift and sag succession or passive margin sequence? Various authors have, somewhat uncritically, referred to the Huronian Supergroup as a rift and passive margin sequence, associated with the breakup of the Superior craton. That implies that the Superior craton broke up along its southern margin sometime around 2.4 Ga following the formation of the lower Huronian rift succession. There is, however, little evidence for this early breakup. Clearly the Superior craton, with its present outline, is just a fragment of a much larger ancestral Archean supercraton, which Bleeker (2003, 2004) has referred to as “Superia”. Breakup of Superia was clearly progressive but may have only started with the very extensive circa 2.22 Ga Nipissing and related mafic magmatic events or large igneous provinces (LIPs), and likely even later along the southern margin of the craton. The specific LIP event that most likely represents initiation of breakup along the southern margin of the Superior is the 2125-2100 Ma Marathon event (Halls et al., 2008; Davey et al., 2020, 2022), with very extensive mafic dyke swarms projecting north into the craton and bimodal magmatism along the southern margin of the craton. The Wyoming and Karelia-Kola cratons are among the cratonic fragments that broke away at that time, as they can be matched to the southern Superior, based on multiple lines of evidence, prior to 2.1 Ga breakup (e.g., Roscoe and Card, 1993; Bleeker and Ernst, 2006; Kilian et al., 2016a, b). This alternative scenario argues for the entire Huronian Supergroup to represent a long-lived intra- and epicontinental rift and sag basin, without a proximal passive margin prism. Consequently, this puts the Huronian Supergroup in a different perspective and has implications for global reconstructions and compilations, such as those of successions thought to represent passive margin sequences through time (Bradley, 2008). Similar arguments very likely apply to other Proterozoic sequences that have been too easily characterized as passive margins. 11 �Proceedings of the 68th ILSG Annual Meeting - Part 2 4. Large Igneous Provinces (LIPs) and Mafic Magmatic Events of the Wider Sudbury Area Mafic magmatic events Here we briefly list and describe some of the large mafic magmatic events in the wider Sudbury area. These events are important as temporal and structural markers, either pre-dating certain structural events, or the impact event itself, or post-dating such events and thus providing minimum age constraints. From old to young, a brief summary of these magmatic events would include (e.g., Krogh et al., 1984; Kamo et al., 1995; Heaman, 1997; Noble and Lightfoot, 1992; Ernst and Bleeker, 2010; Bleeker et al., 2015): • • Major pyroxenite dykes: circa 2507 Ma, perhaps part of the broader Mistassini event? These dykes, which occur just north of the SIC, may indicate the onset of the Huronian rifting event (Bleeker et al., 2015). Matachewan-I: circa 2480 Ma layered intrusions and sills, dykes; the East Bull Lake Suite. Matachewan-II: circa 2460 Ma major diabase swarm, main pulse of Matachewan dykes. May Township sill complex: circa 2250 Ma, large sills to the west of the Sudbury area. Nipissing sills and Senneterre dykes: circa 2217 Ma, volumetrically important in the immediate Sudbury area. Biscotasing dykes: 2167 Ma dykes, mostly north of Sudbury. Marathon dykes: circa 2110 Ma dykes, mostly west of Sudbury. Lauzon Lake dykes: circa 1950 Ma, major NW-trending dykes west of Sudbury. Alkaline dykes, lamprophyres and carbonatite intrusions, circa 1880 Ma; e.g. the Spanish River Complex north of Sudbury. Thin mafic sills/subhorizontal sheets: undated and of yet unknown significance. Lamprophyre dykes: undated and of yet unknown significance; probably more than one event. • Sudbury impact event and associated dykes: 1850-1849 Ma. • Trap dykes: circa 1750 Ma, numerous E-W trending diabase dykes cutting across the Sudbury South Range, still affected by metamorphism. Alkaline intrusions in the wider area, e.g. the Croker Island Intrusion, southwest of Sudbury, Mesoproterozoic in age. Larder Lake dykes: circa 1270 Ma, undeformed olivine diabase dykes similar to Sudbury dykes. Sudbury dykes: circa 1235 Ma, abundant NW-trending undeformed olivine diabase dykes cutting across the area and the Sudbury structure, but themselves cut and truncated by the Grenville Front. Grenville dykes: circa 590 Ma, E-W-trending dykes cutting across the Sudbury structure and Grenville Front, and a manifestation of the extensive Ottawa-Bonnechere Graben system. Kimberlites and associated intrusions in the wider area, Jurassic to Cretaceous. • • • • • • • • • • • • • • Some of the pre-impact mafic magmatic events are rather voluminous in the Sudbury area, and thus a significant component of the overall target rocks. This is particularly true for: 1) the early Matachewan (I) mafic rocks, which form mid-sized layered intrusions in the general area (e.g., James et al., 2002), at or close to the Archean–Paleoproterozoic (i.e. basal Huronian) unconformity (Figure 3), and less voluminous dykes and sills; and 2) the very extensive Nipissing Diabase sill complex, which invaded large parts of the Huronian Supergroup stratigraphy as well as the basement immediately below the unconformity (see Figures 3 and 4). Some of these units have minor, to locally significant, Ni-Cu-PGE magmatic sulphide mineralization; for instance, the Shakespeare Intrusion that is part of the Nipissing magmatic event (e.g., Sproule et al., 12 �Proceedings of the 68th ILSG Annual Meeting - Part 2 2007; Davey et al., 2019). For these reasons (overall volume, sulphur (S), metals (Ni, Cu, PGEs)), they have featured and continue to feature in the debate on metal sources and elemental mass balances in the differentiated Sudbury melt sheet and the associated magmatic sulphide ores. In other words, did certain geological units in the target rocks play an important role in the sulphur and metal budgets of the impact melt sheet, and thus in early sulphur saturation and in the overall metal endowment in the orebodies, or are such roles insignificant in the overall melt sheet evolution and the formation and segregation of the ores? Clearly, immediately following the impact, and the generation of a superheated melt sheet, early sulphide saturation was critical and can be clearly demonstrated (Figure 5); but overall metal budgets may not require enriched source rocks if sulphur saturation was indeed early and given the enormous volume of very hot impact melt available. In the latter scenario, it is simply all about the efficiency of early sulphide saturation, exsolution and segregation, sulphide droplet formation, and the “rain out” of sulphide globules enriched in chalcophile metals. Figure 5: A polished slab of Fe-Cu-Ni magmatic sulphide droplets, millimetres to centimetres in size, in ore from the proximal part of the Copper Cliff “offset dyke”. This dyke was injected into the footwall of the impact crater, during an early stage of the melt sheet evolution. The host rock consists of relatively unfractionated early mafic melt sheet material formed near the base of the melt sheet. Sulphide droplets were actively “raining out” (down in the picture) and physically interacting with mafic inclusions on their way down. Textures such as these provide clear evidence for early sulphide saturation in the melt sheet and that the “rain out” of exsolved sulphide melt droplets was the first-order process collecting magmatic sulphide ores at or near the basal contact of the melt sheet. Note the dumbbell structure of two merging sulphide globules near the top of the sample. 13 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Several of the intrusive events are also of key importance as structural markers. For instance, the extensive Nipissing Diabase sills within the Huronian succession, and locally in basement just below the Huronian unconformity, are folded with the strata of the Huronian Supergroup and, where well-exposed, sill contacts are largely concordant with the sedimentary layering in the immediate host rocks; although locally the sills are cross-cutting and dyke-like due to a “saucer-shaped” sill geometry (see Figure 3). However, the largely concordant nature of the observed contacts of many of the sills, together with a notable absence of clearly documented examples of such sills cutting across both limbs of previously folded Huronian strata, strongly favours an interpretation of the Huronian being largely unfolded, or perhaps only very weakly folded and/or locally tilted, at the time of widespread Nipissing sill emplacement at circa 2217 Ma (Bleeker et al., 2015). This important observation has major implications for the concept of the “Blezardian orogeny” and, ultimately, how to interpret the Sudbury structure (see below), and is fully supported by observations on the orientation of shatter cones. 5. The “Blezardian Orogeny”, Fact or Fiction? The lowermost Huronian Supergroup is intruded by a number of granite plutons in the Sudbury area, one of the major plutons being the Creighton Granite on the South Range, just west of Sudbury (Figure 1). As part of early attempts to systematize the geology of the Canadian Shield and divide this enormous territory into “structural provinces” based on regional deformation patterns and tectono-magmatic events (e.g., Wilson, 1949; Stockwell, 1982), late-stage granites were generally seen, and indeed correctly in many places, as the terminal phase of a major orogenic episode. As it had been clear since the early mapping of the southern Canadian Shield that the folded Huronian Supergroup unconformably overlies Archean basement that had previously been affected by major late Archean deformation and metamorphism, this grew into the concept of the two structural provinces: the Archean Superior Province to the north, affected and shaped by a terminal “Kenoran orogeny”, and a younger Proterozoic Southern Province and fold belt to the south, affected by deformation attributed to a “Blezardian orogeny” dated by the intrusion of the Creighton Granite and similar granite plutons. Early zircon dating attempts of these (shock-deformed!) granite bodies, based on large, multigrain, highly discordant zircon fractions (unabraded), suggested interpreted upper intercept ages of circa 2.35 Ga (Frarey et al., 1982); i.e. a circa 2.35 Ga Blezardian orogeny associated with granite magmatism that terminated the Huronian cycle of sedimentation and associated deformation. Perhaps reasonable at the time, we now know this interpretation is no longer tenable for the following reasons: 1) the granite plutons are lower Huronian rift-related A-type granites, dated at 2455–2460 Ma (Bleeker et al., 2015), not collisional granites as part of a terminal orogenic phase; 2) they only intrude the lowermost Huronian rift volcanics and are deformed with the lower Huronian strata they intrude; and 3) as explained above, the Huronian strata are extensively intruded by Nipissing Diabase sills at 2217 Ma, with most or all of the deformation post-dating the emplacement of the sills. From Bleeker et al. (2015): “A tightly folded Nipissing Diabase sill has been dated at 2215 ±1 Ma. It is fully conformable with surrounding Huronian strata on the South Range, inconsistent with the concept of a pre-Nipissing “Blezardian orogeny”. The main rationale for the Blezardian orogeny was the idea that deformation and intrusion of granite plutons, such as the Creighton Granite, thought to be circa 2.35 Ga in age, terminated the depositional history of the Huronian succession (Frarey et al., 1982; Stockwell, 1982). None of these ideas are supported by present evidence. The Creighton Granite is an early Huronian 2455– 2460 Ma rift-related granite, not an orogenic granite pluton; folding of the Huronian succession did not commence until well after emplacement of Nipissing Diabase sills and sheets with the onset of Penokean accretion and collision events at circa 1860 Ma. Other observations that have contributed to the concept of 14 �Proceedings of the 68th ILSG Annual Meeting - Part 2 a Blezardian orogeny can all be explained without a significant pre-Nipissing deformation event. For instance, saucer-shaped Nipissing sills locally may appear to crosscut Huronian strata and, after superimposed Penokean deformation, could easily lead to confusing field relationships.” With these caveats, does this leave any room for pre-Nipissing deformation of the Huronian? It seems reasonable that there was some local deformation and/or tilting related to movement on rift faults, perhaps local inversion. As shown in Figures 3 and 4, the Huronian succession consists of several major groups, and some of the boundaries between these groups may represent second order sequence boundaries and local, relatively low-angle unconformities. In this context, the basal contact of the aerially extensive Cobalt Group, overlain by coarse conglomerate and diamictite, is particularly relevant. The Cobalt Group far oversteps the lower groups of the Huronian and onlaps onto the Superior craton far to the north. In some areas northeast of Sudbury, careful examination of township geological maps suggests that a low-angle unconformity separates the Cobalt Group from the lower Huronian groups (Hough Lake and Quirke Lake Groups), as depicted in the sections of Figures 3 and 4 (see also, Meyn, 1973). These interesting relationships do not equate with a Blezardian orogeny, however. 6. Post-impact Deformation Beyond the observations discussed above, it is clear that most of the deformation and metamorphism of the Southern Province and the Sudbury area, locally intense, post-dates all of the Huronian Supergroup, the Nipissing Diabase sill emplacement, and also the final settling of the differentiated Sudbury melt sheet, as well as the deposition of the overlying Whitewater Group (Figure 6; see also Figure 3). The major deformation that folds the Sudbury structure, melt sheet, crater fill, and overlying Whitewater Group into a regional-scale, doubly plunging synclinal structure can be largely attributed to the Penokean orogeny (e.g., Card et al., 1972; Brocoum and Dalziel, 1974), amplified and overprinted to varying degrees, particularly on the South Range and toward the Grenville Front, by younger Proterozoic events (Shanks and Schwerdtner, 1991; Bailey et al., 2004; Papapavlou et al., 2017). Figure 6 (next page): Simplified lithologic-stratigraphic column for the Sudbury impact structure, its differentiated melt sheet, and the overlying Whitewater Group. Adapted and compiled from various sources, including Naldrett and Hewins (1984), Grieve et al. (1991), Zieg and Marsh (2005), Ames et al. (2005), Bleeker et al. (2015), and Lightfoot (2016). In the context of the present discussion, note the uppermost units of the Whitewater Group, which represent the foreland depositional wedge of the Penokean foreland basin. These sedimentary formations are only preserved in the regional scale, doubly plunging, Sudbury Basin syncline. The turbiditic Chelmsford Formation is similar in age and character to the Rove (Virginia) Formation of the Lake Superior area. Note the very tight age control on the differentiated melt sheet, including a new age of last crystallizing basal granophyres at 1850.0±0.9 Ma based on chemically abraded, fully concordant zircon data. Numbers highlight the different orebody settings and types: 1) disseminated sulphides in mafic norite near the base of the SIC; 2 and 3) disseminated to semi-massive sulphides in the Sublayer and along the footwall contact; 4) sulphides infiltrated in the footwall breccia; 5) semi-massive to massive sulphides in footwall rocks, variably fractionated and enriched in Cu; 6) massive sulphide sills deeper within the footwall, (sub)parallel to the footwall contact, very Cu-rich; 7) deeper remobilized veins, with a transition to hydrothermal processes; 8) major sulphide concentrations in funnel-like embayments; 9) sulphide ore within inclusion-bearing diorite (IQD) injected into footwall; and 10) globules to semi-massive sulphides, fractionated and Cu-rich, injected deep into footwall dykes. 15 �Proceedings of the 68th ILSG Annual Meeting - Part 2 16 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Early publications, at a time when the Grenville orogenic front was already well recognized, attributed the folding of the Sudbury structure to Grenvillian deformation (e.g., Dietz, 1964). More recent mapping and geochronological data suggest, however, that the actual Grenville Front sensu stricto. is rather sharply defined and that circa 1.0 Ga deformation did not significantly extend far into the immediate foreland and had little effect on the Sudbury structure (e.g., Easton, 1992; Davidson, 1997; Easton et al., 1999). Instead, there is increasing evidence for significant post-Penokean but pre-Grenville deformation, broadly correlated with the Yavapai and Mazatzal (or Labradorian?) orogenic belts/episodes defined in the southwest USA, i.e. broadly in the interval 1.80-1.60 Ga (e.g., Bailey et al., 2004). However, at a somewhat larger scale, Grenvillian deformation, specifically loading of the crust by the Grenvillian thrust stack, is likely to have contributed regional tilting of the crust to the south, causing further uplift of the Sudbury area and the Archean craton to the north. The precisely dated Sudbury impact event (Krogh et al., 1982, 1984; Corfu and Lightfoot, 1996; Davis, 2008; Bleeker et al., 2015; Bleeker and Kamo, in prep.), based on multiple high-precision U-Pb ages on units of the SIC, at 1850 Ma, occurred as a sharply (seconds and minutes!) timed event during the circa 1860–1840 Ma Penokean orogeny. From a broader regional perspective, the onset of the major Penokean orogenic event at circa 1870–1860 Ma predates the impact event (e.g., Holm et al., 2007). In the Sudbury area, this can be demonstrated, perhaps, by Huronian rocks being folded to some degree prior to emplacement of the melt sheet and associated dykes. However, this requires de-convolving the potentially very complex deformation associated with the impact, the immediately post-impact collapse along ring faults, and the large-scale in-flow of material into the transient crater (e.g., see modelling studies on large impact studies and the spectacular deformation, and folding, it induces in the central uplift and the surrounding annulus; Ivanov, 2005). These latter processes would undoubtedly involve tight folding of Huronian strata in places and satisfy the apparent timing relationships. The most pertinent structures in this respect occur to the north of the preserved SIC, where melt sheet injection dykes (known locally as “offset dykes”, e.g. the Hess offset dyke) appear to cut both limbs of fold structures in synclinal outliers of Huronian strata. Is this apparent folding pre-impact (e.g., as assumed by Mungall and Hanley, 2004, using an outdated model of the Blezardian orogeny), or syn-impact and due to large-scale in-flow of material into the annulus around the rebounding central uplift? In any case, these structures were tightened with present dips of Huronian units locally being steep. They were thus likely tightened to some degree by Penokean deformation extending into the foreland. This is the main reason the “Penokean front” in Figure 2 is placed to the north of the preserved SIC, and north of these deformed and folded Huronian outliers. Zooming back out, most of the more intense and final Penokean deformation post-dates the impact structure and also the emplacement of the foreland depositional wedge of the Whitewater Group turbiditic sediments that must have covered the area and which are now preserved only in the keel of the doubly plunging Sudbury Basin syncline (see map of Figure 1). Figure 7 presents a NNW-SSE cross-section through the western half of the Sudbury structure, as constrained by map and outcrop data, as well as a down-plunge projection of the western fold closure of the Sudbury basin (Bleeker et al., 2014). 17 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 7: NNW-SSE cross-section through the western half of the Sudbury structure, illustrating the firstorder regional scale synclinal fold structure, with more intense deformation and tilting of the southern flank (the “South Range”). This structure is largely Penokean in origin, but, particularly on its southern flank, was further amplified and shortened by younger post-Penokean deformation. Overall shortening on the South Range may reach ~50%, with the basal contact of the SIC being steep to subvertical, and locally overturned. Discrete later shear zones, with south-over-north displacement, further imbricated this steepened southern flank of the structure. However, at the scale of the section, there is no large-scale offset on major discrete shear zones. The folded melt sheet must have extended well to the north, to allow injection of the most distal offset dykes, and also to the south. However, all of the southern half of the folded melt sheet has been removed by erosion. The major structural front in the south is the circa 1.0 Grenville Front, with large scale thrust movement. Abbreviations: CG, Creighton Granite; CC, Copper Cliff Rhyolite; EM, Elsie Mt. Formation; MK, McKim greywacke turbidites; MS, Mississagi Formation quartzites; N, Nipissing Diabase sills; R, Ramsay Lake Formation; S, Stobie Formation; SIC, Sudbury Igneous Complex. This cross-section clearly demonstrates the overall north-verging synclinal structure with the younger than 1850 Ma Chelmsford Formation (circa 1840 Ma?) preserved in the core of the syncline. Much of this deformation is likely Penokean, but significantly overprinted by post-Penokean tightening and further shortening on the southern limb of the asymmetric syncline, during deformation associated with the 1.751.65 Ga South Range Deformation Zone. The swarm of mafic “Trap dykes”, which was emplaced at circa 1750 Ma along the South Range (Bleeker et al., 2015; see also the “quartz diabase dykes” of Cochrane, 1984) cuts much of the (Penokean) deformation, but is itself weakly deformed and metamorphosed to upper greenschist facies. It provides an important temporal and structural marker for this post-Penokean interval. The section demonstrates the intensification of structures and overall shortening on the locally steeply dipping southern flank of the large-scale syncline. Here, bulk shortening may locally reach ~40-50%. Much of this deformation was likely Penokean, but was amplified by younger circa 1.80–1.60 Ga deformation that is manifested by a system of southeast-dipping, south-side up reverse shear zones known as the South Range Shear Zone (Shanks and Schwerdtner, 1991; Bailey et al., 2004). Metamorphic titanites in these sheared rocks, which reached epidote amphibolite facies metamorphic grade, date this overprinting deformation as post-Penokean, in the time range of 1.80-1.60 Ga (e.g., Bailey et al., 2004; Papapavlou et al., 2017). Hence, these events have been broadly correlated with Yavapai and Mazatzal orogenic events, as they have in the broader Lake Superior area. Indeed, some of the orebodies on the steep South Range of the SIC are cross-cut and displaced by discrete and well-defined southeast-dipping shear zones. This has been clearly demonstrated in the Thayer Lindsley Mine, and also in the Creighton Mine. This will be discussed during the fieldtrip. 18 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Important as these younger structures are, on the scale of the folded Sudbury structure (i.e. Figure 7), this entire shear zone system only causes minor offsets of the SIC, and the first-order continuity of the SIC at both the western and eastern closures of the folded structure rules out very significant, discrete, fault offsets as has been suggested in some of the published literature (e.g., Milkereit and Green, 1992; Wu et al., 1995; Cowan et al., 1999). The section of Figure 7 is thus best interpreted as a largely Penokean fold structure, moderately overprinted and further shortened by post-Penokean bulk shortening deformation (at the scale of the section), with the final shortening on the South Range reaching perhaps ~50%. 7. Where Was “Ground Zero” and How Big Was the Final Impact Structure? Given the final deformational state of the very large Sudbury impact structure (Figure 7), a first-order question is: where was the geometrical centre of the impact; or in other words, where was “Ground Zero”, relative to the preserved erosional remnants of the structure and those of the melt sheet? An accurate answer to this question has major implications on the interpretation of the size of the impact structure, the volume of impact melt generated, and thus also affects overall mass balance calculations of elements in the melt sheet and the orebodies. Several important datasets may provide an answer to this all-important question: 1) The analysis of preserved ring structures in the relatively undeformed foreland to the north of the deformed Sudbury structure (Butler, 1994; Spray et al., 2004; see also Grieve et al., 1991). 2) An analysis of the statistical focal point of all well-preserved and least deformed (and possibly somewhat re-oriented) shatter cones (Bleeker, in preparation). As presented by Butler (1994), a careful analysis of lineaments and possible impact ring structures in the foreland of the Superior craton places the geometrical centre well to the south of Sudbury by as much as 10-15 km. Butler’s analysis is supported by a statistical analysis of all intersections of shatter cone axes (see also Guy-Bray et al., 1966), which also focus along the South Range, approximately in the Copper Cliff area, or slightly to the northeast in the Frood and Stobie mines area. These findings are summarized in Figure 8. Superimposed on the diagram of Figure 8 is a model of progressive restoration of the preserved outline of the Sudbury structure (i.e. the base of the SIC), in 4-5 steps, including minor deformation in the foreland, 15-20% shortening in the northern half of the SIC syncline, and up to 50% shortening of the southern half of the structure. This restoration of realistic amounts of shortening places the trace of the South Range, and the original focal point of shatter cone axes, just in range of the Butler’s geometrical centre defined by his most robust, lineament-constrained “Ring 3” which is shown in red on Figure 8. A significant conclusion from this work is that all of the preserved SIC represents only a portion of the northern half of the preserved melt sheet. Hence, there is no a priori symmetry between the preserved North Range and the South Range, but rather all kinds of asymmetry: the North Range preserves a thinning northern lobe (but not the edge!) of the original melt sheet, whereas the South Range preserves a more central portion of the melt sheet that was onlapping onto, and partially overlapping the collapsed central uplift. All of the southern part of the melt has been removed by uplift and erosion. These constraints and conclusions are further summarized in Figure 9. These findings can be modified to some extent by increasing the shortening deformation somewhat, moving the restored South Range of the SIC a bit farther south, but without going to unreasonable shortening estimates it will not change the first-order conclusion that the preserved SIC is all from the northern half of the folded melt sheet and thus preserves a relatively small, fundamentally asymmetric sample of the original melt sheet. 19 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 8: Summary map of the size and location of the preserved Sudbury structure relative to the main structural deformation fronts, and the impact ring structures defined by Butler (1994) based on a detailed lineament analysis (here only lineaments related to Ring 3 are shown for illustration, and some to the northeast of the structure where there is some discordance between lineaments and Ring 4). Of these rings, the one shown in red is the most robust (see Butler, 1994) and it, together with the other rings, defines the centre of the ring structures. This is “Ground Zero” and is shown with the small red circle south of Sudbury. This impact centre is pinned to the foreland as it is largely based on Ring 3. The purple star represents the statistical focal point of best-preserved shatter cone axes and is also located south of the preserved structure. Restoring the shortening deformation in 4-5 progressive steps, including 50% shortening of the southern half, less in the northern half, and ~5% in the foreland up to a radius of 50 km, places the original South Range just in contact with Butler’s “Ground Zero” (see the blue trace). Hence, all of the preserved melt sheet represents only part of the northern half of the original melt sheet. Only ~10% of the melt sheet is presently preserved in the erosional remnant of the SIC. 20 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 9: Two critical stages in the formation of the Sudbury impact structure (adapted and simplified after Grieve, 1994): A) early during the impact phase and formation of the transient cavity, as shockwaves are transmitted into the target; and B) after collapse and rebound, and final settling of the melt sheet. Note the estimated position of the North Range (NR) and South Range (SR) in figure B. The diagrams also illustrate the general pattern of shatter cones in the target rocks around an impact crater during its formation and excavation (A) and after rebound (B). A shatter cone (S, outlined by square) forms in response to the shockwaves radiating out from the focus of the impact. When the floor of the transient crater rebounds, the cone and the cone and its axis are rotated up and the rocks may also move inwards due to collapse and large-scale material in-flow into the transient crater. In figure B, the concentric Hess Offset dyke is show, at ~50 km distance from “Ground Zero”. The Foy Offset dyke, more or less parallel to the section, is not shown but extends out to ~65 km (see Figure 8). The Hess Offset dyke cuts small synclinally folded outliers of Huronian rocks that help to define the down-folded and down-faulted annulus surrounding a broad area of central uplift. The overall extent of the final melt sheet, after settling, likely reached the well-defined Ring 3 with a radius of ~67 km. This conclusion is supported by at least one of the injection dykes, the Foy Offset, reaching close to this Ring 3, although the originally overlying melt sheet from which it was injected down has been removed by post-Penokean uplift and erosion. Thus, with an estimated radius of 67 km, and an average thickness of the melt sheet of 2.5 km (see also Grieve et al., 1991 and Naldrett and Hewins, 1984: 2–3 km on the preserved North Range, 3–5 km on the South Range), this suggests a final melt sheet volume (πr2 x d) of ~35 x103 km3. This estimate does not include the considerable melt component preserved in the ~2 km-thick Onaping Formation or within the ejecta that were blown far beyond the crater (tektites?) and into space. Thus, it is likely an underestimate. Nevertheless, reasonable uncertainty estimates put this total impact melt volume in the range of 25–50 x103 km3, which could be used as an input parameter into various scaling models and models of elemental mass balances. 21 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Of this original volume of settled impact melt sheet, the thin edge of which likely extended out to ~67– 70 km radius, only ~2.5–5.0 x103 km3 is preserved in the current extent of the SIC. This estimate can be derived by measuring the cross-sectional area of a central cross-section through the SIC (similar to Figure 7 but a bit farther east), multiplied by 0.5x the preserved width (long axis) of ~60 km. In other words, only on the order of 10% of the original melt sheet is represented by the currently preserved erosional remnant of the SIC, and all of this represents the proximal part of the northwestern half or lobe of the original melt sheet (see Figure 9), perhaps including a thicker “moat” north of the collapsed central uplift. As deformational structures, both large and small, typically nucleate on and amplify original “seed structures” or weaknesses, the curved main synclinal axis of the preserved SIC may be inherited to some degree from such a moat surrounding the core of the central uplift. This is where the melt sheet may be thickest and even larger orebodies could have collected. All of these key points should be considered in any interpretation of the Sudbury structure and its endowment of magmatic sulphide orebodies. Perhaps one of the more robust features of the final impact structure, as presently preserved in the foreland, and in relation to overall size estimates, is the down-folded and faulted annulus of Huronian outliers to the north of the SIC, with a pattern of synclinal fold traces that curves around the entire northern half of the structure, including around the western and eastern first-order fold closures of the structure (Figure 8). This down-folded annulus fits a ring structure with a radius of ~50–55 km. The annulus thus has an apparent diameter of 100-110 km, which is larger than the well-defined annulus of the circa 2023 Ma Vredefort impact structure in South Africa, where it is ~90 km in diameter. This can be scaled to a final crater diameter. The broad, collapsed central uplift must fit within this annulus and therefore has a total radius and diameter of ~45-47 km, and ~90–100 km, respectively, and includes all of the South Range area and the deeply exhumed Huronian to the south, and also the Levack Gneisses in the north. In large complex terrestrial impact craters the diameter of the broad central uplift area is roughly 1/3 of the final crater diameter (Therriault et al., 1997), the scaling relationship being: Dcu = 0.31 Df1.02 where Dcu is the diameter of the central uplift, and Df is the diameter of the final crater rim (Therriault et al., 1997). This would suggest that the final Sudbury structure may have had a diameter closer to 300 km and somewhat bigger than most estimates. Based on the estimate of the annulus diameter alone, it is clear that Sudbury is the largest among known terrestrial impact structures, just slightly larger than the more deeply eroded (no melt sheet preserved) Vredefort structure. Major Sudbury Breccia pseudotachylite occurrences reaching out to a radius of ~120 km supports such a larger estimate based on the scaling relationship of Stöffler et al. (1988): Dpst . 0.8 Df Where Dpst is the diameter of significant pseudotachylite breccia occurrences, i.e. ~240 km. This would result in a similar estimate of ~300 km for the final crater diameter (Df). 22 �Proceedings of the 68th ILSG Annual Meeting - Part 2 8. Conclusions The Sudbury structure represents the deformed erosional remnants of a very large 1850.0±0.9 Ma meteorite (or comet?) impact crater. The current consensus is that the final collapsed crater reached ~200– 260 km (final crater diameter, Df), but it may have been larger based on a robust size estimate of the downfolded and faulted circular annulus, and the most distal observations of significant pseudotachylite occurrences (Sudbury Breccia), some of which occur well beyond the 100 km-radius ring structure (Ring 4 of Butler (1994; see also Thompson and Spray, 1994)). We thus estimate the final crater diameter at ~300 km, which makes it the largest known terrestrial impact crater. The volume of impact melt generated was ~35 x103 km3. Once the crater had collapsed and rebounded, and the melt sheet had settled across a complex peak-ring crater, on a time-scale of mere hours (!) (e.g., see Grieve, 1994), the sheet of impact melt reached out to ~67–70 km, and had an average thickness of ~2.5 km. It may locally have reached 5 km or more (Figure 10). A large proportion of this melt sheet was initially superheated and underwent rapid homogenization and differentiation. Zieg and Marsh (2005; see also Golightly, 1994) point out some of the very complex processes involved as the final melting front burned into the footwall and different blobs of melt or partial melt were generated and may not have fully mixed and homogenized by turbulent convection. There likely was an early separation in i) less dense, more felsic melts, and ii) denser mafic melts, the first floating to the top of the melt sheet pool to form or contribute to the Granophyre of the Main Mass of the SIC, whereas the latter collected towards the base to form the Norite, and Sublayer. The Sublayer represents a complex boundary layer with remnant mafic and ultramafic fragments that collected along the base of the SIC, together with a rain-out of sulphides globules. The Transition Zone Gabbro represent the differentiated top of the lower mafic section, enriched in incompatible elements. The density contrast between the Granophyre and Norite (including the gabbro) is significant, and large enough such that the basal granophyre contact would have equilibrated in a near horizontal position, providing an important reference plane for evaluating the effects of later structural deformation. One of the important implications of this is that it allows a semi-quantitative reconstruction of the footwall topography of the SIC (Figure 10) and aids in general structural reconstructions (e.g., Figure 7). Some of the enormous volume of ejecta fell back into the crater, or was washed back in by various processes and formed the Onaping Formation. In a general sense, the upwards fining stratigraphy of the Onaping Formation and increasing carbon content reflect the waning energy levels and somewhat longer time scales of deposition, with the uppermost and finer-grained “Black Member” material transitioning into post-impact sedimentary and volcano-exhalative processes of the Vermillion member. The impact struck at a target site with complex geology, on the boundary of the rifted Superior craton and the developing Penokean fold belt of the Southern Province. The rebounded and settled crater, and the crater-fill deposits, were then covered by the expanding wedge of foreland basin sediments, first deeper water mud- and siltstones (foredeep?) and finally the greywacke turbidites of Chelmsford Formation at circa 1850–1840 Ma. Soon after, the crater, the melt sheet, and overlying deposits were deformed and shortened by the climax of Penokean deformation, which transformed the area into a doubly plunging regional syncline. The ~3-5 km-thick SIC melt sheet formed a very thick competent layer during this deformation, which resulted in a single large wavelength open fold structure. This syncline, particularly its southern limb, was further modified and shortened by post-Penokean deformation associated with the circa 1.80-1.60 Ga South Range Shear Zone. 23 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 10: A simplified, first-order reconstruction of the South Range melt sheet with its inherent thickness variations and footwall topography. It highlights the embayment structures where most of the significant orebodies collected (see Pye et al., 1984, and contributions therein), as well as the lateral thickness variations which suggest a second-order “peak-ring” on the overall central uplift, where the mafic section of the melt sheet (norites plus gabbro) is much thinner than in the “central puddle” that approximately overlapped “Ground Zero”. The dashed horizontal line indicates the level of the topographic highs of the secondary peak-ring. The section was constructed by first restoring minor second-order folding along strike of the South Range, then correcting from apparent thicknesses to true thicknesses, and finally hanging the section segments from the paleo-horizontal reference plane, i.e. the base of the significantly less dense granophyre upper section of the melt sheet. From west to east, star symbols identify the major magmatic sulphide deposits and producers: Vi, Victoria mine; AK, Aer-Kidd; To, Totten; Lo, Lockerby; Ge, Gertrude; Cr, Creighton; CC, Copper Cliff; CCN and CCS, Copper Cliff North and South mines; Mu, Murray; LS, Little Stobie; Fr, Frood; St, Stobie; Li, Thayer Lindsley; Ga, Garson; and Fa, Falconbridge. Of these Creighton and the Copper Cliff system are among the largest deposits known, and associated with the deepest embayments or funnels, in the floor of what appears to be a deep central puddle in the melt sheet, approximately overlying the centre of the collapsed central uplift. Following various stages of regional uplift and erosion, the end result is that only ~10% to 20% of the original melt sheet is preserved in the current erosional remnant of the SIC. Both the analysis of lineaments and ring structures, and the analysis of shatter cones, suggest that the preserved, folded melt sheet (the SIC) represent parts of the northern half of the original melt sheet and associated crater, with “Ground Zero” being situated south of the main footwall contact of the South Range. This is significant as most studies have implicitly assumed that “Ground Zero” was underneath the preserved SIC and that the preserved SIC (e.g., Golightly, 1994) more or less preserved a symmetrical patch of the original melt sheet and impact structure centered on “Ground Zero”. If the ring structure and shatter cone analysis is correct, this is clearly not the case and there is no inherent basis for symmetry between the North and South Ranges. Rather, the North and South Ranges represent fundamentally different parts of the melt sheet and the collapsed crater; the former represents a more distal northern part of the structure, and the latter a thick, more central part of the melt sheet lapping onto and overlying the rebounded and collapsed central uplift of the final peak-ring crater (see Figure 9). Petrologists had long recognized that something was odd about the Sudbury structure and its igneous rocks, the “Sudbury Irruptive”. In contrast with other large mafic igneous complexes, there was no layered lower section of mafic and ultramafic cumulates near the base, and way too much granitic granophyre near the top (~40% of the volume, rather than ~10% in a fully differentiated mafic intrusive complex). And the 24 �Proceedings of the 68th ILSG Annual Meeting - Part 2 mafic lower part had an unusual SiO2 content (~55–58 wt%), dominantly characterized by norites rather than gabbros 4. The regional deformation was also puzzling and under-appreciated, feeding into interpretations of the whole complex as a concave-upwards igneous lopolith. It took a marine geologist from California, who somehow had solved the mystery of shatter cones in a small number of suspected impact craters in the 1940–1950s (e.g., Dietz, 1947, 1959) 5, and who also was reflecting on the lessons learned from observing the lunar surface (Dietz, 1946), to come and check out Sudbury, and within days confirm his suspicion that it was, fundamentally, a large impact structure with major shock damage in the footwall and no damage in the melt rocks (the igneous rocks of the SIC). Being less familiar with the ores, he still chose to hedge his bets on some of the details, such as the origin of the ores, or the exact nature of the igneous rocks, settling on a hybrid model of a large impact crater that was then intruded by the igneous rocks of the irruptive. As late as 1970, the debate on Sudbury was summarized as follows in a major paper on the structure (see Naldrett et al., 1970): “There are two main theories of origin for the Sudbury structure. These are very different from one another and hinge on the interpretation of the Onaping formation and certain unusual pre-irruptive breccia dikes. Speers (1956, 1957), pointed out that the Sudbury Irruptive lies at the apex of a broad dome some sixty miles in diameter involving Huronian and older rocks. He postulated that uplift of the dome occurred in response to pressure exerted by igneous magma. Successive episodes of uplift, followed by tensional release, gave rise to the breccia dikes and finally resulted in caldera collapse at the apex of the dome. Magma escaping around the rim of the caldera and flowing into the center of collapse produced the Onaping formation. The Irruptive was intruded subsequently, spreading out along the base of this formation. According to this hypothesis, the Nickel Irruptive is a later plutonic manifestation of the igneous activity which previously had given rise to the extrusive Onaping formation. In opposition to this hypothesis, Dietz (1964), suggested that the circularity and brecciation characteristic of the Sudbury structure could best be explained by an explosive meteorite impact, an interpretation for which his own discovery of shatter cones gave support. This theory led French (1967) to find, in inclusions in the Onaping formation, microscopic features characteristic of shock metamorphism. Subsequent work has shown that shock metamorphism, a typical feature of impact sites but unknown in volcanic rocks (French and Short, 1968), is widespread and common in the Onaping formation; it also is found in footwall rocks adjacent to the Irruptive and in fragments in Sudbury breccia. According to the meteorite impact theory the sequence of events was as follows: shock waves radiating from the point of impact produced brecciation, melting, microscopic shock features and shatter cones, and excavated a circular crater; part of the material blasted from this crater fell back as a poorly, sorted See, for instance, the papers by Wilson (1956) and Hamilton (1960), written and published just prior to Dietz’ shatter cone revolution, to appreciate the conceptual struggles that petrologist and geologist were dealing with to explain major aspects of the Sudbury structure. In his paper, Hamilton is inching towards an essentially extrusive interpretation for the Sudbury lopolith. 5 See the amazing review paper by Bourgeois and Koppes (1998) to better understand the historical development of these ideas, and all the players involved, including of course the life and career of Dietz himself. Originally from New Jersey, he was broadly educated with degrees from the University of Illinois and having spent time at the Scripps Institution for Oceanography. He was also well-travelled. A professional posting in Europe during the 1950s had also allowed him to visit the Steinheim and Ries basins in Germany. 4 25 �Proceedings of the 68th ILSG Annual Meeting - Part 2 breccia—the Onaping formation; fracturing and heating of the rocks, and reduction of pressure in the upper mantle below the crater, triggered the evolution of the Nickel Irruptive; the magma was emplaced in the breccia zone beneath the center of the structure.” [End of quote.] It would take another 15–20 years for clarity to emerge. Once the isotopic evidence became available showing that essentially all the igneous rocks, including the more mafic ones, have bulk Nd isotopic signatures that reflect melting of the crust, rather than melting of the mantle, the pendulum finally swung to an impact-only interpretation (Stöffler et al., 1989; Grieve et al., 1991). Bulk sample isotopic values and mixing equations may still hide a very small mantle component into some of the melts but, to date, no conclusive evidence for this has emerged. What lessons can be learned from all of this? One is that it is critical to think “big”, always, and broaden one’s horizon, and to to reflect on new ideas from related or not-so-related sciences. The other is that, as our understanding of the impact record grows, particularly on nearby planetary surfaces, there got to be other Sudbury’s out there, with perhaps less than ~10% preservation: just the odd bit of breccia; or a poorly preserved shatter cone here or there; or some odd dyke of quartz diorite with some sulphides in it, and which was just a little too hot for your average diabase dyke! Remember, every breccia is an interesting breccia! Acknowledgements The first author (WB) was first introduced to some aspects of Sudbury geology in 1987, during a fall fieldtrip of the Canadian Tectonics Group. He has worked, on and off, on Sudbury geology since the early 1990s, first as a researcher for Falconbridge Ltd. at Onaping Mine, and later for the Geological Survey of Canada. The second author (SK) has been involved in precise U-Pb dating of many of the rocks in and around Sudbury. The geology of this unique area never stops to fascinate. WB would like to thank coleaders for helping to put this fieldtrip together, and also those who gave the guidebook a proof-read to catch some of the imperfections. 26 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Excursion Stops, Day 1: Bedding at this locality dips and youngs to the southeast. The shatter cones are well developed and enough of the cone surfaces are visible, such that their axes can be measured accurately. These axes point up towards the north-northwest (~340º/+45º, i.e. up), and they plunge up more or less parallel to the southeast dipping bedding surface, perhaps just a bit shallower than the bedding dip angle (~45– 55º). Although there is plenty of impact brecciation in these footwall rocks, locally chaotic, these outcrops are not chaotic and align with the regional pattern. The shatter cone axes therefore point approximately to “Ground Zero”, the centre of the impact structure. Stop 1: Well-developed shatter cones in Mississagi Formation quartzites, south of Sudbury 46.432448° N, 81.072267° W 494448E 5142100N This is one of the classic shatter cone localities south of the Sudbury structure, and probably one seen by Dietz early on during his 1962 fieldtrip to Sudbury (Figure 11). The conical, striated fracture surfaces can be seen both in the blasted roadside outcrop on the south side of the gravel road (Gibson Road), but also on top on natural outcrop surfaces. Typically, shatter cones are more obvious in blasted outcrops where the rock surface has been opened up, and are more easily missed on natural surfaces. We will see examples of that on this fieldtrip, but this is not a limitation here. Spectacular shatter cones are visible south of the road, on top of the outcrop. In meteorite impacts, the shockwaves travel out from the centre of the transient crater, outwards, and interaction of the high-velocity shock waves with imperfections in the rocks nucleate the conical fracture surfaces. The apparent point of origin of the shockwaves is below ground, as the crater is being excavated and the ground is depressed. Therefore, away from “Ground Zero”, shatter Figure 11: Classical “shatter cones” in the shocked footwall and target rocks of the 1850 Ma Sudbury impact crater, well developed in quartzites of the circa 2.4 Ga Mississagi Formation. With the newly recognized knowledge, in the mid- to late-1950s, that these conical, radiating fracture surfaces represent unique “trace fossils” for high-velocity, very high pressure shock waves associated with meteorite impacts (Dietz, 1959), Sudbury was quickly recognized as an astrobleme—the scar of an ancient impact crater (1962; see Dietz, 1964; and Dietz and Butler, 1964). 27 �Proceedings of the 68th ILSG Annual Meeting - Part 2 cones will fan outward and their axes will be shallow. On rebound of the crater floor, the central uplift moves up and rotates the cones, with their axes now pointing up (see Figure 9). If marginal rocks are turned up or even flipped over during crater formation, the cones may point outwards in rocks that are flipped over (as in the collar around the Vredefort central uplift). Here this is not the case, the rocks are right-way up, dipping moderately to the southeast, and the cones plunge up to the north-northwest. Hence, just from this outcrop alone, it seems “Ground Zero” was to the north, and their dip was subhorizontal when the shockwaves hit. The cones can be examined on top and in section, to determine their axes, as will be demonstrated during the visit. Stop 2: Well-developed shatter cones in blasted Mississagi Formation quartzites, south of Sudbury 46.422030° N, 81.086392° W 493367E, 5140943N This roadside outcrop is again in Mississagi quartzites, somewhat to the southwest of the previous locality. Here beds dip and young towards the north, so we have travelled across one of the many folds in the Mississagi Formation. Several reasonably developed shatter cones are visible in the rocks on the east side of the road (Figure 12). The cone axes here plunge down and to the north, in approximately identical relative orientation to the bedding surface, but now plunging down! Figure 12: Shatter cones plunging down to the north-northwest, on the southern limb of the local syncline. Both bedding and the shatter cone axes have been re-oriented by the Penokean folding. 1) The local Huronian strata were (sub)horizontal at the time of impact; 3) The shockwave traveled from north to south and formed more or less flat lying shatter cones with their apices plunging gently to the north; 4) And both bedding and cones were affected by the folding that affected the south range of the Sudbury structure. Duplicating this exercise at as many as possible localities, and then intersecting all cone axes statistically, looking for a maximum of intersections, can define the origin of the shockwaves, i.e. “Ground Zero”, or perhaps a point below “Ground Zero”. Just from these two outcrops alone, it is clear that there is a correlation between final bedding attitude and final attitude of the cones: both have been affected by moderately tight folding that produced the local synclines and anticlines in the Mississagi Formation. Unfolding of these folds and restoring bedding to approximately horizontal will also align the cones between Stop 1 and Stop 2. This general story is repeated all through the area and the conclusion must therefore be: 28 �Proceedings of the 68th ILSG Annual Meeting - Part 2 This outcrop also demonstrates that shatter cones are rare on naturally weathered outcrop surfaces but are more clearly developed or visible on certain broken rock surfaces. scale of this dyke emplacement, relative to “time zero” could be years to thousand of years? Here the dyke does not show evidence for superheating and melting of adjacent country rocks, but elsewhere melting can be observed along the contact. Stop 3: Southern extent of the Copper Cliff “Offset dyke” cutting across Pecors and Mississagi Formation strata After emplacement and cooling, the dyke was moderately deformed together with the country rocks during the largely Penokean shortening and folding deformation. 46.434569° N, 81.068932° W 494704E, 5142335N The sedimentary rocks on the south side of this Gibson Road locality represent the transition from the finer-grained Pecors Formation to the quartzites (quartz arenites) of the Mississagi Formation. Dips and younging direction are to the south, as at Stop 1, which is just along strike. The Huronian strata are cut by a ~25 m-wide mafic dyke, which at first sight looks not unlike a diabase dyke. This dyke is subvertical and trends south, and its contact relationships look rather typical for a diabase dyke, wandering a little bit and stepping sideways a little bit here and there. There is, however, no known swarm of this trend and/or the right age to explain this dyke. Stop 4: Unsorted matrix-supported conglomerate (reworked diamictite) of the Ramsay Lake Formation 46.433508° N, 81.077564° W 4947041E, 5142218N This glacially polished outcrop is on the north side of the Highway 17 Bypass just south of Kelly Lake. Note that this is a very busy highway with heavy, high-speed traffic. The shoulder of the highway allows cars to stop safely here but make sure to pull well on to the shoulder. This is an outcrop of poorly sorted, matrixsupported, sandy conglomerate typical for the lowermost of the three glacial formations in the Huronian succession. The Ramsay Lake Formation is the basal unit of the Hough Lake Group (see Figure 4), which overlies the rift and rift-fill succession of the Elliot Lake Group, locally with a sharp contact. Petrography shows it to be a medium-grained quartz diorite, with some chilling near the margin. A quartz diorite is rather atypical for a regional diabase swarm, which are essentially all basaltic. It is perfectly along strike of the Copper Cliff Offset dyke to the north, on the north side of Kelly Lake, which can be mapped into the base of the Sudbury Igneous Complex (SIC). This outcrop shows the unsorted nature of the conglomerate and on an outcrop like this one could debate the evidence for a glacial origin. Some bedding surfaces and cross-bedding are visible. Long (2009) describes these rocks as sub-glacial melt-out till. Elsewhere the Ramsay Lake Formation, which typically is thick to very thickly bedded, or even massive, has a finer-grained, darker matrix and looks more like a typical diamictite (Figure 13). Here we are at the top of the Formation, where the till material was reworked and sorted to some degree by sub-glacial processes. It is tempting to think of these “radial offset dykes” as having been emplaced laterally from the north, but that is not the right interpretation. Almost certainly, these dykes of basal melt rock from the SIC were injected downwards from the overlying melt sheet into active fracture planes in the deforming footwall, during the overall crater modification processes and settling and early differentiation of the melt sheet. Overall homogeneity of these quartz diorite dykes indicates that the melt sheet had already mixed and homogenized to some extent. The overall time 29 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The brecciation, locally approaching large-scale chaos, is typical for many areas in the footwall. Dykes and veins, and more irregular domains of fine-grained, dark pseudotachylitic rocks, are what is referred to as “Sudbury Breccia”, indicating intense brecciation and pulverization of the target rocks during impact and intense deformation associated with rebound and collapse of the central uplift. All these rocks are cross-cut by an irregular dyke-like body of quartz diorite emanating from the Copper Cliff funnel. The outer contact of this quartz diorite dyke induced melting and mixed with the adjacent rhyolite (Figure 16). The quartz diorite itself shows typical quench textures: more or less spheroidal structures of radiating, fine acicular pyroxene/amphibole crystals. Together, these observations show that this quartz diorite was injected in a superheated state and then quickly cooled due to interaction with wall rocks. Locally the quartz diorite dyke rock interacted with dark Sudbury Breccia pseudotachylitic material, overall showing that the latter was marginally older but not yet acting fully lithified or brittle. Figure 13: Unsorted more typical diamictite of the Ramsay Lake Formation, with a dark fine-grained matrix (not this outcrop). Pebbles and cobbles are mostly granitoid rocks. Dark patches with striae are a rare example of how shatter cones are exposed on naturally weathered (and polished by Pleistocene ice movement) surfaces. Stop 5: Mineralized Copper Cliff Offset dyke, up from the walking trail at Copper Cliff 46.470977° N, 81.075562° W 494199E, 5146381N The outer phase of the dyke is known as typical quartz diorite, or “QD”, and is generally similar to quartz diorite of many of the offset dykes (e.g., see Stop 3). It generally does not contain sulphides, and does not weather rusty. This outer QD was intruded by one or more phases of dyke injections that are characterized by carrying along inclusions of varying size (inclusion-bearing quartz diorite, or “IQD”), many of which are more mafic, together with variable amounts of sulphides, either disseminated or as conspicuous cm-size globules. A central phase of the dyke carries semi-massive sulphides, which are enriched in Cu. Hence, in these outcrops here, three successive phases of dyke injection can be demonstrated, which mark different stages in the evolution of the melt sheet: This series of outcrops on the side of a hill overlooking the urban area of Copper Cliff exposes parts of the major Copper Cliff Offset dyke which can be mapped to the north into a major “funnel” or narrow embayment structure at the base of the SIC. Major Cu-Ni sulphide mineralization occurs in this funnel structure and into the dyke (Figures 14 and 15). Walking up toward the dyke, one walks across heavily brecciated wackes and arenites at the base of the McKim Formation, and into the top of the Copper Cliff Rhyolite Formation. Bedding features at the base of the McKim Formation, such as graded bedding, scours, and truncated bedding indicate younging is towards the southeast in steeply southeast dipping strata, i.e. stratified sediments that overlie the rhyolites to the north. The latter show quartz phenocrysts and beautiful flow lamination in places. 30 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 14: Map and corresponding longitudinal section, looking west, of the Copper Cliff Offset dyke and its associated orebodies, supporting several active mines. North is to the right in the map figure. Modified after Cochrane (1984) and Farrow and Lightfoot (2002), based on mine and exploration sections. 1) Initial homogenization and injection of unmineralized quartz diorite, probably just prior to sulphide saturation having taken place. 2) Renewed injection from near the base of the melt sheet, with abundant inclusions and sulphide globules being entrained. 3) A final phase of injection, after sulphides had collected and had become enriched in Cu due to Fe-rich and Cu-poor monosulphide solid solution (MSS) having separated out, enriching residual sulphide liquid in Cu and other MSS-incompatible elements (e.g., Craig and Kullerud, 1969). Elsewhere, a fourth and final phase of injection can be recognized, consisting of plagioclasephyric, sulphide-free quartz diorite, representing differentiated norite from the main melt sheet, making it into the footwall. On Figure 14, the mineralized bodies are shown in magenta along the extent of the offset dyke (after Cochrane, 1984). In the corresponding longitudinal section, looking west, the overall extent of the sulphide bodies in the dyke are shown, forming kmscale steeply plunging “fingers”, generally along the centre of the dyke, but sometimes along the margin. These steeply plunging fingers of mineralized IQD clearly indicate emplacement was downwards from the base of the overlying SIC (now eroded away) where sulphides had collected. The vertical plunge of these IQD “dyke in dyke” injections was further amplified by N-S shortening during folding, and vertical extension, but deformation intensity is insufficient to explain the observed aspect ratios. So, the long axis of the orebodies, well defined by drilling and mining, indicates the injection direction: i.e., down from an overlying but now removed melt sheet undergoing critical stages in magmatic evolution: 31 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 15: Longitudinal section similar to but expanded from Figure 14, with full interpretation. Note the full extend of the Copper Cliff Offset dyke with an initial “leading sheet” of relatively homogeneous quartz diorite, followed by later injections of mineralized inclusion-bearing quartz diorite and entrained sulphides, all down from the overlying but now eroded Sudbury melt sheet, which underwent final differentiation into a thick basal norite section (~3 km), a thin transitional quartz gabbro section (~500 m, TZG in blue), and an upper “granophyre” section of broadly granitic composition (~1–2 km, pink). As shown on the figure, the overall ore formation process can be divided into six stages: 1) sulphide saturation, 2) growth and sinking of sulphide globules, 3) collection of sulphides along the basal contact, particularly in topographic low or “embayments”, 4) the onset of liquid fractionation of the sulphide melt due to MSS fractional crystallization, and progressive enrichment of the residual sulphide melt in Cu, 5) episodic injection of dense melts into footwall fractures, and 6) further sub-solidus remobilization of ore components during later deformation. The dip of the SIC is due to subsequent deformation. 1) Initial homogenization due to rapid convection in a superheated impact melt; 2) Rapid sulphide saturation; 3) Collection of basal sulphides and entrained inclusions in a still very hot magma; 4) And repeated injection of basal phases of the evolving melt sheet into footwall fractures forced open by i) on-going movements in the adjusting footwall, and ii) fluid pressure of the magma. The density of the overall “dioritic” (more or less, average crust, well mixed) is about 2.8 g/cm3 and is denser than the average density of the footwall rocks. Figure 15 above shows the completed section and interpretation. In many ways this one section tells much of the story of the Sudbury structure, its evolving melt sheet and its orebodies. 32 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 16: Some key features associated with the quartz diorite offset dykes. A) Margin of the Copper Cliff quartz diorite dyke, against melted rocks of the Copper Cliff Rhyolite country rocks. This degree of melting is highly unusual for normal mafic dykes of this size and clearly indicates the superheated state of the first injections of the Copper Cliff Offset dyke. B) Quench textures in outer (unmineralized) quartz diorite indicating rapid cooling from a superheated states (no crystallization nuclei), followed by rapid crystallization. C) Inner inclusion-bearing quartz diorite with rusty sulphide blebs, cut by a last phase of plagioclase-phyric quartz diorite dyking along the core of the Worthington Offset dyke, Aer-Kidd Mine area. 33 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 6: Extensively developed Sudbury Breccia developed near the top of the Copper Cliff Rhyolite Formation, Lively Arena Robert Dietz, being exposed to this relatively new science of lunar geology in the 1950s, and also having carefully read the early papers/reports by Boon and Albritton (1936, 1937, 1938), quickly broke through this stasis in geological thinking. 6 46.427506° N, 81.144626° W 488888E, 5141558N There is a general lesson here: always broaden your horizon, think out of the box, and listen to what may be totally new views out of left field. This glacially polished outcrop shows spectacular development, essentially chaotic at the ~5–10 m scale, of Sudbury Breccia (often abbreviated to SBX). Metre-size clasts, some rounded, float around in a dark matrix of pulverized rock flower that in some cases may have melted. Many of the fragments can be linked to the surrounding Copper Cliff Rhyolite Formation, the rhyolite that formed the uppermost felsic component of the basal rift succession of the Huronian. However, other fragments are more “exotic”, relative to local outcrops, and illustrate significant movement of breccia material, and injection for some distance into dilating fractures. One final anecdote: One of us remembers, as a 21-year-old undergraduate student back in Amsterdam, in 1980, sitting through a (very) long lecture by an eminent metamorphic petrologist talking about the Vredefort Dome in South Africa, and how it could only be an endogenic cryptoexplosion domal structure based on this or that metamorphic reaction ... . This was 20 years after Hargraves (1961) and Dietz (1961) showed a systematic pattern of shatter cones around the dome, which even then was long known to be riddled with large pseudotachylite bodies and dykes (Shand, 1916). This outcrop illustrates an important truism: any kind of breccia is an interesting breccia! Stop 7: Pillow lavas of the Elsie Mountain Formation, lower Huronian volcanic rocks Seeing this amount of breccia, and the dynamic processes that must have been involved, requires a generative process with sufficient cause and energy, and volcanism clearly is not it. Although this particular outcrop is very spectacular, similar breccia bodies, dykes, and veins occur all over the Sudbury area (in the footwall), and even early workers were familiar with the fact that SBX occurred up ~60–70 km away from the Sudbury Igneous Complex. 46.442417° N, 81.147691° W 488655E, 5143216N In these roadside outcrops along the main road north out of Lively, the mafic volcanic rocks of the lowermost Huronian rift succession are well exposed. Plagioclase-phyric basaltic pillow lavas of the lowermost Elsie Mountain Formation are steeply dipping and facing south (Figure 17). The pillows show well developed rims and, although moderately flattened, show enough asymmetry to determine top directions, to the south. In hindsight, and with our present understanding of Solar System geology, it is easy to see that only an ancient meteorite impact had sufficient energy to do this much damage, and at this scale. As stated in the introduction, it is perplexing how early workers clung to the cryptovolcanic explosion model for that long, perhaps largely due to nonfamiliarity with the emerging knowledge of the geology of the Moon and other planetary bodies. These lavas, together with the Copper Cliff Rhyolite, comprise a bimodal succession typical for continental rifts. Their overall age is 2460–2480 Ma. The mafic lavas, with their prominent plagioclase crystals, were fed by the similarly plagioclase-phyric diabase dykes of the circa 2460 The short and clearly written papers by Boon and Albritton in the 1930s could be characterized as “Sleeping Beauties” (van Raan, 2004; see also Ke et al., 2015; Miura et al., 2021), i.e. papers that were ahead of their time, and went largely unnoticed until they were “discovered” decades later. Boon and Albritton (1937, 1938) add Vredefort to their growing list of suspected impact craters. 6 34 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Ma Matachewan swarm (e.g., Heaman, 1997), which are of the same age and which riddle the Archean basement to the north. None of these dykes cut across the uppermost volcanic unit, the Copper Cliff Rhyolite, where they would be easily noted. Hence, the Copper Cliff Rhyolite marks the final, felsic phase of this rift magmatism. Together with the subvolcanic A-type granite bodies of the Creighton Granite, and similar plutons along strike, this final phase has been dated at circa 2460 Ma (2459±7: Bleeker et al., 2015). Figure 17: Pillow lavas near the top of the Elsie Mountain Formation, steeply dipping and younging to the south. Stop 8: Creighton Mine, among the largest and deepest mines of the structure Overall, the structure of Creighton Mine is typical for the more strongly deformed South Range of the Sudbury structure, with the basal contact of the SIC dipping ~45–50º at surface and steepening at depth, with second-order structures superimposed. At depth, the steep basal contact is cut and offset by a discrete south-dipping shear zone, with south-side up displacement (e.g., Papapavlou et al., 2018), which is part of the postPenokean “South Range Shear Zone” deformation that has further shortened the South Range. 46.461018° N, 81.176757° W 486427E, 51415287N We will make a brief stop here on the access road to the Creighton Mine, one of the largest and deepest mines of the Sudbury structure. Various magmatic sulphide deposits occur at or near the basal contact of the SIC in what is one of the more prominent “embayments” along the basal contact. Creighton Mine geology and structure will be introduced and discussed by means of the composite cross-section shown below (Figure 18). 35 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 18: A) Composite cross-section of the Creighton Mine with its various ore lenses (projected onto a common section). Creighton Mine is one of the largest producers and it is getting very deep. As shown on the section, there is the typical variety of ore types, including more fractionated Cu-rich ore in the footwall, often controlled to some degree by local structures. Section modified after various published sources and original Inco mine sections. Because of the large depth and deep mine infrastructure, the mine also hosts one of the major neutrino labs in the world, the SnoLab, at about 2 km depth. Research at this lab was among the work that has demonstrated some of the fundamental characteristics of these elementary particles. 36 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 18: B) Photo of an active mining face in a Cu-rich orebody in the footwall complex below the SIC at Creighton Mine. The vein-like body is several metres wide and shows sharp, in part structurallycontrolled, contacts. Inclusions of wall rock are suspended in the sulphide matrix and in various states of dismemberment by both physical and magmatic processes. Stop 9: Creighton Granite, gabbro enclaves, and Sudbury Breccia The somewhat coarser grained gabbro enclave at the east end of the road section has been dated at circa 2479 Ma (Bleeker et al., 2015), thus being part of the early Matachewan (Matachewan I) event, which also emplaced larger layered intrusions at the base of the Huronian section, at or near the unconformity with Archean basement. 46.453751° N, 81.186197° W 485700E, 5144481N To the south of Creighton Mine, a ~300 m-long section of the Lively regional road exposes superb outcrops of the Creighton Granite pluton, variably affected by dykes and veins of dark Sudbury Breccia (SBX). The granite hosts major enclaves of early Huronian gabbro/diabase sills or dykes, some with very prominent zone calcic plagioclase megacrysts, which clearly link them to the Matachewan magmatism and large igneous province. Overall, it is rather challenging to get highly precise U-Pb ages on many of units in this area, due mainly to two related reasons: 1) all the pre-1850 Ma zircons are shocked and disturbed, and 2) the zircons are generally altered. Results on the Creighton Granite and the Copper Cliff Rhyolite are shown in Figure 19, showing the scatter and general “pull down” due to shock-induced Pb loss at 1850 Ma, with superimposed younger Pb loss. Many individual analyses, whether single grain or multigrain, are almost meaningless due to these There are also remnants of lower Huronian volcanics with interlayered sandstone layers, which here show graded bedding suggesting tops are to the north, not to the south. 37 �Proceedings of the 68th ILSG Annual Meeting - Part 2 combined Pb loss effects. This resulted in the circa 2350 Ma early age estimates for the Creighton Granite (Frarey et al., 1982), a result based on several un-abraded multigrain zircon fractions that would have averaged out all these effects and plot in the middle of the Pb loss array or triangle (see Figure 19). The current apex of this Pb loss triangle is constrained by analyses of single, small, best- preserved zircon grain fragments pre-treated by chemical abrasion. Collectively the data indicate a minimum age of 2455 Ma for the Creighton Granite and the co-magmatic Copper Cliff Rhyolite, with a most likely upper intercept age of 2459 +7/-4 Ma. Figure 19: U-Pb concordia diagram of the combined results on Creighton Granite and Copper Cliff Rhyolite samples, showing the complex Pb loss array or “triangle” formed by various Pb loss processes. Shock effects of the Sudbury impact have damaged most if not all of the zircon crystals to varying degrees (e.g., Krogh et al., 1996), with results being pulled down to an 1850 Ma lower intercept. The variably altered zircons were then affected by varying stages and degrees of younger Pb loss, including recent Pb loss. Large multigrain fractions, non-abraded, from the early Frarey et al. (1982) study, plot in the middle of the triangle, having averaged out all the various Pb loss processes. Only tiny, best-preserved, single zircon fragments pre-treated by chemical abrasion (CA) from recent studies (Bleeker et al., 2015) plot near the apex of the triangle, constraining both a minimum (2455 Ma) and most likely upper intercept age of 2459+7/-4 Ma for the felsic magmatism. In contrast, comparatively large laser ablation spots on these complicated zircons, without CA pre-treatment, will simply sample all this complexity in Pb loss and result in meaningless upper intercept ages. 38 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 10: Magma mingling structures between Creighton Granite and mafic magmas Stop 11: Basal contact of the Sudbury Igneous Complex 46.459667° N, 81.203009° W 46.450530° N, 81.195107° W 484411E, 5145142N 480515E, 5144125N Roadcuts along the Highway 144 Bypass expose the basal contact of the SIC on underlying Elsie Mountain dark, mafic volcanics and gabbro sills. As was shown at Stop 9, younging in the volcanic section is most likely to the north here, into the base of the SIC. These outcrops show how difficult it is to actually put one’s finger on the lower contact of the SIC, due to the nature of immediate footwall rocks. The section is locally cut by felsic dykes, which are “rheomorphic” dykes of melted footwall that were back-intruded into the base of the SIC. These dykes are of interest in terms of constraining final melt formation and migration in the footwall rocks, but they are very hard to date because they contain essentially all xenocrystic zircons from the underlying rocks (e.g., Creighton Granite), with all their shock damage, and very few newly grown zircon crystals. A couple of hundred metres farther to the north, there are the first outcrops of typical, massive, homogeneous norite from near the base of the Main Mass of the SIC. Just south of the intersection of the Lively regional road with the Highway 144 Bypass, on the east side of the highway, are roadcuts through the Creighton Granite with classic magma mingling structures: rounded blobs of mafic magma suspended in surrounding Creighton Granite. This is significant in the sense that it clearly demonstrates contemporaneous mafic and felsic magmatism in an overall bimodal magmatic system. This is relevant to the interpretation of the Creighton Granite and the confusion about the “Blezardian orogeny”. The Creighton Granite is not a terminal collisional granite, i.e. the interpretation that fed the idea of a Blezardian orogeny, but rather an early A-type granite associated with the final rift-related magmatism at the base of the Huronian Supergroup. Stop 12: Top of the norite section, across the transition zone gabbro, and into the base of the granophyre section 46.487202° N, 81.207053° W 484109E, 5148202N About 3 km north of Stop 11, just north of the powerline and in a lazy curve of the Highway 144 Bypass, occurs the transition zone from uppermost norite, into a ~500 m thick gabbro section where augite becomes the dominant pyroxene, rather than hypersthene, and into the base of the thick granitic granophyre section. Figure 21 shows the typical variation in modal mineralogy across the SIC, as compiled from various sources. The “Transition Zone Gabbro” crystallized oxides and apatite, giving it a much higher magnetic susceptibility. Figure 22 shows some petrographic details of the “black norite”. Figure 20: Typical magma co-mingling structure of rounded blobs of mafic magma interacting with K-feldspar porphyritic granitoid magma of the Creighton Granite. 39 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 21: First-order variations in modal mineralogy, density, and whole-rock chemistry across the Main Mass of the SIC. Figure A and B from Naldrett and Hewins (1984); C from observations by the first author; D, E, F and G from Zieg and Marsh (2005), North Range; F and H from Lightfoot and Zotov (2005), South Range along Highway 144. Profiles resized to a common scale to highlight first-order features. Note the interesting spike in Ni values in the South Range “black norite” (see star in H), where samples also show large scatter in La (see G)), approaching values in basal norite. This area in South Range norites also shows shallow mineral lamination, suggesting a major structure may repeat the basal norites (see section of Figure 18). Grey bars are: in H, initial Ni values in early quartz diorite, highlighting the large Ni depletion in much of the SIC; the dashed line indicates Ni values in glassy melt fragments in the Onaping Formation (Ames et al., 2002); in F and G, upper crustal average values from Rudnick and Gao (2005). Based on normal liquidus and solidus temperatures for these various compositions, one predicts that the base of the granophyre would crystallize last (see Figure 6). Overall heat loss would be highest from the roof of the SIC, and less so into the footwall of the SIC. Crystallization of the norites and gabbro would add latent heat of crystallization into the base of the granophyre section, thus keeping it hot and molten until the upper crystallization front closed in on it from above. Toward the top of the gabbro section occurs a ~50 m-wide zone of very coarse-textured “Crowsfoot Granophyre” (Figure 23). This probably reflects the final accumulation of H2O in an evolved residual gabbro magma, promoting coarse crystal growth in the residual melt. So, although this unit is called “granophyre” it is probably better seen as the top of the gabbro section. Up from here, one enters the base of the “Granophyre” proper, which is of broadly granitic composition and fairly evenly grained and homogeneous. 40 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 22: Thin section image (SEM, back-scattered electron map) of “black norite” showing typical mineral lamination of platy feldspars; typical zircon and baddeleyite crystals are highlighted and shown in close-ups on the right. Note the characteristic dendritic or skeletal habit of the zircons, typical for the norites, suggesting crystallization in a rapidly cooling melt, originally superheated with no nuclei. Zircons from this sample were dated using the Pb evaporation technique, resulting in an age of 1849.7±0.2 (Bleeker et al., 2015). With these expectations, we attempted to precisely date final crystallization of the Main Mass and this resulted in a precise and concordant zircon age of 1850.0±0.9 Ma, which fully overlaps with the best results on ages for the norites. From this we can conclude that crystallization of the entire SIC melt sheet was all within a million years, and likely well within the current resolution of the best available ages. As our ages get better and more precise, this age range of crystallization and cooling may shrink further. The boundary between the top of the Transition Zone Gabbros and the base of the Granophyre represents a major boundary in physical parameters, among them density (Figure 21). The granophyres are on average 1.4 g/cm3 less dense than the norite section and much less dense than the oxide-rich gabbros. The base of the thick granophyre section, also last to crystallize, would thus have equilibrated in an essentially horizontal position after settling of the thick melt sheet. It thus represents an important “paleo-horizontal” reference surface when thinking about the overall structure. 41 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 23: Coarse-grained radiating textures of the “crows foot granophyre”, which probably represents the H2O-rich residual liquid at the top of the Transition Zone Gabbro section. It resembles typical pegmatitic zones at the top of other large differentiated gabbro sill complexes (e.g., Nipissing Diabase sills). If so, this is indeed better interpreted as the top of the gabbros, rather than the base of the granophyre section of the SIC. Stop 13: Typical granophyre, lower half of the Granophyre section Stop 14: Basal section of the Onaping Formation, “Grey Member” rich in angular to rounded Huronian quartzite fragments 46.498015° N, 81.204066° W 46.524529° N, 81.195052° W 484341E, 5149403N 485040E, 5152347N A large pull-out on the west side of the Highway 144 Bypass provides an easy place to examine typical granophyre about 1 km above the transition with the gabbros. Here the granophyre is medium grained, and evenly textured. This spectacular outcrop shows a perfect example of the clast-rich base of the Onaping Formation, with a clast population dominated by Huronian quartzite fragments 1–20 cm in size, somewhat rounded to angular. The fragmental material is tightly compacted into a “suevite breccia”, the latter name used when impact melt clots can be recognized. Early workers on the “Sudbury Irruptive”, among them petrologists, had noted of course that the SIC was unusual in two main ways: 1) the basal norites being more silica-rich (~56–58 wt%) than other large mafic intrusions, and 2) being characterized by a very thick granitic “granophyre” section. Typical large layered intrusion would differentiate into a mafic-ultramafic layered base, and overall perhaps ~10% of evolved granitic granophyric material underneath the chilled roof section. Clearly something was odd about Sudbury! The fragmental rocks are moderately to strongly deformed with a strong, southeast-dipping schistosity/cleavage that is axial planar to the overall syncline of the Sudbury structure, with a stretching lineation that plunges down-dip (on the cleavage plane). Although these structural elements are largely Penokean in origin, they were likely amplified by the younger deformation of the South Range Shear Zone (Shanks and Schwerdtner, 1991). 42 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 15: Upper member of the Onaping Formation (“Black Member”) core of the doubly-plunging Sudbury Basin syncline, which preserved the folded foreland sedimentary wedge from regional uplift and erosion. 46.535044° N, 81.185639° W 485765E, 5153514N Prior to the regional uplift, over the course of the Proterozoic, the Chelmsford and Rove Formation foreland wedges may well have connected. With extensive rock cuts on both sides of the Highway Bypass, this is a good outcrop to examine the upper part of the Onaping Formation, which is finer grained and overall upward fining (see Figure 6), darker, and traditionally was called the “Black Member” of the “fall-back breccias”. The darker colour of the breccias going up-section reflects, in part, a finer-grained matrix, and also an increasing carbon content. Bedding in the Chelmsford Formation is varied, with locally thick sandy turbidite beds. Bedding is folded and overprinted by the regional cleavage that is perfectly aligned with the axial plane of the overall fold structure. Large carbonate concretions are deformed and give an indication of the finite strain. In a general sense this “Grey Member” and “Black Member” terminology is still useful, but more detailed mapping in the last two decades has shown a more complex system of different depositional units including volcanic deposits due to venting of a still active melt sheet, and debris flows (see the work by Ames et al. (2008a,b) and references therein). So only part of the Onaping Formation represents true suevitic fall-back breccias. Other parts were reworked or washed back into the crater by processes other than strict fall back. Here along the highway, and the railway cut above, the Black Member of the Onaping Formation is well exposed in large rock cuts. Micro-diamonds have been reported from these rocks (Masaitis et al., 1999; see also French, 2004). Stop 16: Greywacke turbidites of the uppermost Whitewater Group, the Chelmsford Formation Stop 18: Transition Zone Gabbro, North Range, at Highway 144 – Highway 8 intersection Stop 17: Black Member of the Onaping Formation on the North Range, along Highway 144 at Onaping Falls 46.589851° N, 81.382453° W 470702E, 5159658N 46.617271° N, 81.413896° W 46.575577° N, 81.289467° W 468309E, 5162717N 477820E, 5158042N Outcrops near this intersection expose the Transition Zone Gabbro. The gabbro is medium grained, and characterized by augite being the main pyroxene, with little or no hypersthene. The gabbros also show a spike in oxide (Fe, Ti) and apatite (P) crystallization, typical of the “peak” in relative Fe-Ti concentrations during progressive crystallization, as seen in AFM diagrams of relatively reduced and anhydrous mafic magmas. This is reflected in a spike in magnetic susceptibility, with values increasing by an order of magnitude. The transition zone gabbro is several hundred metres thick and upward transitions into the base of the granophyre section. This roadside outcrop shows the folded and cleaved greywacke turbidites of the Chelmsford Formation, the uppermost formation preserved in the core of the Sudbury Basin syncline. The approximate age of these turbidites is circa 1840 Ma and they represent the foreland depositional wedge of the Penokean orogen. Participants of the fieldtrip who have seen the Rove Formation in the Thunder Bay area will recognize the great similarities and, indeed, these two formations are broadly related. The reason these turbidites are only locally preserved has to do with the fairly high-amplitude down folding in the 43 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 19: Felsic norite towards top of norite section on North Range plagioclase pheno/megacrysts. These dykes are circa 2460 Ma in age and part of the giant Matachewan (II) dyke swarm. They represent the feeders to the plagioclase-phyric basalts of the lower Huronian mafic volcanic rocks seen at Stop 7 north of Lively. They are numerous in this part of the Superior Province and provide an important time marker. As pointed out earlier, they do not cut the Copper Cliff Rhyolite, which represents the terminal phase of the Huronian rift magmatism and volcanism. 46.615361° N, 81.428390° W 467198E, 5162571N These outcrops in the broad curve of the highway at the bottom of the hill expose the top of the North Range norites, referred to as “Felsic Norite” (e.g., Naldrett and Hewins, 1984). The norite is massive and homogeneous, as is typical for most of the SIC rocks. SiO2 contents of these norites are ~58 wt%. Ni values are ~20–30 ppm, which is significantly depleted from values of what are thought to have been primary Ni values in the early undifferentiated impact melts that were perhaps as high as ~100– 200 ppm (Zieg and Marsh, 2004; Lightfoot and Zotov, 2005; Lightfoot, 2006). Both the gneisses and the Matachewan dykes are overprinted by Sudbury Breccia, and shatter cones can be seen at several localities, with the cone axes projecting to the southeast. One well-developed cone in the gneisses has a cone axis of ~120º/+20º (up). A subtle mineral lamination, formed mainly by alignment of platy plagioclase crystals, can be seen and dips moderately to the south, parallel to the attitude of the basal SIC contact in this part of the North Range. Magnetic susceptibility values are ~7.5±1.0 x10-3 SI units. Well-preserved pseudotachylite of the Sudbury Breccia bodies, locally up to ~1 m wide, is very dark in colour here below the North Range, and only weakly recrystallized. This contrasts with pseudotachylite on the South Range (e.g., Stop 9), where it is typically strongly recrystallized and metamorphosed to epidote amphibolite facies and shows cleavage/foliation development due to the more intense deformation of the South Range. Stop 20: Levack Gneisses, cut by Matachewan diabase dykes, all overprinted by Sudbury Breccia and southeast pointing shatter cones 46.624698° N, 81.444722° W Given this relatively good state of preservation of the black pseudotachylite matrix here at these outcrops, it is unlikely that the Levack Gneisses were at a lower crustal level, and thus hot, at the time of impact. The gneisses were likely exhumed to shallower crustal levels during the latest Archean or earliest Paleoproterozoic, well prior to the impact (e.g., James et al., 1992). This point has been debated in the Sudbury literature and is relevant to the question of where “Ground Zero” is located. To the northeast of the Sudbury structure, Levack Gneisses occur in close proximity to the unconformity with the Huronian succession. 465954E, 5163556N Roadcuts on both sides of the highway show typical “Levack Gneisses”, high grade migmatitic gneisses that characterize the Archean basement to the north of the SIC. These rocks reached pyroxene granulite facies grade, before being retrogressed to amphibolite facies in the latest Archean (e.g., Prevec et al., 2005, and references therein). The gneisses contrast with more homogeneous late Archean granites farther north, the Cartier Granites, dated at circa 2640 Ma (Meldrum et al., 1997). The gneisses are cut by large, SSE-trending, subvertical mafic dykes with conspicuous 44 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 24: Shatter cones and Sudbury Breccia in the Levack Gneisses ~0.5 km below the basal contact of the SIC on the North Range. A) Photo of typical, heterogeneous, coarse-textured Levack Gneiss. B) Relatively well-developed partial shatter cone in Levack Gneiss, ~0.5 m in size. C) Shocked zircon with multiple sets of planar deformation features and fractures, from the original study of U-Pb dating of these rocks (see Krogh et al., 1984). D) Well-developed Sudbury Breccia with displaced and variably rounded gneiss fragments floating around in a black pseudotachylite matrix. E) Small shatter cones, ~10-25 cm in size, developed within the Matachewan dyke at this stop, with cone axes focusing toward the southeast (to the right in this picture). 45 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 22: Discovery outcrop along railway cuts, just east of Murray Mine, South Range Stop 21: Cartier Granites, with locally conspicuous Sudbury Breccia, and rare shatter cones 46.521393° N, 81.052983° W 46.668937° N, 81.532490° W 495936E, 5151982N 459268E, 5168513N At this final stop we visit the approximate “discovery outcrop” just east of the Murray Mine. When in 1883 a railway line was cut through this area, Fe-sulphides and some chalcopyrite were noted in the rock cuts. This led to more prospecting and, eventually, the first mining operations in the area. At the time, the main interest was in Cu, as the application of Ni to steel manufacturing had not yet been invented. This final stop or stops, ~10 km out (horizontally) from the footwall contact of the SIC, will examine the more homogeneous late Archean granites of the “Cartier Batholith”, part of the “Algoman granites” to the south of the main Abitibi granite-greenstone terrane. Although somewhat variable and locally showing relict layering, typical parts of the Cartier Granite are homogeneous, relatively massive, pink, and relatively K-feldsparrich late granites with little structure. However, despite locally conspicuous K-feldspar, petrographically the granites are largely monzogranitic to granodioritic in composition, typical for late Archean granites. They contrast with and intrude the older Levack Gneisses. These late-stage granites, reflecting final re-melting of earlier tonalite-trondhjemite-granodiorite (TTG)dominated granite-greenstone crust, have been dated at circa 2640 Ma (Krogh et al., 1984; Meldrum et al., 1997)). They represent a final stage in the Archean crustal evolution of the southern Superior craton prior to stabilization and “cratonization”. The original rock cut was a bit farther to the west, but both the railway and the roads were moved east to allow for the development of the Murray open pit, which is located on the other side of the highway. In the present railway cut, semi-massive sulphide veins anastomose around somewhat deformed mafic fragments, some containing minor disseminated sulphides, others with no sulphides. This kind of material is typical for the basal contact of the SIC and is generally referred to as the “Sublayer” or “Contact Sublayer” (e.g., Pattison, 1979). To the west, these minor sulphide stringers broaden out into the orebodies of the Murray Mine (Figure 25). Occasional shatter cones can be seen in the road cuts, all pointing to the southeast (e.g., GPS waypoint #2686). And there is abundant Sudbury Breccia in places (GPS #2522), forming veins, dykes, and larger breccia bodies with rounded fragments of granite in a black pseudotachylite matrix. Similar Sudbury Breccia occurrences can be mapped radially outwards for another ~75 km, out to ~120 km from “Ground Zero” (see Figure 8; see also Butler, 1994; and Thompson and Spray, 1994) 7. Figure 25 shows a composite section across the basal contact of the SIC in this area, illustrating the SIC dipping to the north at ~40–45º. Relationships are approximately similar to that shown in the Creighton Mine section. The Murray Mine area could be described as yet another minor embayment, just to the east of the major Copper Cliff funnel structure. Confusion is possible with pseudotachylite veinlets formed in relation to regional faults, unrelated in time to the Sudbury structure. However, such veins are typically 0.5–2 cm in width and can mapped along or in proximity to observed fault or slip planes. Almost all Sudbury Breccia, such as discussed here, is developed at a different scale, often as dykes or bodies up to 10-100 cm wide. 7 46 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 25: Composite cross-section, looking west, across the footwall contact of the SIC illustrating the first-order relationships at the Murray Mine and open pit, just to the west of the “Discovery Outcrop” locality along the railway bed, and the Thayer Lindsley Mine farther to the east. Sections of both mines are integrated at the same scale. The Murray Mine section is from old Inco data, as published by Naldrett (1984), and differentiates some of the ore types, all in close proximity to or right along the footwall contact. The outline of the open pit is schematic. The Lindsley Mine section is from old Falconbridge data (see Binney et al., 1994), as published by Bailey et al. (2004). It is one of the localities where the significant offsets due to structures associated with the South Range Shear Zone was first recognized. Ore bodies along this shear zone were highly deformed. A Cu-rich ore body occurred within the footwall complex, dominated by Murray Granite. 47 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Papers by Robert Dietz and related literature (in chronological order) Incorporated into the section of Figure 25, at the same scale, is a section across the Thayer Lindsley Mine. This latter mine is located ~5 km to eastnortheast along strike and represents a next shallow embayment along the footwall contact of the SIC. Papers by Dietz Dietz, R.S. 1959. Shatter cones in cryptoexplosion structures (meteorite impact?). The Journal of Geology, v.67 (5), p.496-505. At the Lindsley Mine, at depth, the basal contact of the SIC is offset by a well-defined, discrete, shear zone that dips to the southeast and shows significant south-side up displacement (Binney et al., 1994). Metamorphic grade in this shear zone is lower amphibolite facies and titanite crystals from sheared norite show two distinct growth phase, brown titanite overgrown by colourless titanite (Bailey et al., 2004). U-Pb data for two brown titanite fractions suggest they grew at circa 1815±15 Ma, whereas the colourless titanites, interpreted to be syntectonic relative to the shear zone fabric, record an imprecise but somewhat younger age, circa 1670±70 Ma (Bailey et al., 2004). Dietz, R.S. 1960. Meteorite impact suggested by shatter cones in rock: Three cryptoexplosion structures yield new evidence of natural hypervelocity shocks. Science, v. 31 (3416), p.1781-1784. Dietz, R.S. 1961. Vredefort Ring structure: meteorite impact scar?. The Journal of Geology, v.69 (5), p.499-516. Dietz, R.S. 1961. Astroblemes. Scientific American, v.205 (2), p.0–59. Dietz, R.S. 1962. Vredefort Ring structure—a reply. The Journal of Geology, v.70, p.502-504. Dietz, R.S. 1963. Astroblemes: ancient meteorite impact scars on earth. In: The Solar System, Volume 4, University of Chicago Press, Chicago. The latter age dates this marked shear zone, which is part of the South Range Shear Zone system that further shortened and imbricated the South Range of the SIC. Dietz, R.S. 1963. Collapsing continental rises: an actualistic concept of geosynclines and mountain building. The Journal of Geology, v.71, p.314-332. Dietz, R.S. 1964. Sudbury structure as an astrobleme. The Journal of Geology, v.72 (4), p.412-434. End of Road Log for Day 1 Dietz, R.S. and Butler, L.W. 1964. Shatter-cone orientation at Sudbury, Canada. Nature, v204 (4955), p.280-281. Dietz, R.S. 1970. Cosmogenic ores at Sudbury astrobleme?. Meteoritics, v.5, p. 91-192. Dietz, R.S. 1971. Shatter cones (shock fractures) in astroblemes. Meteoritics, v.6, p. 58-259. Dietz, R.S. 1971. Sudbury astrobleme: A review. Meteoritics, v.6, p.259-260. Dietz, R.S. 1972. Sudbury astrobleme, splash emplaced sub-layer and possible cosmogenic ores. In: New Developments in Sudbury Geology, J.V. Guy-Bray (ed.), Geological Association of Canada, Special Paper 10, p. 29-40. 48 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Some directly related literature Ames, D.E., Golightly, J.P., Lightfoot, P.C. and Gibson, H.L. 2002. Vitric compositions in the Onaping Formation and their relationship to the Sudbury Igneous Complex, Sudbury Structure. Economic Geology, v.97, p.1541-1562. Shand, S.J. 1916. The pseudotachylyte of Parijs (Orange free State), and its relation to ‘Trap-Shotten Gneiss’ and ‘Flinty Crush-rock’. Quarterly Journal of the Geological Society, v.72 (1-4), p.198–221. Ames, D.E., Buckle, J., Davidson, A. and Card, K. 2005. Sudbury bedrock compilation. Geological Survey of Canada Open File 4570, geology, color map, and digital tables, scale 1:50,000. Boon, J.D. and Albritton, Jr, C.C. 1936. Meteorite craters and their possible relationship to "cryptovolcanic structures". Field and Laboratory, v.5, p.1–9. Ames, D.E., Card, K., Wodicka, N. and Davidson, A. 2008a. 100K Geological map of the Sudbury mining camp and Surrounding area, Ontario, Canada; Supplement to Ames, D.E. and Wodicka, N., 2008, Geology of the Giant Sudbury Polymetallic Mining Camp, Ontario, Canada. Economic Geology, v.103, (5), p.1057-1077. Boon, J.D. and Albritton Jr, C.C. 1937. Meteorite scars in ancient rocks. Field and Laboratory, v.5 (2), p.5364. Boon, J.D. and Albritton Jr, C.C. 1938. Established and supposed examples of meteoritic craters and structures. Field and Laboratory, v.6 (2), p. 4–56. Ames, D.E., Davidson, A. and Wodicka, N. 2008b. Geology of the giant Sudbury polymetallic mining camp, Ontario, Canada. Economic Geology, v.103 (5), p.1057–1077. Boon, J.D. and Albritton Jr, C.C. 1942. Deformation of rock strata by explosions. Science, v.96 (2496), p.402-403. Hargraves, R.B. 1961. Shatter cones in the rocks of the Vredefort Ring. Transactions of the Geological Society of South Africa, v. 4, p.147-161. Bailey, J., LaFrance, B., McDonald, A.M., Fedorowich, J.S., Kamo, S. and Archibald, D.A. 2004. MazatzalLabradorian-age (1.7-1.6 Ga) ductile deformation of the South Range Sudbury impact structure at the Thayer Lindsley mine, Ontario. Canadian Journal of Earth Sciences, v.41, p.1491-1505. DOI: 10.1139/e04-098. Guy-Bray, J.V. and Geological Staff 1966. Shatter cones at Sudbury. The Journal of Geology, v.74, p.243245. French, B.M. 1967. Sudbury structure, Ontario. Some petrographic evidence for an origin by meteorite impact. Goddard S pace Flight Centre, Maryland, Publication X-641-67-67, p.1-56. Bekker, A., Holland, H.D., Wang, P.L., Rumble, D.I.I.I., Stein, H.J., Hannah, J.L., Coetzee, L.L. and Beukes, N.J. 2004. Dating the rise of atmospheric oxygen. Nature, v.427 (6970), p.117-120. Bourgeois, J. and Koppes, S.,1998. Robert S. Dietz and the recognition of impact structures on Earth. Earth sciences history, v.17 (2), p.139-156. Bennett, G., Dressler, B.O. and Robertson, J.A. 1991. The Huronian Supergroup and associated intrusive rocks. In: Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p. 549-592. References related to Day 1 Addison, W.D., Brumpton, G.R., Vallini, D.A., McNaughton, N.J., Davis, D.W., Kissin, S., Fralick, P.W. and Hammond, A.L. 2005. Discovery of distal ejecta from the 1850 Ma Sudbury impact event. Geology, v.33 (3), p. 93-196. Binney, W.P., Poulin, R.Y., Sweeney, J.M. and Halladay, S.H. 1994. The Lindsley Ni–Cu–PGE Deposit and its Geological Setting. In: Proceedings of the Sudbury–Noril’sk Symposium. Ontario Geological Survey, Special Volume 5, p.91-103. 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 basesurge deposits or tsunami deposits?. In: Large Meteorite Impacts and Planetary Evolution IV. Geological Society of America Special Paper 465, p.245-268. DOI: 10.1130/2010.2465(16). Bleeker, W., 2003. The late Archean record: a puzzle in circa 35 pieces. Lithos, v.71 (2-4), p.99-134. Bleeker, W. 2004. Taking the pulse of planet Earth: a proposal for a new multi-disciplinary flagship project in Canadian solid Earth sciences. Geoscience Canada, v.31 (4), p.179-190. 49 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Card, K.D., Gupta, V.K., McGrath, P.H. and Grant, F.S. 1984. The Sudbury structure: its regional geological and geophysical setting. In: The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p. 25-43. 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: Selected Papers of the Fifth International Dyke Conference in Finland, Rovaniemi, Finland, 31 July - 3 Aug 2005 & Fourth International Dyke Conference, Kwazulu-Natal, South Africa 26-29 June 2001, E. Hanski, S. Mertanen, T. Rämö, J. Vuollo (eds.); A.A. Balkema, Rotterdam, p. 3-26. Cochrane, L.B. 1984. Ore deposits of the Copper Cliff Offset. In: The Geology and Ore Deposits of the Sudbury Structure. Ontario Geological Survey, Special Volume 1, p. 347-359. Coleman, A.P.1905. The Sudbury nickel field. Report of Ontario Bureau of Mines, Annual Report for 1905, v.14, pt.3. Bleeker, W., Kamo, S. and Ames, D.E. 2014. Towards a detailed geological cross-section of the deformed Sudbury impact basin: New observations and new geochronology. Presentation at the Ontario Exploration and Geoscience Symposium, 3–4 November, 2014, Sudbury. Corfu, F. and Andrews, A.J. 1986. A U-Pb age for mineralized Nipissing diabase, Gowganda: Canadian Journal of Earth Sciences, vol. 23, p. 107-109. Corfu, F. and Lightfoot, P.C. 1996. U-Pb geochronology of the sublayer environment, Sudbury igneous complex, Ontario. Economic Geology, v.91 (7), p.1263-1269. 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: Targeted Geoscience Initiative 4: Canadian NickelCopper-Platinum Group Elements-Chromium Ore Systems—Fertility, Pathfinders, New and Revised Models, Geological Survey of Canada, Open File 7856, p.151-166. DOI: 10.4095/296686. Cowan, E.J., Riller, U. and Schwerdtner, W.M. 1999. Emplacement geometry of the Sudbury Igneous Complex: Structural examination of a proposed impact melt-sheet. In: Large Meteorite Impacts and Planetary Evolution II; Geological Society of America, Special Paper 339, p.399-418. Bradley, D.C. 2008. Passive margins through earth history. Earth-Science Reviews, v.91 (1-4), p.1-26. Craig, J.R. and Kullerud, G. 1969. Phase relations in the Cu-Fe-Ni-S system and their application to magmatic ore deposits. In Magmatic Ore Deposits— A Symposium, Economic Geology, Monograph 4, p.344–358. Brocoum, S.J. and Dalziel, I.W. 1974. The Sudbury Basin, the southern province, the Grenville Front, and the Penokean orogeny. Geological Society of America Bulletin, v.85 (10), p.1571-1580. Davey, S., Bleeker, W., Kamo, S., Davis, D., Easton, M. and Sutcliffe, R.H, 2019. Ni-Cu-PGE potential of the Nipissing sills as part of the circa 2.2 Ga Ungava large igneous province. In: Targeted Geoscience Initiative: 2018 Report of Activities, N. Rogers (ed.). Geological Survey of Canada, Open File 8549, p.403-419. DOI: 10.4095/313675. Butler, H.R. 1994. 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Dickin, A.P., Artan, M.A. and Crocket, J.H. 1996. Isotopic evidence for distinct crustal sources of North and South Range ores, Sudbury Igneous Complex. Geochimica et Cosmochimica Acta, v.60, p.1605-1613. DOI:10.1016/0016-7037(96)00044-0. Zieg, M.J. and Marsh, B.D. 2005. The Sudbury Igneous Complex: Viscous emulsion differentiation of a superheated impact melt sheet. Geological Society of America Bulletin, v.117, p.1427-1450. Dickin, A.P., Nguyen, T. and Crocket, J.H. 1999. Isotopic evidence for a single impact melting origin of the Sudbury Igneous Complex. In: Large meteorite impacts and planetary evolution II, Geological Society of America, Special Paper 339, p. 361-371. Other relevant references Ames, D.E., Watkinson, D.H. and Parrish, R.R, 1998, Dating of a regional hydrothermal system induced by the 1850 Ma Sudbury impact event. Geology, v.26, p.447–450. Gariépy, C. and Allègre, C.J. 1985. The lead isotope geochemistry and geochronology of late-kinematic intrusives from the Abitibi greenstone belt, and the implications for late Archaean crustal evolution. Geochimica et Cosmochimica Acta, v.49 (11), p. 2371-2383. DOI: 10.1016/0016-7037(85)90237-6. Anders, D., Osinski, G.R., Grieve, R.A.F., Pilles, E.A., Pentek, A. and Smith, D. 2020. Origin and formation of Metabreccia in the Parkin Offset Dike, Sudbury impact structure, Canada. Canadian Journal of Earth Sciences, v.57 (11), p.1324-1336. Bailey, J., McDonald, A.M., Lafrance, B. and Fedorowich, J.S. 2006. Variations in Ni content in sheared magmatic sulfide ore at the Thayer Lindsley mine, Sudbury, Ontario. The Canadian Mineralogist, v.44 (5), p.1063-1077. Holm, D.K., Van Schmus, W.R., MacNeill, L.C., Boerboom, T.J., Schweitzer, D. and Schneider, D. 2005. U-Pb zircon geochronology of Paleoproterozoic plutons from the northern midcontinent, USA: Evidence for subduction flip and continued convergence after geon 18 Penokean orogenesis. Geological Society of America Bulletin, v.117 (3-4), p.259-275. Bleeker, W., Kamo, S. and Ames, D.E. 2013. New field observations and U-Pb age data for footwall (target) rocks at Sudbury: Towards a detailed cross-section through the Sudbury Structure. In: Large Meteorite Impacts and Planetary Evolution V Meeting, 5–8 August, Sudbury, Ontario. Extended abstract, Lunar Planetary Institute contribution no. 1737, p. 13. Ivanov, B.A. and Deutsch, A, 1997. Sudbury impact event: cratering mechanics and thermal history. In: Large Meteorite Impacts and Planetary Evolution, LPI Contribution no. 922, p. 26. Brocoum, S.J. and Dalziel, I.W. 1976. The Sudbury Basin, the Southern province, the Grenville Front, and the Penokean orogeny; Discussion and reply: Reply. Geological Society of America Bulletin, v.87 (6), p.958-958. Kawohl, A., Frimmel, H.E., Bite, A., Whymark, W. and Debaille, V. 2019. Very distant Sudbury impact dykes revealed by drilling the Temagami geophysical anomaly. Precambrian Research, v.324, p.220-235. DOI: 10.1016/j.precamres.2019.02.014. 56 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Keays, R.R. and Lightfoot, P.C., 2004. Formation of NiCu-PGE sulphide mineralization in the Sudbury Impact Melt Sheet. Mineralogy and Petrology, v.82, p.217-258. Mungall, J.E., Ames, D.E, and Hanley, J.J. 2004. Geochemical evidence from the Sudbury structure for crustal redistribution by large bolide impacts. Nature, v.429 (6991), p.546-548. Keays, R.R. and Lightfoot, P.C. 2004. Mafic intrusions in the footwall of the Sudbury Igneous Complex: Origin of the Sudbury impact melt sheet and its associated ore deposits. Ore Geology Reviews, v.120, article 103435. DOI: 10.1016/j.oregeorev. 2020.103435. Prevec, S.A., Lightfoot, P.C, and Keays, R.R. 2000. Evolution of the Sublayer of the Sudbury Igneous Complex: geochemical, Sm-Nd and petrologic evidence. Lithos, v.51, p.271-292. Rousell, H.D. 1972. The Chelmsford Formation of the Sudbury Basin—a Precambrian turbidite. In: New Developments in Sudbury Geology, Geological Association of Canada, Special Paper 10, p.79-91. Ketchum, K.Y., Heaman, L.M., Bennett, G. and Hughes, D.J. 2013. Age, petrogenesis and tectonic setting of the Thessalon volcanic rocks, Huronian Supergroup, Canada. Precambrian Research, v.233, p.144-172. Rousell, H.D. 1975. The origin of foliation and lineation in the Onaping Formation and the deformation of the Sudbury Basin. Canadian Journal of Earth Sciences, v.12, p.1379-1395. Lightfoot, P.C., Keays, R.R. and Doherty, W. 2001. Chemical evolution and origin of nickel sulfide mineralization in the Sudbury Igneous Complex, Ontario, Canada. Economic Geology, v.96, p.18551875. Rousell, H.D. 1984a. Structural geology of the Sudbury Basin. In: The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p. 83-95. Lightfoot, P.C., Keays, R.R., Morrison, G.G., Bite, A. and Farrell, K. 1997. Geologic and geochemical relationships between the Contact Sublayer, inclusions, and the Main Mass of the Sudbury Igneous Complex: A case study of the Whistle Mine embayment. Economic Geology, v.92, p.647-673. Rousell, H.D. 1984b. Onwatin and Chelmsford Formations. In: The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p. 211-218. Shanks, W.S. and Schwerdtner, W.M., 1991. Crude quantitative estimates of the original northwest– southeast dimension of the Sudbury Structure, southcentral Canadian Shield. Canadian Journal of Earth Sciences, v.28 (10), p.1677-1686. Lightfoot, P.C., Keays, RR., Morrison, G.G., Bite, A., and Farrell, K., 1997. Geochemical relationships in the Sudbury Igneous Complex: Origin of the Main Mass and Offset dikes. Economic Geology, v.92, p.289-307. Wieland, F., Gibson, R.L. and Reimold, W.U. 2005. Structural analysis of the collar of the Vredefort Dome, South Africa—Significance for impact‐ related deformation and central uplift formation. Meteoritics & Planetary Science, v.40 (9–10), p.1537-1554. Lightfoot, P.C. and Farrow, C.E.G. 2002. Geology, geochemistry, and mineralogy of the Worthington offset dike: a genetic model for offset dike mineralization in the Sudbury Igneous Complex. Economic Geology, v.97, p.1419-1446. DOI: 10.2113/gsecongeo.97.7.1419. Zolnai, A.I., Price, R.A. and Helmstaedt, H. 1984. Regional cross section of the Southern Province adjacent to Lake Huron, Ontario: implications for the tectonic significance of the Murray Fault Zone. Canadian Journal of Earth Sciences, v.21, p.447-456. Marsh, B.D. and Zieg, M.J. 1999. Melt sheet madness: superheated emulsion differentiation. Geological association of Canada–Mineralogical Association of Canada, Sudbury 1999, Abstracts v.24, p.78 Morrison, G.G. 1984. Morphological features of the Sudbury Structure in relation to an impact origin. In: The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p. 513-522. 57 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Day 2 Sudbury Ore Environments and Offset Dikes – Examples from Whistle and Parkin, NE Sudbury Henning Seibel and Michael Lesher Mineral Exploration Research Centre, Harquail School of Earth Sciences, Laurentian University, 935 Ramsey Lake Road, Sudbury, ON P3E 2C6 1. Organization of Day 2 The second day of pre-meeting Trip 1 will introduce participants to the ore environments at the base of the Sudbury Igneous Complex (SIC) and to the general geology of offset dikes that are exposed on mechanically and hydraulically stripped outcrops in the northeast corner of the Sudbury basin. The day starts adjacent to the former Whistle mine, mined by Inco (now Vale) from 1993 to 1997, where large outcrops highlight the complexity of Superior Province footwall rocks near the contact with the overlying SIC and the transition from mineralized Sublayer through anatectic breccias to traditional offset dike lithologies. The Podolsky North Zone outcrop to the northeast contains footwall-style mineralization on surface, extensions of which were mined underground by FNX/KGHM from 2008 to 2013. In the afternoon, outcrops south (distal Whistle) and north (proximal Parkin) of the Post Creek fault will be compared. The field trip will end with a visit to Rocky’s Restaurant on Lake Wanapitei. In case we are not able to access the Whistle mine outcrops, several interesting and well-preserved outcrops of the Worthington, Trill, and Hess offset dikes to the southwest and northwest of the Sudbury Structure will be visited. Participants will be introduced to typical characteristics of offset dikes, emplacement mechanisms and formation models. 2. Introduction The Sudbury mining camp is the one of the largest magmatic Ni-Cu-PGE mining camps in the world (Fig. 2.1) and has been mined for over 135 years (see review by Lightfoot, 2016). Mineralization is associated mainly with breccias along and near the lower contact of the Main Mass of the Sudbury Igneous Complex (SIC) and within associated offset dikes (Fig. 3.1). Breccias in the Sudbury Structure include 1) pre-impact magmatic breccias (e.g., Levack Breccia), 2) syn-impact pseudotachylitic breccias, locally referred to as Sudbury Breccia (SUBX; e.g., Rousell et al., 2003), 3) syn- to post-impact magmatic breccias directly derived from the SIC, such as Inclusion-Bearing Quartz Diorite (IQD; e.g., Grant and Bite, 1984) and inclusion-bearing Sublayer Norite (SLNR; e.g., Lightfoot and Farrow, 2002; Lightfoot et al., 1997a), and 4) contact metamorphosed and/or partially melted (anatectic) breccias, variably referred to as Footwall Breccia (FWBX; e.g., McCormick et al., 2002) and “Metabreccia” (MTBX; e.g., Lafrance et al., 2014). The ore deposits in the Sudbury Structure occur in two distinct environments (Fig. 3.1): 1) mineralization along or near the basal contact of the Main Mass, including a) disseminated to semi-massive Fe-Ni-Cu sulfides in SLNR and FWBX, and b) veins and disseminations of Fe-Cu-Ni sulfides in underlying SUBX and associated footwall rocks, and 2) disseminated-blebby to semi-massive Fe-Ni-Cu and Fe-Cu-Ni sulfides in offset dikes (e.g., Souch et al. 1969; Naldrett, 2004; Lightfoot, 2016). The Whistle embayment and Whistle-Parkin offset dike at the northeastern corner of the SIC contain elements of both environments, providing an excellent introduction to Sudbury ore systems. 58 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 2.1: Pre-mining Ni resources (past production + current resources) and grades of the world’s largest magmatic Cu-Ni (circles) and PGE (triangle) deposits. Modified from Naldrett (2004). Figure 3.1: Distribution of contact, footwall, and offset dike deposits and occurrences in the Sudbury Impact Structure. Note that locations for deeper ore bodies are projected to surface. Simplified after Ames et al. (2008). 3. Sudbury Ore Environments Ore deposits and occurrences occur all around the Sudbury Structure (Fig. 3.1), but some areas, such as Levack in the North Range and Frood-Stobie, Creighton, and Copper Cliff in the South Range, are much better endowed. 59 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3.2: A) Inclusion-bearing massive sulfide from the Stobie Mine (contact deposit). B) Sharp-walled Cu vein in contact with tonalitic country rock from the Strathcona mine (footwall deposit). C) Disseminated sulfide blebs in variable altered inclusion-bearing quartz diorite from the Copper Cliff North Mine (offset deposit). Courtesy of Harquail School of Earth Sciences (Laurentian University). Contact-Footwall Environment Contact ores occur discontinuously along the SIC-footwall contact and are typically hosted by SLNR and FWBX (also referred to as Granite Breccia or Late Granite Breccia on the North Range). Mineralization typically occurs in funnels (e.g., Whistle, Foy, Copper Cliff), troughs (e.g., Creighton), and embayments (e.g., Levack) (Fig. 3.3) along the basal contact of the SIC, and are subeconomic or absent outside of those features. Mineralization typically grades downward from sparse disseminated sulfides in overlying Main Mass norite through fine and coarse (blebby) disseminated sulfides in Sublayer to semi-massive sulfides in Footwall Breccia (Fig. 3.2A) (see review by Lightfoot, 2016). Contact ores contain a typical magmatic 60 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3.3.: Distribution of contact, footwall, and offset dike deposits and occurrences in the Sudbury Impact Structure. Note that locations for deeper ore bodies are projected to surface. Simplified after Ames et al. (2008). sulfide assemblage of pyrrhotite (containing up to 1% Ni) > pentlandite > chalcopyrite with minor magnetite and platinum-group minerals (PGMs) (see review by Ames et al., 2008). Ore tenors (metals in 100% sulfides) vary widely depending on the composition of the magma, magma:sulfide mass ratio (R factor: Naldrett et al., 1979), and degree of MSS fractionation (e.g., Li and Naldrett, 1994), but typically range from 3.9–6.1% Ni100, 1.3–7.1% Cu100 and 0.7–5.6 ppm (Pd100+Pt100) (Naldrett, 2004). Some contact deposits in the South Range (e.g., Thayer Lindsley: Bailey et al., 2004; Garson: Mukwakwami et al., 2014) are faulted, sheared, and remobilized. Footwall ores appear to occur only below contact deposits, up to 700m (Golightly, 2009) but more typically up to 200–300m (Farrow and Lightfoot, 2002) from the base of the SIC (Fig. 3.3). They are common in the North and East Ranges, but rare in the South Range (Fig. 3.1). Farrow et al. (2005) discriminate three types of footwall mineralization: 1) sharp-walled veins that can reach up to several meters in thickness with predominant chalcopyrite and lesser pentlandite, millerite, and cubanite grading downward and outward into bornite ± millerite veins (Fig. 3.2B), 2) disseminated sulfides often with high PGE/S ratios, and 3) a hybrid type containing both mineralization styles. Ore tenors typically range from 3.5–8.7% Ni100, 28.8–38.3% Cu100, and 13.4–33.5 ppm (Pt100+Pd100) (Naldrett, 2004). A gradual transition from Fe-Co-(Ni)-IPGE-rich contact mineralization to Cu-(Ni)-PPGE-Au-rich footwall mineralization at several deposits (e.g., Frood: Hawley, 1965; Strathcona: Li and Naldrett, 1994; McCreedy East: Gregory, 2006; Levack-Morrison: Nelles, 2012; Nickel Rim South: Glencore Ltd., unpubl.; Podolsky, KGHM unpubl.) is consistent with fractional crystallization and accumulation of Fe-Co-IPGE-rich monosulfide solid solution (MSS) in contact ores and segregation of Cu-rich intermediate solid solution (ISS) and Cu-PPGE-Au-rich residual sulfide liquid in footwall ores (e.g., Mungall, 2007). The formation of bornite-millerite ores across a thermal divide in the Fe-Cu-S system appears to require reactions with wall rocks (Nelles, 2012; see also Lesher, 2017). The formation of distal disseminated PGE-Au rich sulfides appears to require deposition from hydrothermal fluids (e.g., Hanley et al., 2004; Stout, 2009). 61 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3.4: Schematic drawing of typical relationships between marginal QD and interior (mineralized) IQD) in offset dikes as well as several footwall-offset dike relationships in the Worthington offset dike. After Lightfoot and Farrow (2002). Offset Dike Environment Offset dikes are sub-vertical radial or concentric quartz monzodioritic lithologies (historically referred to as quartz diorite) that extend up to 20 km into the underlying footwall rocks from 300–500m-wide funnels at the contact with Main Mass (Fig. 3.1; see review by Lightfoot, 2016). The funnels typically contain SLNR ± MTBX ± IQD-QD pods and grade downward/outward into inclusion- and sulfide-poor Quartz Diorite (QD) margins with inclusion- and sulfide-rich Quartz Diorite cores (IQD; Fig. 3.4; e.g., Pattison, 1979; Grant and Bite, 1984; Lightfoot and Farrow, 2002; Murphy and Spray, 2002; Tuchscherer and Spray, 2002). Inclusions comprise angular to subrounded xenoliths derived from local country rocks, anteliths derived from the offset dike lithologies (i.e., QD clasts in IQD), and ultramafic xenoliths derived from deeper crustal lithologies (Wang et al., 2020). They vary in size from microscopic to tens of meters, and can reach up to 90% in volume (Grant and Bite, 1984). Sulfide contents vary from negligible to massive and are generally linked to the presence of (ultra)-mafic inclusions (Pattison, 1979). Ni-Cu-PGE mineralization ranges from finely disseminated to blebby (Fig. 3.2C), semi-massive, and massive coarse-grained pyrrhotite with variable amounts of pentlandite and chalcopyrite in steeply plunging ore bodies (Cochrane, 1984; Farrow and Lightfoot, 2002). Ore tenors typically range from 3.2–6.5% Ni100, 2.6–12.8% Cu100, and 1.2–28.3 ppm (Pt100+Pd100) (Naldrett, 2004). Ore bodies are often associated with large (ultra)-mafic clasts (e.g., Totten, Podolsky), changes in strike (e.g., Copper Cliff), or cross-cutting faults (Cochrane, 1984). Endowment varies significantly between dikes located in the North Range and South Range, and with proximity to the SIC (Fig. 3.1). Most economic offset deposits (historic and current) are located in South Range dikes (e.g., Frood-Stobie, Copper Cliff). Small PGE-Cu-(Ni) occurrences in the recent discovered Rathbun offset dike indicate potential for untypical (footwall-style) mineralization in distal offset dikes (Kawohl et al., 2020). A gradual transition from Ni-rich contact mineralization to Curich footwall mineralization in some offset deposits (e.g., Frood: Hawley, 1965) is consistent with fractional crystallization of Ni-IPGE-rich MSS to produce residual Cu-PPGE-rich sulfide liquid. 62 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Ore Genesis The ultimate sources of S and metals in the SIC are Fe ± Cu ± Ni-sulfide bearing Archean mafic gneisses, Huronian mafic volcanic rocks, East Bull Lake Suite intrusions, and Nipissing Suite intrusions (Lightfoot et al., 1997, 2001; Keays and Lightfoot, 2004), all of which contain significant amounts of sulfides and the latter two of which contain significant amounts of Ni-Cu-PGE mineralization (e.g., James et al., 2002; Sproule et al., 2007; Holwell and Keays, 2014). However, that leaves two end-member models for the generation of the ores in the SIC (Fig. 3.5): 1) dissolution of Fe ± Cu ± Ni sulfides in the superheated impact melt followed by exsolution and settling during cooling (e.g., Lightfoot et al., 2001; Keays and Lightfoot, 2004) and 2) impact devolatilization of the majority of the S from the impact melt and incorporation of Fe ± Cu ± Ni sulfide xenomelts during thermomechanical erosion of footwall rocks (Lesher, 2019). The latter appears to be more consistent with very consistent Hf isotopic composition of the Main Mass and more heterogeneous Pb-S-Os isotopic compositions of the ores (see review by Wang et al., in press), indicating complete retention and homogenization of refractory Hf (Kenny et al., 2017) but significant loss of more-volatile Pb (O’Sullivan et al., 2016; Kenny et al., 2017; McNamara et al., 2017), Sb (O’Sullivan et al., 2016 GCA), Zn‐Cd‐Rb‐Cs (Kamber and Schoenberg, 2020), and therefore also much/most of the highly volatile S‐Se‐Bi and significant amounts of moderately volatile Ag‐Cu‐Au‐As (Lesher, 2019). Figure 3.5: Schematic representations of end-member sulfide generation and localization models (from Wang et al., in press). Model A: exsolution and convective settling of molten sulfide droplets (T1) followed by gravity flow into embayments, troughs, and funnels (T2). Model B: volatilization of most of the S from the impact melt (T1) followed by generation of sulfide xenomelts by local thermomechanical erosion (T2). 63 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The short interval between emplacement of inclusion- and sulfide-free QD margins and inclusion ± sulfide bearing IQD cores of nested offset dikes (<1–5 days: Wang et al., in press) provides insufficient time for exsolution and settling of sulfide droplets through the 2–5 km-thick impact melt, as dissolution, exsolution, and settling are all inherently slow processes (months: see discussion by Robertson et al., 2015, 2016), whereas the transfer of metals between the magma and sulfide droplets is much faster (hours-days: Yao and Mungall, 2021). Together, available data favor a model involving impact devolatilization of S and other volatile elements, rapid thermomechanical erosion of impact debris by the superheated impact melt, and dynamic upgrading of Fe ± Cu ± Ni sulfide xenomelts (Fig. 3.5). 4. Offset Dike Emplacement Extensive impact melt-bearing dikes are only known from the two largest terrestrial impact structures, Sudbury and Vredefort (Dressler and Reimold, 2004; Osinski et al., 2018, and references therein). Granophyre dikes are the only remnant of an impact melt at the 2023 Ma (Kamo et al., 1996) Vredefort impact crater. They are spatially related to the centrally uplifted dome structure of the deeply eroded impact site (Reimold and Gibson, 2006). Similar to Sudbury offset dikes, the granophyre dikes have a radial and concentrical distribution, widths of 10–50m, lengths of up to 10km, crosscutting relationships with pseudotachylitic breccia, spherulitic textures, and more rarely fragment-poor granophyre margins as well as fragment-rich interiors (Reimold and Gibson, 2006; Osinski et al., 2018; Huber et al., 2022). Granophyre dikes have a homogeneous chemical composition similar to the upper continental crust which could represent the undifferentiated impact melt (Dressler and Reimold, 2004; Huber et al., 2020). Offset Dike Characteristics Quartz Diorite (QD) is predominantly composed of a medium-grained, homogenous matrix with acicular plagioclase, acicular sometimes radiating amphiboles (after pyroxene), variable amounts of quartz and minor biotite, granophyric quartz-alkali feldspar intergrowths, and secondary amphibole. Aphanitic and spherulitic QD margins are common in the intermediate and distal segments of offset dikes and their apophyses (e.g., Hess, Trill). Inclusion-bearing QD (IQD) is mineralogically similar to QD, but typically finer grained with a more granular texture and less pronounced acicular amphibole. Contacts between QD and IQD are often sharp, but sometimes gradational (Fig. 4.3). QD and IQD samples from the same dike are remarkably similar in major and trace element geochemistry, indicating similar melt source and a short succession of injection (Lightfoot et al., 1997a; Pilles et al., 2017). Trace element differences between North Range and South Range dikes appear to reflect incorporation of different amounts of local footwall rocks (Lightfoot et al., 1997a). Mechanisms of Emplacement Two general mechanisms have been proposed to explain the geochemical, petrological, and spatial characteristics of QD and IQD. Some authors (e.g., Lightfoot and Farrow, 2002; Riller, 2005, Prevec and Büttner, 2018) prefer a multi-phase emplacement of two or more melts (Fig. 4.1), whereas others (e.g., Grant and Bite, 1984; Pilles et al, 2018) favor a single injection and flowage differentiation (e.g., Barrière, 1976) to explain the characteristics of some dikes (e.g., Foy). In a multi-phase injection model, an initial phase of sulfide-poor, inclusion-poor QD melt is followed by injection of a second core phase of sulfide-rich, IQD melt in the center of the dikes (Fig. 4.1; e.g., Lightfoot and Farrow, 2002; Prevec and Büttner, 2018). Further injections at later stages to form the more evolved Pele and Cascaden dikes are required, but they do not contain any sulfides (Pilles et al., 2018a). This model is supported by the common presence of inclusions of QD within IQD, common sharp contacts between QD and IQD (Fig. 4.3A), and the spatial relationship between marginal QD and interior IQD. 64 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4.1: Schematic diagram of multi-stage emplacement model. Injection of an initial phase of sulfidepoor, inclusion-poor QD melt is followed by injection of a second phase of sulfide-rich IQD melt in the center of the dikes. Modified after Prevec and Büttner (2018). In a single-stage injection model, IQD melt is injected into the fractured country rock with flowage differentiation producing marginal QD and interior IQD (Fig. 4.2). This model is supported by local gradational transitions between QD and IQD (Fig. 4.3B), rare IQD inclusions in IQD, increasing clast diameters towards the center of the dike, and clast alignments parallel to dike margins in the Foy offset dike (Pilles et al., 2018b). Figure 4.2: Schematic diagram of single-stage model. IQD melt is injected into dilating fractures with flowage differentiation producing marginal QD and interior IQD. Modified after Pilles et al. (2018b) The spatial relationships between marginal QD and internal IQD can be explained by flowage differentiation or by multi-stage emplacement if IQD intruded before QD had completely solidified. However, flowage differentiation and multiple injections should produce different inclusion types and contact relationships. For example, flowage differentiation cannot easily produce the commonly observed sharp contacts between QD and IQD or the frequent inclusions of QD in IQD. In contrast, the common sharp contacts can be produced if IQD intrudes QD, the rare gradational contacts if IQD melts QD (see discussion by Huppert & Sparks, 1985), and the rare inclusions of IQD in IQD by multiple phases of injection or by radial dikes crosscutting/intruding concentric dikes. The weight of evidence presently favors a multiple injection model (Fig. 4.4 and 4.5). 65 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4.3: Sharp (A) and well-defined but gradational (B) contact of medium-grained QD and IQD from two stripped outcrops of the Foy offset dike. Contact is defined by grain size differences as well as abundance of inclusions and blebby sulfides. Image widths are ca. 40cm. Timing of Emplacement Impact cratering can be subdivided into three stages: 1) contact and compression, 2) crater excavation, and 3) crater modification (Gault et al., 1968; Osinski and Pierazzo, 2012). The timing of emplacement of the Sudbury offset dikes proposed by different workers varies from during impact excavation to tens of thousands of years after impact: 1) dilation during transient cavity formation, 1 second to 1 minute after impact (e.g., Lightfoot and Farrow, 2002; QD: Wang et al., in press); 2) dilation during rebound and central uplift, minutes to days after impact (e.g., Tuchscherer and Spray, 2002; IQD: Wang et al., in press); 3) dilation during crater wall collapse, leading to injection of melt into transfer faults (e.g., Scott and Benn, 2002); 4) dilation during isostatic uplift, up to 10,000 years after impact (e.g., Wichman and Schultz, 1993); 5) dilation after melt pressure increase as a result of a coherent roof, 1500–130,000 years after impact (e.g., Prevec and Büttner, 2018); 6) dilation during cooling and subsequent contraction of footwall rocks, >10,000 years after impact (e.g., Riller, 2005); 7) dilation during cooling of the Main Mass, 10,000s to 100,000s of years after impact (Mathieu et al., 2021); 8) dilation during readjustment and late tectonic deformation, >10,000 years after impact (e.g., Therriault et al., 2002). 1) The presence of only sparse local xenoliths and sulfides, and the presence of aphanitic, spherulitic, and radiating “quench” pyroxene textures in the distal parts of the dikes require the impact melt to have been superheated when QD was emplaced, so this precludes all of the models involving emplacement during or after crystallization of the Main Mass. 2) The nested QD/IQD relationships with no occurrences of IQD without QD (with the possible exception of the South Range Breccia Belt) require emplacement while the cores of the QD dikes were still weak, which has been modelled by to have been within 2–5 days (Wang et al., in press). 3) The presence of inclusions and sulfides in IQD require the impact melt to have reached sulfide saturation within that time interval and for IQD to have been forcibly emplaced within that interval. 4) The two events most likely to have caused rapid basin-wide sequential injection of QD and IQD are excavation and impact (Fig. 4.4). 5) The process most likely to have driven the impact melt from a superheated, sulfide-undersaturated state to a liquidus, sulfide saturated state is assimilation of fragments generated by collapse of the peak ring. 66 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4.4: Schematic representation of complex crater-forming and -modification events leading to injection of inclusion- and sulfide-poor marginal QD (T1), generation of inclusion- and sulfide-rich IQD (T3), and melting of inclusions and generation of mineralized FWBX (T4). FWBX – Footwall Breccia, IQD – Inclusion-bearing Quartz Diorite, SLNR – Sublayer Norite, QD – Quartz Diorite. From Wang et al. (in press), as modified after Melosh (1989). Figure 4.5: Inferred geological history of the Sudbury structure, highlighting major events related to the formation of QD, IQD and Sublayer, Footwall Breccia, and associated Ni-Cu-PGE mineralization. From Wang et al. (in press), as modified from Lightfoot, 2016). 67 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Excursion Stops, Day 2: Ore Environments Geology of the Whistle and Parkin Area Footwall Rocks The Whistle funnel and Whistle-Parkin offset dike are located at the northeastern lobe of the Main Mass of the Sudbury Igneous Complex in the Norman and Parkin townships (see regional map from Day 1). The footwall comprises Archean rocks of the Superior Province close to the Whistle embayment and Paleoproterozoic greenschist facies to amphibolite facies metamorphosed volcanic and sedimentary rocks of the Huronian Supergroup in the distal Parkin offset dike (Fig. 5.1; Ames et al., 2008; Lightfoot, 2016 and references therein). With ages of 2725 to 2703 Ma (Nunes and Pykes, 1980) felsic to intermediate volcanic rocks as well as feldspar and quartz-feldspar porphyritic rocks of the Benny greenstone belt are the oldest rocks in the area. (Meyn, 1970). Units of the Levack Gneiss Complex are not shown in Fig. 5.1, but occur as tens of meter size bodies close to the Sudbury Igneous Complex and consist of migmatitic tonalite orthogneiss, biotite paragneiss, mafic to felsic gneiss, and gabbros (Meldrum et al., 1997; Murphy and Spray, 2002). Krogh et al. (1984) established an age of 2711 Ma for leucosomes of a tonalitic orthogneiss. Figure 5.1: Simplified geological map of NE Sudbury after Murphy and Spray (2002). Inset: satellite image of the NE Sudbury area showing the field trip stops on the stripped outcrops at the Whistle funnel and Whistle-Parkin offset dike. MM – Milnet Mine, MMFZ – Milnet Mine Fault Zone, NP – Norman Project, NWP – Norman West Project, PCFZ – Post Creek Fault Zone, PM – Podolsky Mine, WM – Whistle Mine. 68 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The 2642 Ma (Meldrum et al., 1997) Cartier batholith is part of the Algoma plutonic domain and consists of medium- to coarse-grained subporphyritic granite. It intruded and at least partially melted parts of the Levack Gneiss Complex (Langford, 1960; Dressler, 1984b; Ames et al., 2008), resulting in local breccia formation (“Levack breccia”) which can resemble Footwall Breccia. Aphyric to plagioclase glomeroporphyritic diabase and gabbros of the 2473 Ma (Heaman, 1997) Matachewan dike swarm crosscut gneissic and granitic units (Meldrum et al., 1997). North of the Milnet Mine fault, Huronian metasediments of the Quirke Lake and Cobalt Group dominate. In addition, diabase and gabbros of the 2.2 Ga (see Lightfoot, 2016 (p. 88) and references therein) Nipissing mafic intrusive suite occur conformable in the Huronian Supergroup. Whistle Funnel and Contact Mineralization The funnel-shaped Whistle embayment is located at the base of the Main Mass and is overlain by continuous layers of micropegmatite, transition quartz gabbro, and felsic norite, and a thin discontinuous layer of mafic norite (Fig. 5.2; Pattison, 1979; Lightfoot et al., 1997a; Lightfoot et al., 1997b). The dip of the contact changes from ~45 degrees south along the northern side of the funnel to ~70 degrees west along the eastern side of the funnel. First published descriptions of surface and drill core data by Pattison (1979) indicated a zonation of orthopyroxene-rich SLNR at the center of the embayment and gradually more siliceous igneous-textured SLNR matrix towards the margins. Here, the Sublayer transitions into FWBX (Fig. 5.2). In a detailed study of SLNR from the open pit, Lightfoot et al. (1997b) identified several lithologies: The central part of the embayment hosts a two pyroxenite SLNR, whereas orthopyroxene-rich SLNR, olivine norite and leucocratic norite occur at the margins or as centimeter-sized to tens of metersized pods within the others. Sublayer matrix is typically non-poikilitic and hosts disseminated to blebby sulfides. Inclusions can range from millimeters to meters and are described as diabase inclusions, anorthositic to gabbroic inclusions and melanorite inclusions or segregations with gradational contacts (Lightfoot et al., 1997b). The Whistle mine was operated by INCO Limited (now Vale) from 1988–1991 and 1994–1997, producing 5.7 Mt of ore grading 1% Ni and 0.3% Cu (Carter et al., 2009). Inclusion-rich massive sulfides (2–3% Ni, >0.2% Cu, <500ppm Pt+Pd) occur at the Sublayer-Footwall breccia contact and show a fractionation to more Cu-rich ore towards the base of the embayment (Lightfoot et al., 1997b). Whistle Dike and Footwall Mineralization After the Whistle Mine closure, acid-generating waste rocks that had been stored to the northeast of the embayment were mechanically and hydraulically stripped during the backfilling of the open pit, allowing subsequent detailed mapping and studies of the proximal Whistle offset dike (Fig. 5.2). The main outcrop was mapped in 2003 by FNX Mining Company (now KGHM), who were exploring in the proximal Whistle dike for Footwall-type mineralization, as well as locally in more detail by Carter et al. (2009) in 2003, and the Podolsky North Zone by Lafrance et al. (2014) in ~2008–2009 during petrographic-geochemical studies of the “metabreccias” and quartz diorite lithologies. The Whistle dike emerges from the funnel and can be traced for 2km to the northeast where it terminates in a breccia zone at the Post Creek fault (Fig. 5.1; Lightfoot et al., 1997b). MTBX is the dominant lithology with meter- to tens of meters-sized inclusion-poor and inclusion-rich quartz diorite pods more rarely occurring close to the dike margins. Magnetic lineation measurements suggest lateral emplacement with locally downward-directed flow for the dike lithologies and subsequent downward-directed sinking of massive sulfides (Giroux and Benn, 2005). 69 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5.2: Simplified geological map and cross-section of the funnel-shaped Whistle embayment and proximal Whistle dike (after Farrow et al., 2005). Two Footwall-style Cu-(Ni)-PGE ore bodies occur within MTBX and IQD in the proximal segment of the Whistle dike: the Podolsky 2000 Deposit, which occurs at depth near the northwestern margin of the dike, and the Podolsky North deposit, which extends to the surface northeast of the main outcrop (Fig. 5.2). Mineralization in the Podolsky 2000 deposit occurs in the form of disseminated/blebby sulfides and “lowsulfide” stockwork veins in the host rocks, as well as breccia sulfide veins and “sharp-walled” massive sulfide veins in a large gabbroic inclusion (Farrow et al., 2005). The latter mineralization appears to cut the others. Chalcopyrite and millerite are the dominant Cu and Ni ore minerals. 2.24 Mt ore at 4.2% Cu and 0.4% Ni were mined between 2008–2013 (Lightfoot, 2016). A stripped outcrop (field trip stop 2-C1; Fig. 5.1) of the distal Whistle segment located just south of the Post Creek fault indicates that metabreccia with disseminated sulfides is still the dominant lithology. Parkin Dike and Offset Dike Mineralization The Parkin segment of the offset dikes appears north of the Post Creek fault, 2km displaced from the Whistle segment, and can be traced for another 12km to the northeast (Fig. 5.1). It is hosted by units of the Archean Benny greenstone belt for the first 4km and by the metasediments and metavolcanics of the Huronian Supergroup past the Milnet Mine fault. Exploration by Wallbridge Mining in 2014–2015 led to the stripping of several large outcrops of the offset dike close to the Post Creek fault. Here, QD and IQD are the common offset dike lithologies and MTBX occurs mainly as pods or inclusions in the quartz diorite lithologies (Anders et al., 2020) or as parallel bodies adjacent to the IQD (Murphy and Spray, 2002). QD is commonly located at the margins of the dike and in sharp contact with the interior IQD. Disseminated sulfides and sulfide stringers are mainly associated with inclusion-bearing lithologies in the dike center. 70 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Rock Descriptions The most common units/lithologies in the field area are listed in Table 5.1 with short descriptions, the most important of lithologies are discussed in more detail below. Table5.1: Rock units and lithologies to be encountered on the Field Trip. Unit/Lithology Abbreviation Description Levack Gneiss Complex MGN, IGN, and FGN Cartier Granite GR Medium- to fine-grained, weakly to strongly foliated, mafic (MGN), intermediate (IGN) and felsic (FGN) gneisses, gabbros and migmatites Medium- to coarse-grained, alkali feldspar-megacrystic granites to granodiorites and equigranular monzogranites Matachewan Intrusives Gabbro GAB Diabase DIA Impact-Related Rocks Sudbury Breccia SUBX Quartz Diorite QD Medium- to fine-grained, green, magnetic gabbro and diorite bodies intruding other footwall rocks Fine- to medium-grained, northwest striking, plagioclase glomeroporphyritic dikes intruding other footwall rocks Polymictic, matrix-supported breccia with aphanitic to very finegrained, black to dark grey groundmass that supports subrounded heterolithic footwall fragments (millimeter to tens of meters) Medium-grained, leucocratic, homogeneous, igneous textured granodioritic matrix with tabular to acicular and sometimes radiating amphibole (after pyroxene) laths; interpreted to represent variably contaminated impact melt Magmatic Impact-Related Breccias Inclusion Quartz Diorite IQD Polymictic, homogeneous, matrix-supported breccia with finegrained, grey to “salt-and-pepper” igneous-textured groundmass; local and exotic (mainly ultramafic) inclusions; often with <2% disseminated sulfides Sublayer Norite SLNR Polymictic, homogeneous, matrix-supported breccia with fine- to medium-grained, subophitic, noritic groundmass; abundant local and exotic (mainly ultramafic) inclusions; often with <5% disseminated sulfides Metamorphic-Anatectic Impact-Related Breccias Footwall Breccia FWBX Polymictic, heterogeneous, matrix-supported breccia with finegrained, pinkish white, granitic groundmass; subrounded inclusions (Granite Breccia) (GRBX) of local footwall rocks and SUBX Metabreccia MTBX Polymictic, heterogeneous, matrix- to clast-supported breccia with fine-grained dark grey to pinkish-grey, recrystallized groundmass; >50% inclusions of local footwall rocks ± pods of QD and IQD Late Dikes Olivine diabase UM East-west trending, up to 7m wide, dark, fine-grained dikes crosscutting all other lithologies in the area 71 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Breccia Classification Pre-impact Archean Breccia (“Levack Breccia”) Intrusion of the 2642 Ma Cartier Batholith brecciated many parts of the Levack Gneiss Complex (Card and Innes, 1981; Dressler, 1984b; Meldrum et al., 1997). Levack Breccia is characterised by a medium- to coarse-grained, locally alkali feldspar-megacrystic, granodioritic to monzogranitic matrix with angular to rounded, cm- to tens of meter-sized mafic, gneissic, and migmatitic inclusions (Fig. 5.3A). The distribution and abundance of Levack Breccia is not well known, but there are several areas on the Whistle outcrop where this breccia is present. Sudbury Breccia Pseudotachylitic SUBX appears to be the earliest formed impact breccia as it is crosscut by all other impact-related lithologies. Shock compression and cataclasis during crater formation led to shattering, pulverization, and frictional melting of footwall rocks, forming centimeter- to meter-sized veins and more irregular bodies of intense brecciation, often at structurally weakened zones and at lithology boundaries (e.g., Speers, 1957; Dressler, 1984a; Rousell et al., 2003; Lafrance et al., 2008; Lafrance and Kamber, 2010). It is characterized by a black to dark grey, aphanitic to very fine-grained matrix with (sub)-rounded centimeter- to meter-sized clast of immediate country rock lithologies (Fig. 5.3B). Sudbury breccia occurs frequently in the vicinity of the SIC but has been described up to distances of 50 km or more (Dressler, 1984a). Close to the SIC, contact metamorphism led to a grain size increase (often accompanied by lighter matrix colours), growth of biotite porphyroblasts, and partial melting of felsic mineral clasts. Inclusion-Bearing Quartz Diorite IQD is a polymictic breccia constraint to the offset dikes. It typically consists of a homogeneous, equigranular, fine- to medium-grained groundmass with a “salt-and-pepper” appearance (Fig. 5.3C). Inclusion sizes and abundancies vary from millimeter to tens of meters and <10% up to 90%, respectively. The matrix typically consists of tabular plagioclase with lesser alkali feldspar, quartz, and amphiboles. Radial amphiboles, which are typical for QD, occur less frequent in IQD. Sublayer Norite SLNR forms a discontinuous layer at the base of the Sudbury Igneous Complex normally occurring within funnels, troughs, and embayments. It is a variably mineralized polymictic breccia with centimeterto meter-size inclusions of Ol melanorite anteliths, local xenoliths, and exotic ultramafic xenoliths set in a fine- to medium noritic matrix (Fig. 5.3D; e.g., Pattison, 1979; Naldrett et al., 1984; Wang et al., 2018, 2020). Footwall Breccia FWBX (also termed “leucocratic breccia” and “late granite breccia”) is a polymictic breccia containing inclusions of local footwall rocks, SUBX, Main Mass norite, and exotic ultramafic inclusions (e.g., Pattison, 1979; Coats and Snajdr, 1984; Dressler, 1984a; Lakomy, 1990; McCormick et al., 2002; Wang et al., 2020). Grain sizes, compositions, and textures of the matrix are highly variable and dependent on proximity to the SIC contact and mineralization: the matrix is coarser-grained, igneous-textured and of dioritic to tonalitic composition closer to the SIC (Fig. 5.3E) and finer-grained, metamorphic-textured (recrystallized) and of granitic composition further into the footwall (Lakomy, 1990; McCormick et al., 2002). In proximity to mineralization, FWBX is often characterized by a grey colour (Greenman, 1970). 72 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Several modes of formation have been proposed, including a) contact metamorphism of impactbrecciated footwall during cooling of impact melt (Dressler, 1984a), b) partial melting and anatexis of felsic layers in the Levack Gneiss Complex due to heat conduction of cooling impact melt (Coats and Snajdr, 1984), and c) contact metamorphism during injection of Sublayer norite (Pattison, 1979). Metabreccia The term “metabreccia” was first used by INCO mining geologists in the 1970’s to describe thermal metamorphism of SUBX in proximity to the superheated melt sheet (E.F. Pattison, 2019, pers. comm.), which has been adopted by some of the other mining companies (e.g., Poulin et al., 2009). Other workers have interpreted MTBX to represent variable amounts of impact melt (QD/IQD) and partially melted footwall rocks (FWBX) mobilized into the dike structure (Murphy and Spray, 2002; Giroux and Benn, 2005; Lafrance et al., 2014; Carter et al., 2009; Anders et al., 2020). Farrow et al. (2005) were the first to introduce “metabreccia” as a term in the published literature, but different terminology exists by further authors, such as diatexite, metatexite (Lightfoot, 2016), radial breccia, mafic sulfide-bearing breccia (both Murphy and Spray, 2002), metamorphic leucocratic breccia/Footwall Breccia (Carter, 2005; Carter et al., 2009), and recrystallised Footwall Breccia (Grant and Bite, 1984). Research on metabreccia (or its synonyms) is restricted to the Whistle funnel and WhistleParkin offset dike, Ministic, Foy and Trill offset dikes – all located in the North Range. Metabreccia is a heterogeneous, grey pinkish, polymictic breccia with abundant millimeter- to centimeter-sized and lesser meter-sized footwall inclusions (Fig. 5.3F). Textures and mineralogy vary from igneous-textured with similar modal abundancies as QD and IQD (Lafrance et al., 2014) to dynamically recrystallized with a fine-grained quartz-feldspar-rich matrix (Anders et al., 2020). MTBX is the dominant lithology in the Whistle dike, whereas it occurs as pods in quartz diorite in Parkin. 73 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5.3. A) Levack breccia near Whistle funnel. B) Metamorphosed SUBX near Whistle funnel (field trip stop 2-A2. C) IQD in Whistle funnel (field trip stop 2-A4). D) SLNR in Whistle funnel (field trip stop 2-A3). E) FWBX near Coleman Mine. F) MTBX in Whistle offset dike. 74 �Proceedings of the 68th ILSG Annual Meeting - Part 2 A. Whistle Embayment and Proximal Offset Dike The stripped outcrops of the Whistle funnel provide excellent exposure and a wealth of interesting field relationships and textures. Key aspects are highlighted on the satellite image (Fig. 5.4) and on the geological map (Fig. 5.5), and are described in more detail below. Figure 5.4: Satellite image of the filled and reclaimed Whistle open pit (prior to re-greening stage), showing field trip stops on the associated stripped outcrops. The Podolsky Mine is in the upper right part of the image. 75 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5.5: Map of the Whistle outcrop (simplified from FNX Mining Company) showing field trip stops. Dashed lines delineate the footwall-funnel/dike contact. 76 �Proceedings of the 68th ILSG Annual Meeting - Part 2 2-A1) SLNR-GRBX contact (NAD83 17N, 509053, 5179920) The first stop is located at the western flank of the embayment near the reclaimed open pit. SLNR is in contact with FWBX over a few centimeters. SLNR has a fine- to medium-grained, dark grey to black matrix with minor rugged sulfide blebs and some subrounded inclusions (Fig. 5.6A–C). FWBX has a pinkish grey, fine- to medium-grained, heterogenous matrix with abundant strongly intergrown quartz-feldspar and variable amounts of interstitial green amphibole. Inclusions are typically subangular and of mafic to intermediate composition (Fig. 5.6D–F). Figure 5.6: High-resolution sample scans, plane-polarized and cross-polarized images of SLNR (A-C) and FWBX (D-F) from the eastern funnel flank. 77 �Proceedings of the 68th ILSG Annual Meeting - Part 2 2-A2) SUBX and footwall rocks (NAD83 17N, 509289, 5180141) This 170m x 120m stripped outcrop is located in the footwall just west of the former open pit. Meter- to tens of meters-scale blocks of various country rocks are enveloped and truncated by cm- to dm-wide pseudotachylitic SUBX veins and bodies (Fig. 5.7A). Walking from the NW end of the outcrop towards the embayment (ESE direction), a slight gradational change of color, grain size and clast angularity are noticeable in the SUBX veins and might reflect a baking of the breccia closer to the former impact melt. The matrix changes from a black/dark grey to a medium grey, coarsens slightly, biotite porphyroblasts are observable and small felsic clasts are getting wispier closer to the former impact melt. Small SUBX bodies are intruded by a pinkish medium- to coarse-grained feldspar-rich granitoid in some areas, which might indicate the formation of localized feldspar-rich FWBX (Fig. 5.7B). Figure 5.7: A) Clast of a pre-impact Levack Breccia consisting of sub-angular to -rounded dm-sized gabbro in a granitic matrix, surrounded by SUBX. B) SUDBX fragment (?) wrapped around subrounded, 30cm large gabbro inclusion itself intruded by feldspar-rich granitoid (FWBX?). 2-A3) Sublayer norite (NAD83 17N, 509540, 5180011) SLNR is only exposed at the SW end of the 300m x 300m main outcrop. It typically displays a rustybrown weathering surface sulfide oxidation (Fig. 5.3). The noritic matrix is typically homogeneous, fineto medium-grained, and displays equigranular textures. It contains abundant country rock inclusions, minor dark green/brown ultramafic inclusions, and QD-type inclusions. Globular and ragged sulfide blebs up to 2cm in size are common. 2-A4) IQD-MTBX (NAD83 17N, 509563, 5180046) To the NE, the SLNR transitions into FWBX/MTBX in the central part of the main outcrop, indicating the beginning transition from an embayment environment to an offset environment. At the margins of the FWBX/MTBX, bodies or pods of leucocratic QD and IQD occur (Fig. 5.5), often in sharp contact with FWBX/MTBX as shown here. The IQD is characterized by a homogeneous fine-to medium-grained matrix with abundant amphiboles needles and less than 20% (sub)-rounded inclusions. FWBX/MTBX is appearing more heterogenous and rich in inclusions. 78 �Proceedings of the 68th ILSG Annual Meeting - Part 2 2-A5) SLNR-IQD-MTBX-QD (NAD83 17N, 509600, 5179940) This area on the southeastern flank of the embayment, which was mapped in detail by Carter et al. (2009), features excellent exposure of contact relationships between leucocratic QD, MTBX/FWBX, IQD, and SLNR. It also highlights the complexity and variability in breccia matrix composition in a small area (Fig. 5.8 and Fig. 5.9). Footwall gabbros and granitoids are crosscut by decimeter wide SUBX veins (e.g., location 4 in Fig. 5.9). Abundant subrounded inclusions from local footwall rocks are set in an aphanitic to very fine-grained groundmass which is, independent of the host unit, composed of predominantly chlorite and epidote. Leucocratic QD occurs as a large body in contact with the footwall (location 1 in Fig. 5.9). It is mediumgrained with abundant amphibole and plagioclase laths as well as oikocrystic quartz, feldspar, and granophyric quartz-feldspar intergrowth (Fig. 5.8B–C). Figure 5.8: A) Sharp contact between LQD (left) and polymictic MTBX/FWBX (right). See Fig. 5.9 for location. Plane- (B) and cross-polarized (C) images of the leucocratic QD matrix. 79 �Proceedings of the 68th ILSG Annual Meeting - Part 2 MTBX/FWBX (depending on the map) is fine-grained, equigranular, polymictic, and in sharp contact with leucocratic QD at location 1. The polymictic breccia is typically inclusion-rich (>50%) with inclusion sizes ranging from millimeters to several meters. Feldspar with lesser quartz and interstitial amphibole are the dominant minerals in the dark grey, fine-grained breccia groundmass. In contrast, MTBX/FWBX at location 3 (Fig. 5.9), 5m away from location 1, displays typical grey-pinkish colors, is inclusion-rich and appears recrystallized on fresh surfaces. The poikilitic groundmass consists of highly altered equant plagioclase and tabular amphibole chadacrysts in coarser, oikicrystic quartz-(feldspar). IQD at location 2 (Fig. 5.9) is characterized by a homogeneous matrix with a typical “salt-and-pepper” color on fresh surfaces. It is polymictic with <25% inclusions and minor sulfide blebs. The matrix is fineto medium-grained and poikilitic. Variably altered, equant plagioclase and rare amphibole needles are set in oikocrystic quartz with minor feldspar and granophyric intergrowth. Macroscopically, the different breccia units appear to be quite different and easily distinguishable. SUBX adjacent to the funnel/offset dike displays distinct grain size differences, is strongly altered and of a more mafic matrix composition. Groundmass compositions of the other breccia units are remarkably similar (see also Lafrance et al., 2014). 80 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5.9: Detailed map of the southeastern flank with SUBX, SLNR, IQD, MTBX/FWBX, leucocratic QD, and footwall rocks. Locations of breccia samples are shown in the map. The leucocratic QDMTBX/FWBX contact is shown in Fig. 5.8. Images of slabs and thin sections are described in the text. 81 �Proceedings of the 68th ILSG Annual Meeting - Part 2 82 �Proceedings of the 68th ILSG Annual Meeting - Part 2 2-A6-1 GRBX-Cartier Batholith contact (NAD83 17N, 509538, 5179867) Located at the SE flank of the embayment just meters away from the covered and re-greened open pit. FWBX is in a well-defined contact over 1–2cm with Cartier granite. 2-A6-2 IQD with QD Inclusions (NAD83 17N, 509552, 5179875) Sharp contact between QD and IQD (SLNR?). QD is leucocratic, medium-grained and displays well developed amphibole needles <1cm. IQD is characterized by a finer-grained, homogeneous matrix with salt-and-pepper texture, 10–40% inclusions and disseminated blebby sulfides. Several subrounded QD pods or inclusions are visible and highlighted by their more leucocratic appearance (Fig. 5.10A). 2-A7 Leucocratic QD with local footwall clasts (NAD83 17N, 509693, 5180036) Coarse-grained, leucocratic QD with up to 2cm long amphibole needles and angular, greenish-altered gabbro inclusions (Fig. 5.10B). Located just 1m away from the SE dike margin right at the contact between Cartier batholith and gabbro, this area indicates only a short lateral transport of the gabbro inclusions. Country rock inclusions at QD margins is also documented in other offset dikes, such as Foy, Hess, and Worthington. Figure 5:10: A) IQD with three (leucocratic) QD pods/inclusions in sharp contact with (leucocratic) QD at 2-A6-2. B) QD with large amphibole needles and local clasts at 2-A7. Image width is approximately 75 cm. 2-A8 IQD-MTBX (NAD83 17N, 509799, 5180205) This area has been mapped in detail by Carter et al. (2009) and shows some excellent IQDMTBX/FWBX features on glacially polished surfaces. The IQD-MTBX/FWBX contact can be traced over several meters and is generally sharp. MTBX/FWBX is dark grey on fresh surfaces, inclusion-rich (ca. 50%), and closely associated with large (>5m) Cartier granite clasts. Small (<3cm), rounded, red feldspar inclusions are common with mafic inclusions typically being smaller and less common. IQD is leucocratic, medium-grained with well-developed amphibole needles and typically less than 10% inclusions. 83 �Proceedings of the 68th ILSG Annual Meeting - Part 2 B. Intermediate Whistle Dike 2-B1 Podolsky IQD-MTBX-Sulfides (NAD83 17N, 509937, 5180490) The Podolsky outcrop located 500m NE of the Whistle embayment (Fig. 5.4), features the surface exposure of sharp-walled Cu-PGE-rich sulfide veins from the Podolsky North Zone. Steeply dipping chalcopyrite-rich veins are locally up to 3m wide but range more commonly on a cm- to dm-scale. They follow lithological boundaries as well as crosscuts IQD and MTBX (Fig. 5.11B–C). Flow laminations can be observed in the veins (Lafrance et al., 2014). IQD predominates in the southern part of the outcrop and MTBX in the northern part. QD pods or inclusions (typically <3m) are more abundant in contact with IQD than with MTBX and display several centimeter long radiating amphibole needles (after pyroxene). Contacts between QD and the other units can be sharp or gradational over a few centimeters (Fig. 5.11A). IQD and MTBX have similar clast compositions, but the latter is generally more enriched. Detailed descriptions of the lithologies can be found in Lafrance et al., (2014). 84 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5.11: Geological map of the Podolsky outcrop (simplified after Carter et al., 2009). A) Sharp contact between MTBX and QD. B) Chalcopyrite-rich sample with inclusions. C) Drone image of the Cu-vein with gneissic clast. 85 �Proceedings of the 68th ILSG Annual Meeting - Part 2 C. Distal Whistle Dike 2-C1 (NAD83 17N, 510293, 5181352) This stripped outcrop 1.5km northeast of the Whistle embayment (Fig. 5.1) belongs to claims of North American Nickel Inc. and displays the northern most segment of the Whistle dike before it is offset by the Post Creek fault. The eastern part of the outcrop consists of Cartier granite. MTBX is the dominant lithology in the central part. It is very fine-grained, grey pinkish with abundant small feldspar inclusions and <2% disseminated sulfides (Fig. 5.12A–B). Meter-sized inclusions of subangular to subrounded diabase, granite and QD inclusions are common (Fig. 5.12C). Figure 5.12: A) close-up image of MTBX matrix. B) polished slab of MTBX with abundant feldspar clasts. C) QD clast in contact with MTBX at the distal Whistle outcrop 2-C1. D) Well defined contact between inclusion-poor and inclusion-rich quartz diorite at the proximal Parkin outcrop 2-D1. Hammer head for scale 86 �Proceedings of the 68th ILSG Annual Meeting - Part 2 D. Proximal Parkin Dike Wallbridge Mining´s proximal Parkin properties are located just north of the Post Creek Fault (Fig. 5.1). Mechanical stripping, mapping, and drilling of 63 drill holes as well as geophysical surveys were conducted in 2015 and 2016. Outcrops maps are shown in Fig. 5.13 and Fig. 5.14. 2-D1 Northern Part (NAD83 17N, 509199, 5183212) The northern stripped outcrops show large sections of the central portion of the Parkin offset dike. IQD is the dominant lithology and incorporates millimeter to tens of meter sized country rock inclusions as well as disseminated and stringer sulfides. QD can locally be observed at the margins of the dike and the outcrops. There are several areas where a sharp contact between QD and IQD can be observed (Fig. 5.12D). MTBX can be observed as inclusions and irregular pods in QD and IQD, often with well-defined contacts over <1mm and more rarely, with gradational contacts as indicated by changes in inclusion abundance (see Anders et al., 2020). 2-D2 Southern Part (NAD83 17N, 509104, 5183025) The southern stripped outcrop displays similar features as the northern outcrops. In addition, two prominent large (˃10m) felsic gneiss inclusions occur in the central portion of the dike. These potential Levack Gneiss inclusions are not from the immediate country rocks, thus indicating the high energy of the quartz dioritic melt during injection. Similar observations were made by Murphy and Spray (2002) 1km to the northeast of this outcrop. 87 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5:13: Map of Wallbridge Mining´s proximal Parkin stripped outcrops (northern part, from Wallbridge Mining Assessment Report on the Parkin Property 2015-16). 88 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5:14: Map of Wallbridge Mining´s proximal Parkin stripped outcrops (northern part, from Wallbridge Mining Assessment Report on the Parkin Property 2015-16). 89 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Alternative Field Trip Stops: Worthington, Trill East, and Hess Ermatinger Worthington The Worthington offset dike is located at the SW lobe of the SIC and is connected to the Main Mass of the SIC through the Victoria embayment (Fig. 2.1). The dike crosscuts and incorporates inclusions of metasediments of the Huronian Supergroup as well as Nipissing diabase. Several historical mines targeted small ore bodies close to the surface over the last century. Vale´s Totten mine, in operation since 2014, is the latest addition to Sudbury operations and the only active mine at the Worthington dike. Nevertheless, there are several exploration projects in the area. KGHM´s Victoria project is located close to the embayment and SPC Nickel is conducting an exploration and drilling program on their AER-KIDD property (Fig. 6.1), located between Victoria to the northeast and Totten to the southwest. Figure 6.1: Satellite image with outcrop locations of SPC Nickel’s AER-KIDD exploration project. Outcrop 1 is located in the southwest, outcrop 2 in the center and outcrops 3 and 4 in the northeast of the map. 90 �Proceedings of the 68th ILSG Annual Meeting - Part 2 AER-KIDD Outcrop 1 Outcrop 1 displays the eastern contact between the Worthington offset dike and Footwall metasediments of the McKim Formation, which are in sharp contact with very fine-grained to spherulitic QD (Fig. 6.2A– B). QD matrix grain size gradually increases over 2–3m towards the center where it is in a well-defined contact with finer grained IQD (Fig. 6.2A–B). The contact is characterized by a) the grain size difference, b) inclusions, and c) disseminated sulfides. Centimeter- to decimeter-sized inclusions are generally aligned parallel to the strike of the dike and are comprised of footwall rock assemblages, i.e., gabbros and amphibolites presumably of the Nipissing mafic suite and McKim Formation metasediments. Subrounded, meter-sized inclusions of QD in IQD are common. AER-KIDD Outcrop 2 Outcrop 2 is adjacent to the former small Robinson mine and displays similar field relationships as outcrop 1, although less well defined due to abundant sulfide oxidation. In addition to QD and IQD, the most central part of the dike (and most westerly part of the outcrop before the fenced-off mine) is comprised of an amphibolite inclusion-rich quartz diorite (AIQD) which can grade into amphibolite-bearing sulfide matrix breccia. Amphibolite inclusions are typically larger in AIQD than in IQD. AER-KIDD Outcrop 3 and 4 Outcrops 3 and 4 are located 300m northeast of outcrop 2 and display similar features as outcrop 1 (Fig. 6.3). QD-IQD contacts are generally well defined over a few millimeters and characterized by grain size differences as well as inclusion and sulfide abundances (Fig. 6.3A). In some areas, rounded inclusions of McKim Formation are incorporated into QD close to the footwall contact. Also, “mushroom-like” bulging of McKim Formation into QD can be observed (Fig. 6.3B). 91 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 6.2: Geological map, measurements, and field images of the AER-KIDD outcrop 1. Most country rock clast in the IQD are rotated parallel to the IQD-QD and QD-Footwall contact. A-B) QD-IQD contact is well-defined and characterized by grain size differences as well as inclusion and sulfide content. 92 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 6.3: Geological map and field images of AER-KIDD outcrops 3 and 4. A) Sharp QD-IQD contact defined by grain size differences, inclusions and disseminated blebby sulfides. B) McKim Fm. “intruding” QD. 93 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 6.4: Geological map of Trill East outcrop (simplified after Wallbridge Mining Assessment Report on the Trill Property 2014). A) SUBX vein crosscut by QD. MTBX might be present at the contact. B) Close-up of MTBX inclusion/pod in QD. 94 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Trill East The 3–20m wide, radial Trill offset dike is located southwest of the Sudbury Igneous Complex, just north of the boundary between Superior province and Southern province (Fig. 2.1). Exploration by Wallbridge Mining in 2005 led to the discovery of a 65m x 5m long sulfide mineralization 4km west of the SIC. Mechanical stripping of several outcrops revealed several offset dike units: QD, IQD, spherulitic QD and glassy QD. Common QD is typically found at the dike margin and IQD in the center. Glassy and spherulitic QD are restricted to thin apophyses in the footwall (Klimesch, 2009). The Trill East outcrop, ca. 500m west of the SIC, features the northern dike-footwall contact (Fig. 6.4). SUBX crosscuts the Cartier granite in fine veinlets as well as in a thicker zone at the eastern end of the outcrop and is itself crosscut by the offset dike (Fig. 6.4A). Dark grey QD is the only offset dike unit present. Several inclusions or pods of MTBX occur within QD (Fig. 6.4B) or between the QD-footwall contact. MTBX is fine-grained, has a light grey appearance and is enriched in small (mm–cm) plagioclase-quartz, alkali feldspar mineral, and mafic clasts (Anders, 2016). Hess Ermatinger The concentric Hess offset dike follows the northern outline of the SIC in ca. 15km distance. It has been traced from E of the Foy-Hess intersection to the Ermatinger township WNW of Sudbury (Fig. 2.1). Exploration efforts by Wallbridge Mining in 2010–2012 led to the mechanical stripping and drilling of several outcrops with one in particular displaying excellent exposures of offset dike-footwall interactions, QD and IQD. The footwall contact of the steeply SE dipping offset dike can be traced for 40m and is characterized by a decimeter-wide zone of QD chilled against Cartier granite (Fig. 6.5). The contact is often sharp but can be irregular in zones of QD bulging into the footwall (Fig. 6.5C). Granite inclusions are abundant in the chilled QD, vary in size from cm-dm and are sometimes partially digested (Fig. 6.5D). Several smaller and larger apophysis are cutting through the footwall granite and diabase dikes. QD grain sizes gradationally increase from chilled to medium-grained over 5m. The contact between medium-grained QD and fine-grained IQD is predominantly defined by the abrupt change in grain size (Fig. 6.5A–B). IQD is characterized by a fine-grained, sulfide-poor (<2%), inclusion-poor (<10%) IQD. In comparison to the local inclusions in the chilled QD, inclusions in IQD are of gabbroic, dioritic and ultramafic composition not directly associated with the local footwall. Disseminated blebby sulfides occur in patches and are composed of a typical magmatic sulfide assemblage. Acknowledgements The authors for Day 2 would like to thank the mining companies for access to their properties: Whistle mine (Vale); Podolsky (KGHM); Parkin, Trill and Hess (Wallbridge Mining); Worthington (SPC Nickel); distal Whistle (North American Nickel). We are grateful to Wouter Bleeker for organizing both field days and to Mike Easton for editorial handling. Safety gear was kindly provided by the Harquail School of Earth Sciences at Laurentian University. 95 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 6.5: Geological map of the Hess Ermatinger outcrop. A) Northern contact (dashed lines) between medium-grained QD and fine-grained IQD. Rusty spots in IQD indicate presence of sulfide blebs. B) Southern QD-IQD contact with fresh and weathered surfaces. Gabbro clast in IQD is not from the local footwall. C) Chilled QD intrudes into Cartier granite. D) Dm-sized granite clast showing partial digestion. Diabase dikes are crosscut by QD. 96 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Literature related to Day 2 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.383-386. Ames, D.E., Davidson, A. and Wodicka, N. 2008. Geology of the giant Sudbury polymetallic mining camp, Ontario, Canada: Economic Geology, v.103, p.1057-1077. Anders, D. 2016. The Sudbury Impact Structure – new insights into the origin and emplacement of the Basal Onaping Intrusion and the Parkin, Trill and Foy offset dykes of the North Range: Unpublished PhD thesis, The University of Western Ontario, 183p. Dressler, B.O. 1984a. 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Poulin, R., Dunlop, S. and Everest, J.O. 2009. Podolsky field guide - Podolsky mine property and Whistle pit, Norman Twp., Sudbury mining district, Ontario, FNX Mining Company Ltd., 8p. Naldrett, A.J., Hoffman, E.L., Green, A.H., Chou, C.L., Naldrett, S.R. and Alcock, R.A. 1979, The composition of Ni-sulfide ores, with particular reference to their content of PGE and Au: The Canadian Mineralogist, v. 17, p. 403-415. Prevec, S.A. and Büttner, S.H. 2018. Multiphase emplacement of impact melt sheet into the footwall: offset dykes of the Sudbury Igneous Complex, Canada: Meteoritics & Planetary Science, v.53, p.1301-1322. Nelles, E.W. 2012. Genesis of Cu-PGE-rich Footwalltype mineralization in the Morrison Deposit, Sudbury: Unpublished MSc thesis, Laurentian University, 87p. Reimold, W.U. and Gibson, R.L. 2006. The melt rocks of the Vredefort impact structure–Vredefort Granophyre and pseudotachylitic breccias: implications for impact cratering and the evolution of the Witwatersrand Basin: Geochemistry, v.6, p.135. Nunes, P. and Pyke, D. 1980. Geochronology of the Archean metavolcanic belt, Timmins-Matachewan area—Progress report; in Summary of Field Work and Other activities, Ontario Geological Survey, Miscellaneous Paper 92, p. 34-39. Riller, U. 2005. Structural characteristics of the Sudbury Impact Structure, Canada: impact-induced versus orogenic deformation—a review: Meteoritics & Planetary Science, v.40, p.1723-1740. Osinski, G.R., Grieve, R.A.F., Bleacher, J.E., Neish, C.D., Pilles, E.A. and Tornabene, L.L. 2018. Igneous rocks formed by hypervelocity impact: Journal of Volcanology and Geothermal Research, v.353, p.2554. Robertson, J., Ripley, E.M., Barnes, S.J. and Li, C. 2015. Sulfur liberation from country rocks and 100 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stout, A.E. 2009. Geology, mineralogy, and geochemistry of the McCreedy East 153 Cu-Ni-PGE Deposit, Sudbury, Ontario: Unpublished MSc thesis, Utrecht University, 39p. incorporation in mafic magmas: Economic Geology, v.110, p.1111-1123. Robertson, J.C., Barnes, S.J. and Le Vaillant, M. 2016. Dynamics of magmatic sulphide droplets during transport in silicate melts and implications for magmatic sulphide ore formation: Journal of Petrology, v.56, p.2445-2472. Therriault, A.M., Fowler, A.D. and Grieve, R.A.F. 2002. The Sudbury Igneous Complex: a differentiated impact melt sheet: Economic Geology, v.97, p.15211540. Rousell, D.H., Fedorowich, J.S. and Dressler, B.O. 2003, Sudbury Breccia (Canada): a product of the 1850 Ma Sudbury event and host to footwall Cu–Ni– PGE deposits: Earth-Science Reviews, v.60, p.147174. Tuchscherer, M.G. and Spray, J.G. 2002. Geology, mineralization, and emplacement of the Foy offset dike, Sudbury Impact Structure: Economic Geology, v.97, p.1377-1397. Wang, Y., Lesher, C.M., Lightfoot, P.C., Pattison, E.F. and Golightly, J.P. 2018. Shock metamorphic features in mafic and ultramafic inclusions in the Sudbury Igneous Complex: Implications for their origin and impact excavation: Geology, v.46, p. 43446. Scott, R.G. and Benn, K. 2002. Emplacement of sulfide deposits in the Copper Cliff offset dike during collapse of the Sudbury crater rim: evidence from magnetic fabric studies: Economic Geology, v.97, p.1447-1458. Souch, B E., Podolsky, T. and Wilson, H.D.B. 1969. The sulfide ores of Sudbury: Their particular relationship to a distinctive inclusion-bearing facies of the Nickel Irruptive, Magmatic Ore Deposits, 4, Society of Economic Geologists. Wang, Y., Lesher, C.M., Lightfoot, P.C., Pattison, E.F. and Golightly, J.P. 2020. Geochemistry and petrogenesis of mafic and ultramafic inclusions in Sublayer and Offset Dikes, Sudbury Igneous Complex, Canada: Journal of Petrology, v.61. Speers, E.C. 1957. The age relation and origin of common Sudbury Breccia: The Journal of Geology, v.65, p.497-514. Wang, Y., Lesher, C.M., Lightfoot, P.C., Pattison, E.F. and Golightly, J.P. (in press). Genesis of Sublayer, Footwall Breccia, and associated Ni-Cu-PGE Mineralization in the Sudbury Igneous Complex: Economic Geology. Spray, J.G., Butler, H.R. and Thompson, L.M. 2004. Tectonic influences on the morphometry of the Sudbury Impact Structure: implications for terrestrial cratering and modeling: Meteoritics & Planetary Science, v.39, p.287-301. Wichman, R.W. and Schultz, P.H. 1993. Floor-fractured crater models of the Sudbury Structure, Canada: implications for initial crater size and crater modification: Meteoritics, v.28, p.222-231. Sproule, R.A., Sutcliffe, R., Tracanelli, H. and Lesher, C. M. 2007. Palaeoproterozoic Ni–Cu–PGE mineralisation in the Shakespeare intrusion, Ontario, Canada: a new style of Nipissing gabbro-hosted mineralisation: Applied Earth Science, v.116, p.188200. Yao, Z.-s. and Mungall, J.E. 2021. Kinetic controls on the sulfide mineralization of komatiite-associated Ni-Cu-(PGE) deposits: Geochimica et Cosmochimica Acta, v. 05, p.185-211. 101 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Location of stops for ILSG Field Trip 2. Trip starts and leaves from Science North (upper left). Stops 1 to 5 are accessed along Highway 17 and the Highway 17 bypass. Stops 6 is on Highway 537. Stops 7 to 14 are accessed from Estaire Road (formerly Highway 69). 102 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Field Trip 2 – Geology of the Grenville Front and the Grenville Front Tectonic Zone in the Sudbury area R.M. Easton Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, Ontario P3E 6B5 Introduction For metamorphic rocks, mineral prefixes are listed in order of relative abundance, starting with least abundant first. Mineral abbreviations follow Whitney and Evans (2010). The following conventions are used regarding descriptive adjectives. A gneissic granite is a meta-igneous rock of granitic composition. A granitic gneiss, a granite gneiss, or a gneiss of granitic composition may be either a meta-igneous or a metasedimentary rock. Similarly, a tonalitic gneiss or a tonalite gneiss is a gneiss of tonalite modal composition but may be of either meta-igneous or metasedimentary origin. A gneissic meta-arkose is a metasedimentary gneiss of overall granitic composition. The term metamorphic grade is used where bulk-rock composition or other factors prevent a more detailed assignment of metamorphic conditions. Where metamorphic conditions can be outlined, metamorphic facies terminology is used. The field trip uses road accessible outcrops. All of the road stops can be accessed using a 2-wheel drive vehicle. Unless otherwise stated, all UTM coordinates are in Zone 17, datum NAD 83. Safety Many of the field trip stops are located on highways that are especially busy during the summer season. Care should always be exercised when parking, exiting vehicles, and crossing the roads. Use of safety vests and/or bright clothing is recommended, in order to improve your visibility to motorists. Most of the trip routes are on Crown land or public roadways, but access is on or near private property in some cases. As in all such situations, please respect the property rights of others, so as to maintain good relationships, so that future access for geologists is not adversely affected. Many rocks in the Grenville Province were subjected to extreme ductile deformation and subsequently recrystallized, and can be described either as tectonites or gneissic mylonites. Several field-based terms have been proposed to describe these gneissic mylonites including the terms straight gneiss, block gneiss, and porphyroclastic gneiss (e.g., Davidson et al. 1982; Hanmer and Ciesielski 1984). Terminology A number of terms used in this report are outlined below. Rock Classification Layering thickness terms used in this report are listed below. These terms apply to bedded, layered and gneissic rocks. Very thinly layered Thinly layered Medium layered Thickly layered Very thickly layered Extremely thickly layered A migmatite is a heterogeneous rock composed of two or more components, one generally quartzofeldspathic in composition (leucosome or neosome) and the other more mafic in composition (paleosome or mesosome). Within the field trip area, such rocks are commonly layered, and in many instances, are formed by partial melting during high-grade regional metamorphism. <3 cm 3 to 10 cm 10 to 30 cm 30 to 100 cm 1 to 3 m >3 m Terminology for plutonic rocks follows that of Streckeisen (1976) and LeMaitre et al. (2002). 103 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Geological Setting Descriptive terminology for these rocks follows Sawyer (2008) and Mehnert (1971). Migmatites collectively display a wide variety of features depending on the degree of partial melting and deformation during development. The first-order division of migmatites, based on morphology and proportion of leucosome, results in 2 types: metatexite and diatexite. The division between the 2 is based on the relative amount of melt (leucosome) in the rock. The Ontario Geological Survey uses a boundary of 20% leucosome between metatextite and diatextite, which is near the minimum value suggested by Sawyer (2008) but does not require the same precision in estimating leucosome content as the use of 16% would require. The 20% boundary also accounts for the fact that initial bulk-rock composition of the protolith is a factor in the amount of partial melt that can be produced, and thus is better suited for a wide range of bulk-rock compositions. Proterozoic Rocks in the Sudbury area Proterozoic rocks in the Sudbury area are assigned to either the Paleoproterozoic Southern Province or the Mesoproterozoic Grenville Province (cf. Wynne-Edwards 1972; Easton 1992). The Southern Province in Ontario comprises Paleoproterozoic metasedimentary and metavolcanic rocks of the Huronian Supergroup and gabbroic intrusions of the Nipissing gabbro suite. Also included in the Southern Province are the Sudbury Igneous Complex and the Whitewater Group; plutonic and minor volcanic rocks of the Killarney Magmatic Belt; and rocks of the Sudbury diabase dike swarm (Figure 1) (Bennett, Dressler and Robertson 1991). The Huronian Supergroup (Figure 2) was deposited unconformably on Archean plutonic and supracrustal rocks of the Superior Province. The lowest unit, the Elliot Lake Group, consists of both metavolcanic and metasedimentary rocks (Figure 2). In the Sudbury area, the metavolcanic units include tholeiitic basalts of the Elsie Mountain Formation, evolved tholeiitic basalts, dacites, and metasedimentary rocks of the Stobie Formation, and dacites and rhyolites of the Copper Cliff Formation (circa 2460 Ma). The latter are likely coeval with the Murray and Creighton granites (Bennett, Dressler and Robertson 1991; Bleeker et al. 2015). The metavolcanic units interfinger with, and are overlain, by the Matinenda Formation in the west and the McKim Formation in the east. Geochemical data reported by Innes (1972, 1977), Easton (1998), Gordon (2021) indicate a tholeiitic affinity for the Stobie Formation, whereas felsic metavolcanic rocks of the Copper Cliff Formation and the Murray and Creighton granites show calcalkalic signatures. Purpose The purpose of the trip is two-fold. First, is to examine the nature of the Grenville Front and the associated Grenville Front tectonic zone using a variety of exposures in the Sudbury area. Second, is to examine outcrops mapped in the fall of 2021 between Wanup and Estaire. This new mapping indicates that Nepewassi domain rocks occur closer (by approximately 5-10 km) to the Grenville Front then previously recognized. Also, although not specifically part of the field trip route, the new field data also suggests the presence of a major north-northwest-trending structure along the Wanapitei River near Wanup, with the character of the rocks in the Grenville Front tectonic zone being very different in lithology and structural character on either side of the structure (see section on “The GFTZ zone near Wanup”). 104 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 1. Geology of the northern Central Gneiss Belt of the Grenville Province and the Grenville Front region in Ontario. Locations of mapping areas described by Easton (2014) and Van de Kerckhove (2014) are indicated by dashed-line boxes. Abbreviations: C, Cosby pluton; SF, Sturgeon Falls batholith; WB, West Bay batholith; and WC, Wanapitei complex. Figure modified from Easton (1992, p.755). At the base of the Huronian Supergroup in the Elliot Lake, Agnew Lake and Sudbury areas are several layered gabbro to anorthosite intrusions referred to as the East Bull Lake intrusive suite (Peck et al. 1993; James et al. 2002a, 2002b). These bodies have ages of circa 2475 Ma (Clough and Hamilton 2017; Krogh et al. 1984) and appear to be slightly older than the rocks of the Elliot Lake Group. and sandstone units are interpreted to represent deposition during warmer intraglacial or postglacial periods in either fluvial or marine environments (Junnila and Young 1995; Fralick and Miall 1989). Huronian Supergroup deposition was complete by 2217 Ma, the age of the Nipissing gabbro (Davey et al. 2019; Corfu and Andrews 1986; Noble and Lightfoot 1992). The Huronian Supergroup has been interpreted to represent a Wilson cycle, starting from a rifting phase represented by the Elliot Lake, Hough Lake and Quirke Lake groups; followed by a passive margin sequence (Cobalt Group); and concluded by a continent-arc collision between the SuperiorSouthern provinces and the Wisconsin Magmatic Arc Terrane (e.g., Young 1983; Hoffman 1989; Bennett et al. 1991). Each of the 3 groups overlying the Elliot Lake Group consists of sedimentary cycles of conglomerate, mudstone, siltstone or carbonate, capped by crossbedded sandstone (Bennett et al. 1991) (Figure 2). Conglomerate units (e.g., Ramsey Lake, Bruce and Gowganda formations) in each of the cycles have been interpreted as being glaciogenic in origin, likely deposited in a marine environment adjacent to an ice shelf. The siltstone 105 �Proceedings of the 68th ILSG Annual Meeting - Part 2 This collisional event at 1870 to 1835 Ma, termed the Penokean Orogeny, is believed to be responsible for most of the metamorphism and deformation present in the Huronian Supergroup. The scale and intensity of the Penokean Orogeny remains a subject of debate (Davidson et al. 1992; Card 1992; Raharimahefa et al. 2014; Holm et al. 2018; Zi et al. 2022), in part because the Penokean Orogeny has no associated plutonism in Ontario. In contrast, Riller et al. (1999) attributed deformation and peak metamorphism of the Huronian Supergroup to the Blezardian Orogeny (2470-2220 Ma), with subsequent transpressional deformation during the Penokean. These divergent views reflect the lack of constraints on the age of Huronian Supergroup metamorphism and deformation. The Sudbury Igneous Complex was emplaced at 1850 Ma (Krogh et al. 1984; Davis 2008) and consists of a lower, ore-bearing sublayer, a main mass of norite, and an upper granophyre (e.g., Dressler et al. 1991). Associated with the Sudbury Igneous Complex are brecciated rocks, termed the Sudbury breccias (e.g., Dressler et al. 1991), consisting of randomly oriented blocks of country rock in a fine-grained, pseudotachylite matrix. The breccias occur up to 200 km from Sudbury but are most abundant near Sudbury. The Sudbury Igneous Complex and related rocks have been variously interpreted as originating from meteorite impact, impact-induced plutonism and volcanism, and volcanism (see reviews in Pye et al. 1984). The southern part of the Sudbury Igneous Complex was weakly metamorphosed by an event that also retrograded metamorphosed rocks of the Huronian Supergroup. Regional sodium and potassium metasomatism and silicification have intensely altered rocks locally within the Huronian Supergroup, especially along faults, at circa 1700 Ma (Meyer et al. 1990; Gates 1991; Schandl et al. 1994; Easton et al. 1996; Fedo et al. 1997). Significant magmatism occurred again at 1750 to 1730 Ma and at 1500 to 1450 Ma in the Killarney Magmatic Belt (van Breemen and Davidson 1988; Davidson and van Breemen 1994; Krogh 1994). Figure 2. Idealized Huronian Supergroup stratigraphy as utilized by the Ontario Geological Survey based on Robertson, Card and Frarey (1969). Yellow units are sandstone dominated, blue units are carbonate rocks (limestone or dolostone), brown units are mudstones and/or turbidites (lined) or conglomerates. Green units are volcanic rocks. 106 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 1. Timing of major geological events and summary of age constraints on the main rock units present in the Sudbury area. Event and/or Map Unit Age Constraint (Ma) Comment and/or Source Grenville dike swarm 586±4 Kamo, Krogh and Kumarapeli (1995) Pegmatite vein emplacement 989±2 Corfu and Easton (2000) Age of peak metamorphism in the hangingwall of the Grenville Front tectonic zone 1000 to 990 Corfu and Easton (2000), this study Age of peak Grenvillian metamorphism in the Central Gneiss Belt 1040 to 1030 Carr et al. (2000) Sudbury dike swarm 1238±4 emplaced in or along northwest-trending faults in the Southern Province, deformed and metamorphosed in the Grenville Province, Krogh et al. (1987). Killarney magmatic belt second-stage magmatism, coincident with magmatism in the Eastern Granite Rhyolite Province and in the Central Gneiss Belt 1471±3 van Breemen and Davidson (1988) Regional albitization metasomatic event 1701±4 U/Pb monazite, Schandl, Gorton and Davis (1994); fluid focussed along northwest faults Killarney magmatic belt volcanism and high-level plutonism 1740, 1747±3, 1749±12 van Breemen and Davidson (1988); Sullivan and Davidson (1993); Davidson and van Breemen (1994) Northwest-trending regional faults Pre-1700, post-1850 Faults cut Sudbury Structure Penokean orogeny (folding and metamorphism of Huronian Supergroup rocks?) 1775±10 ~1835 Peak deformation. Zi et al. (2022) Peak metamorphism. Holm et al. (2001) Impact event and formation of Sudbury breccia 1850±1 Krogh, Davis and Corfu (1984); Davis (2008) Penokean arc formation 1880-1870, 1845-1830 Zi et al. (2022) Thrust faulting post-F2 pre-regional faulting Sudbury breccia localized along these faults, suggesting they are pre-Sudbury Structure F2 folding post-2200, pre-1700, pre 1850? Pre-regional faulting, Nipissing sills axial planar to folds F1 folding pre-2200 Nipissing sills folded or intruded into early folds Emplacement of Nipissing gabbro sills 2217±4 Davey et al. (2019); Corfu and Andrews (1986); Noble and Lightfoot (1992) Huronian Supergroup sedimentation >2220 but <2460 Youngest detrital grains in Bar River Fm are 2306 Ma (Hill, Davis and Corcoran 2018) Huronian Supergroup felsic volcanism and related plutonic rocks, including the Matachewan dike swarm ~2477 to 2375 (2450±25, 2460±20, 2477±9, 2415±5 Krogh, Davis and Corfu (1984), Heaman (1997); Corfu and Easton (2000), Krogh, Kamo and Bohor (1996), Smith (2002); Bleeker et al. (2015) Emplacement of East Bull Lake intrusive suite rocks 2475±2 Heaman (geochronologist, University of Alberta, personal communication, 1999); Clough and Hamilton (2017) Emplacement of orthopyroxene hornblendite bodies (East Bull Lake suite) 2468±5 Corfu and Easton (2000) Emplacement of alkali feldspar granite and megacrystic granodiorite near River Valley 2660 to 2665 Bodies intrude Crerar and Pardo gneiss, Easton (2003) High-grade Archean metamorphism and migmatization 2647±4 Krogh, Davis and Corfu (1984); Wodicka and Card (1995); Ames et al. (2005) Emplacement ages of Archean units in the Sudbury area 2711±7 to 2642±1 see also Table 5 Krogh, Davis and Corfu (1984); Wodicka and Card (1995); Chen, Krogh and Lumbers (1995); Meldrum et al. (1997); Ames et al. (2005) 107 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The last major magmatic activity in the Southern Province occurred at circa 1240 Ma with the emplacement of the northwest-trending Sudbury diabase dike swarm (Krogh et al. 1987). This event is noteworthy, as rocks of this dike swarm can be traced across the Grenville Front into the Grenville Front tectonic zone, providing an important marker horizon (e.g., Bethune 1997). (Figure 3). movement would be necessary to juxtapose granulites of the Levack gneiss complex against higher-crustal-level rocks to the east, however, if the Upper Wanapitei River fault re-activated an older listric fault system, then considerably less relative uplift across the fault may be present. Murray fault system Significant changes in stratigraphic thickness within the Huronian Supergroup occur across the Murray fault system (e.g., Bennett et al. 1991, and references therein). North of the Murray fault, the McKim Formation is tens of metres thick, whereas south of the fault, it probably exceeds 1000 m. The thickness and facies variations across the Murray fault system suggest that the faults represent south side down, syn-sedimentary, growth faults that were reactivated during compression attributed to the Penokean Orogeny (e.g., Zolnai et al. 1984). Significant changes in thickness within the Huronian Supergroup occur east and west of a line roughly coincident with the trace of the northtrending Upper Wanapitei River fault. Debicki (1990) estimated the thickness the Huronian Supergroup at Sudbury to be approximately 10,350 m, 85% of which consists of the lower 3 groups. In contrast, east and northeast of Wanapitei Lake, the thickness of the Huronian Supergroup is approximately 6,250 m, 75% of which consists of the Cobalt Group. The Elliot Lake, Hough Lake and Quirke Lake groups are all considerably thinner east of Wanapitei Lake, and the McKim and Ramsey Lake formations are apparently absent (Easton and Murphy 2002). In addition to stratigraphic thickness variations, the Murray fault system also marks profound changes in structural style, metamorphic grade and magmatic associations (Card et al. 1972). Deformation is more complicated and of greater intensity south of the fault. Likewise, metamorphic grade is higher immediately south of the fault (amphibolite facies transitional southward to greenschist facies) than to the north (greenschist to subgreenschist facies). South of the fault, there are several 1750 Ma and younger granitoid complexes (e.g., Cutler batholith) (Davidson and van Breemen 1994), whereas, north of the fault, there are no such intrusions. Major Fault Systems Upper Wanapitei River fault The Upper Wanapitei River fault has had a protracted deformation history, exhibiting at least 7 to 8 km of left-lateral movement between 2170 and 1850 Ma (Buchan and Ernst 1994), and at least 3 km of left-lateral movement post-1040 Ma (Easton and Murphy 2002). According to Easton (2000), the north-trending Upper Wanapitei River fault apparently divides the Archean rocks in the Elliot Lake to North Bay area into two domains, with the boundary between these domains passing through Street Township. The eastern domain, which includes the River Valley–Hagar area consists of supracrustal and metaplutonic rocks, with deeper levels in the crust being exposed to the south, likely due to Grenville orogenesis. In contrast, the western domain is pluton-dominated, with deeper levels of the crust, being exposed to the east. The amount of vertical movement across the fault is unknown. Significant vertical Northeast of Coniston, the Grenville Front boundary fault and the Murray fault system are thought to merge into the Wanapitei fault (Davidson 1997), which can be traced into Street Township. This fault is then offset to the north by the Upper Wanapitei River fault and continues eastward as the Ess Creek and Grenville Front boundary faults along the trend of the Kabikotitwia and Sturgeon rivers (Easton and Murphy 2002). Thus, Huronian Supergroup strata located west and northwest of the Wanapitei and Ess Creek faults occur north of the Murray fault system, whereas 108 �Proceedings of the 68th ILSG Annual Meeting - Part 2 any Huronian Supergroup rocks preserved within the Grenville Province would have originally been deposited south of the Murray fault system. These rocks are considered parautochthonous in the sense that they were once part of the autochthon, having been subsequently transposed and uplifted northwestward. It is difficult to recognize such rocks directly in many places along the Grenville Front This is due not only to the effects of intense reworking within the Grenville orogen, but also to juxtaposition, on opposing sides of the front, of rocks that were originally at different crustal levels, thus exhibiting different states of deformation and metamorphism, and not necessarily representing the same lithologic units. Geochronology has been of inestimable value in making broad correlations across the front. Lack of identification of deformed and metamorphosed equivalents in the Grenville Province of the flatlying supracrustal rocks (e.g., Huronian Supergroup) northwest of the front is interpreted to be due to uplift and erosion of these successions, so that only their substrate is preserved. This field trip specifically examines these issues. The Grenville Province Introduction Rocks of the Grenville Province in Ontario range in age from circa 2690 to 990 Ma. All rocks older than 1300 Ma are pre-Grenvillian, whereas those younger than 1300 Ma are Grenvillian. With respect to nomenclature, a variety of subdivisions are in use for the Grenville Province in Ontario and fall into 2 broad groups: those that are lithologically based, commonly with a long history of usage (e.g., Wynne-Edwards 1972); and those that are more tectonic or interpretative in character, generally of more recent vintage (e.g., Rivers et al. 1989; Carr et al. 2000). Geological domains and their boundaries between the different types do not always coincide from one scheme to another (e.g., the Central Gneiss Belt contains paraautochthonous and allochthonous rocks), however, both approaches are valid, and usage is based on needs (e.g., lithologic- and historic-based terminology may be used more on detailed maps (<1:50 000 scale), tectonic-based terminology may be used on regional maps and in academic literature). Key divisions of the Grenville Province are listed in Table 2. The extent of Superior and Southern province rocks within the Grenville orogen can be documented by Nd depleted mantle model ages. Dickin and McNutt (1989) found that Archean and Paleoproterozoic model ages of gneisses are restricted to northwest of a line extending from Key Harbour to Timiskaming, well southeast of the GFTZ and some 60 km from the Grenville Front. It can be argued, however, that meta-sedimentary rocks younger than the Huronian Supergroup (>2.2 Ga) could equally well have had Archean provenance. Gneisses southeast of this line have distinctly younger Nd model ages (circa 1.9 Ga). The field trip route mostly lies within the northernmost part of the Grenville Front tectonic zone (GFTZ), but also includes some rocks in northernmost Nepewassi domain. Both are part of the parautochthonous belt or the Laurentian Margin, which is defined as that part of the Grenville Province in which the rocks, although thoroughly reworked during the Grenvillian and/or earlier orogenies, can be reasonably equated with rocks of older Shield provinces to the northwest (Rivers et al. 1989; Carr et al. 2000). 109 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3. Distribution of 1.24 Ga Sudbury diabase and metadiabase, -1.17–1.15 Ga coronitic olivine metagabbro, and eclogitic rocks in the Central Gneiss Belt, Ontario and westernmost Quebec. The broken line near the Grenville Front is the southeast margin of the Grenville Front tectonic zone (Wynne-Edwards 1972). Figure from Ketchum and Davidson (2000). 110 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 2. Key divisions and boundaries within the Grenville Province in Ontario. Historic/Lithologic Regional Tectonic Local Tectonic/Historic Grenville Front Tectonic Zone (GFTZ) Para-autochthonous belt (Rivers et al. 1989) or Laurentian margin 1 (Carr et al. 2000) Segments 1, 2, 3 Central Gneiss Belt (CGB) (WynneEdwards 1972; Easton 1992) Para-autochthonous and/or allochthonous belt (Rivers et al. 1989), Laurentian margin 2 and 3 (Carr et al. 2000) Parry Sound, Algonquin, Tomiko, Beaverstone terranes Britt, Fishog, Go Home (lower), Go Home (upper), Huntsville, Kiosk, McCraney, McClintock, Moon River, Nepewassi, Novar, Powassan, Shawanaga, Sequin, Tilden Lake domains Central Metasedimentary Belt (CMB) (Wynne-Edwards 1972; Easton 1992) Composite Arc Belt (CAB) and Frontenac-Adirondack Belt (FAB) (Carr et al. 2000) Bancroft, Elzevir, Frontenac terranes (Elzevir contains Anstruther, Belmont, Grimsthorpe, Mazinaw, Sharbot Lake domains), Adirondack Lowlands and Highlands Grenville Front (Wynne-Edwards 1972; Easton 1992) North limit of Grenville metamorphism and penetrative deformation (locally migmatite front) Grenville Front boundary fault (GFBF) Allochthon Boundary Thrust (ABT) (Rivers et al. 1989) Separates para-autochthonous and allochthonous rocks (Rivers et al. 1989; Carr et al. 2000) a.k.a. central Britt shear zone, Shawanga shear zone Laurentian Margin - Composite Arc Belt boundary (Carr et al. 2000) Composite Arc boundary zone (CABZ) (Carr et al. 2000) Central Metasedimentary boundary zone (CMBBZ), a.k.a Central Metasedimentary Belt boundary thrust zone (CMBbtz) Composite Arc Belt – FrontenacAdirondack Belt boundary (Carr et al. 2000) Frontenac-Adirondack boundary zone (FABZ) (Carr et al. 2000) a.k.a. Maberly shear zone, Sharbot LakeFrontenac boundary Important Boundaries Nepewassi domain eastward continuation into the Grenville Province of the main igneous components of the Killarney Magmatic Belt, which is straddled by the Grenville Front in the Killarney area. Table 3 summarizes the ages from plutons of both suites in the Killarney Magmatic Belt, the Grenville Front tectonic zone, and the Nepewassi domain. Leucogabbro to anorthosite of the St. Charles and Mercer intrusions cut the West Bay batholith and were emplaced at circa 1225 Ma (Prevec 2004). The Nepewassi domain (Easton 1992) was discriminated from its neighbours on the basis of structural trends as well as rock types. The area underlain by the Nepewassi domain in the field trip area was mapped by Lumbers (1975) at 1:126 720 scale. The Nepewassi domain is underlain by compositionally heterogeneous migmatitic gneisses which have a polycyclic history. Plutonic rocks in the domain form 2 suites which are less deformed than their typically migmatitic host rocks (Lumbers 1975): an older, granite-monzogranite suite circa 1740 Ma that includes the migmatitic West Bay and the Sturgeon Falls batholiths, and a younger, non-migmatitic suite, circa 1450 to 1420 Ma, that includes the Cosby pluton. Near Alban, the Cosby pluton intruded a thick sequence of quartzite, known as the French River quartzite (cf. Lumbers 1975). Both plutonic suites represent an Limited geochronological data are available from the Nepewassi domain (Table 3). Tonalitetrondhjemite gneisses exposed between Hagar and Warren yielded U/Pb zircon ages of 2678 to 2683 Ma, with titanite indicating that regional metamorphism of these same gneisses occurred between 996 and 975 Ma (Chen et al. 1989). A grey gneiss located on Highway 535 north of Noelville yielded a similar age of 2680±11 Ma by laser 111 �Proceedings of the 68th ILSG Annual Meeting - Part 2 ablation on zircon (Van de Kerckhove 2016). These zircon ages are consistent with the Nd/Sm and Pb/Pb model ages of Dickin (1998a, 1998b) which indicate the presence of Archean crust throughout much of Nepewassi domain. (Lumbers 1975, p.98), but neither the details on the age determination nor the location of the sample are available. Aldis (2016) reported a laser U/Pb zircon age of circa 1434 Ma from the Cosby pluton near Noelville, similar to the age reported by Lumbers (1975). Along the western margin of the Nepewassi domain and the Grenville Front tectonic zone, a layered gneiss unit, the French River “paragneiss”, yielded a zircon population with an age of 1744±11 Ma (Krogh 1989). The homogeneous nature of the zircon population (Krogh 1989) suggested that the French River “paragneiss” is not a typical clastic metasedimentary rock, but rather that it may have been derived from metamorphosed volcanic and/or volcaniclastic rocks. Alternatively, it may be a highly strained orthogneiss. In contrast, the French River quartzite contained only Archean zircons, with monazite giving a metamorphic age of 1062±15 Ma (Krogh 1989). Quartzites examined by Van de Kerckhove (2016) northeast of Noelville had detrital zircon populations ranging from 2563 to 2962, with peak populations between 2686-2702 Ma, consistent with detrital zircon populations from Huronian Supergroup rocks in the Southern Province (see summary in Easton 2019). Van de Kerckhove (2016) also reported metamorphic zircon and monazite ages of 1755±11 and 1761 Ma, respectively, from the same area, suggesting a regional metamorphic event coincident with Killarney belt magmatism. New observations from the northwestern Nepewassi domain, which will be seen during the field trip. include the identification of granulitefacies, green and pink, garnet-bearing, potassium feldspar megacrystic granodioritic gneiss of the Estaire pluton and incorporated pods of hypersthene-bearing gneissic diorite, a likely comagmatic phase (Stop 7, 8). The Estaire pluton granodiorites are characterized by high Ba (>2000 ppm) and high Zr (>500 ppm) contents similar the those found in the West Bay pluton, south of Verner. In addition, quartzite (Stop 9), possibly correlative with the French River quartzite, occurs as a 2.2 km long, up to 300 m wide, belt on the south side of the Wanapitei River, only 2.2 km south of the boundary with the GFTZ. The Grenville Front The Grenville Front itself is a zone of southeastdipping faults and mylonites and has generally been placed at the southeast limit of recognizable Southern Province rocks (e.g., Lumbers 1975; Davidson 1997). Locally there are complications that have led to many debates concerning the identity of the Grenville Front and its distinction from other faults that intersect, merge with or are parallel to the Front (see discussion in Davidson 1997). In central Street Township the Grenville Front (which is coincident with the Wanapitei fault) has been displaced to the north by the younger, north-trending, Upper Wanapitei River fault by at least 850 m of sinistral and west-side-up movement (Easton and Murphy 2000, 2002). This displacement likely occurred after circa 590 Ma, as a Grenville swarm diabase dike in northern Henry and Loughrin townships is also displaced by northtrending faults. The age of the major plutonic units in Nepewassi domain is poorly known, but many probably have affinities to the Killarney magmatic suite (circa 1740) (Easton 2014). The migmatitic, French River granite, located along the western boundary between the Nepewassi domain and the Grenville Front tectonic zone, and intruded into the French River “paragneiss”, gave a robust Rb/Sr age as well as a U/Pb zircon age of circa 1700 Ma (Krogh and Davis 1969, 1972). The age of the texturally similar West Bay batholith near Lavigne has not been determined reliably, although a poor-quality laser U/Pb zircon age of circa 1255 was reported by Aldis (2016). A U/Pb zircon age of circa 1420 Ma has been reported from the Cosby pluton 112 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 3. Summary of geochronological data for Killarney area magmatic rocks, the Wanapitei complex and the Nepewassi domain. Age (in Ma) Unit Comment Killarney Area Magmatic Rocks 1749+12/–8 biotite granodiorite North of Chief Lake granite 1747±3 granodiorite Eden Lake complex 1744±29 granodiorite Eden Lake complex 1704±13 granitic dike 1596±39 granitic dike 1746+16/–6 granodiorite Cuts fabric and shear zone in Huronian Supergroup rocks near McFarlane Lake Cuts fabrics in Eden Lake complex Cutler batholith 1742±1.4 1464±2 1467±18 granite granite granite Killarney granite Chief Lake granite Chief Lake granite 1429 1447 1464 pegmatite dike pegmatite dike pegmatite dike South of GF, Chief Lake area South of GF, Chief Lake area South of GF, Chief Lake area Wanapitei Complex 1746+12/–6 quartz monzonite dike 1707±17 quartz monzonite dike 1746+6/–5, hornblende 996 metanorite Wanapitei complex, cuts metagabbro Wanapitei complex, cuts metagabbro Wanapitei complex, lower intercept age of metamorphism Wanapitei complex Method Source U/Pb TIMS zircon U/Pb TIMS monazite U/Pb LA-ICP–MS zircon U/Pb LA-ICP–MS zircon Davidson and van Bremen (1994) Sullivan and Davidson (1993) Raharimahefa, Lafrance and Tinkham (2014) Raharimahefa, Lafrance and Tinkham (2014) U/Pb LA-ICP–MS zircon U/Pb TIMS zircon Raharimahefa, Lafrance and Tinkham (2014) Davidson, van Breemen and Sullivan (1992) van Breemen and Davidson (1988) Davidson and van Bremen (1994) Raharimahefa, Lafrance and Tinkham (2014) Krogh (1994) Krogh (1994) Krogh (1994) U/Pb TIMS zircon U/Pb TIMS zircon U/Pb LA-ICP–MS zircon U/Pb TIMS zircon U/Pb TIMS titanite U/Pb TIMS monazite U/Pb TIMS zircon U/Pb LA-ICP–MS zircon U/Pb TIMS zircon Davidson, p.39 in Easton, Davidson and Murphy (1999) Rousell et al. (2012) Prevec (1993, 1992) U/Pb LA-ICP–MS zircon U/Pb LA-ICP–MS zircon Rousell et al. (2012) Nepewassi Domain French River 1744±11 “paragneiss” U/Pb TIMS zircon Krogh (1989) circa 1700, 1689±16 U/Pb TIMS zircon, Rb/Sr whole rock Krogh and Davis (1969, 1972) U/Pb TIMS zircon Lumbers (1975) U/Pb TIMS monazite Krogh (1989) U/Pb TIMS zircon Prevec (2004, 1992) U/Pb SHRIMP zircon Prevec (2004, 1993) U/Pb TIMS zircon Chen, Krogh and Lumbers (1995); Van de Kerckhove (2016) Chen, Krogh and Lumbers (1995) 1735±3 garnet metagabbro 1694±7, 1640±10 garnetiferous mafic Wanapitei complex, dike cuts other units, 2 populations 1420 1062±15 1245±48 1244±100 2678 to 2683 975 to 996 GFTZ near western boundary of Cosby subdomain, 2 sample sites, single population GFTZ near western boundary of French River Cosby subdomain, migmatitic, “granite” cuts French River “paragneiss” Cosby subdomain, Cosby pluton no location given French River Cosby subdomain, quartzite quartzite only had Archean zircons Mercer anorthosite Southern subdomain, 1222±2 Ma in Prevec (1992) St. Charles Southern subdomain, anorthosite 1206±36 Ma in Prevec (1993) tonalite, Northern subdomain, granodiorite range from 7 sample sites tonalite, Northern subdomain, age of metamorphism, granodiorite range from 6 sample sites U/Pb TIMS zircon, lower intercept Rousell et al. (2012) Abbreviations: GF, Grenville Front; GFTZ, Grenville Front tectonic zone; LA-ICP–MS, laser ablation inductively coupled plasma mass spectrometry; SHRIMP, sensitive high-resolution ion microprobe; TIMS, thermal ionization mass spectrometry. 113 �Proceedings of the 68th ILSG Annual Meeting - Part 2 At the time of formation, the Grenville Front was probably equivalent to the Main Boundary thrust of the current Himalayan orogen, marking the boundary between lowlands to the north and high-standing mountains to the south. afield, perhaps contributing to the fill of late Mesoproterozoic basins in the western and northern parts of North America (e.g., Hoffman and Grotzinger 1993; Rainbird et al. 1997). In the field, the Grenville Front has generally been placed at the southeastern limit of recognizable Southern Province rocks (Lumbers 1975; Davidson 1997, 1998). Locally, however, there are complications that have led to many debates concerning the identity of the Grenville Front and its distinction from other faults that intersect, merge with, or are parallel to it (see discussion in Davidson (1995, 1997, 1998)). In the west, the Grenville Front can be traced from Georgian Bay across the Killarney Magmatic Belt until Coniston where it intersects the Murray and the Creighton faults. Along most of its length in Ontario, the Grenville Front is characterized by a major, intense, moderately southeast-dipping mylonite zone a few metres to tens of metres thick. The mylonitic rocks have a dip-parallel lineation and kinematic indicators show a northwestward thrust sense. Rocks in the immediate foreland adjacent to front-parallel faults, which are generally steeper than the front mylonite zone, show cataclastic deformation. Gneissic and protomylonitic rocks southeast of the front show penetrative ductile deformation with the same northwest-directed thrust sense as the front mylonites. Thus, the Grenville Front marks a transition from brittle to ductile deformation toward the southeast. The timing of isotopic closure in a variety of mineral systems adjacent to the Grenville Front in the Sudbury area was examined by Corfu and Easton (2000). Zircon closed at 995 to 987 Ma, similar to ages reported all along the Grenville Front from Killarney to Labrador by Krogh (1994); this age represents some of the youngest Grenvillian activity in Ontario. Titanite and monazite from the same localities record only slightly younger ages between 989 to 977 Ma, consistent with the slightly lower closure temperatures for these minerals. Rutile and apatite ages from the same samples record ages of 973 to 971 Ma and 959-932 Ma, respectively, consistent with slow cooling along the Grenville Front after last movement at circa 995 Ma. In many places, front-parallel mylonite zones occur within the gneisses southeast of the front, and both these and the front mylonitic rocks exhibit local, superimposed cataclasis. This demonstrates the changing nature of deformation, from ductile to brittle, during the time taken for the orogen to rise and, with accompanying exhumation, to cool. With respect to the time taken for uplift, it is noteworthy that nowhere along the length of the Grenville Front is there any evidence that the foreland was ever the site of a basinal depression that received detritus from an elevated Grenvillian mountain belt (it is probably significant that sediments of suitable age that are part of the Midcontinent Rift fill have Archean or Penokean and not Grenvillian provenance). This can be explained through a combination of factors such as the length of time taken for exhumation, and the effect of crustal thickening within the orogen that may have allowed the foreland to remain isostatically buoyant (e.g., Jamieson and Beaumont 1989). Lack of a foreland basin would have allowed only ephemeral deposition of coarse detritus adjacent to the orogen, and finer detritus to bypass the foreland and to be spread widely farther As mentioned, the Murray fault (Figure 1, 4, 5) is a major west-trending lineament, which locally separates weakly metamorphosed Huronian Supergroup rocks to the north from strongly metamorphosed and deformed, Mesoproterozoicgranite-bearing, Huronian Supergroup rocks to the south. Metamorphic contrast across the Murray fault is most pronounced in proximity to plutonic rocks (e.g., the Eden Lake and Cutler plutons). Northeast of Coniston (Figure 4, 5), the Grenville and Murray faults are thought to merge into the Wanapitei Fault (Davidson 1997). In central Street 114 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4. A) The juncture between the Grenville Front mylonite zone and the Murray–Wanapitei fault south of Coniston. Stop 2-5 is the same location as Stop 1 in this guide, Stop 2-3 is the same as Stop 3. Abbreviations: GFMZ, Grenville Front mylonite zone; L, lake. B) Regional relationship between the Grenville Front, the Murray fault, and faults extending westward from the Ottawa-Bonnechere rift system. Dashed lines are Neoproterozoic Grenville swarm dikes, inverted triangles in Lake Nipissing are alkalic complexes associated with Neoproterozoic rifting. Both figures from Davidson (1995). 115 �Proceedings of the 68th ILSG Annual Meeting - Part 2 intrusive rocks of both the East Bull Lake and the Nipissing intrusive suites, as well as several types of migmatitic gneisses, likely of Neoarchean age (Easton 2000). The eastern segment, between River Valley and the Ottawa River, includes rocks mainly derived from the adjacent Superior Province. The field trip area straddles the boundary between the central and eastern segments. The GFTZ can be envisaged broadly as an anastomosing network of higher-strain rocks, moderately inclined and shallowing southeastward, surrounding lower-strain pods and lenses that are presumably elongate parallel to the prevalent southeast-plunging stretching lineation. The same style of structure can be seen in many places at outcrop scale and, if one considers porphyroclasts in mylonite, also at microscopic scale. Its southeast margin is ill-defined; as the front-parallel layered structure becomes shallower to the southeast; it also becomes progressively warped (buckled) about gentle, predominantly southeast-plunging axes so that the structural grain expressed at the surface changes from northeast to southeast. Map-scale enveloping surfaces, however, maintain a generally northeast trend. Figure 5. Fault interpretations in the Coniston area. A) after Lumbers (1975); B) modified after Dressler (1984). Stop 6 on the figure is located just southwest of Stop 3 in this guidebook. Figure from Davidson (1997). Township, the Grenville Front (Wanapitei Fault) is displaced to the north along the younger, northtrending, Upper Wanapitei River fault by at least 850 m of sinistral and west-side-up movement (Easton et al. 1996; Easton and Murphy 2002). In the Street Township area east of Sudbury, a rapid southeastward increase in metamorphism is present, with garnet-staurolite-kyanite developed within 800 m of the front and sillimanite-potassium feldspar rocks within 2 km. (Easton and Murphy 2002; Easton et al. 1999). Metamorphism and progressive disruption of the Sudbury swarm dikes occurs much closer to the front than to the southwest, and successive orthopyroxeneclinopyroxene-garnet coronas are fully developed in non-deformed cores of dismantled lenses within a kilometre of the front. Northwest of the front, shaly interbeds in the Mississagi Formation (predominantly cross-bedded feldspathic arenite) and shales of the underlying Pecors Formation are low-grade phyllites (muscovite-chlorite-albitequartz) 800 m from the Grenville Front. Nipissing gabbro contains the assemblage epidote-actinolitechlorite-albite ± quartz. The Grenville Front tectonic zone (GFTZ) The Grenville Front tectonic zone (GFTZ) is a region up to 30 km across lying between the Grenville Front and the Central Gneiss Belt of the Grenville Province (Figure 1). Easton (1992) divided the Grenville Front tectonic zone in Ontario into 3 lithologic segments. The western segment between Killarney and Wahnapitae comprises rocks equivalent in age, geophysical signature, and rock type to the adjacent Killarney Magmatic Belt. The central segment stretches from Wahnapitae to River Valley and contains mafic 116 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Despite its proximity to Sudbury, no rocks that can be related to the Sudbury impact event (e.g., metamorphosed Sudbury breccia; Offset dikes, etc.) have been reported from the GFTZ. the Sudbury dike swarm. There is no relative age constraint on the leucogranite pegmatites, although they are metamorphically recrystallized. East of the Wanapitei River The GFTZ zone near Wanup In contrast, east of the river, several large roundish or ovoid plutonic bodies are present, likely related to the Killarney Magmatic Belt, including the Wanapitei Complex (6 x 2 km in size) and the Cleland stock (3 x 3 km in size). East Bull Lake intrusive suite rocks former larger, more continuous, and better-preserved bodies, including the Red Deer Lake intrusion (11 km long, up to 1 km wide) and a body present along the southeast margin of the Wanapitei Complex. Host gneisses are dominantly quartzofeldspathic and contain. garnet-in both melanosomes and leucosomes (Photos F, G, H, I), and are similar to many of the gneisses present in Street Township only a few kilometres to the north-northeast. Also present are deformed, lenticular granitoid bodies (“metaarkoses” of Lumbers 1975), which in map pattern appear to define broad folds (see maps of Lumbers 1975; Dressler 1984). Similarly, the kyanitebearing paragneiss units west of Wahnapitae (Grant et al. 1962; Pearson 1959; Easton and James 1997) are more continuous, and better preserved, than possible correlative schistose rocks to the west of the Wanapitei River (Stop 13). Calc-silicate gneisses and impure marbles have not been reported from the area east of the river. Sudbury swarm dikes are pod-like, with minimal strike lengths. Late granite pegmatites seem to be less abundant in the area east of the river. Finally, the boundary with the Nepewassi domain is approximately 12-15 km from the Grenville Front. A key observation from the 2021 mapping program by the author is that the Grenville Front tectonic zone in the Sudbury area displays different structural styles whether one is west, or east, of the Wanapitei River. Most of the stops on the field trip are in the area west of the Wanapitei River, primarily for logistical reasons. West of the Wanapitei River West of the river, lithological units consist mainly of highly-strained gneisses, typically migmatitic (Photo A, B), that form thin, near continuous belts interlayered with migmatitic amphibolite, amphibolite and garnet amphibolite. Possible metasedimentary units, including schists containing aluminosilicate minerals, calc-silicate gneisses, and minor impure marble (Photo C), form thin, lenticular, discontinuous units that occur locally within the package of highly-strained gneissic and amphibolitic units. Rocks of the East Bull intrusive suite are locally present, but form thin, discontinuous units. Sudbury swarm dikes are large, with strike lengths of 50 to 100 m. In addition, the boundary with the Nepewassi domain is only 8 km from the Grenville Front. Pegmatite dikes are common west of the Wanapitei River and are predominantly granitic. Some are deformed and concordant or nearconcordant, with gneissosity, whereas others are highly discordant. Large, late, niobium-yttriumfluorine (NYF), variably-zoned, discordant, granitic pegmatite dikes (e.g., Stop 6a, 6b) are common in the area west of the river, and many have been quarried in the past, mainly for feldspar and/or mica (Vos et al. 1981). Narrow (0.5 to 3.0 m wide), garnet-baring, fine-grained leucogranite pegmatites are abundant in the GFTZ near the boundary with Nepewassi domain (Photo D, E) .and were not observed by the author east of the river. The late NYF pegmatites are younger than Structure along the Wanapitei River The feature causing the observed lithological differences across the Wanapitei River has a linear, north-northwest trend and appears as a weak linear magnetic feature in the low-resolution magnetic data available for the area (Figure 6, upper). The course of the Wanapitei River coincides with this structure from Coniston to Estaire, which obscures direct examination of the rocks immediately 117 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Photo C. Rocks west of the Wanapitei River. Impure dolomitic marble and calc-silicate layers in southerly of two calc-silicate and marble bands on the west side of Highway 69/400, northside of Old Wanup Road overpass. Scale card is 9 cm long (UTM 510571E 5140422N). Photo A. Rocks west of the Wanapitei River. Folded granitic leucosome in migmatitic, garnetbearing, gneissic diorite at north end of the roadcut. West side of Highway 69/400, 1.8 km from the Grenville Front (UTM 509315E 5141309N). Photo D. Rocks west of the Wanapitei River. Most of rock face is relatively massive, leucosome-poor migmatitic, garnet-bearing, gneissic granodiorite, however there is a zone that is much more leucosome-rich in the lower, centre part of the photo. West side of Highway 69/400 (UTM 5111057E 5137014N). Photo B. Rocks west of the Wanapitei River. Most of rock face is relatively massive, leucosome-poor migmatitic, garnet-bearing, gneissic granodiorite, however there is a zone that is much more leucosome-rich in the lower, centre part of the photo. West side of Highway 69/400, 1.8 km from the Grenville Front (UTM 509366E 5141236N). Photo E. Close-up of white, garnet-bearing, near concordant white pegmatite dike shown in Photo D. Scale card is 9 cm long. 118 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Photo F. Rocks east of the Wanapitei River. Closeup view of layered felsic metatexite gneiss at this station. Outcrop is on the west side of St. Cloud Road. Scale card is 9 cm long (UTM 515775E 5139250N). Photo H. Rocks east of the Wanapitei River. Layered garnet-bearing felsic metatexite gneiss at this station. Outcrop is on the west side of St. Cloud Road. Scintillometer for scale, instrument is 23 cm long, back of scintillometer is 10 cm square (UTM 515544E 5138949N). Photo G. Rocks east of the Wanapitei River. Layered felsic metatexite gneiss at this station. Outcrop is on the west side of St. Cloud Road. Hammer for scale, handle is 33 cm long (UTM 515544E 5138949N). Photo I. Rocks east of the Wanapitei River. Layered felsic metatexite gneiss at this station. Outcrop is on the west side of St. Cloud Road. Hammer for scale, handle is 33 cm long. 119 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Wanapitei River are consistent with higher temperatures. Regardless, more work needs to be done on rocks from both sides of the northnorthwest-trending structure to understand its origin and tectonic history. Granite Pegmatites in the GFTZ Ercit (1999) demonstrated that most granitic pegmatites in the northern Grenville Province, regardless of age, are classified as niobiumyttrium-fluorine (NYF) type pegmatites with no parental granite, indicating an anatectic origin and post-kinematic emplacement. Lumbers (1975) recognized two types of granitic pegmatite intrusions in the field trip area. Gneissic, deformed, granitic pegmatite intrusions composed mainly of quartz and potassium feldspar with minor mica and amphibole that occur as discontinuous dikes and sills. Zircon from one of these dikes in Cleland Township, east of the Wanapitei River, yielded an age between 1600 and 1700 Ma (Krogh and Davis 1970; Lumbers 1975), suggesting an affiliation with the Killarney magmatic suite. Figure 6. Upper. First vertical derivative of the residual magnetic intensity of the Wanup area showing the location of Wanup, the Grenville Front, and the north-northwest structure along the Wanapitei River. Northwest linear magnetic highs west of the Grenville Front are Sudbury diabase dikes. Lower. Bouguer gravity field. More common are late, post-metamorphic granitic pegmatite intrusions (circa 1000 Ma) that are distinctly zoned, commonly with quartz-rich cores (Stop 6b). Some are locally radioactive because of the presence of allanite (Stop 6a). Many have been quarried in the past, mainly for feldspar and/or mica (Vos et al. 1981). adjacent to the structure. The feature is also apparent in the Bouguer gravity data, separating a broad gravity high on the east side of the river from an area of lower density rocks to the west (Figure 6, lower). Which is interesting, as many of the mafic and granulite facies rocks west of the river have specific gravity values of 3.0 and 3.2 g/cm3. Acknowledgements Field work related to this guidebook was conducted in September to October 2021, with Julie Chartrand of the Ontario Geological Survey providing excellent field assistance. Dr. Manuel Duguet of the Ontario Geological Survey provided a technical review of the manuscript prior to publication. The north-northwest-trending structure is located where there is a flexure in the trace of the GFTZ, from north-northeast from Killarney to Wanup, to northeast from Wanup eastward. A preliminary explanation for the lithological and structural differences across the northnorthwest-tending structure would be that different structural levels are exposed on either side, with a likely deeper, and possibly hotter level on the west compared to the east. The greater degree of migmatization, and the presence of granulite facies rocks in the Nepewassi domain west of the 120 �Proceedings of the 68th ILSG Annual Meeting - Part 2 FIELD TRIP DETAILS we are on the same limb of the Coniston syncline that we will be on at Stop 3 (see Figure 4a). Geological Maps Cross the road and walk slightly west to the roadcut on the north side. The roadcut and rounded outcrops north of the road consist of Nipissing metagabbro that is massive but metamorphosed, containing secondary green amphibole, epidote and chlorite. The metagabbro also has low magnetic susceptibility. This metamorphic assemblage is typical of Nipissing gabbro throughout the Southern Province and thus does not appear to be a function of proximity to the Grenville Front. Geological compilation maps covering all or parts of the area of the field trip include Ames et al. (2005), Dressler (1984), Lumbers (1975) and Card and Lumbers (1973). ROAD LOG Note: Caution should be taken when parking vehicles on the shoulder of the highway and when examining outcrops located along Highway 17 and on other roads along the field trip route. All UTM co-ordinates are given in NAD 83 datum, zone 17. Cross the road again back to the south side. Here we see the northeast contact of the Sudbury diabase dike. Blocks containing its chilled contact with Mississagi sandstone can be found at the roadside. Near this contact the diabase is fine grained and contains xenocrysts of plagioclase which in turn include earlier-crystallized olivine crystals — a common feature at the margins of Sudbury dikes (Bethune 1997; Bethune and Davidson 1997). There is absolutely no evidence in thin section of any metamorphic reaction between these two minerals at this location. Note that this dike is unmetamorphosed and has high magnetic susceptibility. The Sudbury dikes are chemical distinct compared to other dike swarms in the region, and are characterized by TiO2 >2.5 wt.%, >700 ppm barium and >300 ppm Zr (Ketchum and Davidson 2000). This distinctive chemistry allows for these dikes to be recognized south of the Grenville Front, where they serve as important markers of deformation and metamorphic history (Figure 3). Leave from Science North at the junction of Paris Street and Ramsey Lake Road in Sudbury. Head south on Paris toward highway 69. 0.0 km — Junction of Highway 69 and the southeast and southwest bypass. Turn right onto the eastbound ramp and proceed east on Highway 17 after merging on to the Highway. 6.3 km — pull over on the right shoulder by the 1552.0 kilometre sign. Walk ahead (east) onto the roadcut on the south (right) side of the road. This locality is Stop 2-3 in Davidson (1995), Stop 1 in Davidson (1997); Stop 1-1 in Davidson et al. (2002). Stop 1. Mississagi Formation sandstone, Nipissing gabbro sill and a Sudbury swarm olivine-diabase dike UTM co-ordinates 508085E, 5144868N This stop lies within the Southern Province 900 m northwest of the Murray fault and approximately 3 km west-southwest of Stop 3. It is only two kilometres from the Grenville Front (see Figure 4a). Here a southeast-trending, vertical, 65m-thick olivine diabase dike of the 1235 millionyear-old Sudbury dike swarm cuts across the contact between Mississagi sandstone (south side of the road) and Nipissing metagabbro (north side). In the same dike that we see here, but farther to the south on the southeast side of the Murray fault, plagioclase xenocrysts become clouded with fine epidote, and rims of fine actinolite appear between olivine and plagioclase grains. Where Sudbury dikes have been identified in the immediate hanging wall of the Grenville Front, olivine in plagioclase xenocrysts has reaction coronas of pale orthopyroxene with outer rims of pargasite-spinel symplectite, and Ti-Fe oxide grains are surrounded Crossbedding in the sandstone indicates that the steeply dipping beds face north-northwest. In fact, 121 �Proceedings of the 68th ILSG Annual Meeting - Part 2 by Ti-biotite and garnet symplectite. Near the Grenville Front to the southwest, Sudbury dikes cut across and are chilled against a pre-existing mylonitic fabric that is developed in granitoid rocks as young as 1470 Ma (Bethune 1997; Davidson and Ketchum 1993), pointing to the existence of some kind of pre-Grenvillian tectonic front roughly coincident with the Grenville Front sensu stricto. In this regard it is pertinent that, in the hanging wall within a few kilometres of the front, metamorphic monazite in pelitic gneiss (Dudas et al. 1994) and zircon from pegmatitic leucosomes (Krogh 1994) record an age of circa 1445 Ma. produced small (cm-size) poorly defined cones (Spray et al. 2007). Return to vehicles and continue east on the bypass. 11.2 km — Junction of the bypass and Highway 17, take the right merge lane onto Highway 17 and head east toward Coniston. 14.1km — Junction to Coniston (traffic light), continue east on Highway 17. 16.8 km — Junction with Highway 17 and the Coniston Hydro Dam Road just past the overpass over the railway tracks. Turn right onto Coniston Hydro Dam Road. Between Highway 17 and the parking area, the road crosses feldspathic sandstone beds of the Mississagi Formation. Return to vehicles and continue east on the bypass. 9.0 km — pull over on the right shoulder by just after the curve on the way up the hill. Examine outcrops on the east (right) side of the road. 18.4 km — Park in pullout area opposite the gates to the Coniston Hydro Dam. We will walk to the first series of outcrops which are on the north side of the railway tracks (UTM 513570E, 5146857N). This is Stop 2-5 of Davidson (1995), Stop 7 of Davidson (1997), Stops C-1 and C-2 of Easton et al. (1999), and Stop 1-2 of Davidson et al. (2002). Stop 2. Shattercones in Mississagi Formation sandstone UTM co-ordinates 509418E, 5147073N Sudbury is famous for its shatter cones, which are well exposed in units of the Huronian Supergroup, especially the Mississagi Formation. The roadcut exposes numerous large shatter cones developed in quartz arenite of the Mississagi Formation. This outcrop is best visited in late afternoon, where the evening sun provides excellent lighting. Note that the shatter cones are not isolated individuals. The whole outcrop is full of shatter cones, something that is not revealed on a polished glaciated surface. Stop 3a. Mississagi Formation sandstone, Southern Province side of the Grenville Front Two major faults of the Murray fault system in the Southern Province, the Creighton fault and the Murray fault itself (Card 1978), converge eastward and meet just north of Alice Lake, 2 km west of here (see Figure 4, 5). East-northeast of this juncture, a narrow valley in line with the Murray fault marks the Grenville Front. North of this valley are well-preserved Huronian Supergroup sandstones (Mississagi Formation) and Nipissing gabbro (not observed at the stop) at low metamorphic grade, and south of it, high-grade migmatitic gneisses of the Grenville Province. The covered interval between the two conceals the Wanapitei fault, and is as little as 25 m wide in places between Alice Lake and the village of Wahnapitae, 6 km to the northeast. There is no field evidence to suggest that the Murray and Wanapitei faults are not one and the same, contrary to The shatter cone collar around the Sudbury structure forms a near continuous ring extending up to 20 km distant from the contact between the footwall rocks and the Sudbury Igneous Complex. Prior to regional deformation and folding, most shatter cones pointed upward, as their impact source origin was from above. Despite statements to the contrary, no volcanic blast has ever formed a shatter cone collar, and nuclear blasts have at best 122 �Proceedings of the 68th ILSG Annual Meeting - Part 2 published maps of this area (Lumbers 1975; Dressler 1984); the two interpretations are illustrated in Figure 5. by Corfu and Easton (2000) indicates that the metamorphism is indeed Grenvillian and culminated at circa 995 Ma. Mississagi sandstone at this stop displays obvious primary sedimentological features — crossbedding is well preserved and indicates that the beds face the same way as they dip, namely steeply to the northwest. Cleavage in the silty interbeds dips steeply southward and is axial planar to a major, southwest-plunging syncline whose surface trace lies about 2 km to the northwest (see Figure 4a). Metamorphic grade is low: cleavage in the silty interbeds is given by aligned sericite; thin sections show the presence of minor greenish biotite, indicating that the grade is probably no higher than middle greenschist facies. To the east, the trace of the Wanapitei fault lies in the valley along which a power line runs; to the west it passes just south of the slag heaps that can be seen in the distance. Return to vehicles. Retrace route back to Highway 17 and continue east on 17. 20.2 km — Highway 17 and Coniston Hydro Dam Road, turn right and continue east on Highway 17. 22.7 km — Bridge over the Wanapitei River in Wahnapitae village. The Wanapitei fault which, as at Stop 3, marks the Grenville Front in this area, lies in the river valley. Bare hills on the north (left) side of the river are underlain by well-bedded Mississagi sandstone that faces northwest, away from the front, similar to what we observed at Stop 3. The large roadcut to the right (south), just past the bridge and the variety store exposes kyanite-bearing metasedimentary gneiss and mafic gneiss (garnet-bearing amphibolite) derived from gabbro; both of which are cut by coarse-grained pegmatite, itself deformed. Walk back to the outcrop by the road on the opposite side of the valley and railway line. 27.2 km — Junction Highway 17 and 537, continue east on Highway 17. Stop 3b. Gneisses on the Grenville Province side of the Grenville Front 31.5 — Junction Highway 17 and Sunset Road. UTM co-ordinates 5135633E 5146768N 32.6 — Pull off onto gravel area on the south side of the Highway. This locality is Stop 1, Day 3 in Easton, James and Jobin-Bevans (2010). The outcrop on the south side of the railway crossing is composed of migmatitic quartzofeldspathic, mafic and minor pelitic gneiss, and includes narrow mylonite zones that diverge southwestward from the Murray-Wanapitei fault line. The pelitic gneiss contains kyanite and sillimanite, and amphibolite contains garnet, attesting to middle to upper amphibolite facies. Stop 4 (Optional). Shear-Zone Hosted Orthopyroxene Hornblendite Body UTM co-ordinates 525451E 5152056N Examine the outcrop and large blasted boulders present on the west side of the pullout. They belong to an orthopyroxene hornblendite body of the East Bull Lake intrusive suite that is present within a high-strain zone that extends subparallel to the highway. Examples of these highly strained felsic gneisses can be examined in outcrops at the base of the hill east of the pullout. The top of the ridge south of the road and above the stop consists of layered leucogabbronorite of the East Bull Lake intrusive suite. Thus, although proximal to rocks of the suite here, the orthopyroxenite body is not This outcrop exposes a highly deformed mix of granitoid and hornblende gneiss, some with garnet, cut by mylonite zones whose rotated feldspar porphyroclasts indicate south-side-up displacement. These rocks clearly represent an entirely different crustal level to that exposed just 70 m to the north, which implies several kilometres of vertical displacement along the Wanapitei fault, provided that the metamorphism in these rocks is younger than that in the Mississagi Formation. In Street Township to the northwest, geochronology 123 �Proceedings of the 68th ILSG Annual Meeting - Part 2 directly in contact with other rocks of the suite, although that relationship has been observed elsewhere in Street and Awrey townships (Easton and Murphy 2002). Note the large, equant, orthopyroxene crystals, and the fine-grained amphibole matrix. Mineral chemistry indicates that the amphiboles are magnesium hornblende, tremolite and cummingtonite (Buckley et al. 1997; Easton and Murphy 2002). The amphiboles occur as individual grains and complexly exsolved and intergrown grains. Olivine grains from a body located closer to the Grenville Front are Fo73. Orthopyroxene compositions lie in the bronzite field and are magnesium-rich (En75-80). (Easton and Murphy 2002). closely and attempt to assess possible protoliths. Are these metasedimentary or metavolcanic rocks or are they highly deformed orthogneisses? Once at the northeast end of the outcrop, the protolith of some of these gneisses will become readily apparent, due to an area of lower-strain present in a macroscopic fold nose (Photo 2). This locality illustrates the perils of trying to identify protolith in many gneissic terranes, especially in areas of poor, incomplete, and lichencovered exposures. Remember, we are only a kilometre south of the Grenville Front at this locality, and only at upper amphibolite conditions. What cannot be ascertained at this stop is whether these gneisses reflect a rather simple metamorphic history, for example an Archean high-grade metamorphic event followed by reworking during the Grenville, or multiple metamorphic episodes throughout the Archean, the Paleoproterozoic and the Mesoproterozoic. Turn around an retrace route westward on Highway 17 32.7 km — Junction Highway 17 and Sunset Road. Turn right (north) onto Sunset Road and proceed for 400 m. Pull over and park on the left shoulder of the road at the entrance to MTO gravel pit 402002 (phone 705-4775478). Walk to the well exposed outcrops in the central part of the pit. This stop was used by Davidson (1997) but was not included in the guidebook descriptions. Return to vehicles and retrace route back to Highway 17. 34.4 km — Turn right and head west on Highway 17 toward Sudbury. 43.3 km — Junction Highway 17 and 537 in Wahnapitae just east of the bridge. Turn left onto 537 and head south. Stop 5. Gneisses in the Grenville Province 1100 metres south of the Grenville Front 48.8 km — Metamorphosed Sudbury swarm diabase dike is exposed on the west side of the roadcuts which consist mainly of Grenville gneisses. UTM co-ordinates of gate 524338E 5152230N A variety of predominantly migmatitic gneisses are exposed in the well-exposed outcrops in the floor and northwestern wall of this pit. We are approximately 1 km south of the Grenville Front at this stop The grey, garnet-rich, gneisses may have been metasedimentary rocks. In contrast, a variety of mafic rocks, some garnet-bearing, some not, are interlayered with the grey gneisses. Some of these mafic rocks are large rafts in the gneisses (Photo 1), whereas others are thinner, boudinaged and aligned, and may represented dismembered dikes. Garnet is abundant and occurs in both the melanosome and leucosomes of the gneisses. 60.5 km — Junction in Wanup, continue west (straight) toward Highway 400-69. Proceed to the most northeastern exposed outcrop. While doing so, examine the gneisses 124 �Proceedings of the 68th ILSG Annual Meeting - Part 2 69 and is located just to the south of the part of Dill Township that was mapped in detail by Kwak (1968) and Davidson (1997, 1998). This road cut is illustrative of what rocks look like in the Grenville Front tectonic zone only 7 km south of the Grenville Front. The key takeaway is that is near impossible to easily relate the rocks that we see here to any of the rock units that we see on the north side of the front in either the Southern or the Superior provinces. In contrast to Stop 5, which appears to be dominated by an abundance of metasedimentary gneisses, the large roadcuts to the west and the east are more representative of much of the Grenville Front tectonic zone in the Sudbury area, where mafic gneisses predominant, albeit with slivers and layers of rocks that may have originally been metasedimentary. In examining the mafic gneisses in the two roadcuts, pay attention to features such as degree and style of leucosome formation; the presence or absence of garnet, and the abundance of garnet in some of the mafic gneisses, which far exceeds what would normally be generated in a mafic rock during a single-stage metamorphic event. We will examine the west roadcut first. Photo 1. Grey migmatite with mafic pods at Stop 5. Stop 6A. West Roadcut. From west to east the roadcut consists of: Photo 2. Fold in grey migmatite at Stop 5 (524200E 5152280N). In the nose of the fold, even though the rock is still deformed and metamorphosed, it is clear here that the protolith of the rock was a matrix-supported conglomerate of unknown stratigraphic affinity.  Approximately 100 m of interlayered grey to locally rusty, thin layered paragneiss and weakly layered, variably migmatitic mafic gneiss and deformed granitoid layers  13 m wide outcrop gap  Approximately 25 m of variably migmatitic, texturally varied mafic gneiss. This unit hosts a thin sub-horizontal pegmatite. 61.8 km — Pull over and park in the pullout area on the right side of the road. This stop will examine the two large roadcuts to the west and east of the parking area.  Approximately 55 m of gneissic gabbro, with distinct east and west contacts. Along the western contact, relict large plagioclase crystals occur in the groundmass as dark equant to lath shaped crystals and as isolated crystals up to 20 mm long (Photo 4c). Energy dispersive X-ray analysis indicates that the large crystals are andesine, with a composition of An46. Plagioclase xenocrysts are common occurrence in Sudbury swarm dikes in the Sudbury area (Stop 1; Davidson 1997, p.16). The presence of large plagioclase crystals only along the western contact Stop 6. Gneisses in the Grenville Province 6 kilometres southeast of the Grenville Front Pullout 512055E 5137085N The pullout splits a near continuous road cut, approximately 800 m long, and up to 15 m high, into eastern and western halves. The road cut was created during the process of four-laning Highway 125 �Proceedings of the 68th ILSG Annual Meeting - Part 2 may simply reflect better preservation at that locality, as the gneissic gabbro at the contact has moderate magnetic susceptibility (1.1 to 2.3 x 10-3 SI units) compared with the rest of the body (<0.75 x 10-3 SI units). The fact that the gneissic gabbro is likey a metamorphosed Sudbury swarm gabbroic dike is confirmed by geochemistry, as these samples have the high TiO2 (>2.5 wt.%), barium (>700 ppm) barium and Zr (>300 ppm) contents typical of the Sudbury swarm (analyses 1-4, Table 4). The gneissic gabbro hosts a near-vertical feldsparrich pegmatite dike (Photo 3). The eastern part of the pegmatite dike contains bluish apatite crystals (Photo 4a) and allanite (Photo 4b). Scintillometer readings from the eastern part of the dike range from 50 to 98 ppm U and 130 to 208 ppm Th (Easton, unpublished data). The age of the pegmatite is not known, but is younger than circa 1240 Ma, the age of the host Sudbury dike. Photo 4. A) Blue apatite in “trains” along albite crystal boundaries in granitic pegmatite. B) Single allanite crystal in granitic pegmatite. C) Andesine xenocrysts in metamorphosed mafic dike adjacent to pegmatite. Photos from Péloquin et al. (2020). Photo 3. Near-vertical pegmatite dike cutting gneissic gabbro of the Sudbury dike swarm at Stop 6a. Dike is unevenly zoned, with the right (west) side dominated by potassium feldspar, and the left (east) side being more radiogenic and containing more quartz, allanite and apatite. 126 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 4. Summary of geochemical and mineral chemistry data from mafic rocks at Stop 6. Work was performed at the OGS Geoscience Laboratory. Co-ordinates in NAD83, Zone 17. Analysis 4 from Easton (2003), analysis 6 from Peck et al. (1995). Analysis Number Sample Number Easting (m) Northing (m) Rock Name SiO2 (wt %) TiO2 Al2O3 Fe2O3total MnO MgO CaO Na2O K2O P2O5 CO2 S LOI Total Ba (ppm) Rb Sr V Pb Th U Nb Y Zr Ni Cr Cu Au (ppb) Pd Pt amphibole pyroxene biotite plagioclase oxide sulphide 1 2 3 4 5 19RME-2001 19RME-2003 19RME-2005 99RME-0337 19RME-2003 511908 5137095 gneissic gabbro, cgr 46.37 2.87 17.01 15.51 0.193 5.56 7.26 3.19 1.40 0.639 <0.023 0.135 0.65 100.74 769 30 389 206 5.9 3.7 2.2 15.1 36.4 229 90 83 84 0.7 0.4 0.3 n/a n/a n/a n/a n/a n/a 511949 5137093 gneissic gabbro, east contact 46.28 3.18 15.88 16.41 0.208 5.23 7.45 3.39 1.38 0.710 0.038 0.211 0.42 100.63 805 31 373 257 7.3 4.8 <1.6 16.1 41.4 259 67 97 96 0.7 0.6 0.4 H none 10% An27-31 ilmenite pyrite 511898 5137080 gneissic gabbro, west contact 46.98 3.01 15.61 16.08 0.201 5.35 7.63 3.22 1.38 0.666 0.042 0.113 0.41 100.63 780 27 390 240 5.6 4.0 <1.6 15.4 38.6 241 80 90 66 1.o 0.6 0.4 H none 10% An27-31 ilmenite pyrrhotite 6 7 90DCP-226 19RME-2004 556111 5109966 Sudbury dike 511951 5137095 leucogabbroic gneiss 402140 5143082 gabbro. East Bull Lake 512003 5137095 migmatitic mafic gneiss 45.53 2.96 16.61 16.92 0.200 5.95 7.76 3.49 1.25 0.59 <0.03 0.08 <0.05 99.73 700 1 363 233 6 2.5 0.7 16 39 243 93 45 58 25 <8 <5 n/a n/a n/a n/a n/a n/a 44.61 1.38 18.11 13.26 0.125 7.63 11.03 2.36 1.02 0.038 0.116 0.249 0.57 100.16 204 31 360 439 3.5 <1.5 <1.6 2.2 18.8 42 12 35 70 1.5 <0.14 0.1 MH none trace An52-65 ilmenite pyrrhotite 48.82 1.55 17.70 15.13 0.20 2.54 8.74 2.76 1.32 0.11 0.09 n/a 1.50 100.37 399 59 263 261 nr nr nr nr 17 78 20 nr 282 nr nr nr n/a n/a n/a n/a n/a n/a 51.10 0.42 13.98 8.61 0.160 10.29 12.38 1.44 0.57 0.034 0.159 0.044 1.00 100.06 64 42 96 212 1.9 <1.5 <1.6 0.7 10.7 29 164 531 65 10.2 24.7 30.6 FH diopside trace An74-77 none pyrite Notes: Major element oxides are in weight %; trace element data are in parts per million, except for Au, Pd, Pt which are in parts per billion. Plagioclase composition of xenocrysts in sample 19RME-2005 is An46-55. Abbreviations: cgr, coarse-grained; FH, ferrohornblende; H, hastingsite; LOI = loss-on-ignition; MH, magnesiohastingsite; n/a = not applicable; nr, not reported. Amphibole classification of Leake et al. (1997). 127 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 4 — continued. Analysis Number Sample Number Easting (m) Northing (m) Rock Name 8 9 10 11 12 19RME-2011 19RME-2008 19RME-2014 19RME-2009 19RME-2010 512277 5137074 garnet amphibolite 41.91 0.99 19.86 16.90 0.154 7.30 11.31 0.95 0.52 0.024 0.312 0.306 0.58 512313 5137069 garnet amphibolite 41.72 1.00 18.98 18.00 0.165 8.16 9.92 0.79 0.59 0.023 0.133 0.239 1.10 512290 5137073 ultramafic dike 512312 5137069 paragneiss SiO2 (wt %) TiO2 Al2O3 Fe2O3total MnO MgO CaO Na2O K2O P2O5 CO2 S LOI 511971 5137100 gneissic granodiorite 73.31 0.27 13.98 2.61 0.022 0.81 1.87 3.15 3.70 0.024 0.035 0.011 0.41 46.39 0.56 10.52 10.99 0.166 17.60 9.77 1.33 0.51 0.105 0.239 0.050 1.61 90.32 0.03 5.78 0.47 0.09 0.86 0.39 0.87 0.82 0.012 <0.023 0.019 0.46 Ba (ppm) Rb Sr V Pb Th U Nb Y Zr Ni Cr Cu Au (ppb) Pd Pt TREE 1861 126.6 267 15 39.3 34.7 4.1 7.4 5.2 136 6 25 <9 0.6 0.16 0.19 160.91 131 14.6 551 774 <1.7 0.28 0.16 <0.7 3.9 11 40 58 167 2.3 0.23 0.28 18.73 106 13.6 454 797 <1.7 0.19 0.07 <0.7 4.3 8 31 37 263 2.6 0.15 0.22 17.21 162 8.1 217 230 <1.7 1.11 0.38 1.6 11.7 45 593 1891 30 1.1 4.94 4.36 47.78 1683 13.7 96 <3 <1.7 0.63 0.34 <0.7 4.8 104 5 23 21 <0.6 <0.14 <0.06 76.26 Notes: Major element oxides are in weight %; trace element data are in parts per million, except for Au, Pd, Pt which are in parts per billion. Abbreviations: LOI = loss-on-ignition; n/a = not applicable; nr, not reported; TREE = total rare earth elements. 128 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 5. Summary of geochemical data from felsic and mafic intrusive rocks at Stops 7 and 8. Work was performed at the OGS Geoscience Laboratory. Co-ordinates in NAD83, Zone 17. Analysis Number Sample Number Easting (m) Northing (m) Rock Name SiO2 (wt %) TiO2 Al2O3 Fe2O3total MnO MgO CaO Na2O K2O P2O5 CO2 S LOI Ba (ppm) Rb Sr V Pb Th U Nb Y Zr Ni Cr Cu Au (ppb) Pd Pt TREE 1 2 3 4 21RME-021 Stop 7 514042 5129295 gneissic megacrystic granodiorite 21RME-022 Stop 7 5129235 5129037 gneissic megacrystic granodiorite 21RME-024 >0.75 >0.75 0.71 >0.75 >0.75 0.15 0.14 0.08 0.16 0.14 0.20 0.059 0.14 0.167 0.054 0.33 0.190 0.017 0.39 0.430 0.103 0.66 0.196 0.137 0.82 2265 60.0 370 28 17 4.1 <1.3 15.7 32.0 743 5 12 25 1996 56.8 384 46 15 4.3 <1.3 24.4 31.7 541 6 11 17 <2700 95.5 363 33 18 4.7 <1.3 16.3 20.3 544 3 <7 13 666 44.4 434 218 12 3.4 <1.3 15.1 47.2 46 37 22 46 324 22.5 327 185 8 <1.9 <1.3 5.3 31.7 85 85 129 46 >129 >129 >99 >120 >53 514123 5129524 gneissic megacrystic granodiorite 5 21RME-020 21RME-023 Stop 7 Stop 8 514046 514151 5129293 5129037 fine-grained fine-grained gneissic gneissic diorite gabbro Notes: Major element oxides are in weight %; trace element data are in parts per million, except for Au, Pd, Pt which are in parts per billion. Abbreviations: LOI = loss-on-ignition; n/a = not applicable; nr, not reported; TREE = total rare earth elements. 129 �Proceedings of the 68th ILSG Annual Meeting - Part 2  Approximately 16 m of non-migmatitic, gneissic leucogabbro (Photo 5). In thin section, the gneissic leucogabbro has a texture characterized by 120° grain boundaries and consists predominantly of brown amphibole and andesine. Brown amphibole is common under granulite facies conditions (Froese 1973). This is consistent with metamorphic conditions in this part of Dill Township identified by Kwak (1968). Geochemical data suggests the leucogabbro may be part of the East Bull Lake intrusive suite (emplaced circa 2475 Ma; compare analyses 5 and 6, Table 4).  Approximately 10 m of migmatitic mafic gneiss.  4 m of migmatitic granodiorite gneiss (Photo 6), similar to the host rock to samples C96-2 and C96-1 in Corfu and Easton (2001). This granodiorite gneiss has elevated U and Th contents (analysis 8, Table 4) typical of early Paleoproterozoic felsic rocks in the Sudbury area such as the Creighton granite and the Copper Cliff rhyolite (both circa 2460 Ma).  10 m wide outcrop gap  Approximately 35 m of migmatitic mafic gneiss, locally with deformed granitoid veins and layers present. Photo 5. Non-migmatitic, leucogabbroic gneiss at Stop 6a. This rock is cut by the gneissic gabbro of the Sudbury dike swam, and thus is older than 1240 Ma. Scale card is 10 cm long. Stop 6B. Roadcut. From west to east the roadcut consists of:  Approximately 125 m of interlayered grey to locally rusty, thin layered paragneiss, complexly folded, and weakly layered, variably migmatitic mafic gneiss and deformed granitoid layers.  Approximately 45 m of migmatitic mafic gneiss, locally with deformed granitoid veins and layers present. This gneiss hosts a 30 m long, discordant zoned non-radiogenic pegmatite vein.  Approximately 200 m of garnet amphibolite (25-50% garnet) (analysis 9, 10, Table 4) (Photo 7). At one point, it is cut by a near-vertical ultramafic dike (analysis 11, Table 4). A thin (1 m thick), near vertical quartzose gneiss band is likely of metasedimentary origin given its high silica content (analysis 12, Table 4).  Approximately 100 m of interlayered grey to locally rusty, thin layered paragneiss and weakly layered, variably migmatitic mafic gneiss and deformed granitoid layers. Photo 6. Gneissic granodiorite to monzogranite at Stop 6a, possibly related to the Creighton granite. Scale card is 10 cm long. 130 �Proceedings of the 68th ILSG Annual Meeting - Part 2 All of the units are steeply-dipping (70-85°) to the southeast. The age of the paragneiss and mafic gneiss is not well constrained, although Easton and Murphy (2002) and Easton and James (1997) observed that migmatitic mafic gneiss in the Street Township area to the northeast were Archean, whereas non-migmatitic mafic gneisses and amphibolite units were Proterozoic. If the migmatitic granodiorite gneiss in the west roadcut is indeed correlative with similar granitoid rocks in the Street Township and the Sudbury area that are circa 2460 Ma, this would suggest that the migmatitic mafic gneisses in the roadcut are Archean. Photo 7. Close-up of garnetite that constitutes much of the roadcut at Stop 6b. Scale card is 1 cm long. Analyses 9 and 10, Table 4, are from this unit. The garnet-rich mafic gneisses (25-50% garnet) present in the eastern third of the roadcut (Photo 6) are of potential economic interest. Garnet-rich mafic gneisses occur as discontinuous lenses in the Grenville Front tectonic zone from Sudbury to River Valley and have been mined locally in Street Township as a source of garnet (Easton and Murphy 2002). Kwak (1968) shows the presence of other outcrops of garnet-rich gneiss to the northeast of the roadcut, suggesting that they may be more abundant in Dill Township than previously suspected. Easton and Murphy (2002) and Easton (1996, 2003) suggested that these garnet-rich rocks may represent metamorphosed hydrothermal altered rocks similar to those present in volcanogenic massive sulphide (VMS) systems. 63.1 km — Optional Stop – Large roadcuts dominated by grey orthogneiss occur on both sides of the highway (UTM co-ordinates north side 511060E 5136470N, south side 511010E 5136460N), with an approximately 15 m wide Sudbury swarm dike similar to that observed at Stop 6A present in the north roadcut (UTM 511032E 5136448N). 63.3 km — Junction 537 and Estaire Road, turn south onto Estaire Road. 63.6 km — Cross from Grenville Front tectonic zone into Nepewassi domain. 72.3 km — Junction Estaire Road and Nelson Road, turn right onto Nelson Road. 73.7 km — Pull over and stop by roadcut on north side of road. Return to vehicles and continue west on Highway 537. 131 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 7 – Estaire pluton, granulite, Nepewassi domain UTM co-ordinates514045E 5129295N At this location, we are approximately 15 km south of the Grenville Front, and approximately 6.5 km south of the northern boundary of Nepewassi domain. This roadcut consists predominantly of green (Photo 8a) and pink (Photo 8b), garnetbearing, potassium feldspar megacrystic granodioritic gneiss. The gneiss incorporates irregular pods of fine- to medium-grained gneissic diorite (Photo 8c), which could be a co-magmatic phase (analysis 4, Table 5). The green coloration is a typical effect of granulite-facies metamorphism. In thin section, orthopyroxene occurs in the gneissic diorite, along with clinopyroxene- garnet [Alm58Adr2Grs18Prp19Sps4]-hastingsite-biotite (4.5 wt. % TiO2) and andesine. Opaque minerals are pyrrhotite and pyrite. Orthopyroxene is locally altered to calcite and is the only mafic phase that shows alteration. The presence of orthopyroxene indicates that this rock reached temperatures of at least 800°C (Pattison et al. 2003). The overall mineralogy of the green and pink granodioritic gneiss is similar, but there are a few key differences. In the green rock, garnet [Alm66Adr3Grs21Prp7Sps4] is abundant, and occurs primarily along the margins of potassium feldspar megacrysts and large plagioclase laths. Amphibole is hastingsite, plagioclase is oligoclase, potassium feldspar contains 1.3-1.9 wt. % BaO, and biotite contains (3-5 wt. % TiO2). Both magnetite and ilmenite are present, with minor amounts of Al in magnetite and Mn in ilmenite; no sulphide minerals were observed. Fluorapatite and 50-200 micron-size zircon are also present. In contrast, in the pink rock, garnet [Alm67Adr2Grs20Prp10Sps2] is sparser, and plagioclase phenocrysts have andesine cores and oligoclase rims. Opaque minerals are ilmenite (stoichiometric) and minor pyrrhotite and pyrite, and allanite are also present. Neither sulphide mineral nor allanite are in the green rock. Photo 8. A) Fresh surface of garnet-bearing potassium megacrystic gneissic granodiorite at Stop 7 showing green coloration suggestive of granulite facies metamorphism. B) Fresh surface of garnet-bearing potassium megacrystic gneissic granodiorite at Stop 7 showing pink coloration. C) Fine-grained mafic gneissic diorite raft hosted in potassium megacrystic gneissic granodiorite. Scale card in all images is 10 cm long. 132 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 9 – Quartzite, Nepewassi domain The presence of potassium megacrystic granodiorite and fine-grained dioritic rocks is typical of rocks of the Killarney intrusive suite (circa 1740 Ma), which are common in the eastern part of Nepewassi domain (cf. Easton 2014). Preliminary geochemical data for granodiorite samples of the Estaire pluton are characterized by high Ba (>2000 ppm) and high Zr (>500 ppm) contents (analyses 1-3, Table 5.) These geochemical characteristics resemble those of the West Bay pluton in Nepewassi domain, located south of Verner (Van de Kerckhove and Easton 2016). A sample of this unit, from the roadcut at the road junction 100 m from the stop was collected for geochronology, but results were not available at the time of the trip. UTM co-ordinates 513333E 5134250N This roadcut consists of medium layered quartzite, compositionally a quartz arenite, with thin, boudinaged mafic layers (dikes?) (Photo 9). The stop is near the southern end of a 2.2 km long, up to 300 m wide, belt of quartzite, part of which was quarried for silica flux between 1910 and 1924, first by the Canadian Copper Company and later by the International Nickel Company of Canada (MDI 41I07SW00002). Quartzite units occur throughout the Nepewassi domain, most notably in the French River area but also as thin slivers in southeastern Nepewassi domain (Easton 2014; Van de Kerckhove 2016). A sample from this roadcut was collected for geochronology, but results were not available at the time of the trip. Return to vehicles and continue to road junction. 73.8 km — Junction, Nelson, McVittie, Secord roads, turn left onto McVittie Road. 74.1 km — Pull over and stop, roadcuts are present on both sides of the road. We are interested in the north end of the roadcut on the west side of the road. Stop 8 – Estaire pluton, intrusion breccia, amphibolite, Nepewassi domain UTM co-ordinates 514150E 5129080N. From north to south, the roadcut exposes a spectacular intrusion breccia with rafts of finegrained mafic material (gabbro to diorite) hosted by medium-grained granodiorite. The centre part of the outcrop is medium-grained gneissic granodiorite to monzogranite. The south end of the outcrop is a fine- to medium-grained gneissic gabbro to diorite (analysis 5, Table 5). As at the last stop, well-developed intrusion breccias and intercalation of mafic and felsic magmatic phases are typical of the Killarney intrusive suite. Photo 9. Quartzite at stop 9. Note thin amphibolite layer above the end of the hammer handle (which is 40 cm long). Krogh (1989) reported a predominantly Archean population (≥2650 Ma) with a metamorphic age of circa 1060 Ma. Quartzites in southeastern Nepewassi domain studied by Van de Kerckhove (2016) also had predominantly Archean populations (≥2650 Ma, but generally <2700 Ma). Van de Kerckhove (2016) suggested that the populations in the quartzites were suggestive that they may be correlative with the Lorrain or Bar River Formations of the Huronian Supergroup, however, there is nothing unique in the zircon 74.4 km — Turn around and return to Nelson Road, right onto Nelson Road. 75.8 km — Nelson Road and Estaire Road, turn left and head north. 81.6 km — Pull over and stop, examine roadcut on east (right) side of road. 133 �Proceedings of the 68th ILSG Annual Meeting - Part 2 populations of any of the Nepewassi domain quartzites that is diagnostic that they are part of the Huronian Supergroup versus representing another stratigraphic package. Davidson (1997, p.29-32) has an extensive discussion regarding whether or not Huronian Supergroup rocks are present in the Grenville Front tectonic zone. Biotite from near this Stop show evidence for excess argon, yielding ages between 950 to 1020 Ma (Fairbairn, Hurley and Pinson 1960; Hanson and Gast 1967), which are similar to zircon ages from the GFTZ in the Sudbury area (Krogh 1994; Corfu and Easton 2000). Return to vehicles and continue north on Estaire Road. 82.5 km — Enter Wanup pluton 83.3 km — Pull over on right shoulder, outcrop opposite on west side of road. Stop 10 – Wanup pluton, flattened megacrystic granodiorite, Nepewassi domain UTM co-ordinates 512357E, 5135635N This roadcut exposes rocks of another megacrystic granite pluton in Nepewassi domain, the Wanup pluton. We are at the north end of the pluton at this stop, and overall, the Wanup pluton exhibits a more intense gneissic fabric than does the Estaire pluton that we saw at Stops 7 and 8. There are 2 rock types present in the roadcut, a matrix-rich, gneissic megacrystic diorite to granodiorite (Photo 10a, 10b) and a matrix-poor, megacrystic granodiorite to monzogranite (Photo 10c). As was the case for the Estaire pluton, it is likely that the Wanup pluton may be a Killarney intrusive suite body. Photo 10. A) weathered surface of potassium feldspar megacrystic gneiss granodiorite of the Wanup pluton from an outcrop 200m to the southeast of Stop 10. B) fresh surface of matrixpoor potassium feldspar megacrystic gneiss granodiorite of the Wanup pluton from Stop 10. C) fresh surface of matrix-rich potassium feldspar megacrystic gneiss granodiorite of the Wanup pluton from Stop 10. Scale card in all images is 10 cm long. Return to vehicles, continue north on Estaire Road. 84.7 km — Pull over on right shoulder, outcrop on east side. 134 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 11 – Nepewassi terrane boundary? UTM co-ordinates 511070E 5136160N This 100 m long, up to 3 m high, roadcut exposes a variety of tectonites related to the boundary between the Nepewassi domain to the south and the Grenville Front tectonic zone to the north. The south end of the roadcut consists of irregularly layered migmatitic paragneiss with boudinaged mafic layers (Photo 11). Continuing northward, is a transition into a zone of flattened grey granodioritic gneisses and amphibolite (Photo 12). Present within this zone are extremely deformed, porphyroclastic granitic pegmatite dikes that are parallel to the near-vertical gneissic fabric present throughout the roadcut (Photo 13). Photo 12. Flattened grey granodioritic gneiss (left) and flattened amphibolite (right) at Stop 11. Hammer handle is 33cm long. Near the north end of the roadcut, large potassium feldspar porphyroclasts are present locally in the tectonites (Photo 14). Isoclinal folds are also present in the tectonites (Photo 15). Return to vehicles, continue north on Estaire Road. Photo 13. Flattened, migmatitic amphibolite (left) and porphyroclastic granitic pegmatite dike (right) at Stop 11. Hammer handle is 33cm long. Photo 11. Irregularly layered migmatitic paragneiss with boudinaged mafic layers at the south end of Stop 11. Hammer handle is 33cm long. Photo 14. large potassium feldspar porphyroclast at Stop 11. Scale card is 10 cm long. 135 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Return to vehicles and continue north on Estaire Road. 87.7 km — Pull over on right shoulder (UTM 510020E, 5137675N), outcrop on right (east) side. Optional – mafic pod (migmatitic amphibolite) with straight gneiss on north side. 88.5 km — Pull over on right shoulder, outcrop on the right (east) side of Estaire Road. Stop 4.5 of Davidson, Carmichael and Pattison (1990), Stop 2-10 of Davidson (1995). Stop 13 – Garnet amphibolite and schists UTM co-ordinates 509410E 5139350N This approximately 80 m long roadcut consists of several different phases. The southern end of the roadcut consists of weakly migmatitic, garnet amphibolite (analysis 8, Table 6), which is in sharp contact with the pelitic and semi-pelitic schists that constitute the bulk of the outcrop. There are at least 3 schist units in the roadcut, from south to north these are. Photo 15. Folding of leucosome layer in grey, dioritic gneiss, north end of Stop 11. Scale card is 10 cm long.  Biotite-garnet schist with boudinaged quartz layers (quartz veins?) 85.0 km — Junction with 537, continue north on Estaire Road.  Muscovite-garnet schist with andesine porphyroblasts and intercalated with thin layers of semi-pelite to psammite (analysis 5, Table 6). Garnet [Alm75Adr0Grs7Prp14Sps5]. 86.6 km — Pull over on right, outcrop on right (east) side. Stop 12 – Migmatitic amphibolite  Muscovite-staurolite-garnet schist (analysis 6, Table 6) with andesine porphyroblasts. Staurolite is black on the weathered surface and is best seen on slabbed surfaces. Garnet [Alm76Adr1Grs6Prp9Sps5]. UTM co-ordinates 510020E 5137675N This roadcut is representative of much of the fine- to medium-grained amphibolite units with 1% to 5% thin, stringy feldspathic leucosome present in the GFTZ west of the Wanapitei River. This rock was described as “common amphibolite” by Kwak (1969). The affinity of these amphibolite units has not been firmly established. Both Lumbers (1975) and Dressler (1984) considered them as possible equivalents of the Nipissing intrusive suite, however, the presence of leucosome suggests that they might be older. Geochemical data (unpublished) obtained by the author indicates that the amphibolite bodies are more primitive than typical Nipissing intrusive suite rocks. The schist unit can be traced almost continuously over approximately 3.5 km and is near its widest extent at this stop. Only 900 m to the north, on Highway 69 (Photo 16), the unit is only about 15 m thick. Davidson, Carmichael and Pattison (1990) report that at this stop, in thin section, kyanite forms single crystals in contact with all the other minerals (garnet-biotite-muscovite-plagioclasequartz) and that it is commonly grown across large muscovite flakes. Sillimanite occurs as bundles of needles, generally associated with, or replacing, biotite, particularly at the edges of garnet 136 �Proceedings of the 68th ILSG Annual Meeting - Part 2 porphyroblasts. Kwak (1971) reported kyanite and potassium feldspar together as well. Figure 7 presents thermobarometric curves which indicate equilibration at ~8 kbar and 710°C. This is ~1 kbar and 40°C higher than the result obtained from a sample collected only 1.2 km to the north at the golf course, suggesting a thermal gradient of ~50°C/km (Davidson et al. 1990). Figure 7. TWEEQ plot showing 22 thermobarometric curves (5 independent) for pelitic schist with the main-stage assemblage garnet-kyanite-sillimanite-muscovite-biotiteplagioclase-quartz-ilmenite-rutile from a sample collected on the hill above Stop 13, as reported by Davidson, Carmichael and Pattison (1990). Photo 16. Muscovite-potassium feldspar schist on the east side of Highway 69, 900 m north of Stop 13 (analysis 7, Table 6). Schist is bounded to the south by garnet amphibolite (not in the photo) and to the north by irregularly layered grey, migmatitic, tectonite of intermediate composition (lower left of the photo). Rare earth element (REE) patterns for these rocks are inconclusive. The schist samples from this Stop have patterns (Figure 8) consistent with the post-Archean Australian shale composite (Taylor and McLennan 1985), but also with the pattern found in FIII rhyolites (Lesher et al. 1978) associated with volcanogenic massive sulphide (VMS) systems. It is tempting to think that these schists might be metamorphosed equivalents of the Huronian Supergroup, mostly likely the McKim Formation which has the bulk-rock major element composition capable of forming the observed mineral assemblages, most notable a highaluminum content (compare analyses 1-4, Table 6). Although the major element geochemistry supports this possibility, trace element data (analyses 5-7, Table 6) are less conclusive, as none of the samples analyzed have Ni/Co or Cr/Zn ratios typical of post-Archean fine-grained sedimentary rocks, such as the McKim Formation (between 1 and 2.5 and 1 and 1.6, respectively, Tang, Chen and Rudnick 2016). In fact, the schist sample from the site shown in Photo 16 (analysis 7, Table 6), is similar to that of the Nepewassi domain quartzite sample from Stop 9 rather than the McKim Formation. If these schists are part of the McKim Formation, this correlation is only possible if the McKim Formation has undergone extreme tectonic thinning (from approximately 1,000 m thick immediately north of the Grenville Front to approximately 100 m or less here. In addition, the adjacent Huronian Supergroup units, such as the Mississagi Formation, which are sandstonedominated, are nowhere to be seen. It could be that some of the neighbouring mafic host rocks, such as the garnet amphibolite at this stop, be metamorphosed Huronian Supergroup volcanic 137 �Proceedings of the 68th ILSG Annual Meeting - Part 2 rocks. Alternatively, as discussed at Stop 6b, the schists could represent metamorphosed hydrothermally altered rocks, similar to those present in VMS systems. Return to vehicles and continue north on Estaire Road. 88.5 km — Junction Estaire Road, Gladu Road, and Bentley Avenue. Park vehicles and walk to older roadcuts on the south side of Gladu Road. Stop 14 (Optional). Mylonitic rocks near the Grenville Front UTM co-ordinates 508520E 5141335N We are approximately 800 m south of the Grenville Front at Stop 14. The two old roadcuts south of Gladu Road consist of thin-layered, compositionally varied, straight gneisses, with some near vertical, gneissic fabric parallel, porphyroclastic granite pegmatite dikes. As we saw at Stop 5, protolith of these gneisses is not easily determined. Figure 8. Rare earth elements for selected samples normalized to the post-Archean Australian shale composite (Taylor and McLennan 1985). Schist samples from Stop 13 (open and filled triangles, analyses 5, 6, Table 6) and a McKim Formation sample from the Southern Province (filled diamonds) are straight lines centred around 1, suggesting that they could be metasedimentary rocks of the McKim Formation. Similarly, a sample of quartzite from Stop 9 (open diamonds, analysis 4, Table 6) parallels the composite, but at 1/10 the REE content, likely due to the abundance of quartz. Another schist sample (photo 16, open squares, analysis 7, Table 6) and the felsic rock from Stop 6B (open triangles, analysis 8, Table 4) have unusual concave heavy rare earth patterns, suggesting that they may have been affected by alteration. Return to vehicles and continue north on Estaire Road. 91.5 km — Optional. follow Bentley Avenue to turnaround at end for a view across the Grenville Front. Retrace route to Estaire Road, take Highway 69 back into town to Science North. End of road log. 138 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 6. Summary of geochemical from rocks at Stop 13 as well as average samples of the McKim Formation. Analysis 1 from Card, Innes and Debicki (1977, Table 8, p.40); analyses 2, 3, from Kwak (1968, 1971), with samples 3 from near Stop 13; analyses 4-8 from Easton (unpublished data, 2022). Analysis Number Sample Number 1 2 Average of Average of 8 Samples, 14 samples McKim Fm 3 4 5 6 7 8 232C 21RME102 Stop 9 513333 5134250 quartzite 21RME029 Stop 13 509443 5139333 schist 21RME030 Stop 13 509401 5139388 schist 21RME063 21RME027 Stop 13 509443 5139333 garnet amphibolite 0.0 >0.75 >0.75 0.12 >0.75 0.00 0.09 0.05 0.00 0.21 0.074 0.035 0.90 0.09 0.174 2.98 0.10 0.105 2.93 0.11 <0.03 1.93 0.06 0.131 0.79 95 47.7 39 5.6 11 3 <1.9 <1.3 0.8 4.8 82 9 37 3 2.97* 5.97* 1802 197.0 57.7 31 283 16 13.8 4.4 15.8 45.9 258 92 184 257 0.42 0.26 1968 196.8 68.7 30 261 15 11.2 3.2 13.8 34.5 200 81 150 166 0.49 0.22 1288 18.8 50.0 14 24 5 <1.9 <1.3 1.9 3.0 54 9 17 2 4.55* 1.21 123 17.4 116 16 268 9 <1.9 <1.3 3.5 21.9 60 127 221 106 n/a n/a 32 202 206 27 <27 Easting (m) Northing (m) Rock Name semi-pelite schist schist SiO2 (wt %) TiO2 Al2O3 Fe2O3total MnO MgO CaO Na2O K2O P2O5 CO2 S LOI/H2O+ 57.90 0.83 21.94 8.10 0.06 2.93 1.01 1.29 2.77 0.12 0.11 nr 3.20 59.20 1.11 20.75 8.54 0.08 2.90 1.34 1.49 3.25 0.15 nr nr 1.56 60.95 1.23 19.90 5.48 0.05 2.60 1.49 2.27 3.74 0.10 nr nr 1.94 Ba (ppm) Rb Sr Ga V Pb Th U Nb Y Zr Ni Cr Cu Ni/Co Cr/Zn Au (ppb) Pd Pt TREE 509866 5140422 schist Notes: Major element oxides are in weight %; trace element data are in parts per million, except for Au, Pd, Pt which are in parts per billion. * indicates Ni/Co or Cr/Zn with Archean ratios. Abbreviations: LOI = loss-on-ignition; n/a = not applicable; nr, not reported; TREE, total rare earth element. 139 �Proceedings of the 68th ILSG Annual Meeting - Part 2 References ——— 1992. Circa 1.75 Ga ages for plutonic rocks from the Southern Province and adjacent Grenville Province: what is the expression of the Penokean orogeny?: Discussion; in Radiogenic Age and Isotopic Studies: Report 6; Geological Survey of Canada Paper 92-2, p.227-228. Aldis, C.H. 2016. Geochronology and geochemistry of Mesoproterozoic granitoids in the Nepewassi domain, Grenville Orogen; unpublished BSc thesis, University of Waterloo, Waterloo, Ontario, 82p. Ames, D.E., Davidson, A., Buckle, J.L. and Card, K.D. 2005. Geology, Sudbury bedrock compilation, Ontario; Geological Survey of Canada, Open File 4570, scale 1:50 000. Card, K.D. and Lumbers, S.B. 1975. Sudbury–Cobalt; Ontario Geological Survey, Map 2361, scale 1:253 440. 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Clough, C.E. and Hamilton, M.A. 2017. Matachewan LIP revisited: a revised, high-resolution U-Pb age for the East Bull Lake intrusion and associated units; Geological Association of Canada—Mineralogical Association of Canada, Abstracts, v.40, p.65. Buchan, K.L. and Ernst, R.E. 1994. Onaping fault system: age constraints on deformation of the Kapuskasing structural zone and units underlying the Sudbury structure; Canadian Journal of Earth Sciences, v. 21, p.1197-1205. Corfu, F. and Andrews, A.J. 1986. A U-Pb age for mineralized Nipissing diabase, Gowganda, Ontario; Canadian Journal of Earth Sciences, v.23, p.107109. Buckley, S.B., Easton, R.M., Ford, F.D., and Lobanok, L. 1997. Analysis of P-T conditions along the Grenville Front east of Sudbury, Ontario; Geological Association of Canada-Mineralogical Association of Canada-Canadian Geophysical Union, Program with Abstracts, v.22, p.A-20. Corfu, F. and Easton, R.M. 2000. U-Pb evidence for polymetamorphic history of Huronian rocks underlying the Grenville Front Tectonic Zone east of Sudbury, Ontario; Chemical Geology, v.172, p.149171. Davey, S., Bleeker, W., Kamo, S[.L.], Davis, D.[W], Easton, M.[R.] and Sutcliffe, R.H. 2019. Ni-Cu-PGE potential of the Nipissing sills as part of the ca. 2.2 Ga Ungava large igneous province; in Targeted Geoscience Initiative: 2018 report of activities; Card, K.D. 1978. Geology of the Sudbury-Manitoulin area, districts of Sudbury and Manitoulin; Ontario Geological Survey, Report 166, 238p. 140 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Geological Survey of Canada, Open File 8549, p.403-419. zircon: Application to crystallization of the Sudbury impact melt sheet; Geology, v.36, p.383-386. Davidson, A. 1995. Tectonic history of the Grenville Province, Ontario, Precambrian’95 Field Trip Guidebook A5; Geological Survey of Canada, Open File 3142, 149p. Debicki, R.L. 1990. Stratigraphy, paleoenvironment and economic potential of the Huronian Supergroup in the southern Cobalt embayment; Ontario Geological Survey, Miscellaneous Paper 148, 154p. ——— 1997. New information on the Grenville Front near Sudbury, 43rd Annual Institute on Lake Superior Geology, Proceedings, v.43, pt.3, 38p. Dickin, A.P. 1998a. Nd isotope mapping of a cryptic continental suture, Grenville Province of Ontario; Precambrian Research, v.91, p.433-44. ——— 1998. Questions of correlation across the Grenville Front east of Sudbury, Ontario, in Current Research, 1998-C, Geological Survey of Canada, p.145–154. ——— 1998b. Pb isotope mapping of differentially uplifted Archean basement: a case study from the Grenville Province, Ontario; Precambrian Research, v.91, p.445-454. Davidson, A. and Ketchum, J.W.F. 1993. Grenville Front studies in the Sudbury region, Ontario; in Current Research Part C, Geological Survey of Canada, Paper 93-1C, p.271-278. Dickin, A.P. and McNutt, R.H. 1989. Nd model age mapping of the southeast margin of the Archean foreland in the Grenville Province of Ontario; Geology, v.17, p.299-302. Davidson, A., and van Breemen, O. 1994. U-Pb ages of granites near the Grenville Front, Ontario, in Radiogenic age and isotopic studies: report 8, Geological Survey of Canada, Current Research 1994-F, p.107–114. Dressler, B.O. 1984. Sudbury geological compilation; Ontario Geological Survey, Map 2491, scale 1:50 000. Dressler, B.O., Gupta, V.K. and Muir, T.L. 1991. The Sudbury Structure; in Geology of Ontario, Chapter 15, Ontario Geological Survey, Special Volume 4, pt.1, p.593-625. Davidson, A., Easton, R.M., Corriveau, L. and Martignole, J. 2002. Transect of the southwestern Grenville Province; Geological Association of Canada, Saskatoon’02, Fieldtrip B6 Guidebook, 114p. Dudas, F.O., Davidson, A. and Bethune, K.M. 1994. Age of the Sudbury diabase dykes and their metamorphism in the Grenville Province, Ontario; in Radiogenic Age and Isotopic Studies: Report 8; Geological Survey of Canada, Paper 94-F, p.97-106. Davidson, A., Carmichael, D.M. and Pattison, D.R.M. 1990. Metamorphism and Geodynamics of the southwestern Grenville Province, Ontario; IGCP Project 235-304, Field Trip #1 Guidebook, 123p. Easton, R.M. 1992. The Grenville Province; in Geology of Ontario, Chapter 19, Ontario Geological Survey, Special Volume 4, pt.2, p.713-904. Davidson, A., Culshaw, N.G. and Nadeau, L. 1982. A tectono-metamorphic framework for part of the Grenville Province, Parry Sound region, Ontario; in Current Research, Part A, Geological Survey of Canada, Paper 82-1A, p.175-190. ——— 1996. Geology of garnetiferous gneisses in Street Township, District of Sudbury; in Summary of Field Work and Other Activities, 1996, Ontario Geological Survey, Miscellaneous Paper 166, p.7073. Davidson, A., van Breemen, O. and Sullivan, R.W. 1992. Circa 1.75 Ga ages for plutonic rocks from the Southern Province and adjacent Grenville Province: what is the expression of the Penokean orogeny?; in Radiogenic Age and Isotopic Studies: Report 6; Geological Survey of Canada Paper 92-2, p. 107118. ——— 1998. New observations related to the mineral potential of the Southern Province and the Grenville Front tectonic zone east of Sudbury; Ontario Geological Survey, Open File Report 5976, 28p. ——— 2000. Variation in crustal level and large-scale tectonic controls on rare-metal and platinum-group element mineralization in the Southern and Grenville provinces; in Summary of Fieldwork and Davis, D.W. 2008. Sub-million-year age resolution of Precambrian igneous events by thermal extraction– thermal ionization mass spectrometer Pb dating of 141 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Ercit, T.S. 1999. 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U-Pb age constraints on Early Proterozoic mafic magmatism from the southern Superior and western Grenville provinces, Ontario; in Radiogenic Age and Isotopic Studies, Report 6, Geological Survey of Canada, Paper 92-2, p.97-106 Sawyer, E.W. 2008. Atlas of migmatites; The Canadian Mineralogist, Special Publication 9, NRC Research Press, Ottawa, 371p. ——— 1993. An isotopic, geochemical and petrographic investigation of the genesis of early Proterozoic mafic intrusions and associated volcanics near Sudbury, Ontario; unpublished PhD. thesis, University of Alberta, Edmonton, Alberta, 223p. Schandl, E.S., Gorton, M.P. and Davis, D.W. 1994. Albitization at 1700+/-2 Ma in the SudburyWanapitei Lake area, Ontario; Canadian Journal of Earth Sciences, v.31, p.597-607. ——— 2004. Basement tracing using Mid-Proterozoic anorthosites straddling a palaeoterrane boundary, Ontario, Canada; Precambrian Research, v.129, p.164-184. Smith, M.D. 2002. The timing and petrogenesis of the Creighton pluton, Ontario: an example of felsic magmatism associated with Matachewan igneous events; unpublished MSc thesis, University of Alberta, Edmonton, Alberta, 123p. Pye, E.G., Naldrett, A.J., and Giblin, P.E, eds. 1984. Geology and Ore Deposits of the Sudbury Structure; Ontario Geological Survey Special Volume 1, 603p. Spray, J., Federowich, J.S. and Kontak, D.J. 2007. An introduction to the geology of the Sudbury impact structure; Intra-Meeting Field Trip, NUNA 2007 Meeting, Sudbury, Ontario, 38p. Raharimahefa, T., Lafrance, B. and Tinkman, D.K. 2014. New structural, metamorphic, and U-Pb geochronological constraints on the Blezardian Orogeny and the Yavapai Orogeny in the Southern Province, Sudbury, Canada; Canadian Journal of Earth Sciences, v.51, p.750-774. Streckeisen, A. 1976. To each plutonic rock its proper name; Earth-Science Reviews, v.12, p. 1-33. Sullivan, R.W. and Davidson, A. 1993. 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Abbreviations of names for rock-forming minerals; American Mineralogist, v.95, p.185-187. 146 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Field Trip 3 – Magmatism and Brecciation in the Footwall Rocks of the Southwestern Sudbury Structure Caroline Gordon Ontario Geological Survey, Sudbury, Ontario P3E 6B5 Carol-Anne Généreux Mineral Exploration Research Centre, Harquail School of Earth Sciences, Laurentian University, Sudbury, Ontario P3E 2C6, Terrane Geoscience Inc., Canada Brad Clarke SPC Nickel Corp., Sudbury, Ontario P3E 5P5 Introduction 2019) and the Ramsey–Algoma granitoid complex (Card 1979), which includes the Cartier (2642 Ma: Meldrum et al. 1997) and Birch Lake (2651 Ma: Kamo 2006; Easton and Heaman 2008; Gordon, et al. 2018a) batholiths (Figure 2). The southern boundary of the Superior Province is overlain unconformably by the supracrustal rocks of the Huronian Supergroup. The Huronian Supergroup was deposited in a continental rift and on a continental platform between 2450 and 2219 million-years ago (Krogh et al. 1984; Bennett et al. 1991) and has been interpreted to represent a partial Wilson cycle (Young 1983; Hoffman 1989; Bennett al. 1991; Young et al. 2001). This one-day field trip presents geological highlights from the Ontario Geological Survey (OGS) Southwest Sudbury Structure bedrock mapping project. This project is part of a collaborative program with the OGS, the Mineral Exploration Research Centre at the Harquail School of Earth Sciences, Laurentian University, and the private sector. Regional Geology The Sudbury Igneous Complex (SIC) is interpreted to represent a melt sheet produced by the impact of a meteorite at 1850 Ma (Dietz 1964; Krogh et al. 1984; Davis 2008). The SIC is part of the Southern Province of the Canadian Shield and is located north of the Grenville Front astride the southern contact of the Archean Superior Province (Figure 1). The term Sudbury Structure, which is used throughout this field guide, refers to the SIC, the Sudbury Basin that contains rocks of the Whitewater Group, and the outer zone of brecciated footwall rocks. The Sudbury Structure is geographically subdivided into north, south and east ranges (Figure 2). The Sudbury area has been intruded by numerous dikes, sills and plutons of various ages (Figure 2). Known intrusions and intrusive suites include, in chronological order, the 1) Joe Lake gabbro (2660 Ma: Bleeker et al. 2015); 2) Matachewan dike swarm (2480-2460 Ma: Heaman 1997; Bleeker et al. 2012); 3) Drury Township, Falconbridge and Frood intrusions of the East Bull Lake Intrusive Suite (2480 Ma: Krogh et al. 1984; James et al. 2002; Keays and Lightfoot 2020); 4) Creighton and Murray plutons (2460 Ma: Bleeker et al. 2015); 5) Nipissing Intrusive Suite (2219 to 2210 Ma: Davey et al. 2019; Noble and Lightfoot 1992; Corfu and Andrews 1986: Bleeker et al. 2015); 6) Trap dike swarm (1750 Ma: Bleeker et al. 2015); 7) Sudbury dike swarm (1238 Ma: Krogh et al. 1987), and; 8) Grenville dike swarm (590 Ma: Kamo et al. 1995). In the Sudbury area, Archean rocks of the Superior Province are part of the Abitibi Subprovince and consist of the Levack Gneiss Complex (2711 to 2647 Ma: Krogh et al. 1984; Wodicka and Card 1995), the Benny greenstone belt (2680-2700 Ma: Ontario Geological Survey 147 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 1. Sketch map showing the regional setting of the Sudbury Igneous Complex and the Huronian Supergroup (modified from Young et al. 2001). The Sudbury area has been affected by several episodes of deformation and metamorphism. Regional metamorphism is thought to have reached mid-greenschist to lower-amphibolite facies and generally increases to the south (Card 1978; Card et al. 1984; Fox 1971). Country rocks adjacent to the SIC were thermally metamorphosed but have since been overprinted by regional metamorphism (Dressler et al. 1991; Jørgensen et al. 2019; Généreux et al. 2021). Ductile deformation of the Southern Province has historical been interpreted to have started prior to or concurrent with emplacement of the Nipissing Intrusive Suite during the Blezardian Orogeny (2415-2219 Ma; Raharimahefa et al. 2014; Stockwell 1982), and continued during the Penokean Orogeny (18901830 Ma; Dressler 1984a; Bennett et al. 1991), the Yavapai and Mazatzal orogenies (1770-1600 Ma; Bailey et al. 2004; Raharimahefa et al. 2014; Papapavlou et al. 2017), and the Grenville Orogeny (1120-980 Ma; Carr et al. 2000). Recent studies have attributed most of the deformation in the Sudbury area to the Yavapai–Mazatzal orogenies (Bailey et al. 2004; Raharimahefa et al. 2014; Papapavlou et al. 2017). The lower age limit of ductile deformation is constrained by the age of the undeformed Sudbury dike swarm (1238±4 Ma; Krogh et al. 1987). Three types of ore environments are associated with the SIC: 1) contact deposits, which are hosted in depressions at the base of the SIC; 2) offset deposits, which occur in quartz diorite dikes that extend for several kilometres from the SIC into the footwall rocks, and; 3) footwall deposits, which are found in the brecciated footwall rocks directly underlying the SIC (cf. Lightfoot 2017). It is generally accepted that nickel-copper-platinum group element (Ni-Cu-PGE) deposits in Sudbury are primarily magmatic and formed by differentiation of a sulphide melt during crystallization of the SIC (Keays and Crocket 1970 148 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 2: Geological map of the Sudbury area (modified from Ames and Farrow 2007). The location of Drury and Denison townships are shown in the lower left of the figure. 149 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 1. Mineral deposits in Drury and Denison townships. Table includes names for mines and prospects. Occurrences and discretionary occurrences are not included here. Reference numbers in table correspond with mineral occurrence symbol on Figure 3. Naldrett et al. 1994; Naldrett 1999; Ames and Farrow 2007), with some involvement of hydrothermal fluids (Farrow et al. 1994; Jago et al. 1994; Morrison et al. 1994; Molnár and Watkinson 2001; Péntek et al. 2008). Many deposits in the South Range of the SIC were modified during post 1850-Ma tectonic events (Lightfoot 2017). Mining and Exploration History The southwestern portion of the Sudbury Structure has been an important mining and exploration area since the discovery of Ni-Cu-PGE mineralization related to the SIC. The first deposit discovered in the area was the Worthington deposit in 1884 (cf. Lightfoot 2017). Since then, the southwestern Sudbury Structure has supported numerous past-producing mines and currently hosts one active mine (Totten Mine), advanced prospects (Victoria Project) and undeveloped mineral occurrences (Figure 3; Table 1). The majority of mineral deposits are associated with SIC-related rocks and occur as contact deposits or as offset deposits within the Worthington and Vermilion offset dikes. Commodities include NiCu-PGE, gold (Au), cobalt (Co), silver (Ag), iron (Fe), tin (Sn), selenium (Se), tellurium (Te) and arsenic (As) (Table 1). Several contact and offset deposits were modified by post-SIC deformation and ore is now hosted within shear zones (i.e. Vermilion and Chicago mines). In addition to SIC-related mineralization, known prospects, occurrences and discretionary occurrences include: 1) Ni-Cu±PGE mineralization hosted by the Nipissing gabbro, volcanic and sedimentary rocks of the Huronian Supergroup, and within shear zones; 2) Au-Cu bearing quartz veins; and, 3) uranium-thorium (U-Th) mineralization in pyritic quartz-pebble-rich conglomeratic arenites in the lower Matinenda Formation of the Huronian Supergroup (Figure 3). Quartz veins and quartzites in the area have also been quarried for silica. Ref No. Name Commodity* (Primary/Secondary) 1 Totten Mine Ni, Cu / PGE, Co, Au, Ag 2 Totten #1 Ni, Cu / PGE, Co, Au, Ag 3 Worthington Mine Ni, Cu, PGE / Co, Au, Ag 4 Worthington #2 Ni, Cu, PGE / Co, Au, Ag 5 Howland Pit Ni, Cu / PGE, Co 6 Robinson Mine Ni, Cu / PGE, Co 7 Aer Mine (Rosen and Gersdorffite mines) Ni, Cu / PGE, Co 8 Victoria Mine Ni, Cu / PGE, Au 9 Vermilion Mine Ni, Cu / PGE, Au, Ag, Co, Fe, Sn, Se, Te, As 10 Crean Hill Mine Ni, Cu, PGE / Au, Co, Fe, Ag, Se, Te 11 Lockerby Mine Ni, Cu, Co / PGE, Au, Ag 12 Ellen Pit Ni, Cu / Au, Ag, PGE, Fe, Co, Se, Te 13 Chicago Mine Ni, Cu / PGE, Au, Ag, Co, Fe, Se, Te 14 Sultana Nickel Mine Ni, Cu 15 Delta Occurrence Ni, Cu / PGE, Au 16 McIntyre Mine Ni, Cu 17 Victoria Project Ni, Cu, PGE 18 Alanaen and Maki West (Kerr Addison Prospect) U, Th, Cu *Primary and secondary commodities as listed in the Ontario Mineral Inventory Database (Ontario Geological Survey 2022). 150 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3. Geological map of Drury and Denison townships (after Gordon et al. 2018a; Gordon and Généreux 2017; Gordon 2018). Mineral occurrences are from Ontario Geological Survey (2022) and Gordon et al. (2018a). Universal Transverse Mercator coordinates are in North American Datum 1983, Zone 17. CCF = Cameron Creek Fault, VLF = Vermilion Lake Fault, FLF = Flack Lake Fault, CAF = Chicago Fault, CF = Creighton Fault, VF = Victoria Fault, CHF = Crean Hill Fault, MF = Murray Fault, WRTH-W = Worthington Offset dike, western limb, WRTH–E = Worthington Offset dike, eastern limb, VO = Vermilion Offset dike, CVDZ = Creighton-Victoria deformation zone. 151 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Geological Overview of the Southwestern Sudbury Structure In the southwestern Sudbury Structure, the magmatic breccia along the base of the SIC (Figure 3) has been subdivided into 4 types based on matrix composition and clast abundance (cf. Gordon et al. 2018a). Geological mapping in the southwest Sudbury Structure was completed between 2015 and 2018 and focussed on Drury and Denison townships (Figure 3) located approximately 50 km west of the City of Greater Sudbury (Figure 2). 1. Clast-rich Sublayer Norite (>30% clasts): heterolithic breccia with abundant gabbroic and lesser amounts of granitoid clasts in a noritic to leuconoritic matrix. For sake of clarity, the authors have omitted the prefix “meta” for the rock names in this field guide (e.g. gabbro versus metagabbro). 2. Clast-poor Sublayer Norite (<30% clasts): heterolithic breccia with a noritic matrix. Sudbury Igneous Complex (SIC) 3. Sublayer Granite Breccia: Heterolithic breccia (>35% clasts) with abundant granitoid clasts and few mafic clasts in a pink-weathering matrix. The SIC is elliptical in shape and approximately 30 km x 60 km in size (Figure 2). It is made-up of three main components: 1) Main Mass, which is a differentiated igneous body; 2) Contact Sublayer, a basal magmatic breccia; and, 3) Offset dikes, quartz diorite dikes emplaced in the footwall of the SIC (Giblin 1984; Dressler et al. 1991; Ames et al. 1997, 1998, 2002). All three components are exposed in the southwest Sudbury Structure. 4. Heterolithic breccia with gabbroic and anorthositic gabbro clasts in a variably textured, leucogabbroic matrix. Clast-rich and clast-poor Sublayer Norite occurs at the SIC–Archean granitoid and SIC–Huronian Supergroup contacts. The Sublayer Granite Breccia unit, which likely represents a variation of the classic Sublayer Granite Breccia, occurs exclusively at the SIC–Archean granitoid contact. Breccia type 4 occurs exclusively at the SIC–Drury Township intrusion contact and, although it has been tentatively classified as sublayer, it may represent Footwall Breccia. Main Mass The SIC, as exposed in the southwestern Sudbury Structure, exhibits the complete Main Mass stratigraphic sequence (Figure 3). From top to bottom, it consists of: 1. Granophyre: leucocratic monzogranite and upper plagioclase-rich phase of granodiorite. Offset Dikes 2. Transition zone quartz gabbro: melanocratic to mesocratic quartz gabbro and monzogabbro with cumulus magnetite and apatite. Offset dikes of the SIC, also known as Offset Sublayer, are radial, concentric, and discontinuous segmented bodies that were emplaced into the footwall of the SIC (cf. Lightfoot 2017). Offset dikes consist of two phases: quartz diorite (QD) and a sulphide-enriched inclusion-bearing quartz diorite (IQD) (Grant and Bite 1984). 3. South Range norite: leucocratic to mesocratic quartz monzogabbro and norite. Contact Sublayer The main mass of the SIC overlies an extensive zone of magmatic breccia known as the Contact Sublayer. The Sublayer is heterolithic in matrix composition and clast type, variably gossanous and laterally discontinuous. It is generally classified into two groups based on the composition of its matrix: Sublayer Norite consists of breccia with a noritic matrix, whereas Sublayer Granite Breccia contains a granitic matrix (cf. Lightfoot 2017). The Worthington Offset is a branching radial offset dike that trends southwest from the SIC contact (Figure 3). The proximal part, occupying a possible embayment in the Victoria Mine area north of Ethel Lake, occurs as numerous faulted segments (Grant and Bite 1984). Southward from the SIC, the offset dike narrows and then broadens before bifurcating into eastern and western limbs. 152 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The eastern limb tapers gradually over a distance of 1500 m at the surface, trends southeast, and broadens at depth (Grant and Bite 1984). The western limb extends to the southwest for at least 15 km, with a thickness between 30 and 100 m. The dike truncates volcanic and sedimentary rocks of the Huronian Supergroup as well as Nipissing gabbro and is crosscut by the Creighton Fault (Figure 3). The Worthington Offset contains both the inclusion-poor QD and mineralized IQD phases. The contact between phases is sharp and inclusions of QD are present within IQD. formed by cataclasis (Lafrance and Kamber 2010; O'Callaghan et al. 2016; Rousell et al. 2003) and/or frictional melting of the target rocks during cratering (Dressler 1984a; Lafrance and Kamber 2010; O'Callaghan et al. 2016; Rousell et al. 2003; Thompson and Spray 1994). Within the southwestern Sudbury Structure, Sudbury Breccia occurs as fine-grained, dark green to greenish-yellow veins, dikes and irregular shaped bodies of various thicknesses and orientations. Thicker breccia veins, typically several metres in width, are usually clast-rich and form corridors along major lithologic contacts and structures. Two major heterolithic Sudbury Breccia belts were identified in Drury Township: 1) at the contact between the Drury Township intrusion and the Archean granitoid, and 2) following the folded contacts between Huronian sedimentary rocks and Nipissing gabbro in south-central Drury Township (Figure 3). These breccia belts range from a few metres to several hundred metres wide. In the belts, breccia matrix typically makes-up more than 40% of the outcrops. The Vermilion Offset dike crops out east of Ethel Lake, south of Crean Hill (Figure 3). The Vermilion Offset occurs within a 200 m long, northwest-striking zone of discontinuous QD and IQD pods at the contact of Sudbury Breccia and sedimentary and volcanic rocks of the Stobie Formation (Grant and Bite 1984; Szentpeteri et al. 2003). Breccias in the Footwall Rocks Two types of breccia occur within the footwall of the SIC: Footwall Breccia and Sudbury Breccia. Footwall rocks of the SIC Footwall Breccia Archean Ramsey-Algoma granitoid complex Footwall Breccia is a parautochthonous breccia that occurs in discontinuous lenses and sheets between the Contact Sublayer and underlying footwall rocks. It consists mainly of brecciated and partially melted footwall rocks, and has an igneous to granoblastic matrix (Lakomy 1990; McCormick et al. 2002). Within the southwestern Sudbury Structure, Footwall Breccia has been identified at the Crean Hill Mine area (Figure 3) (Généreux et al. 2021), where it consists of breccia dikes and pods of various compositions hosted in volcanic rocks of the Stobie Formation. Monzogranite, granite and granodiorite belonging to the Archean Ramsey–Algoma granitoid complex are the oldest rocks in the southwestern Sudbury Structure (Figure 3). The contact between the Archean basement and rocks of the Huronian Supergroup is known to be unconformable (Stockwell 1964; Card 1990) but is highly sheared in the Sudbury area. The transition between the Cartier and Birch Lake batholiths of the Ramsey-Algoma granitoid complex is reportedly within Drury Township, but the exact location of the contact has not been determined (Tolman 1929; Meldrum et al. 1997). U/Pb zircon geochronology on a granitoid sample collected in western Drury Township yielded a minimum age of 2645±1 Ma (Gordon et al. 2018a), which suggests that the granitoid rocks in Drury Township are older than the Cartier batholith and more similar in age to the Birch Lake batholith. Sudbury Breccia All rock types in the Sudbury area that are older than circa 1850 Ma contain various amounts of impact breccia, locally called Sudbury Breccia. It is a parautochthonous breccia with an aphanitic to microcrystalline matrix, and occurs as discontinuous veins or tabular bodies within the footwall rocks. Sudbury Breccia is thought to have 153 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Elliot Lake Group The Elliot Lake Group is the lowermost unit in the Huronian Supergroup and the only unit to contain volcanic rocks. The basal volcanic rocks are interpreted as fissure eruptions related to deeppenetrating crustal faults during rifting (Card 1978). In the Sudbury area, the volcanic rocks are up to 3 km thick and are subdivided into the Elsie Mountain, Stobie and Copper Cliff Formations (Card 1978) (Figure 4). The felsic volcanic rocks of the Copper Cliff Formation are not exposed in Drury and Denison townships and will not be discussed further in this guide. Paleoproterozoic Huronian Supergroup The Huronian Supergroup extends from the east shore of Lake Superior across the north shore of Lake Huron, and northeastward across the Cobalt Embayment to the Noranda area in northwestern Quebec (Bennett et al. 1991) (Figure 1). It consists of a southward thickening, up to 12 km thick, package of sedimentary and volcanic rocks that are subdivided, from oldest to youngest, into the Elliot Lake, Hough Lake, Quirke Lake and Cobalt groups (Bennett et al. 1991; Robertson et al. 1969). The minimum age of the Huronian Supergroup is constrained by the age of the intruded Nipissing Intrusive Suite (2210-2219 Ma: Davey et al. 2019; Corfu and Andrews 1986; Noble and Lightfoot 1992; Bleeker et al. 2015). Its maximum depositional age is constrained by the age of the felsic volcanic rocks of the Copper Cliff Formation (2452-2460 Ma: Krogh et al. 1984; Ketchum et al. 2013; Bleeker et al. 2015). The Huronian Supergroup volcanic rocks in Drury and Denison townships exhibit significant lateral variation from east to west and will be discussed as 3 separate segments: western, central and eastern. In the western segment, the majority of rocks previously identified as volcanic have been reclassified as mylonites of the CreightonVictoria deformation zone (Figure 3) (Gordon et al. 2015; Simard et al. 2016; Généreux et al. 2016; Généreux et al. 2017; Gordon et al. 2018a). In the central segment, which is east of the CreightonVictoria deformation zone, volcanic rocks of the Elsie Mountain Formation are dominated by massive and pillowed basaltic flows that are locally amygdaloidal and porphyritic, with minor amounts of intercalated arenite and siltstone. South of the Elsie Mountain Formation, bimodal volcanic rocks of the Stobie Formation are intercalated with arenite of the Matinenda Formation (Figure 3). In this area, the Stobie Formation consists predominantly of massive and pillowed basaltic to rhyolitic flows with interbedded arenite, siltstone, and minor amounts of pyroclastic rocks. The central and eastern segments are separated by an unnamed northwest-trending fault in Denison Township (Figure 3). The eastern segment is dominated by the bimodal Stobie Formation, which is bound to the north and east by the SIC and Creighton pluton, respectively. The Elsie Mountain Formation is largely absent in the eastern segment, except for a thin sliver of basalt exposed adjacent to the Creighton pluton. The oldest and lowermost Elliot Lake Group consists of an intercalated sequence of sandstone, conglomerate, siltstone, mudstone, and local volcanic rocks (Card 1978). With the exception of the carbonate-bearing Serpent Formation (Quirke Lake Group), the overlying Hough Lake, Quirke Lake and Cobalt groups each contain cyclical repeating sequences of lower conglomeratic units, middle siltstone and mudstone units, and upper sandstone units (Roscoe 1969). Rocks of the lower Huronian Supergroup, specifically the Elliot Lake and Hough Lake groups, are represented in the southwestern Sudbury Structure. In the Sudbury area, most of the Huronian Supergroup strata are subvertical and are approximately west-northwest- to east-trending. Reversals of facing direction define the synclines, anticlines and thrust faults in the area. Despite folding and faulting within formations, the overall younging direction of the stratigraphy is southward. Unit thicknesses stated herein are apparent thicknesses. 154 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4. Generalized stratigraphy and facies relationships of the Elliot Lake Group in the Southwest Sudbury Structure (modified from Card 1978). The Matinenda Formation, which hosts the uranium-rich pyritic quartz pebble conglomerates in the Elliot Lake area (Figure 1), is interpreted to have been deposited in a braided fluvial environment (Fralick and Miall 1989). In the southwestern Sudbury Structure, the Matinenda Formation is up to 1-km thick and thins eastward where it eventually disappears from the Stobie– McKim Formations contact but crops out as discontinuous layers within the Stobie Formation (Figure 3). The Matinenda Formation consists of subfeldspathic arenite and quartz arenite with quartz-pebble conglomeratic beds. The quartzpebble conglomeratic beds locally contain pyrite and elevated concentrations of U and Th. The arenites, which constitute the bulk of the formation, are massive to crudely bedded, locally displaying graded beds and cross-bedding. The McKim Formation, which is interpreted to represent a marine transgression that gradually drowned the Matinenda fluvial plain (Fralick and Maill 1989), is one of the most aerially extensive Huronian Supergroup units in the southwestern Sudbury Structure. The preserved sequence was significantly thickened (up to 1.5 km) by folding (Figure 3). The McKim Formation consists of interbedded mudstone, siltstone and minor sandstone. The amount of interbedded sandstone increases significantly eastward through Denison Township. The sandy and silty turbidites are 155 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Proterozoic Intrusive Rocks typically thickly laminated to thinly bedded and commonly display cross-bedding, graded beds, ripples and scour marks. Staurolite, chloritoid and rutile porphyroblasts are present, predominantly in the muddier beds. Drury Township Intrusion, Matachewan Dike Swarm and Huronian Synvolcanic Intrusions The Drury Township intrusion, Matachewan dike swarm, and volcanic rocks of the Huronian Supergroup were all emplaced during continental rifting associated with the Matachewan Igneous Event (Heaman 1997). The Drury Township intrusion and Matachewan dike swarm are also tentatively genetically linked to eruption of the basal Huronian Supergroup mafic volcanic rocks (Fahrig 1987; Vogel et al. 1998). Hough Lake Group The Hough Lake Group contains sedimentary rocks of the Ramsay Lake, Pecors and Mississagi Formations. It is the lowest of the 3 groups that display the cyclical repetition of conglomerate— siltstone-mudstone—–sandstone. Each cycle is interpreted to represent a sequence of glaciogenic —marine—fluvial and/or shallow marine deposition (Roscoe 1969; Robertson 1976; Fralick and Miall 1989). The Drury Township intrusion is interpreted as one of several leucogabbro-anorthosite sills of the East Bull Lake Intrusive Suite (Prevec 1993; Prevec and Baadsgaard 2005). Intrusions of the East Bull Lake Suite were emplaced between ~2491 and 2475 Ma and occur in a discontinuous east-northeast-trending belt along the Archean– Proterozoic contact between Elliot Lake and the Ottawa River (Krogh et al. 1984; James et al. 2002; Bleeker et al. 2012; Bleeker et al. 2015). The Drury Township intrusion is up to 2 km thick and is exposed at the Ramsey–Algoma granitoid complex – Huronian Supergroup – SIC contact in Drury Township (Figure 3). The intrusion is composed of medium- to coarse-grained, locally pegmatitic, vari-textured anorthositic gabbro with a marginal gabbroic phase. The Ramsay Lake Formation is up to 300 m thick and consists of 2 conglomerate units: the “beige” and “grey” members, and, an upper sandstone unit, the “sandy” member (Gordon et al. 2018a). The “beige” member is a bilithic conglomeratic subfeldspathic arenite with a beige, quartz-rich matrix and clasts of granite and quartz. The “grey” member is consistent with the classic description of the Ramsay Lake Formation (cf. Young 1991; Bennett et al. 1991). It is a heterolithic conglomeratic wacke or sandstone with a grey, quartz-rich matrix. The “sandy” member is composed of crudely bedded, subfeldspathic arenite, wacke and quartz arenite. The Pecors Formation is 100 to 300 m thick and exposed between the Ramsay Lake and Mississagi Formations (Figure 3). The formation consists of siltstone and mudstone. The siltstone and mudstone are thickly laminated and locally exhibit cross-beds and load structures. The Matachewan dike swarm is an extensive radial dike swarm consisting of north- and northwest-trending mafic dikes, which intruded granitoids of the Superior Province and crop out over 300,000 km2 in Ontario and southwestern Quebec (Halls and Bates 1990). The Matachewan dike swarm was emplaced in 2 main pulses. The first, earlier pulse at circa 2480 Ma is believed to have been coincident with emplacement of the East Bull Lake Intrusive Suite (Krogh et al. 1984; James et al. 2002; Bleeker et al. 2012; Bleeker et al. 2015). The second, and “main pulse” of the Matachewan dike swarm. occurred at circa 2460 Ma (Heaman 1997; Bleeker et al. 2012; Bleeker et al. 2015). In the southwestern Sudbury Structure, northwest- and northeast-trending Matachewan The Mississagi Formation is at least 2 km thick, including thickening by folding (Figure 3). Sandstones of the Mississagi Formation consist of well-sorted, fine- to medium-grained subfieldspathic arenite and quartz arenite. They are thinly to thickly bedded, locally contain beds of siltstone, and commonly display cross-beds. 156 �Proceedings of the 68th ILSG Annual Meeting - Part 2 dikes crosscut the Drury Township intrusion and the Archean Ramsey-Algoma granitoid complex (Figure 3). These mafic dikes are fine- to mediumgrained and locally plagioclase-phyric. The northeast trend for the Matachewan dike swarm in the Sudbury area is atypical; it is possible that their trend represents a small-scale, concentric dike swarm (Gordon et al. 2018a). emplaced between 2219 to 2210 Ma (Davey et al. 2019; Corfu and Andrews 1986; Noble and Lightfoot 1992; Bleeker et al. 2015) as part of the Ungava large igneous province during continental rifting (Ernst and Bleeker 2010; Davey et al. 2019). In the southwestern Sudbury Structure, numerous Nipissing sills intruded the Huronian Supergroup and adjacent Archean basement rocks (Figure 3). Individual sills can reach up to 400 m thick and extend over 2 km in length. The larger Nipissing sills typically occur near or at the contact between the Ramsay Lake and McKim Formations. These larger sills are crudely differentiated, with gabbro and melagabbro phases. Pegmatoidal and/or anorthositic pods occur within the thicker portions of the sills. Smaller, narrower sills, on the other hand, are undifferentiated. Mafic sills that are geochemically similar to the Elsie Mountain and Stobie formations intrude the volcanic rocks of the Elsie Mountain and Stobie Formations as well as sedimentary rocks of the Matinenda and McKim Formations (Figure 3) (Gordon et al. 2018a; Gordon 2021, 2022). These mafic sills likely represent a combination of synvolcanic intrusions and intercalated mafic flows. Locally, peperite textures are preserved where the mafic sills intruded the McKim Formation, suggesting that Huronian Supergroup volcanism continued (or resumed) during deposition of the McKim Formation (Gordon et al. 2018a; Gordon 2021). Most of these mafic sills are 50-100 m wide, but a few larger intrusions are up to 200 m wide and 2 km long (Figure 3). All are roughly east-trending and fine- to medium-grained. Trap Dike Swarm In the Sudbury area, east- to east-northeasttrending mafic dikes, interpreted to belong to the Trap dike swarm (circa 1750 Ma; cf. Bleeker et al. 2015), crosscut the Worthington Offset dike, Sudbury Breccia, folded Nipissing sills and Huronian Supergroup stratigraphy (Figure 3). These dikes are typically less than 50 in width and consist of undeformed quartz diabase. Field relationships suggest that these mafic dikes postdate the SIC event and they are tentatively linked to a post-Penokean rifting event (cf. Bleeker et al. 2015). Creighton Pluton The Creighton pluton is a 2455-2460 Ma subvolcanic sill that was emplaced into the Huronian Supergroup mafic volcanic rocks during rifting (Bleeker et al. 2015). It is interpreted as the high-level magma chamber to rhyolites of the Copper Cliff Formation (Bleeker et al. 2015). The western extent of the Creighton pluton crops out in the northeastern corner of Denison Township (Figure 3). It intruded the mafic volcanic rocks of the Stobie and Elsie Mountain Formations and is truncated to the north by the SIC. The pluton consists of leucocratic granite and porphyritic quartz monzonite. Inclusions of porphyritic quartz monzonite are locally found in the granite phase. Dikes of Unknown Affinity East-northeast-trending mafic dikes of similar appearance to the Trap dikes, but which are geochemically distinct, crosscut the Archean basement, folded Nipissing sills and Huronian Supergroup strata (Gordon et al. 2018a). The mafic dikes are fine- to medium-grained, locally plagioclase-phyric and are variably foliated. In the southern central portion of Drury Township, northwest-trending felsic dikes crosscut folded Nipissing sills. The felsic dikes are up to 5 m wide, quartz-rich, massive and contain inclusions of Nipissing gabbro (Gordon et al. 2018a). Nipissing Intrusive Suite The voluminous Nipissing Intrusive Suite is exposed for over 400 km from the Ontario–Quebec border to Sault Ste. Marie. The intrusive suite was 157 �Proceedings of the 68th ILSG Annual Meeting - Part 2 These mafic and felsic dikes have not been assigned to any known intrusive suite and likely represent one or more magmatic events that occurred after regional folding. They are not included on Figure 3; readers are referred to Gordon et al. (2018a) for more information. McKim Formation are juxtaposed against rocks of the Mississagi, Pecors and Ramsay Lake Formations. Northwest-trending faults produced significant offsets of the Huronian Supergroup strata. In Denison Township, rocks of the Stobie and McKim Formations are juxtaposed against rocks of the Elsie Mountain Formation along a northwesttrending fault (Figure 3). This configuration of Huronian Supergroup strata is truncated to the north by the SIC and crosscut by the CreightonVictoria deformation zone. Reactivation of the northwest-trending faults is indicated by sheared sublayer along the fault. These structures may represent southward extension of the similarly oriented Onaping Fault system which, like the Murray Fault, is thought to have originated as extensional faults during deposition of the Huronian Supergroup and reactivated during subsequent orogenic events (cf. Card 1978; Zolnai et al. 1984). Sudbury Dike Swarm The Mesoproterozoic Sudbury dike swarm (circa 1238 Ma: Krogh et al. 1987) consists of northwest-trending olivine diabase dikes that crosscut the Superior and Southern provinces, and their deformed and metamorphosed equivalents occur within the northwestern Grenville Province (Ketchum and Davison 2000). The Sudbury dike swarm extends ~300 km west and northwest from the Sudbury area. Different tectonic settings for their emplacement have been proposed, which include continental rifting (Shellnut and MacRae 2012), a back arc setting (Ernst and Bleeker 2010), or mantle-plume upwelling (Easton et al. 2021). In the southwestern Sudbury Structure, undeformed, northwest-trending dikes of the Sudbury dike swarm crosscut the Ramsey-Algoma granitoid complex, the Drury Township intrusion, rocks of the Huronian Supergroup and the SIC (Figure 3). The Sudbury dikes consist of undeformed olivine diabase, are fine- to mediumgrained and locally plagioclase-phyric. Pre-impact Foliation An early, southeast-trending foliation is locally preserved within basaltic rocks of the Elsie Mountain Formation adjacent to the Creighton Pluton contact. This foliation is locally crosscut by Sudbury Breccia, which also contains randomly oriented clasts of the foliated basalt and granitoid indicating that deformation started before the Sudbury impact event (Gordon 2018). Structure The southwestern Sudbury Structure displays multiple generations of structural fabrics and major structures, which are described below in chronological order. Bedding Parallel Thrust Faults Bedding-parallel thrust faults are recognized on the basis of stratigraphic repetition and/or absence of specific Huronian Supergroup formations. The thrust faults are folded along with Huronian Supergroup strata (Gordon et al. 2018a; Généreux et al. 2018) and, thus, predate regional folding. In Drury Township, a bedding-parallel foliation is associated with these thrust faults and is locally preserved in the matrix of Sudbury Breccia, indicating that these faults formed after the impact event (Généreux et al. 2018). Murray Fault and Northwest-Trending Faults The Murray Fault is a major east- to eastnortheast-trending structural feature in the Southern Province (Figure 1). It is thought to have originated as an extensional fault during Huronian sedimentation and has been periodically reactivated during subsequent tectonic events (Card and Hutchinson 1972; Card 1978; Zolnai et al. 1984). The Murray Fault truncates the Huronian Supergroup strata in the southeastern area of Denison Township (Figure 3), where rocks of the 158 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The CVDZ is characterized by a strong westnorthwest- to east-trending subvertical foliation, and a steeply plunging eastward-trending stretching lineation. Shear sense indicators correspond to dextral (horizontal) shearing in the west, and south-over-north dextral (oblique) shearing to the east (Généreux et al. 2016; Généreux, et al. 2017). The development of the CVDZ is interpreted as coeval with regional folding (Généreux et al. 2018). Regional Folds, Foliation and Lineation The present structural configuration of the Huronian Supergroup strata in Drury Township is strongly controlled by kilometre-scale, northeasttrending isoclinal folds. Adjacent to the Superior– Southern provinces contact, smaller scale folds are west-northwest to east-trending, following the orientation of the Superior–Southern provinces contact. Throughout Drury Township, Nipissing sills are folded along with the Huronian Supergroup strata or follow their axial trace. Folding decreases in intensity eastward, into Denison Township, and appears to be restricted to the McKim Formation (Figure 3). Northeast-trending Faults This fault group includes the Cameron Creek, Fairbank Lake, Chicago and Vermilion Lake faults, as well as unnamed faults of similar orientation (Figure 3). These northeast-trending brittle-ductile faults produce significant offset of SIC-related rocks, folded Huronian Supergroup strata and Nipissing sills. A moderate to strong east- to east-northeasttrending regional foliation is ubiquitous within the Huronian Supergroup stratigraphy and Sudbury Breccia. The regional foliation overprints Sudbury Breccia, and locally crenulates the beddingparallel foliation associated with the early thrust faults (Généreux et al. 2018). It is axial planar to folds on outcrop and contains a southeast- to southwest-plunging stretching lineation that parallels fold axes. Both fabrics also rotate along with fold orientations adjacent to the SouthernSuperior provinces contact (Généreux et al. 2018). Minor Structures Shattercones, which are impact-related conical fractures with distinctive cone or fan-shaped features, are found within arenites and conglomerates of the Mississagi and Ramsey Lake Formations, respectively. The regional foliation and mylonitic fabric are locally overprinted by north-northwest-trending crenulation cleavage and kink bands. Most kink bands observed are S-shaped and locally form conjugate sets of centimetre-scale box-folds (cf. Généreux et al. 2016; Gordon et al. 2018a). Local, brittle, northwest-trending faults also crosscut all rock types, including Sudbury dikes. Creighton-Victoria Deformation Zone The Creighton-Victoria deformation zone (CVDZ) consists of an east-southeast-trending, 200 to 400-m wide mylonite zone that occurs along the contact between the Archean Superior Province and Paleoproterozoic Southern Province across Drury Township (Figure 3) (Gordon et al. 2015; Simard et al. 2016; Généreux et al. 2016; Généreux et al. 2017; Gordon et al. 2018a). Eastward into Denison Township, the CVDZ extends into a 1.5 km wide deformation corridor bound by the Creighton and Victoria faults (Généreux et al. 2017). The corridor consists of discrete, 5 to 20-m wide, shear zones that follow internal contacts within weakly to moderately foliated Huronian volcanic and sedimentary rocks. 159 �Proceedings of the 68th ILSG Annual Meeting - Part 2 FIELD TRIP DETAILS The Trap dike is massive, grey in colour, homogeneous and fine-grained. It is at least 20-m wide. Along the contact with the Trap dike, the Nipissing gabbro is strongly amphibolitized. Both the Nipissing sill and the Trap dike were metamorphosed to greenschist facies. Plagioclase is partially sausseritized and pyroxene is completely replaced by amphibole and, locally, biotite. Geological Maps Geological compilation maps covering all or parts of the area of the field trip include Ames et al. (2005), Dressler (1984b) and Card and Lumbers (1973). A detailed geology map is available for Drury Township (Gordon et al. 2018b). ROAD LOG Approximately 50 m south of the Trap dike, the Nipissing sill is intruded by olivine diabase of the Sudbury dike swarm (Figure 6). This Sudbury dike can be traced for over 8 km and truncates the folded Huronian Supergroup stratigraphy and the Worthington Offset dike (Figure 3). The contact between the Sudbury dike and Nipissing sill is sharp, strikes northwest and dips steeply northward. The Sudbury dike is at least 25 m wide and has a distinctive brown weathered surface. It is magnetic, massive, fine- to medium-grained and plagioclase-phyric. It is also relatively unaltered, its primary mineralogy consisting of plagioclase, clinopyroxene, orthopyroxene and olivine. The Sudbury dike is truncated by late, northwesttrending brittle faults adjacent to the southern contact with the Nipissing gabbro and within the central area of the exposure (Figure 6). Note: Caution should be taken when parking vehicles on the shoulder of the roads and when examining outcrops along any road along the field trip route. All UTM co-ordinates are given in NAD 83 datum, zone 17. Figure 3 shows the location of the field trip stops. The mileages in the road log represent the distance from one stop to another. 38.5 km (30 minutes) – Starting at the Willet Green Miller Centre in Sudbury, head west toward Ramsey Lake Road. Turn left onto Ramsey Lake Road and continue west for 1.9 km. Use the left 2 lanes to turn left (south) on Paris Street and continue for 4 km. Use the right lane to take the Highway 17W ramp to Sault Ste. Marie. After 800 m, continue straight to merge onto Hwy 17. Drive west on Hwy 17 for 30 km. Turn right (north) onto Fairbank Lake/Totten Mine Road and continue for 1.1 km. Stop 1 will be on the right (east) side of the road. 6.4 km (~10 minutes) – Head north on Fairbank Lake Road toward Bay Street and drive for 1.4 km. Turn right (north) onto Crean Hill Road and continue for 2.1 km. Turn right at the fork to stay on Crean Hill Road and continue for 2.9 km. Stop 2 is past the gates and will be on the right (east) side of the road. Stop 1 Nipissing sill, Trap and Sudbury dikes UTM coordinates 0471368E 5136502N Exposed on the east side of Fairbank Lake/Totten Mine Road is a large outcrop of Nipissing gabbro intruded by mafic dikes of the Trap and Sudbury dike swarms (Figure 3). The Nipissing gabbro is part of a northwest-trending sill that is up to 300-m wide and 7-km long (Figure 3). The Nipissing gabbro is green-grey, massive and medium-grained. Along its northern margin, it is intruded by a Trap dike (Figure 5). The contact between dikes is sharp, strikes east-northeast and dips steeply southward. Stop 2 Crean Hill - Footwall Breccia UTM coordinates 0473051E 5141707N Just beyond the gates and on both sides of the access road are stripped outcrops that expose variably mineralized Crean Hill Footwall Breccia. The exposure is within 200 m of the verticallydipping SIC contact. The open pit of the pastproducing Crean Hill Mine is located just beyond the fence at the northern edge of the outcrops (Figure 7). 160 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5. Nipissing gabbro crosscut by a northeast-trending Trap dike (Stop 1). Figure 6. Sudbury dike in contact with Nipissing gabbro and crosscut by northwest-trending brittle faults (Stop 1). 161 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 7. Partial view of the Crean Hill outcrops with the Crean Hill open pit in the background (Stop 2). Crean Hill mine, currently owned by Vale Canada Limited, operated sporadically from 1909 to 2000, producing 744,747 tons grading 2.14% Ni and 2.9% Cu to the end of 1916 (Card 1968). The main orebody consisted of sulphide breccia hosted in brecciated Huronian Supergroup volcanic rocks along the SIC contact (Coleman 1913; Knight 1917; Card 1968), with disseminated low-sulphide PGE mineralization occurring in the underlying brecciated footwall rocks. Monolithic breccias occur as irregular veins in dacite (felsic breccia) and basalt (mafic and quartzofeldspathic breccias). Partial melt patches make up 5–20 vol.% of the host rocks, and consist of quartz and plagioclase, with hornblende porphyroblasts. They terminate en biseau and cut across mafic breccia veins. Heterolithic dioritic breccias occur as pods (bilithic breccia), dikes (breccia dikes), and anastomosing veins (mixed breccia) within basalt (Figure 9). They formed as melts with contactparallel flow textures, which are defined by elongate wispy clasts that wrap around basalt clasts. Partial melt textures are not observed within the dioritic breccias, but narrow (<1 cm) quartzplagioclase leucosomes occur along their contact with the host basalt. The Crean Hill stripped outcrops expose brecciated basalt with minor dacite and quartzofeldspathic arenite of the Huronian Supergroup, which are crosscut by dikes and pods of Footwall Breccia (Figure 8). Four compositions of breccia are found in the outcrop: monolithic felsic, mafic, and quartzofeldspathic breccias, and heterolithic dioritic breccias (Généreux et al. 2021). 162 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 8. Geological map of the Crean Hill outcrops (from Généreux et al. 2021). UTM coordinates are in NAD 83, Zone 17 (Stop 2). 163 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Mineralized mixed breccia appears gossanous on outcrop, containing <5 vol.% of disseminated pyrrhotite, pentlandite, chalcopyrite, gersdorffite and precious metal minerals (PMM), which are commonly associated with epidote-quartz(-calcite) alteration patches. Shear zones appear to be the main control on the distribution of PGE (Gibson et al. 2010), with strongly foliated mixed breccia generally enriched in Au, Pt and Pd. 1.6 km (~3 minutes) – Head southwest on Crean Hill Road toward Fairbank Lake Road and drive for 1.4 km. Turn left (southeast) on the gravel road. Continue east on the gravel road for 200 m, keep left at the fork. Stop 3 is on the north side of the road. Figure 9. Mixed breccia unit crosscut by a heterolithic dioritic breccia dike (from Généreux et al. 2021) (Stop 2). Stop 3 Vermilion Mine – Vermilion Offset Dike, Sudbury Breccia and Stobie Formation A detailed study of the Crean Hill breccias by Généreux et al. (2021) showed that they have significantly lower SiO2 and higher TiO2 than the SIC, suggesting that the breccias likely did not form by injection of SIC melt into the fractured target rocks. Modeling of partial melt compositions during contact metamorphism showed that the dioritic breccia matrices are too mafic to have formed by anatexis during contact metamorphism. Instead, their composition mirrors that of their host rocks, thus they are best interpreted as locallyderived shock melts that formed during shock compression and which were trapped in the basement rocks during cooling of the SIC (Généreux et al. 2021). The breccias were subsequently modified by contact metamorphism (T ≥ 750°C) during cooling of the melt sheet, and by later syn-tectonic regional metamorphism at upper greenschist to amphibolite conditions. UTM coordinates 0472251 E 5140280 N Francis L. Sperry discovered sperrylite (PtAs2) at the Vermilion mine (Wells 1889), which is also the type locality for arsenohauchecornite (Ni18Bi3AsS16), michenerite (PdBiTe) and violarite (FeNi2S4). The mine operated from 1887 to 1916 and produced over 4000 tonnes of ore with exceptionally high grade of 6.64% Ni and 6.89% Cu, including 180 tonnes at 20-25% Cu-Ni, 125 g/t Ag, 125 g/t Pd, 46.9 g/t Pt and 10.3 g/t Au (Holloway et al. 1917). The stripped outcrop exposes the Vermilion Offset quartz diorite (Figure 10 and 11). This offset dike is not connected to the Main Mass of the SIC (Grant and Bite 1984), but instead forms a lens at the contact between Sudbury Breccia and mafic volcanic rocks of the Stobie Formation (Figure 10). At the northernmost part of the outcrop are slightly deformed basaltic volcanic rocks of the Stobie Formation, where bedded lapilli tuff is interlayered with vesicular basalt, pillow breccia and possible flow-top breccia (Figure 12). Volcanic textures such as hyaloclastite, amygdules, and lapilli are generally well-preserved but are locally overprinted by coarse acicular amphiboles. Low-sulphide PGE mineralization is hosted in anastomosing veins of ‘mixed’ dioritic breccia, which contain a medium-grained dioritic matrix intermingled with a fine-grained basaltic matrix. This mixed breccia is locally crosscut by irregularly shaped dioritic breccia dikes (Figure 9) and displays a strong mottled texture that is further complicated by the presence of leucosomes. 164 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 10. Geology of the Vermilion Offset (modified from Grant and Bite 1984). Detailed geology of the Vermilion Mine surface outcrop modified from Lightfoot et al. (1997) (Stop 3). 165 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 11. Overview of the stripped outcrop at the Vermilion mine (Stop 3). The central portion of the main outcrop exposes the Vermilion quartz diorite, which is generally medium-grained and contains up to 20% of small (<5 cm), partially digested, felsic and mafic clasts. At the top of the main outcrop is a 10-m wide section of finer-grained and foliated inclusionbearing quartz diorite (IQD), which contains 2040% of dacite, basalt and amphibolite clasts that are up to 10 m in size. IQD is in sheared contact with basaltic rocks to the north and Sudbury Breccia to the South. Figure 12. Basaltic pillow breccia with wellpreserved hyaloclastite (Stop 3). 166 �Proceedings of the 68th ILSG Annual Meeting - Part 2 1.5 km. Park at marked trail. Walk north on trail for 150 m. Trail leads to outcrops exposed along the powerline. Stop 4 Creighton Fault and Stobie Formation UTM coordinates 474656E 5140835N Along the powerline is an excellent exposure of ductile deformation observed within the deformation corridor bound by the Creighton and Victoria Faults. The Creighton Fault can be traced across most of Denison Township as a prominent, relatively continuous, topographic low with periodic outcrops. Along the powerline, a 20-m wide flow-banded dacite is in sheared contact with a relatively massive, locally amygdaloidal, basaltic flow. Both the intermediate and mafic volcanic rocks are part of the Stobie Formation. Figure 13. Matrix-supported Sudbury Breccia with elongated clasts (Stop 3). Sudbury Breccia is best exposed south and southwest of the main quartz diorite outcrop, closest to the parking area. It occurs as variably foliated, matrix-supported, heterolithic breccia (Figure 13) within brecciated mafic volcanic rocks. Ductile deformation is expressed in the intermediate unit as a strong east-northeasttrending foliation that is consistently oriented counter-clockwise to flow banding, suggesting apparent dextral shearing (Figure 14). An intersection lineation between the flow banding and the foliation steeply plunges to the southeast. The same kinematic indicators are observed in other shear zones in the area, including the Creighton-Victoria mylonite zone farther west (Généreux et al. 2017). The Vermilion ore occurred in shear-hosted sulphide veins and irregular lenses ranging from a few centimetres to 40 cm in diameter. The shear zones are found mainly in Sudbury Breccia and strike parallel to the quartz diorite–Sudbury Breccia contact (Szentpéteri et al. 2003). Disseminated sulphides and platinum-group minerals (PGM) also occur in quartz diorite and foliated Sudbury Breccia, and are associated with irregularly distributed epidote-albite-chlorite alteration patches that range from a few centimetres to several metres in size. The close spatial association of PGM with alteration minerals, their finely disseminated nature, the presence of sulphides-PGM in secondary hydrothermal veins, and the occurrence of sulphide veins within shear zones all suggest a complex multistage magmatic-hydrothermal and metamorphic-hydrothermal origin of the sulphidePGM assemblages (Szentpéteri et al. 2003). 4 km (15 minutes) – Return to Crean Hill Road (~200 m). Turn left (southwest) toward Fairbank Lake Road for 1.3 km. Turn right (northwest) on Fairbank East Road and continue north for approximately 1 km. Turn right (east) on the gravel road. Drive east for Figure 14. Strong pervasive foliation (S) in felsic volcanic rock consistently oriented counterclockwise to flow banding (S0), suggesting apparent dextral shearing (from Généreux et al. 2017) (Stop 4). 167 �Proceedings of the 68th ILSG Annual Meeting - Part 2 13 km (25 minutes) – Walk south 150 m along trail to return to the gravel road. Travel west for 1.5 km. Turn left (south) on Fairbank East Road and continue for 1.1 km. Continue south onto Crean Hill Road for 2.1 km. Turn right (west) onto Fairbank Lake Road. Fairbank Lake Road turns into Spanish River Road after 7 km, continue along Spanish River Road for another 4 km. Where Spanish River Road turns southward, continue straight (west) onto High Falls Road for 200 m. Stop 5 is at the bend along High Falls Road. The outcrops are exposed north of the road. Use the small gravel road adjacent to the outcrop for parking. Stop 5 Synvolcanic mafic sill and McKim Formation UTM coordinates 0460806E 5135978N On the north side of Spanish River Road, a large outcrop contains turbidites of the McKim Formation and a synvolcanic mafic sill (Figure 3). The McKim Formation consists of thickly laminated mudstone with siltstone layers, and contains porphyroblasts of chloritoid and staurolite (Figure 15A). Bedding is subvertical and trends northeast. There is a strongly developed beddingparallel foliation overprinted by regional eastnortheast-trending foliation (Figure 15B). Figure 15. A) Laminated mudstones of the McKim Formation with chloritoid and staurolite porphyroblasts. B) Bedding parallel foliation in thickly laminated mudstones of the McKim Formation overprinted by regional east-northeasttrending foliation, vertical face (Stop 5). North of the McKim Formation a 50-60 m wide mafic sill is exposed. The sill trends parallel to bedding and exhibits distinctly rounded and lobate, mafic enclaves that are enclosed in a felsic, micarich and garnet-bearing matrix (Figure 16). This texture has been interpreted as a peperite, formed as a result of a mafic sill intruding what were originally unconsolidated wet sediments of the McKim Formation. The mafic enclaves are geochemically similar to that of the mafic volcanic rocks of the Elsie Mountain Formation (Gordon et al. 2018; Gordon 2021, 2022). Figure 16. Peperite with fine-grained, amoeboid mafic lobes encompassed by a felsic matrix (Stop 5). 168 �Proceedings of the 68th ILSG Annual Meeting - Part 2 10 km (15 minutes) – Drive east on High Falls Road and continue straight onto Spanish River Road for 4 km, turn left (north) on Fairbank Lake Road, drive north for 5.5 km. Stop 6 is on the left (west) side of Fairbank Lake Road. Stop 6 Drury Township Intrusion and Matachewan dike swarm UTM coordinates 0464241E 5142157N On the west side of Fairbank Lake Road, outcrops of the Drury Township intrusion are exposed. The Drury Township intrusion consists of anorthositic gabbro that is medium- to coarsegrained, locally pegmatitic and vari-textured (Figure 17A). Mineralogy is characterized by greenschist grade assemblages. Plagioclase is almost entirely saussuritized and pyroxenes are replaced by amphibole and chlorite. On the east side of the road, the anorthositic gabbro varies from undeformed to mylonitized (Figure 17B). Where deformed, shear zones and foliation trend northeast and are parallel to the adjacent Chicago Fault (Figure 3). The anorthositic gabbro is also crosscut by a narrow, plagioclase-phyric mafic dike of the Matachewan dike swarm. The Matachewan dike is fine-grained, massive and metamorphosed to greenschist facies. Figure 17. A) Coarse-grained to pegmatitic anorthositic gabbro of the Drury Township intrusion. B) Mylonitized anorthositic gabbro of the Drury Township intrusion (Stop 6). 8 km (10 minutes) – Drive south on Fairbank Lake Road for 5.5 km. Turn left (east) to stay on Fairbank Lake Road and continue for 1.8 km. Turn left (north) on Kidd Copper Mine Road and drive north for 700 m. Stop 7a is on the left (west) side of Kidd Copper Mine Road. The Worthington Offset dike has been mined periodically since 1885 and is host to Sudbury’s most recently developed deposit, Totten Mine, with grades of 1.42% Ni, 1.9% Cu and 4.8 g/t PGM (Lightfoot 2017; Lightfoot and Farrow 2002). Three past producing mines can be found on the Aer-Kidd property (Figure 18). The first is the Howland Pit, which has been filled in and partially reclaimed. The second is the Robinson Mine (Stop 7b), which includes a shaft cap, a large, fenced hole, and small adit (Figure 19). The third pastproducing site is the Rosen and Gersdorffite mines (also known as Aer Mine), which are located northeast of the old mill site and contains a shaft cap. Stop 7 Aer Kidd Property - Worthington Offset Dike The beginning of Kidd Copper Mine Road is Vale Canada Limited property and is gated. The property changes ownership to SPC Nickel Corp. at the crest of the hill just before the Howland pit. The Worthington Offset dike crops out on the west side of the road and is intermittently exposed within the Aer-Kidd Property from the Howland pit to Perch Lake (Figure 18). 169 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 18. Geological map of the Aer-Kidd Property along the Worthington Offset dike. Past-producing mines are located within the dike where amphibole- and inclusion-bearing quartz diorite (AIQD) pods are observed. Stop 7a and 7b are located between the Howland Pit and Robinson Mine. UTM coordinates are in NAD 83, Zone 17 (Stop 7). 170 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 19. Geological map of the stripped outcrop between the Howland Pit and the Robinson Mine sites (Stop 7a). The Stop includes siltstones of the McKim Formation, as well as the quartz-diorite (QD) and inclusion-bearing quartz-diorite (IQD) phases of the Worthington Offset dike. UTM coordinates are in NAD 83, Zone 17. 171 �Proceedings of the 68th ILSG Annual Meeting - Part 2 There are many stripped and forested outcrops of the Worthington Offset dike over the length of the property. The two best surface exposures are: 1) the large, stripped outcrop on the left side of the road between the Robinson Mine and Howland Pit (Stop 7a) (Figure18), and 2) the Robinson Mine site (Stop 7b) (Figure 18, 19). foliated siltstone of the McKim Formation. QD is found farther up the hill, but it is much thinner than that at Stop 7a. Further north is the sharp contact between the QD and IQD. The mineralized matrix of the IQD at this site has a slightly higher grade than at Stop 7a. Massive sulphides are found further north, just below the fence. Stop 7a: Worthington Offset Dike and McKim Formation UTM coordinates 0466620E 5137925N On the west side of Kid Copper Mine Road, past the Howland Pit and core farm, is a large, stripped outcrop sloping north. This outcrop shows the phase separation of the dike well and is representative of the weakly mineralized to unmineralized portions of the Worthington Offset. The outcrop contains a small section of McKim Formation siltstone (near the road) in sharp contact with Worthington quartz diorite (QD) (Figure 19). The siltstone is thickly laminated and exhibits a moderate east-northeast-trending foliation. QD is medium-grained, massive, with local veins and jointing, and contains several rounded sedimentary enclaves adjacent to the contact with the siltstone (Figure 20A). Farther north, the QD phase is crosscut by inclusion-bearing quartz diorite (IQD). The IQD phase is heterolithic and contains inclusions of QD, amphibolite, basalt and siltstone that are enclosed in a massive, medium-grained quartz diorite matrix (Figure 20B). Most inclusions are subrounded and range from a centimeter to submeter in diameter. Sulphide blebs are visible throughout the matrix of the IQD. Figure 20. A) Sharp and linear contact between the Worthington Offset dike and the McKim Formation. Note the rounded sedimentary enclaves in QD near the contact. B) Heterolithic, massive IQD with sub-rounded inclusions derived from local host rocks. Note the pervasive sulphide burns throughout the quartz diorite matrix (Stop 7a). Stop 7b: Mineralization in the Worthington Offset Dike UTM coordinates 0466765E 5138015N 200 m (1 minute) – Continue northeast along Kidd Copper Mine Road for 200 m. Stop 7b is on the left (north) side of the road. Northward up the hill, toward the fenced hole and behind the shaft cap, is the Robinson Mine outcrop. The hill exposes a large portion of the Worthington Offset dike, but the best exposure is the stripped outcrop located south of the fence adjacent to the Robinson pit (Figure 21). The southern edge of the outcrop consists of weakly 172 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 21. Geological map of the stripped outcrop found on the Robinson Mine site at Stop 7b. The Stop includes siltstones of the McKim Formation, and quartz diorite (QD), inclusion-bearing quartz diorite (IQD) and amphibole- and inclusion-bearing quart diorite (AIQD) phases of the Worthington Offset dike. Massive sulphide mineralization is hosted within the AIQD phase present along the northern edge of the exposure. A sub-unit of IQD is found at this location, and locally is called amphibolite-bearing IQD (AIQD). The contact between IQD and AIQD is transitional, which is why the latter is considered a sub-unit of IQD rather than a distinct phase of the dike. AIQD contains almost exclusively large, rounded, amphibolite inclusions, with massive sulphides wrapping around them (Figure 22A). The amphibolite inclusions are dark green, massive, coarse-grained, and can range from a few centimetres to several metres in diameter. These inclusions are thought to have been derived from the nearby Nipissing sills. Similar inclusions have been reported at the Totten Mine, where they are referred to as “Sudbury Gabbros” (Lightfoot 2017). 173 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Massive to semi-massive sulphide mineralization is associated with AIQD and occurs solely within the quartz diorite matrix. The amphibolite inclusions themselves are not mineralized, thus they are quite dilutive to the mineralization of this sub-unit. This style of mineralization is representative of the mineralization observed in the Worthington Offset dike. Chalcopyrite, pyrrhotite, and pentlandite are the dominant sulphide minerals (Figure 22B), and gersdorffite and niccolite have been identified locally in drill core. Sulphide mineralization is not a necessary feature of AIQD, but the latter is always present where mineralization occurs. The distribution of mineralized AIQD is complex and varied. On this property, 4 modeled vertical shoots are known to contain significant sulphide mineralization (Howland, Robinson, Rosen, and Perch Lake). Elsewhere along the Worthington Offset, mineralization is structurally controlled, occurring in bends, folds, and boudins along the dike. A narrow diabase dike is also found at this stop. The dike crosscuts the Worthington Offset dike and has been tentatively assigned to the Trap dike swarm. It is massive, fine-grained, weakly magnetic and can be followed for several metres. Diabase dikes are commonly observed in drill core. Two large olivine diabase dikes of the Sudbury dike swarm also cut through the property. Shear zones are observed throughout the Aer Kidd property, including at this outcrop, where an east-northeast-trending shear zone displaced the offset dike by less than a metre. Such localized displacement is observed along several other shear zones observed on surface and in drill core, and has been reported on many historical mine maps. The orientation of the shear zone is similar to other easttrending shear zones in Denison and Drury townships, including the CVDZ (Figure 3), and likely formed during the same deformation event. Figure 22. A) Mineralized outcrop on the surface exposure of the Robinson Mine, showing AIQD with amphibolite inclusions. The QD matrix between the inclusions hosts massive to semimassive sulphide mineralization. B) Mineralized AIQD in drill core. Two massive sulphide stringers are hosted within quartz diorite matrix and wrap around rounded amphibolite inclusions. Note the smaller amphibolite inclusions the sulphide stringer vein. Dominant sulphide minerals are pyrrhotite and pentlandite, with lesser chalcopyrite (Stop 7b). A narrow diabase dike is also found at this stop. The dike crosscuts the Worthington Offset dike and has been tentatively assigned to the Trap dike swarm. It is massive, fine-grained, weakly magnetic and can be followed for several metres. Diabase dikes are commonly observed in drill core. Two large olivine diabase dikes of the Sudbury dike swarm also cut through the property. 174 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Acknowledgments References The authors would like to thank R-L. Simard for her collaboration and contributions to the Southwest Sudbury Structure bedrock mapping project. R.M. Easton is thanked for his ongoing guidance and assistance on this project, and M. Duguet, S. Evers, J.E. Chartrand, S.J. McIlraith, and P. Gervais for all the helpful discussions, as well as assistance with databases, software logistics, sample preparation and drafting. B. Lafrance and D.K. Tinkham from Laurentian University are thanked for their ongoing advice and helpful discussions in building the geological interpretation. Special thanks are extended to Vale Canada Limited, Lonmin PLC, Wallbridge Mining Company Limited, SPC Nickel Corp., KGHM International, Sudbury Integrated Nickel Operations (a Glencore Company) and private landowners for providing access to their properties, sharing information and knowledge, and for all additional support and assistance that has been provided. Last but not least, this project would not have been possible without our hard-working field assistants, J. Ménard, J. Enright, L. Lebeau, S. Robillard, C. Stone, E. Campbell, G. Broughton, B. Clarke, J. Creppin, M. Jacques and J. AndersonButcher. Ames, D.E. and Farrow, C.E.G. 2007. Metallogeny of the Sudbury mining camp, Ontario; in Mineral Deposits of Canada: A Synthesis of Major DepositTypes, District Metallogeny, the Evolution of Geological Provinces, and Exploration Methods, Geological Association of Canada, Mineral Deposits Division, p.329-350. Ames, D.E., Bleeker, W., Heather, K.B. and Wodicka, N. 1997. Timmins to Sudbury transect: new insights into the regional geology and setting of mineral deposits; Geological Association of Canada– Mineral Association of Canada, Joint Annual Meeting, Ottawa, Field Trip Guidebook B6, 133p. Ames, D.E., Davidson, A., Buckle, J.L. and Card, K.D. 2005. Geology, Sudbury bedrock compilation, Ontario; Geological Survey of Canada, Open File 4570, scale 1:50 000. Ames, D.E., Golightly, J.P., Lightfoot, P.C. and Gibson, H.L. 2002. Vitric compositions of the Onaping Formation and their relationship to the Sudbury Igneous Complex, Sudbury Structure; Economic Geology, v.97, p.1541-1562. Ames, D.E., Watkinson, D.H. and Parrish, R.R. 1998. Dating of the regional hydrothermal system induced by 1850 Ma Sudbury impact event; Geology, v.26, p.447-450. Bailey, J., Lafrance, B., McDonald, A.M., Federowich, J.S., Kamo, S. and Archibald, D.A. 2004. MazatzalLabradorian-age (1.7–1.6 Ga) ductile deformation of the South Range Sudbury impact structure at the Thayer Lindsley Mine, Ontario; Canadian Journal of Earth Sciences, v.41, p.1491-1505. Bennett, G.B., Dressler, B.O. and Robertson, J.A. 1991. The Huronian Supergroup and associated intrusive rocks; in Geology of Ontario, Ontario Geological Survey, Special Volume 4, Part 1, p.549-591. Bleeker, W., Hamilton, M.A., Ernst, R.E. and Söderlund, U. 2012. Resolving the age structure of the Matachewan event: Magmatic pulses at ca. 24452452 Ma, 2458-2461 Ma, and 2475-2480 Ma; unpublished CAMIRO Reports A96, A97, and A98, 17p. 175 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Large Igneous Province; in Targeted Geoscience Initiative: 2018 report of activities, Geological Survey of Canada, Open File 8549, p.403-419. 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 Targeted Geoscience Initiative 4: Canadian nickelcopper-platinum group elements-chromium ore systems – Fertility, pathfinders, new and revised models, Geological Survey of Canada, Open File 7856, p.151-156. Davis, D.W. 2008. 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Coexisting chloritoid and staurolite and the staurolite-chlorite isograd from the Agnew Lake area, Ontario, Canada; Geological Magazine, v.108, p.205-219. ——— 2021. Geology and geochemistry of the Elsie Mountain and Stobie formations, Huronian Supergroup: Developing a chemostratigraphy to address challenges with the current subdivision; abstract in Geoscience Canada, v.48, no.4, p.178. Fralick, P. and Miall, A. 1989. Sedimentology of the lower Huronian Supergroup (early Proterozoic), Elliot Lake area, Ontario, Canada; Sedimentary Geology, v.63, p.127-153. ——— 2022. Geology and geochemistry of the Elsie Mountain and Stobie Formations, Huronian Supergroup: Developing a chemostratigraphy to address challenges with the current subdivision; in 68th Institute on Lake Superior Geology, Proceedings, v.68, pt.1, p.18-19. Généreux, C-A., Lafrance, B. and Gordon, C.A. 2017. The Creighton Fault and its relation to the mylonite zone in Drury Township, Southwest Sudbury Structure; in Summary of Field Work and Other Activities, 2017, Ontario Geological Survey, Open File Report 6333, p.16-1 to 16-10. Gordon, C.A. and Généreux, C-A. 2017. Preliminary Results from Geological Mapping in Denison Township, Southwest Sudbury Structure; in Summary of Field Work and Other Activities, 2017, Ontario Geological Survey, Open File Report 6333, p.15-1 to 15-15. Généreux, C-A., Lafrance, B., Simard, R-L. and Gordon, C.A. 2016. Mylonite zone and regional deformation in Drury Township, southwest Sudbury Structure; in Summary of Field Work and Other Activities, 2016, Ontario Geological Survey, Open File Report 6323, p.19-1 to 19-9. Gordon, C.A., Simard, R-L. and Généreux, C-A. 2015. Geology and low-sulphide platinum group element mineralization of Drury Township, South Range, Sudbury Igneous Complex; in Summary of Field Work and Other Activities, 2015, Ontario Geological Survey, Open File Report 6313, p.21-1 to 21-18. Généreux, C-A., Lafrance, B., Tinkham, D.K. and Gordon, C.A. 2018. Polyphase deformation in the southwest Sudbury impact structure; Geological Association of Canada–Mineralogical Association of Canada, Joint Annual Meeting, Vancouver, abstract. ——— 2018a. Precambrian geology of Drury Township, Southwest Sudbury Structure: Explanatory notes for Preliminary Map P.3823; Ontario Geological Survey, Open File Report 6346, 49p. Généreux, C-A., Tinkham, D. and Lafrance, B. 2021. On the role of shock melting and anatexis in breccia formation: Southern Sudbury impact structure, Canada; Precambrian Research, v.363. 106346, 25p. ——— 2018b. 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Mafic intrusions in the footwall of the Sudbury Igneous Complex: Origin of the Sudbury impact melt sheet and its associated ore deposits; Ore Geology Reviews, v.120. 103435, 31p. Holloway, G.T., Miller, W.G., Young, M. and Gibson, T.W. 1917. Report of the Royal Ontario Nickel Commission with appendix; King’s Printer, Toronto; available on-line as Ontario Geological Survey, OP01, 1006p. Ketchum, K.Y., Heaman, L.M., Bennett, G. and Hughes, D.J. 2013. Age, petrogenesis and tectonic setting of the Thessalon volcanic rocks, Huronian Supergroup, Canada; Precambrian Research, v.233, p.144-172. Jago, B.C., Morrison, G.G. and Little, T.L. 1994. Metal zonation patterns and microtextural and micromineralogical evidence for alkali- and halogen-rich fluids in the genesis of the Victor Deep and McCreedy East footwall copper orebodies, Sudbury igneous complex; in Proceedings, Sudbury–Noril’sk Symposium, Ontario Geological Survey, Special Volume 5, p.65-76. Knight, C. 1917. Geology of the Sudbury area and description of Sudbury ore bodies; in Report of the Royal Ontario Nickel Commission, p. 104-211. 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 UPb 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. James, R.S., Easton, R.M., Peck, D.C. and Hrominchuk, J.L. 2002. The East Bull Lake intrusive suite: Remnants of a ~2.48 Ga large igneous and metallogenic province in the Sudbury area of the Canadian Shield; Economic Geology, v.97, p.15771606. Krogh, T.E., Davis, D.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, Ontario Geological Survey, Special Volume 1, p.431-446. Jørgensen, T.R.C., Tinkham, D.K. and Lesher, C.M. 2019. Low-P and high-T metamorphism of basalts: Insights from the Sudbury impact melt sheet aureole and thermodynamic modelling; Journal of Metamorphic Geology, v.37, p.271-313. Lakomy, R. 1990. Implications for cratering mechanics from a study of the Footwall Breccia of the Sudbury impact structure, Canada; Meteoritics, v.25, p.195207. Kamo, S.L. 2006. Report on U-Pb geochronological data from the southern Abitibi Subprovince, Bannockburn–Montrose and Vernon townships, and the Grenville Front region, Thistle–Sisk townships, Ontario; internal U/Pb age report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, Department of Geology, University of Toronto, Toronto, Ontario, 20p. Lafrance, B. and Kamber, B.S. 2010. Geochemical and microstructural evidence for in situ formation of pseudotachylitic Sudbury breccia by shock-induced compression and cataclasis; Precambrian Research, v.180, p.237-250. Lightfoot, P.C. 2017. Nickel Sulfide Ores and Impact Melts; Elsevier, 1130p. Kamo, S.L., Krogh, T.E. and Kumarapeli, P.S. 1995. Age of the Grenville dyke swarm, Ontario-Quebec: implications for the timing of Iapetan rifting; Canadian Journal of Earth Sciences, v.32, p.273280. Lightfoot, P.C. and Farrow, C. 2002. Geology, Geochemistry, and Mineralogy of the Worthington Offset Dike: A Genetic Model for Offset Dike Mineralization in the Sudbury Igneous Complex; Economic Geology, v.97, p.1419-1446. Ketchum, J.W.F. and Davidson, A. 2000. Crustal architecture and tectonic assembly of the Central Gneiss Belt, southwestern Grenville Province, Canada: a new interpretation; Canadian Journal of Earth Sciences, v.37, p.217-234. Lightfoot, P.C., Doherty, W., Farrell, K., Keays, R.R., Moore, M. and Pekeski, D. 1997. Geochemistry of 178 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Geochronology Inventory of Ontario (July 2019 update), online database. the main mass, sublayer, offsets, and inclusions from the Sudbury Igneous Complex, Ontario; Ontario Geological Survey, Open File Report 5959, 231p. ——— 2022. Ontario Mineral Inventory database; Ontario Geological Survey, Mineral Inventory database (March 2022 update), online database. McCormick, K.A., Fedorowich, J.S., McDonald, A.M. and James, R.S. 2002. A textural, mineralogical, and statistical study of the Footwall Breccia within the Strathcona Embayment of the Sudbury Structure; Economic Geology, v.97, p.125-143. Papapavlou, K., Darling, J.R., Storey, C.D., Lightfoot, P.C., Moser, D.E. and Lasalle, S. 2017. Dating shear zones with plastically deformed titanite: New insights into the orogenic evolution of the Sudbury impact structure (Ontario, Canada); Precambrian Research, v.291, p.220-235. Meldrum, A., Abel-Rahman, A.M., Martin, R.F. and Wodicka, N. 1997. The nature, age and petrogenesis of the Cartier Batholith, northern flank of the Sudbury Structure, Ontario; Precambrian Research, v.82, p.265-285. Péntek, A., Molnár, F., Watkinson, D.H. and Jones, P.C. 2008. Footwall-type Cu-Ni-PGE mineralization in the Broken Hammer area, Wisner Township, North Range, Sudbury Structure; Economic Geology, v.103, p.1005-1028. Molnár, F. and Watkinson, D.H. 2001. Fluid-inclusion data for vein-type Cu-Ni-PGE Footwall ores, Sudbury Igneous Complex and their use in establishing an exploration model for hydrothermal PGE-enrichment around mafic-ultramafic intrusions; Exploration and Mining Geology, v.10, no.1-2, p.125-141. Prevec, S.A. 1993. An isotopic, geochemical and petrographic investigation of the genesis of early Proterozoic mafic intrusions and associated volcanism near Sudbury, Ontario; unpublished PhD thesis, University of Alberta, Edmonton, Alberta, 223p. Morrison, G.G., Jago, B.C. and White, T.L. 1994. Footwall Mineralization of the Sudbury Igneous Complex; in Proceedings, Sudbury–Noril'sk Symposium, Ontario Geological Survey, Special Volume 5, p.57-64. Naldrett, A.J. 1999. Summary: Development of ideas on Sudbury geology, 1992-1998; Geological Society of America, Special Paper 339, p.431-442. Prevec, S.A. and Baadsgaard, H. 2005. Evolution of Paleoproterozoic mafic intrusions located within the thermal aureole of the Sudbury Igneous Complex, Canada: Isotopic, geochronological and geochemical evidence; Geochemica et Cosmochimica Acta, v.69, no.14. p.3653-3669. Naldrett, A., Pessaran, A., Asif, M. and Li, C. 1994. Compositional variation in the Sudbury ores and prediction of the proximity of footwall copper-PGE orebodies; in Proceedings, Sudbury–Noril’sk Symposium, Ontario Geological Survey, Special Volume 5, p.133-146. Raharimahefa, T., Lafrance, B. and Tinkham, D.K. 2014. New structural, metamorphic, and U–Pb geochronological constraints on the Blezardian Orogeny and Yavapai Orogeny in the Southern Province, Sudbury, Canada; Canadian Journal of Earth Sciences, v.51, p.750-774. Noble, S.R. and Lightfoot, P.C. 1992. U–Pb baddeleyite ages for the Kerns and Triangle Mountain intrusions, Nipissing diabase, Ontario; Canadian Journal of Earth Sciences, v.29, p.1124-1129. Robertson, J.A. 1976. The Blind River uranium deposits: The ores and their setting; Ontario Division of Mines, Report 147, 73p. Robertson, J.A., Card, K.D. and Frarey, M.J. 1969. The Federal–Provincial Committee on Huronian stratigraphy progress report; Ontario Department of Mines, Miscellaneous Paper 31, 26p. O'Callaghan, J.W., Osinski, G.R., Lightfoot, P.C., Linnen, R.L. and Weirich, J.R., 2016. Reconstructing the geochemical signature of Sudbury Breccia, Ontario, Canada: Implications for Its formation and trace metal content; Economic Geology, v.111, p.1705-1729. Roscoe, S.M. 1969. 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VIII.--Sperrylite, a new Mineral; American Journal of Science (1880-1910) v.37, issue 217, p.67. Simard, R-L., Gordon, C.A. and Généreux, C-A. 2016. Geology of Drury Township, southwest Sudbury Structure: An overview; in Summary of Field Work and Other Activities, 2016, Ontario Geological Survey, Open File Report 6323, p.16-1 to 16-19. Wodicka, N. and Card, K.D. 1995. Late Archean history of the Levack gneiss complex, southern Superior Province, Sudbury, Ontario: New evidence from UPb geochronology; in Precambrian’95, Program with Abstracts, p.191. Stockwell, C.H. 1964. Fourth report on structural provinces, orogenies, and time-classification of rocks of the Canadian Precambrian Shield. Part II. Geological Studies; Geological Survey of Canada, Paper 64, p.1-7. Young, G.M. 1983. Tectono-sedimentary history of early Proterozoic rocks of the northern Great Lakes area; in Early Proterozoic Geology of the Great Lakes Region; Geological Society of America, Memoir 160, p.15-32. ——— 1982. 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The Birch Lake Batholith, Ontario; American Journal of Science, series 5, v.17, no.101, p.403-424. 180 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Field Trip 4 – Overview of the Sudbury Structure Sudbury District RGP Office Resident Geologist Program, Ontario Geological Survey, Sudbury, Ontario P3E 6B5 Introduction Geology of the Sudbury Area – Brief Overview This field trip is based on the “Sudbury Structure Field Trips” that have been given for over 20 years by the Sudbury District Office of the Ontario Geological Survey, Resident Geologist Program. There are, therefore, numerous contributors, not all of which have been recorded over the years. The Rousell and Brown (2009) Field Guide is also heavily leaned upon. Formation of the Sudbury Structure It is generally accepted that the Sudbury Structure and its associated abundant metal endowment were the result of a collision of a large meteorite (or comet; e.g. Petrus et al. 2015) with the Earth approximately 1850 million-years-ago (Krogh et al. 1984; Davis 2008). The impact happened near the contact of the Archean Superior and the Paleoproterozoic Southern provinces. Greater detail on the Sudbury Structure is presented in the introductory section of Field Trip 1 at this meeting “A Traverse Across the Sudbury Impact Structure” (Bleeker et al. 2022). Furthermore, Lightfoot (2017) gives a very indepth treatment of the Sudbury Igneous Complex and its associated nickel deposits. The entire meteorite, as well as a tremendous mass of earth’s crust, was volatilized on impact. A complex crater (Figure 1) developed and was filled with a ‘melt sheet’ of molten crustal material that segregated and crystallized to become the Sudbury Igneous Complex (SIC); which was then overlain by fragmental “fall-back” material and younger, more typical, sedimentary rocks that constitute the Whitewater Group. This one-day field trip will provide an overview of the Sudbury Structure, one of the most prolific nickel camps in the world and the remnant of one of the world’s largest impact craters. The sites visited provide a cross-section of the Sudbury Structure, including the footwall rocks, the Sudbury Igneous Complex and the crater-fill sedimentary rocks of the Whitewater Group. At the time of the impact the Penokean Orogeny (1870-1835 Ma) was underway. It is posited that the impact occurred in the foreland marine basin at the leading edge of the orogen. The Penokean Orogeny continued after the collision and resulted in modification to the impact crater. The subsequent nearly 2 billion years of tectonic and erosional history also changed the shape and size of the Sudbury Structure, resulting in its current configuration. In the footwall rocks, the evidence for the impact origin of the structure include the Sudbury Breccia (pseudotachylite) and shatter cones. The nickel deposits formed at the base of the Sudbury Igneous Complex, and in the associated quartz-diorite “Offset” dikes. The first of the crater-fill units, the Sandcherry Member of the Onaping Formation, is a fallback breccia from the impact. The subsequent crater-fill units, the Onwatin and Chelmsford Formations, will also be visited. Geology of the Sudbury Structure After the impact, the rocks and structure generated by the impact were subjected to varying degrees and orientations of deformation. The Sudbury Structure, therefore, has been divided into three distinct “ranges” in order to distinguish sections with different footwall rock affinities and The Stops are not in stratigraphic order, but in the order that provides the best safety for the participants (e.g., avoidance of crossing major highways whenever possible). 181 �Proceedings of the 68th ILSG Annual Meeting - Part 2 different degrees of structural over-print. Figure 2 outlines the approximate location of the “North Range”, “South Range” and “East Range”. The East Range is distinguished primarily on its distinct and complicated structural overprint. SIC Footwall Rock Types Sudbury Breccia (Stops 3 and 8) Sudbury Breccia resulted from the shattering of the rocks. that were at or near the surface of the Earth. due to the shock from the impact of the meteorite. This network of in-situ breccias is relatively abundant in zones that occur as far as 15 km from the Sudbury Structure contact, whereas more isolated occurrences have been identified up to 80 km from the exposed footwall contact. Breccia contacts are typically sharp and have a range of configurations from straight and regular through complex, anastomosing, riverine patterns. Sudbury Breccia dikes preferentially conform to pre-Sudbury Event structures and lithological contacts. Breccia Zones range from millimetres to hundreds of metres wide and may be traced for tens of kilometres (e.g., the South Range Breccia Belt). A stratigraphic section (Figure 3A) summarizes the rocks and features that developed as a result of the impact. These features are found in the country “Footwall” rocks as well as in the SIC and the Whitewater Group. A second cross-section (Figure 3B) summarizes the various ore forming environments within the entire Sudbury Structure. In the vicinity of the Sudbury Structure, the Superior Province rocks comprise the Levack Gneiss Complex; a suite of 2711 to 2642 millionyear old metamorphosed and intimately intercallated supracrustal and felsic and mafic intrusive rocks (Krogh et al. 1984; Wodicka and Card 1995). The gneisses are cut by the metagranitoid intrusive rocks of the Cartier batholith (2642 Ma; Meldrum et al. 1997). The Southern Province rocks are characterized by shallow marine and terrestrial metavolcanic and metasedimentary rocks of the Huronian Supergroup (2460 to 2300 Ma; see Trip 5, this volume; Easton and Bennett 2022) and associated mafic-ultramafic intrusions associated with the East Bull Lake intrusive suite (circa 2480 Ma; James et al. 2002) and the Nipissing gabbro suite (circa 2217 Ma; Davey et al. 2019). The character of the distinct Superior versus Southern province rocks is reflected in the Sudbury Structure rocks in proximity with either subprovince (i.e. the nature of Sudbury Breccia in the North Range (Superior Province) is distinct from that of the South Range (Southern Province)). Sudbury Breccia is typically a clast-supported breccia with a pseudotachylitic (glassy) to finely comminuted matrix that has been variably recrystallized. In proximity with the SIC, the matrix displays partial melt in pods and encased lithic fragments. Clasts are generally locally derived, with some exotic material, and are commonly equant with sub-rounded to sub-angular shapes. Economic mineralization can occur in proximity to, and associated, with Sudbury Breccia. The Frood–Stobie and Broken Hammer mines are examples of deposits hosted in, or associated, with Sudbury Breccia. Footwall Breccia (Stop 5: Access requires permission from City of Greater Sudbury) Footwall Breccia is an economically significant unit that hosts deposits such as at the Levack Mine (Figure 4). This style of mineralization is most prevalent in the North Range. Footwall Breccia is not uniformly distributed around the Sudbury Structure contact environment and is absent at the contact in many places. Unmineralized Footwall Breccia is present at Stop 5. Outside the Sudbury Structure proper (in what is referred to as the “Footwall” environment), rock types that resulted from the Sudbury Event include the Sudbury Breccia, Footwall Breccia and Offset Dikes (or Quartz Diorite - QD). Rocks within the Sudbury Structure include the Sudbury Igneous Complex (SIC) and the overlying Whitewater Group supracrustal rocks. 182 �Proceedings of the 68th ILSG Annual Meeting - Part 2 This enigmatic rock type occurs as “Megabreccia” and discontinuous sheets at the SIC contact, as well as part of some Offset Dikes (i.e. Whistle and Foy), and as intrusions into Felsic Norite. It is thickest in “embayments” (interpreted slump features along the over-steepened crater wall) where they provided suitable sites for sulphide accumulation (Figure 4A). The SIC, from bottom to top (or from Footwall contact to the base of the Whitewater Group), includes the Contact Sublayer/Offset Sublayer and the Main Mass (Norite, Quartz Gabbro and Granophyre) units. Contacts between the units are gradational and subject to interpretation. The character of the units in different regions within the Sudbury Structure can be distinct. Footwall Breccia is a heterolithic breccia with angular to sub-rounded fragments of varied sizes. Fragments are mostly locally derived and may include gabbro, diabase, mafic gneiss, and Huronian Supergroup sandstones – dependent upon the rock types of the footwall environment. Fragments of Sudbury Breccia have been identified in this unit, constraining the relative age of these two breccia types. The matrix is what makes this such an enigmatic lithology; it is crystalline rather than fragmental in nature. The matrix is subigneous at the SIC contact and has a metamorphic character with increased distance from the SIC. High temperatures at the SIC contact melted the breccia matrix resulting in the sub-igneous texture (Lakomy 1990). With increased distance from the contact, the temperature gradient resulted in a transition from an igneous-textured breccia matrix to a metamorphic-textured breccia matrix (Fedorowich et al. 2009). Contact Sublayer/Offset Sublayer (Stops 1, 5 and 9) Sublayer rock types are inclusion-bearing, igneous-textured, gabbronoritic or quartz dioritic bodies of varied compositions. The two main Sublayer environments are contact and offset dike. Both environments have potential to host economic Ni-Cu±PGE (Platinum Group Element) mineralization. The Discovery Outcrop (Stop 1) is an example of mineralized Contact Sublayer, and the recently re-opened Totten Mine is an example of economic mineralization in the Offset Sublayer environment. In the contact environment, rocks of the Sublayer are discontinuous lenses or sheets at the contact between the “Main Mass” of the SIC and the Footwall rocks. The presence and thickness of the Contact Sublayer appears to be controlled by the three-dimensional topography of the Footwall contact, with the Sublayer preferentially occurring in troughs (embayments and terraces) at the contact. The Contact Sublayer is a varied mixture of igneous silicate matrix, inclusions of silicate rock material, and magmatic Cu-Ni-Fe sulphides. Fragments include • footwall country rocks that can be identified and correlated with those directly observed in the Footwall, • xenoliths related to the SIC (similar to the basal Norite unit) and • exotic inclusions (anorthosite to dunite) with no known affiliation or source. Sudbury Igneous Complex (SIC) The Sudbury Igneous Complex (SIC) represents the crystalline rocks generated from the impact melt sheet. The current theory suggests that the melt sheet was created on the floor of the evolving impact crater within moments of the impact of the meteorite. This dynamic fluid evolved with the crater. As the crater stabilized the melt sheet differentiated and eventually solidified to form the SIC. The fall-back material (the Onaping Formation of the Whitewater Group) was originally ejected from the crater as it was excavating itself, but then was deposited on top of the melt sheet and was likely partially consumed and incorporated into it. In the Offset environment (Stop 9), the Sublayer is commonly referred to as Quartz Diorite or QD. This type of Sublayer occurs in thin, dike-like intrusions into the Footwall that occur in three morphologies 183 �Proceedings of the 68th ILSG Annual Meeting - Part 2 • • • predominate in this unit and intact crystalline plagioclase laths and prisms are present. radial offsets that extend out nearly orthogonally from the SIC, concentric offsets that are oriented subparallel to but outboard of the SIC, and discontinuous or breccia-hosted offsets. Whitewater Group (Stops 2, 6 and 7) The Whitewater Group encompasses an accumulation of nearly 3 km of supracrustal clastic and chemical “sedimentary” basin-fill rocks. The lowermost Onaping Formation is an atypical clastic rock of contentious origin; however, it is overlain by more recognisably sedimentary units that include the Vermilion, Onwatin and Chelmsford Formations. The Whitewater Group has been recognized only within the confines of the Sudbury Structure. However, it is likely that related clastic sediments (most particularly of the Onaping Formation) were deposited well outside the structure, but were either • not lithified and were dispersed and/or redeposited in other forms or • were eroded away or • have not yet been recognized or • a combination of the above scenarios. Most of the offset dikes are composite intrusions consisting of a central core of inclusion-bearing (commonly sulphide-rich) quartz diorite flanked by relatively inclusion- and sulphide-poor quartz diorite. Main Mass (Stop 4) The Main Mass of the SIC is a gradational series of igneous rocks that can be interpreted as having crystallized from a single differentiating magma (Pattison 2009; Lightfoot et al. 1997). At the base of the Main Mass, the Norite unit is a massive, medium- to coarse-grained, cumulate textured, two-pyroxene gabbronorite with varied amounts of quartz (up to 15%) and numerous accessory minerals. In the South Range, the unit is black as a result of ilmenite growth in plagioclase grains. In the North Range, the plagioclase grains do not typically contain ilmenite in their crystal structure and the rock colour is generally grey. “Mafic Norite” can locally be distinguished at the base of the Main Mass as modally more orthopyroxenerich and quartz-poor than the rest of the Norite, and geochemically by a sharp increase of magnesium. Onaping Formation (Stop 6) Originally described as a tuff, this contentious formation is a thick (1.4 km) accumulation of heterolithic breccias with igneous textures. The base of the formation is intruded by Granophyre of the underlying SIC and the top grades up into the carbonate and mudstone rocks of the overlying Vermilion and Onwatin Formations. The intermediate unit of the SIC Main Mass is a greenish-grey, two-pyroxene Quartz Gabbro. Contacts between the Quartz Gabbro and the flanking Norite and Granophyre are gradational. Quartz contents range from 15-60%. Passing stratigraphically up through the Quartz Gabbro there is an increase in the abundance of granophyric-textured quartzofeldspathic material. At Stop 4 this can be observed as an increase in the pink component of the rock relative to the green component. Recent detailed work by Ames et al. (2009) has led to the interpretation that the Onaping Formation formed in a dynamic and changing environment with an internal stratigraphy that captures the early evolution of the impact crater. “The formation represents a succession of glassrich breccias and coeval hypabyssal intrusions that have been hydrothermally altered to a variable degree.” (Ames et al. 2009) The most compelling evidence that distinguishes the rocks of the Onaping Formation from the volcaniclastic rocks that they texturally resemble is the presence of The upper and most felsic of the SIC Main Mass units is the Granophyre (formerly called micropegmatite). Granophyric intergrowth of quartz-plagioclase-potassium feldspar modally 184 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 1. Complex crater development after a meteorite impact (from Stoffler et al. 1980 in Taylor 1982). Figure 2. Sudbury Structure showing the surrounding geologic provinces and the 3 “ranges” of the SIC. Note the Grenville Province only formed in the area after the impact event (from Rousell and Card 2009). 185 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3. Schematic sections illustrating the A) stratigraphic and B) ore deposit environments associated with the Sudbury Structure. Note the difference in scale and the absence of the uppermost stratigraphy in B. SIC = Sudbury Igneous Complex, OF = Onaping Formation. Figure from Ames et al. (2008). 186 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4. A) An idealized representation of the embayment structures and their relationship to slump terraces of the Sudbury Structure. B) A section through the Levack No. 4 ore body (after Morrison 1984). Onwatin Formation (Stop 7) The Onwatin Formation consists of massive to laminated, carbonaceous and sulphidic argillite and siltstone with minor greywacke (Ames et al. 2009). This formation conformably overlies the Vermilion Formation. Its thickness has been interpreted as between 600 m and 1410 m. The formation is interpreted to have developed in a restricted, anoxic basin. shock metamorphosed quartz in the lithic fragments (French 1967; Peredery 1972). This is further supported by the presence of shock induced diamonds, and fullerenes (Becker et al. 1994; Masaitis et al. 1999), shatter cones in lithic fragments (Peredery 1972) and an iridium anomaly (Mungall et al. 2004). At Stop 6, we encounter the chaotic fragmental rocks of the Sandcherry Member of the Onaping Formation. “Bombs” with lithic cores and welldeveloped, banded glass rims are distinctive features of these outcrops. Chelmsford Formation (Stop 2) The extensively exposed Chelmsford Formation is the uppermost preserved unit of the Whitewater Group and occupies the elliptical core of the Sudbury Structure. This unit comprises turbiditic greywackes with a minor argillic component (Ames et al. 2009). The contact between the Chelmsford Formation and the underlying Onwatin Formation is gradational. Well-defined sedimentary structures including channels, flute casts, ripples, convolute laminations and iron-carbonate concretions can be observed in outcrops of the Chelmsford Formation. Vermilion Formation (no stop on this trip) The Vermilion Formation is a carbonaceous argillite unit that was subjected to significant syndepositional hydrothermal alteration. This alteration resulted in economic accumulations of zinc, lead, copper and silver mineralization (Errington and Vermilion mines). This unit, traceable around the Sudbury Structure in drillcore, is only exposed at surface in the southwest part of the basin in the vicinity of the pastproducing Errington Mine. The average thickness of the unit is 13.5 m (Stoness 1994). 187 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 5. MAP OF FIELD TRIP STOPS (geology from Ames et al. 2005) . Red stars indicate active mines A Copper Cliff North Mine B Creighton Mine C Totten Mine D McReedy West Mine E Fraser Mine F Coleman (Lower Coleman) Mine G Nickel Rim South Mine N Garson Mine 188 �Proceedings of the 68th ILSG Annual Meeting - Part 2 FIELD TRIP DETAILS Geological Maps Geological compilation maps covering all or parts of the area of the field trip include Ames et al. (2005), Dressler (1984b), and Card and Lumbers (1977). ROAD LOG Note: Caution should be taken when parking vehicles on the shoulder of roads and highways and when examining outcrops located along major roads along the field trip route. UTM co-ordinates are given in NAD 83 datum, zone 17; latitude and longitude are also provided in decimal degrees). 10.1 km — There are several routes to get to the Discovery Site from Science North. The “easiest” is: Turn left (west) out of Science North onto Ramsey Lake Road. Turn right (north) onto Paris Street (650m). Turn left onto Elm Street (2.75km) and continue on Elm Street/Regional Road 35 (6.7km) to turn off for Discovery Site on right (NE). Trail to site. Stop 1. Sudbury Discovery Site: A) Gossaneous outcrop of Sublayer along strike from original Discovery Outcrop. B) Close-up of stringer sulphides. Ruler marked in inches on the top, centimeters on the bottom. Stop 1 – Discovery Site UTM co-ordinates 495917E, 5151952N latitude-longitude 46.5211243N, 81.0532319W Protected site: NO HAMMERS Caution: Outcrop is adjacent to the active transcontinental line of the Canadian Pacific Railway (CPR). This discovery led to one of Canada’s most active staking-rushes. In just over 100 years the Sudbury Structure has generated tremendous wealth and produced huge masses of critically important raw materials for Ontario’s and the world’s manufacturing. Nickel-copper mineralization was first identified in the Sudbury area by Alexander Murray of the Geological Survey of Canada in 1856. However, it wasn’t until 1883 that the economic significance of the area was appreciated by the public. At this time construction of the Canadian Pacific Railway exposed rich coppernickel mineralization at a site close to the spot commemorated here. In the 1970s the actual discovery outcrop was mined and is now occupied by the water-filled Murray Pit, a few hundred metres from here. At this site, in line with the original discovery outcrop, gabbro-peridotite inclusion-bearing Contact Sublayer is exposed. Mineralization, though perhaps not as rich, is similar to that found in 1883. The Clarabelle No. 2 open-pit and the head frame of the inactive Murray Mine are visible from the Stop. 189 �Proceedings of the 68th ILSG Annual Meeting - Part 2 16.1 km — Northeast on Regional Road 35. Turnoff on right (north). Outcrop 90m west of parking area on north side of highway. Caution: Turn-off is a driveway. Stop 2 – Chelmsford Formation, Whitewater Group UTM co-ordinates 481688E, 5157337N latitude-longitude 46.5693493N, 81.2389532W Well-defined Bouma sequences can be identified in these outcrops. These turbidite beds display excellent sedimentary structures indicating tops, including fining upward sequences, crossbedding, flame structures and rip-ups. Distinctive, large, elongate “concretions” overprint bedding and can be traced within beds over many metres along the outcrop face. A well-developed, steeply dipping cleavage associated with the post-Sudbury event South Range Shear Zone (SRSZ) cuts bedding and appears to stretch the concretions into upright oval cross-sections. 23.5 km — Highway 144. Turn-off on right (NE). Outcrop 50m southeast of parking area on northeast side of the highway. Caution: Outcrops adjacent to Highway 144. Stop 3 – Levack Gneiss Complex with Matachewan dike and Pseudotachylite (Sudbury Breccia) Stop 2. A) Bedded Chelmsford Formation with flame structures. B) ‘Concretions’ (outlined in dashed red line) in Chelmsford Formation. UTM co-ordinates 464642E, 5164239N latitude-longitude 46.6307803N, 81.4619083W This location is part of what is considered the North Range footwall. It is outside the Sudbury Structure. Compositionally heterogeneous, complex gneiss of the Archean Levack gneiss complex is cut by pegmatitic dikes associated with the Archean Cartier Granite and a Paleoproterozoic Matachewan diabase dike. All these rocks are cut by narrow, impact-generated Sudbury Breccia veinlets. These outcrops display a complexity that is inherent in the footwall of the Sudbury Structure. Stop 3. Sudbury Breccia hosted in Levack gneiss. Ruler marked in inches on the top, centimeters on the bottom. 190 �Proceedings of the 68th ILSG Annual Meeting - Part 2 4.8 km — Return south on Highway 144 (4.7km). Turn left on onto Regional Road 8 (Onaping/Levack). Turn-off 160m from intersection on right. Outcrop 30m south of parking area; trail to outcrop. Stop 4 – Quartz Gabbro and Granophyre UTM co-ordinates 468456E, 5162737N latitude-longitude 46.6174539N, 81.4119820W This outcrop shows the gradational contact between the Quartz Gabbro and the Granophyre units of the Sudbury Igneous Complex. Rousell and Brown (2009) describe the granophyre on the southern part of the outcrop as being “approximately three parts micrographic intergrowth (potassium -feldspar and quartz) to one part tabular, plagioclase phenocrysts”. The contact between the granophyre and the quartz-gabbro to the north is described as “a gradual change in the micrographic intergrowth to plagioclase ratio. The contact is arbitrarily placed where the modal plagioclase exceeds that of intergrowth”. 1.9 km — North on Regional Road 8 (450m). Turn left on onto Onaping Drive (1.3km). Turn right onto unnamed road to gate for rehabilitated Onaping Landfill (150m). Outcrop is 250m north of parking area; walk on road to outcrop. Stop 4. A) Sudbury Igneous Complex, quartz gabbro with pink felspars. B) micropegmatite (granophyre) dike. of partial melt pods. Differences are best seen from a distance. Stop 5 – SIC Contact: Levack gneiss, Sublayer and Footwall Breccia (access requires permission from the City of Greater Sudbury) The outcrop on the west exemplifies barren Contact Sublayer. The exposure at the Discovery Site would look like this if it were not strongly mineralized. This rare, nearly barren outcrop of Sublayer shows fragments of norite > ultramafic >> felsic gneiss in an igneous crystalline matrix. UTM co-ordinates 467459E, 5163351N latitude-longitude 46.5892202N, 81.4298083W This stop is at the west end of the highly productive Levack–Onaping embayment structure that hosts the currently operating Fraser, Coleman, and McCreedy West Cu-Ni-PGE mines (Figure 5). Two large outcrop faces expose Contact Sublayer rocks (western exposure) and Footwall Breccia rocks (eastern exposure). The differences between them are subtle and are primarily based on the fragment composition, and the presence or absence The outcrop on the east has been interpreted as unmineralized to weakly mineralized Footwall Breccia. Here, unlike the Sublayer outcrop, the fragments are all Levack gneiss. The Footwall Breccia is locally cut by zones of partial melting characterized by acicular amphibole crystals (that may originally have been pyroxenes) in a pink matrix. 191 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 6 – Onaping Formation, Whitewater Group UTM co-ordinates 470724E, 5159607N latitude-longitude 46.583893N, 81.3821632W Please watch for highway traffic at this site! Caution: Stop requires walking across Highway 144. Traffic can be very heavy and fast-moving. This is the A.Y. Jackson Lookout, a scenic spot to enjoy High Falls. The outcrops to be examined are on the far side of Highway 144 and with the curves in the road and the speed of the traffic it is a dangerous crossing. Stop 5. A) Levack Gneiss Complex. B) Sublayer (dark norite breccia). Stop 6. Breccia of the Onaping Formation, Sandcherry Member. A) General aspect. B) Fragment within a streamlined glass rim. 8.0 km — Return along unnamed road to Onaping Drive (150m). Turn left. Turn right (south) onto Regional Road 8 (1.3km). Turn left (south) on onto Highway 144 (500m). Turn left into A.Y. Jackson Lookout (6.2km). Outcrop is along west side of highway; 230m north of turn-off; 160m from parking area. The Onaping Formation is the lowermost formation of the Whitewater Group. These rocks belong to the Sandcherry Member, a chaotic fragmental basin fill created from the fall-back of ejecta from the impact. Originally described as volcanic in origin, a few features identified in this 192 �Proceedings of the 68th ILSG Annual Meeting - Part 2 unit clearly distinguishes these rocks from the tuffs they closely resemble. The most compelling evidence is geochemical or microscopic: presence of an iridium anomaly (Mungall et al. 2004), presence of “impact diamonds” (Masaitis et al. 1999), and fragments containing shock metamorphosed quartz (French 1967; Peredery 1972). The planar deformation features in the quartz have only ever been identified in rocks affected by an impact and rocks affected by an atomic blast. Look for lithic cored “bombs” and fragments with glassy rims. These rocks do look very much like volcaniclastic material. Stop 7. Onwatin Formation slate. Bedding is near horizontal (white dashed line); cleavage is vertical. 31.4 km — East on Highway 144 to Highway 144S (Lively/Highway 17; 11.7km). Turn right to stay on Highway 144S. Turn left onto Regional Road 24 (Lively; 13.6km). Turn right onto Anderson Drive (5.9km). Turn left into Tom Davies Community Centre (150m). NOTE: the entrance to the community centre is after the building (one-way). Outcrop is on the east side of the building next to the highway. 5.4 km — Turn right (south) out of A.Y. Jackson Lookout onto Highway 144. Continue 5.4 km to stop on right (south). Park on the highway shoulder. Stop 7 – Onwatin Formation, Whitewater Group UTM co-ordinates 475298E, 5159222N latitude-longitude 46.5861084N, 81.3224356W The Onwatin Formation is approximately 600 m thick, comprising carbonaceous and pyritic argillite and minor wacke. The formation is thought to have been deposited in a deep restricted basin with stagnant and anoxic bottom waters. Estimates of the range of total carbon content of the Onwatin Formation vary, and the carbonaceous material may have originated as floating algal mats (Rousell 1984; Arengi 1977). Coleman (1905) estimated a range of 6.8 to 10% carbon in the Onwatin slate, whereas Arengi (1977) calculated 0.26 to 4.05% free carbon in the Onwatin Formation. Arengi (1977) concluded that the carbon occurs as elongated segmented structures, either as individuals or in clots, which resemble modern and fossil algal and fungal filaments. Stop 8 – Pseudotachylite: Sudbury Breccia UTM co-ordinates 488791E, 5141810N latitude-longitude 46.4297697N, 81.1458923W The glaciated outcrop shows fragment-rich pseudotachylite that is part of the South Range Breccia Belt. Fragments consist of lower Huronian Supergroup metasedimentary and metavolcanic rocks, which form the host rocks to the pseudotachylite dike. More specifically, the fragments consist of mafic metavolcanic rocks of the Elsie Mountain Formation, metarhyolite of the Copper Cliff Formation, and metapelite and metaquartzite of the McKim Formation. Fragments may show reaction rims and incipient marginal fragmentation. This suggests that the matrix was a melt. 193 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stop 9 – Copper Cliff Offset: Copper Cliff No. 1 Mine site UTM co-ordinates 494819E, 5146946N latitude-longitude 46.476050N, 81.0674912W Stop 8. Polymictic Sudbury Pseudotachylite. (GPS is 15 x 6.5 cm). The Copper Cliff Offset is one of the Sudbury Offset Dikes (also referred to as Quartz Diorite Dikes or QD) that host much of the economic copper-nickel-PGM mineralization in the Sudbury Camp. The offset dikes are part of the noritic sublayer, and the Copper Cliff Offset merges with the Main Mass norite in a funnel-shaped embayment to the north. This segment of the offset is about 8 km in length. At this site, the dike lies along the contact between the supracrustal rocks of the Elsie Mountain and Stobie Formations to the east and the Creighton Granite to the west (Cochrane 1984). The dike is cored by the subunit “IQD” or Inclusion Quartz Diorite. This subunit is the host of economic mineralization in the offset environment. Breccia– 10.0 km — Turn right onto Main Street from Anderson Drive (130m). Turn left onto Old Highway 17 (Regional Road 55 to Sudbury; 1.3km). Turn left onto Power Street (7.2km). Turn left onto Godfrey Drive (950m). Parking for stop on right (400m). Stop 85m east along powerline. The Copper Cliff No. 1 Mine was the first underground mine in Sudbury. Production began in 1886. The Copper Cliff North Mine continues production on the same Offset. Of other historical interest, the park on the other side of Godfrey Street was a roast bed. The area has been re-greened. 9.5 km — Return south on Godfrey Drive and turn left onto Balsam Street (130m). Turn left onto Regional Road 55 (1.3km). Continue on Regional Road 55, keeping left at cloverleaf for Big Nickel Drive (Regional Road 34). Regional Road 55 becomes Lorne Street after Big Nickel Drive cloverleaf. Turn left onto Martindale Road (3.6km). Left turn to remain on Martindale (450m). Slight left onto Walford Road at Regent Street intersection (1.4km). Turn left onto Paris Street (750m). Turn right onto Ramsey Lake Road (700m), Stop on right (1.5km). Outcrop begins 50m back (SW) from parking spot along footpath. Stop 9. Copper Cliff No. 1 Mine site. 194 �Proceedings of the 68th ILSG Annual Meeting - Part 2 workers, however, have noted that at other locations around the SIC, shatter cone orientations are in opposing directions (Dressler 1984a). The shatter cones, together with planar features in quartz, are believed to be evidence of a meteorite impact origin for the Sudbury Structure (e.g., Dietz 1964, 1972; French 1972). End of road log. References Ames, D.E., Davidson, A., Buckle, J. and Card, K., 2005. Geology, Sudbury bedrock compilation, Ontario; Geological Survey of Canada, Open file 4570, scale 1:50 000. Stop 10. Shattercones. Pen is 13 cm long. Stop 10– Shatter Cones Ames, D.E., Davidson, A. and Wodicka, N., 2008. Geology of the giant Sudbury polymetallic mining camp, Ontario, Canada; Economic. Geology, v.103, p.1057-1077. UTM co-ordinates 501535E, 5146083N latitude-longitude 46.4683165N, 80.9800069W Protected site: NO HAMMERS This outcrop has some of the best preserved and abundant shatter cones found in the Sudbury area. The host rock here is quartzite of the Mississagi Formation. The shatter cones appear as conical striated features whose surfaces are often micaceous and shiny. They range in length from a few centimetres to about a metre. Large cones may have numerous small cones along their flanks. Cones are exposed only where they control the outcrop surface. On other surfaces intersecting crescent-shaped fractures give the characteristic shattered appearance to the rock. They are best seen here when obliquely illuminated by the late afternoon sun. Ames, D.E., Stoness, J.A. and Rousell, D.H. 2009. Whitewater Group; in A Field Guide to the Geology of Sudbury, Ontario; Ontario Geological Survey, Open File Report 6243, p.37-44. Arengi, J.T. 1977. Sedimentary evolution of the Sudbury Basin; unpublished MSc thesis, University of Toronto, Toronto, Ontario, 141p. Becker, L., Bada, J.L., Winans, R.E. Hunt, J.E., Bunch, T.E. and French, B.M. 1994. Fullerenes in the 1.85billion-year-old Sudbury Impact Structure; Science, v.265 p.642-645 (Erratum v.265 p.1644) Bleeker, W. Kamo, S.L., Henning, S. and Lesher, M. 2022. A traverse across the Sudbury Impact Structure; in 68th Institute on Lake Superior Geology, Proceedings, v.68, pt.2, Guidebook, Field Trip 1. The distinctive fractures, termed shatter cones, form by passage of shock waves through rock, and are found at many astroblemes and “cryptoexplosion” structures. They have also been reported from localities with no known explosive associations. Geological mapping has shown that the SIC is surrounded by a belt of rocks more than 16 km wide containing shatter cones. In some locations, if the rocks are returned to their hypothetical orientation during the Sudbury Event, the apices of the shatter cones appear to point inward toward the basin (French 1972). Other Card, K.D. and Lumbers, S.B. 1977. Sudbury-Cobalt; Ontario Geological Survey, Map 2361, scale 1:253 440. Cochrane, L.B. 1984. Ore Deposits of the Copper Cliff Offset; in The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p 97-136. Coleman, A.P. 1905. The Sudbury Nickel Region; Report of the Ontario Bureau of Mines, v.14, pt.3, 183p. 195 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Canadian Shield; Economic Geology, v.97, p.15771606. Davey, S., Bleeker, W., Kamo, S.L., Davis, D.W, Easton, M.R. and Sutcliffe, R.H. 2019. Ni-Cu-PGE potential of the Nipissing sills as part of the ca. 2.2 Ga Ungava large igneous province; in Targeted Geoscience Initiative: 2018 report of activities; Geological Survey of Canada, Open File 8549, p.403-419. Krogh, T.E., Davis, D.W. and Corfu, F. 1984. Precise U-Pb Zircon and Baddleyite ages for the Sudbury areal in The geology and ore deposits of the Sudbury Structure, Ontario Geological Survey Special Volume 1; p. 431-446. 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.383-386. Lakomy, R. 1990. Implications for cratering mechanics from a study of the Footwall Breccia of the Sudbury impact structure, Canada; Meteorics, v.25, p 95-207. Lightfoot, P.C. 2017. Nickel sulfide ores and impact melts: Origin of the Sudbury Igneous Complex; Elsevier Inc., 662p. Dietz, R.S. 1964. Sudbury Structure as an astrobleme; Journal of Geology, v.72, p.412-434. Lightfoot, P.C., Doherty, W., Farrell, K., Keays, R.R., Moore, M. and Pekeski, D. 1997. Geochemistry of the Main Mass, Sublayer, Offsets, and Inclusions from the Sudbury Igneous Complex, Ontario; Ontario Geological Survey, Open File Report 5959, 231p. ——— 1972. Sudbury Astrobleme, splash emplaced sublayer and possible cosmogenic ores; in Geological Association of Canada, Special Paper 10, p.754-756. Dressler, B.O. 1984a. The effects of the Sudbury Event and the Intrusion of the Sudbury Igneous Complex on the Footwall Rocks of the Sudbury Structure; in The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p 97-136. Masaitis, V.L.; Shafranovsky, G.I.; Grieve, R.A.F.; Langenhorst, F.; Peredery, W.V.; Therriault, A. M.; Balmasov, E.L.; Fedorova, I.G.; Dressler, B.O. and Sharpton, V.L. 1999. Impact diamonds in the suevitic breccias of the Black Member of the Onaping Formation, Sudbury Structure, Ontario, Canada; in Large meteorite impacts and planetary evolution; II, Geological Society of America, Special Paper 339, p. 317-320. ——— 1984b. Sudbury geological compilation; Ontario Geological Survey, Map 2491, scale 1:50 000. Easton, R.M. and Bennett, G. 2022. A cross-section through the Huronian Supergroup at Elliot Lake, Ontario; in 68th Institute on Lake Superior Geology, Proceedings, v.68, pt.2, Guidebook, Field Trip 5, 57p. Meldrum, A., Abdel-Rahman, A.F., Martin, R.F. and Wodicka, N. 1997. The nature, age and petrogenesis of the Cartier Batholith, northern flank of the Sudbury Structure, Ontario; Canada; Precambrian Research, v.82, p.265–285. Fedorowich, J.S., Golightly, J.P. and Rousell, D.H. 2009. Breccias in the Footwall; in A Field Guide to the Geology of Sudbury, Ontario; Ontario Geological Survey, Open File Report 6243, p.45-55. Morrison, G.G. 1984. Morphological Features of the Sudbury Structure in Relation to an Impact Origin; in The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1; p. 513-520. French, B.M. 1967. Sudbury structure, Ontario: some petrographic evidence for origin by meteorite impact; Science, v.156, p.1094–1098. Mungall, J.E.; Ames, D.E. and Hanley, J.J. 2004. Geochemical evidence from the Sudbury Structure for crustal redistribution by large bolide impacts; Nature, v.429, p.546-548. ——— 1972. Shock-metamorphic features in the Sudbury Structure, Ontario: a review; in Geological Association of Canada, Special Paper 10, p.19-28. Pattison, E.F. 2009. Sudbury Igneous Complex; in A Field Guide to the Geology of Sudbury, Ontario; Ontario Geological Survey, Open File Report 6243, p.56-74. James, R.S., Easton, R.M., Peck, D.C. and Hrominchuk, J.L. 2002. The East Bull Lake intrusive suite: remnants of a ~2.48 Ga large igneous and metallogenic province in the Sudbury area of the 196 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Stöffler D., Gault D.E. and Reimold W.U. 1980. Cratering experiments in non-cohesive and weakly cohesive sand: Excavation mode and ejecta characteristics (abstract); in Papers Presented to the Conference on Multi-ring Basins: Formation and Evolution Lunar and Planetary Institute; p.89-91. Peredery, W.V. 1972. Chemistry of fluidal gases and melt bodies in the Onaping Formation, in New Developments in Sudbury Geology, Geological Association of Canada, Special Paper 10, p.49-59. Petrus, J.A., Ames, D.A. and Kamber, B.S. 2015. On the track of the elusive Sudbury impact: geochemical evidence for a chondrite or comet bolide; Terra Nova, v.27, p.9-20. Stoness, J.A. 1994. The stratigraphy, geochemistry and depositional environment of the Paleoproterozoic Vermilion and Onwatin formations, and their relationship to the Zn-Cu-Pb massive sulphide deposits in the Sudbury Basin; unpublished MSc thesis, Laurentian University, Sudbury, Ontario, 205p. Rousell, D.H. 1984. Onwatin and Chelmsford formations; in The Geology and Ore Deposits of the Sudbury Structure, Ontario Geological Survey, Special Volume 1, p.211-218. Rousell, D.H. and Brown, G.H., editors. 2009. A Field Guide to the Geology of Sudbury, Ontario; Ontario Geological Survey, Open File Report 6243, 200p. Taylor, S.R. 1982. Planetary Science: A Lunar Perspective. Lunar and Planetary Institute. 508p. Wodicka, N. and Card, K.D. 1995. Late Archean history of the Levack gneiss complex, southern Superior Province, Sudbury, Ontario: New evidence from UPb geochronology; in Precambrian ’95, Program with Abstracts, p.191. Rousell, D.H. and Card, K.D. 2009. Geological Setting; in A Field Guide to the Geology of Sudbury, Ontario; Ontario Geological Survey, Open File Report 6243, p.1-6. 197 �Proceedings of the 68th ILSG Annual Meeting - Part 2 198 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Field Trip 5 – An Overview of the Huronian Supergroup in the Elliot Lake area R.M. Easton Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, 933 Ramsey Lake Road, Sudbury, Ontario P3E 6B5 with contributions by G. Bennett Retired, formerly Resident Geologist, Sault Ste. Marie, Ontario Geological Survey Introduction The Huronian Supergroup is one of the Earth’s most studied sequences of rocks. Since the turn of the century the results of hundreds of studies of Huronian rocks have been published in scientific journals and government publications. These studies have led geoscientists, to present evidence for the Earth’s earliest glacial periods, the development of free oxygen in the atmosphere of the early Earth, the deposition of paleoplacer deposits of uranium, and evidence for plate tectonic activity during the Paleoproterozoic. Much of the evidence is based on rock exposures which will be visited during this field trip. The field trip uses road accessible outcrops. All of the road stops can be accessed using a 2-wheel drive vehicle. Unless otherwise stated, all UTM coordinates are in Zone 17, datum NAD 83, which is essentially equivalent to NAD WGS84. Safety Many of the field trip stops are located on highways that are especially busy during the summer season. Care should always be exercised when parking, exiting vehicles, and crossing the roads. Use of safety vests and/or bright clothing is recommended, in order to improve your visibility to motorists. Proterozoic rocks of the Canadian Shield in the Sudbury to Elliot Lake area are assigned to either the Paleoproterozoic Southern Province or the Mesoproterozoic Grenville Province (cf. Card et al. 1972; Wynne-Edwards 1972). The Southern Province in Ontario comprises Paleoproterozoic metasedimentary and metavolcanic rocks of the Huronian Supergroup and gabbroic intrusions of the Nipissing gabbro suite. Also included in the Southern Province are the Sudbury Igneous Complex (SIC), the Whitewater Group; plutonic and minor volcanic rocks of the Killarney Magmatic Belt; and rocks of the Sudbury mafic dike swarm (see Figure 2; Bennett et al. 1991). Table 1 summarizes the major geological events affecting the Superior, Southern and Grenville provinces in the Sault Ste. Marie to Sudbury area. Most of the trip routes are on Crown land or public roadways, but access is on or near private property in some cases. As in all such situations, please respect the property rights of others, so as to maintain good relationships, so that future access for geologists is not adversely affected. Purpose The transect through the Huronian Supergroup at Elliot Lake used by this field trip provides an opportunity to examine nearly all of the major units of the supergroup in a single day. These first pages are intended to give participants new to the Huronian Supergroup of Ontario a summary of what we think we know of these ancient rocks. Much of the material is borrowed from prior ILSG guidebooks (Bennett 2006; Bennett et al. 1997), but with updating and the inclusion of additional stops by R.M. Easton. The Huronian Supergroup (Robertson et al. 1969a; see Figure 3) is a sequence of variably metamorphosed Paleoproterozoic sedimentary and minor volcanic rocks that lie unconformably upon 199 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Archean rocks of the Superior Province. The Huronian rocks extend eastward from Lake Superior, along the north shore of Lake Huron to Sudbury, and then northward to the Noranda area of Quebec; a distance of approximately 450 km (Figures 1, 2). dominated phase at 2724.9±1.4 Ma (Optional Stop C) and a younger, calc-alkaline phase at 2686.5±1.1 Ma (Easton 2010, 2013a). Volcanism had ceased by 2674.8±0.8 Ma, the emplacement age of a granodiorite intrusion into the greenstone belt (Stop 1). The Huronian Supergroup attains its greatest thickness of 12,000 metres southeast of Sudbury (Debicki 1990). The sequence thins northward due to the wedging out of basal units, the thinning of the siliciclastic units, and erosion within the sequence (Roscoe 1969; Frarey and Roscoe 1970). The ages of both the greenstone belts and the Ramsey-Algoma granitoid complex are predominantly in the range of 2695 to 2650 Ma, in contrast to ages from the Abitibi greenstone belt which are typically in the range of 2740 to 2690 Ma. This distinction is important, as it allows for discrimination between potential source regions for the Huronian Supergroup strata, as described in detail in the section on “What Detrital Zircon studies tell us about the Source Region for the Huronian Supergroup”. The circa 2310 Ma U/Pb age from zircons in tuffaceous layers in the Bar River Formation (Hill et al. 2018; Rasmussen et al. 2013) places an upper age limit on the age of deposition of the bulk of Huronian Supergroup, with all sedimentation being completed well before emplacement of the Nipissing gabbroic intrusions at circa 2217 Ma (Davey et al. 2019; Corfu and Andrews 1986; Noble and Lightfoot 1992). The age of the rhyolites of the Copper Cliff Formation (circa 2450 Ma; Bleeker et al. 2015; Ketchum et al. 2013) is probably close to the start of initial deposition of the Huronian Supergroup. Huronian Magmatism Introduction Four distinct, more-or-less coeval (2480 to 2460 Ma), post-Kenoran igneous rock sequences are associated with the Huronian Supergroup:  Mafic dikes in the basement rocks but which do not cut the Huronian Supergroup (Matachewan and Hearst dike swarms). The Archean Basement  Igneous complexes, typically layered, of gabbro, gabbronorite and anorthosite (East Bull Lake intrusive suite). The basement to the Huronian Supergroup consists predominantly of rocks of the Ramsey– Algoma granitoid complex (Card 1979), which includes several large felsic batholiths, including the Cartier (2642 Ma: Meldrum et al. 1997) and Birch Lake (2651 Ma: Kamo 2006) granite batholiths. The batholiths were emplaced into a slightly older, granodiorite and quartz diorite intrusive and gneissic complex, which have ages of 2700 to 2675 Ma (Easton 2013a; Prevec 1993; Ontario Geological Survey, unpublished data).  Mafic to felsic volcanic flows (Elliot Lake Group).  Felsic plutons in the Southern and Grenville Provinces in the Sudbury area (2475 to 2460 Ma). Basement Dikes The granitoid rocks of the Ramsey-Algoma granitoid complex are intruded by mafic dikes of the Matachewan–Hearst swarm, which was likely emplaced in 2 main pulses, the first, earlier pulse at circa 2480 Ma is believed to have been coincident with emplacement of the East Bull Lake intrusive suite of layered intrusions (Krogh, et al. 1984; James et al. 2002a; Bleeker et al. 2015; Heaman 1997). The second and “main pulse” of the Matachewan–Hearst dike swarm occurred at circa 2460 Ma (Bleeker et al. 2015). These intrusive phases are younger than the few greenstone belts present in the Sudbury to Elliot Lake area; the largest of which is the Whiskey Lake greenstone belt, located south and southeast of Elliot Lake. Limited geochronology from the Whiskey Lake greenstone belt indicates 2 main periods of volcanism, an earlier, tholeiitic mafic 200 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 1. Sketch map showing the regional setting of the Sudbury Igneous Complex and the Huronian Supergroup (modified from Young et al. 2001). Layered Gabbro, Gabbronorite and Anorthosite Complexes et al. 2002a, 2002b; Easton et al. 2010). The three largest bodies contain platinum-group element mineralization near their basal contacts (cf. Easton et al. 2010). At the base of the Huronian Supergroup in the Elliot Lake, Agnew Lake and Sudbury areas are several layered gabbro to anorthosite intrusions referred to as the East Bull Lake intrusive suite (Peck et al. 1995; James et al. 2002a, 2002b). These bodies have ages of circa 2475 Ma (Clough and Hamilton 2017; Krogh et al. 1984) and appear to be slightly older than the rocks of the Elliot Lake Group. U/Pb chemical-abraded thermal-ionizationmass-spectrometry (CA-TIMS) zircon ages from the River Valley, Agnew and East Bull Lake intrusions cluster at 2475 Ma (Clough and Hamilton 2017; Easton et al. 2010). Although similar in age to the older phase of the Matachewan dike swarm, there are numerous occurrences of Matachewan dikes cutting intrusive rocks of the East Bull Lake intrusive suite (cf. Easton 2003, 2009; Easton et al. 2010) indicating that emplacement of all these mafic rocks was coeval. All the East Bull Lake intrusive suite bodies found to date have been emplaced into the Archean basement at, or just below, the Archean–Huronian Supergroup boundary. Intrusions of the East Bull Lake intrusive suite are characterized by the presence of anorthositic phases, and locally by a well-developed, primary rhythmic layering of alternating anorthositic and gabbroic layers (cf. James et al. 2002a, 2002b; Easton et al. 2010). Major intrusions of the East Bull Lake intrusive suite occur at Agnew Lake and East Bull Lake, with the largest body, the River Valley intrusion being found in the Grenville Province adjacent to the Grenville Front (cf. James 201 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 2. A time-rock chart for the southeast Lake Superior region (from Bennett 2006). 202 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 3. A stratigraphic column for the Elliot Lake fieldtrip transect (from Bennett 2006). 203 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 1. Timing of major geological events and summary of age constraints on the main rock units present in the Sudbury to Elliot Lake area. Event and/or Map Unit Age Constraint (Ma) Comment and/or Source Grenville dike swarm 586±4 Pegmatite vein emplacement 989±2 1000 to 990 Corfu and Easton (2000) 1040 to 1030 Carr et al. (2000) Age of peak metamorphism in the hangingwall of the Grenville Front tectonic zone Age of peak Grenvillian metamorphism in the Central Gneiss Belt Sudbury mafic dike swarm Killarney magmatic belt second-stage magmatism, coincident with magmatism in the Eastern Granite Rhyolite Province and in the Central Gneiss Belt Regional albitization metasomatic event Killarney magmatic belt volcanism and high-level plutonism Northwest-trending regional faults Penokean orogeny (folding and metamorphism of Huronian Supergroup rocks) Impact event and formation of Sudbury breccia Penokean arc formation and magmatism Thrust faulting F2 folding F1 folding Emplacement of Nipissing gabbro sills Huronian Supergroup sedimentation 1238±4 1471±3 Kamo, Krogh and Kumarapeli (1995) Corfu and Easton (2000) emplaced in or along northwest-trending faults in the Southern Province, deformed and metamorphosed in the Grenville Province. Krogh et al. (1987). van Breemen and Davidson (1988) U/Pb monazite, Schandl, Gorton and Davis (1994); fluid focussed along northwest faults 1740, 1747±3, 1749±12 van Breemen and Davidson (1988); Sullivan and Davidson (1993); Davidson and van Breemen (1994) Pre-1700, post-1850 Faults cut Sudbury Structure 1775±10 Peak deformation. Zi et al. (2022) ~1835 Peak metamorphism. Holm et al. (2001) 1701±4 1850±1 1890-1860, 1845-1830 post-F2 pre-regional faulting post-2200, pre-1700, pre 1850? pre-2200 2217±4 >2220 but <2460 Huronian Supergroup felsic volcanism and related plutonic rocks, including the Matachewan dike swarm ~2477 to 2375 (2450±25, 2460±20, 2477±9, 2415±5 Emplacement of East Bull Lake intrusive suite rocks 2475±2 Emplacement of orthopyroxene hornblendite bodies (East Bull Lake suite) 2468±5 Emplacement of alkali feldspar granite and megacrystic granodiorite near River Valley High-grade Archean metamorphism and migmatization Emplacement ages of Archean units in the Sudbury area 2660 to 2665 2647±4 2711±7 to 2642±1 204 Krogh, Davis and Corfu (1984); Davis (2008) Zi et al. (2022) Sudbury breccia localized along these faults, suggesting they are pre-Sudbury Structure Pre-regional faulting, Nipissing sills axial planar to folds Nipissing sills folded of intruded along folds Davey et al. (2019); Corfu and Andrews (1986); Noble and Lightfoot (1992) Youngest detrital grains in Bar River Fm are 2306 Ma (Hill, Davis and Cochran 2018) Krogh, Davis and Corfu (1984), Heaman (1997); Corfu and Easton (2000), Krogh, Kamo and Bohor (1996), Smith (2002); Bleeker et al. (2015) Heaman (geochronologist, University of Alberta, personal communication, 1999); Clough and Hamilton (2017) Corfu and Easton (2000) Bodies intrude Crerar and Pardo gneiss, Easton (2003) Krogh, Davis and Corfu (1984); Wodicka and Card (1995); Ames et al. (2005) Krogh, Davis and Corfu (1984); Wodicka and Card (1995); Chen, Krogh and Lumbers (1995); Meldrum et al. (1997); Ames et al. (2005) �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 4. The cyclicity of Huronian Supergroup rocks (from Bennett 2006). Figure 5. Paleocurrent directions in the Matinenda and the Mississagi Formations (from Bennett 2006). 205 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The Huronian Supergroup Marie area and between 110-300 m thick in the Thessalon area. It consists of 2 distinctive rock types: an upper, well-sorted, grey sandstone and a clast-supported polymictic conglomerate (Bennett et al. 1991). (Figures 3, 6, 7). The Livingstone Creek Formation has not been recognized east of the Quirke Lake Syncline (Bennett 2006). Introduction The Huronian Supergroup is subdivided into 4 groups (Robertson et al. 1969a, 1969b), which in ascending stratigraphic order are: the Elliot Lake Group, Hough Lake Group, Quirke Lake Group and Cobalt Group (Figure 3). Formations of the 3 upper groups, with the exception of the Serpent Formation of the Quirke Lake Group, show regional stratigraphic continuity, and display a remarkable cyclicity of lithological units (Figure 4). Each cycle begins with matrix-supported conglomerate (diamictite), followed by mudstone, siltstone and/or limestone, and capped by a thick sequence of crossbedded, coarse sandstone (Bennett et al. 1991). Paleocurrent studies (cf. Long 1976, 1978; McDowell 1957) have shown flow to the south to southeast, with southeast being the predominant direction (Figure 5). In most areas, clast-supported, polymictic conglomerate is predominant in the lower sections of the Livingstone Creek Formation. Cobble- to boulder-sized clasts, generally of grey granitic rocks and minor mafic plutonic and metamorphic rocks, are set in a sparse matrix of grey coarse arkose or arkosic grit. Bennett (2006) notes that he had not observed clasts of Huronian Supergroup volcanic rocks in these conglomerates. Locally, thin units of crossbedded, grey arkose are interbedded with the conglomerate (Frarey 1977, Bennett et al. 1991). The granitic mega-clasts of the conglomerate member are predominately pale grey in contrast with the predominantly reddish hues of the underlying Archean basement rocks. Some of the granitic megaclasts in the predominately grey conglomerate near the south end of Pine Ridge Road near Thessalon show the distinct texture of the typical Archean, massive, pink, potassium feldspar-megacrystic granite – but with only a hint of the pink color in the phenocrysts. The grey colour of the Livingstone Creek Formation conglomerates appears to be due to the reduction of ferric iron in the feldspars of the granitic clasts, and not a result of differing provenance as some have suggested. This conclusion is supported by Bennett’s (2006) observation that granitic rocks in a “paleosol zone” a few metres to a few tens of metres below the base of the Livingstone Creek Formation commonly are grey as well. Conglomerate units (e.g., Ramsey Lake, Bruce and Gowganda formations) in each of the cycles have been interpreted as being glaciogenic in origin, likely deposited in a marine environment adjacent to an ice shelf. The siltstone and sandstone units are interpreted to represent deposition during warmer intraglacial or post-glacial periods in either fluvial or marine environments (cf. Junnila and Young 1995; Fralick and Miall 1989). The Elliot Lake Group The Elliot Lake Group differs from the overlying Huronian groups in that:  its internal stratigraphy is generally discontinuous and less extensive.  it does not have the diamictite-mudstone-sandstone sequence of the overlying groups.  it contains the only important uranium deposits and the only volcanic rocks of the Huronian Supergroup. The grey, sandstone member of the Livingstone Creek Formation can be distinguished from most Huronian Supergroup sandstones by its uniform grain size (fine- to medium-sand). In addition, mudstone and pebbly units are lacking in the sandstone member of the Livingstone Creek  st formations have disconformable surfaces. The Livingstone Creek Formation Conglomerates and sandstones of the Livingstone Creek Formation (Frarey 1967, 1977) form the lowermost Huronian Supergroup unit The formation is at least 400 m thick in the Sault Ste. 206 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Formation, and carbonate occurs along the foreset beds of the well-developed trough crossbeds. The Huronian Supergroup Volcanic Rocks of the Sault Ste Marie-Elliot Lake area In addition to the well-known exposures in the Thessalon and Sault Ste Marie areas, other possible occurrences of the Livingstone Creek Formation include: Frarey (1967) named the Huronian volcanic rocks in the Thessalon and Sault Ste. Marie areas that overlie the Livingstone Creek Formation as the “Thessalon Formation” (Figures 3, 6).  grey sandstone and conglomerate near Crazy Lake in Nicholas Township (Bennett 1978; Bottrill 1971). Bennett (2006), based on his examination of all known exposures of Huronian Supergroup volcanic rocks as well as all available drill-core and drill-hole logs reporting volcanic rocks in the Elliot Lake–Sault Ste Marie area, concluded that there is no credible evidence for more than one period of Huronian Supergroup volcanism in the Elliot Lake–Sault Ste. Marie area and that all the Huronian Supergroup volcanic rocks west of the nose of the Quirke Lake Syncline are stratigraphically correlative with the Thessalon Formation (Figure 6) (Bennett 1978, 2006; Bennett et al. 1991).  a basal grey sandstone unit directly underlying the Matinenda Formation in Haughton Township (Bennett et al. 1991).  an area of clast-supported, grey granite-cobble conglomerate near Samried Lake (Jackson 2001). The clast size, local source and low stratigraphic position of the Livingstone Creek Formation conglomerates are consistent with deposition as an alluvial fan(s). The uniform, fine- to mediumgrained sand of the trough cross-bedded sandstone member suggests a different, although likely related, more distal depositional environment than that of the conglomerate. The sandstone member may represent deposition by median streams flowing in a fault-bounded valley with walls of Archean rocks partly covered by alluvial fans (Bennett et al. 1991). The well-sorted nature of the sandstone suggests an aeolian component or even aeolian deposition as proposed by Meyer (1983). Unfortunately, none of the many attempts to obtain an absolute age determination from rocks of the Thessalon Formation have been successful. Nonetheless, there is no reason to think that the age of the Thessalon Formation differs significantly from that of the Copper Cliff Formation (circa 2460 Ma). The maximum thickness of the Thessalon Formation in the Sault Ste. Marie area is approximately 650 to 820 m (Frarey 1977). Diamond drilling has indicated at least 670 m of Thessalon Formation volcanic rock under Lake Huron south of the town of Thessalon, and the formation may be up to 1080 m thick north of Bass Lake in Aberdeen Township (Bennett 2006). The Huronian Supergroup Volcanic Rocks – Overview As noted by Easton (2013a), the transition between dominantly subaerial and dominantly submarine deposition of metavolcanic rocks of the Huronian Supergroup occurs in the Elliot Lake area, and this change in depositional environment may have been significant with respect to sedimentary depositional environments in the Elliot Lake area itself. Thus, it may be no coincidence that Huronian Supergroup mafic metavolcanic rocks are found in proximity to all the past-producing uranium mines and current prospects in the Elliot Lake area. In the Sault Ste. Marie, Thessalon and Aberdeen Lake areas, the Thessalon Formation can be subdivided into an upper tholeiitic basalt unit and a lower complex or “mixed member” (Bennett et al. 1991), which includes fractionated rocks, including basaltic andesite, tholeiitic andesite, mugearite, hawaiite and rhyolite. Magnesium-rich basalts with some of the chemical characteristics of komatiites are present in the Dollyberry Lake, Pecors Lake and Thessalon 207 �Proceedings of the 68th ILSG Annual Meeting - Part 2 areas. The lower flow sequences show much lower concentrations of Ni, Cr, and contain higher amounts of Ti and P than do the upper flows (cf. Ketchum et al. 2013). In the Dollyberry Lake area, the upper basalts appear to be missing, possibly due to erosion (Bennett 2006). In most areas the metamorphic grade of the Thessalon Formation flows is lower greenschist facies, although the presence of albite and primary pyroxene, along with elevated sodium contents, indicates sub-greenschist facies at the northern end of the Duncan volcanic belt near Sault Ste. Marie (Bennett 2006). A comprehensive analysis of the geochemistry of the Thessalon Formation volcanic rocks between Sault Ste. Marie and Thessalon, concluded that the lavas are divisible into 7 distinct units based on mapping, petrography and major and trace element geochemistry (Figure 8) (Tomlinson 1996; Ketchum et al. 2013). The 7 units were grouped into 2 “lava series”. The upper lava series (unit 6) is equivalent to the upper tholeiitic basalt sequence of Bennett et al. (1991). The lower lava series (units 1-5, of Tomlinson 1996 and Ketchum et al. 2013) consists mainly of basaltic andesite with subordinate, local rhyolite, mugearite, andesite and high magnesium basalt flows; and corresponds to the “diverse member”of Bennett et al. (1991). Amygdules of epidote, chlorite, calcite, quartz and stilpnomelane in complex zonal arrangements are common. Flattened chlorite-filled amygdules a centimetre or less across are a distinctive feature of most mafic flows of the Thessalon Formation. Pillow structures are rare but are present in most areas. Scoriaceous flow-tops and crosscutting breccias are commonly filled with a fine-grained mixture of quartz and grey to red secondary albite (Bennett 2006). With regard to the geochemistry and tectonic setting of the Thessalon Formation, Tomlinson (1996) stated “that the source of the lavas was metasomatized upper mantle rather than a deep mantle or plume component. Structural subsidence patterns in the Archean basement (Zolnai et al. 1984) are thought to be responsible for lithospheric stretching, in-turn causing mantle upwelling, episodic partial melting and volcanism.” Bennett et al. (1991) proposed that the upper, basaltic flows of the Thessalon Formation (upper lava series) probably represent part of a continental flood basalt sequence, whereas the diverse member (lower lava series) appears to have erupted from central vents. The volcanic rocks of the Quirke Lake Syncline display lithological and geochemical similarities to the lower lava series of the Thessalon Formation west of the Quirke Lake Syncline (Bennett 2006). Syndepositional features present where sedimentary rocks of the Livingstone Creek Formation infill fractures in the underlying Archean basement indicate that initially volcanism was a consequence of rifting. In active rifts, a single uplift and melting event occurs as a plume impacts the lithosphere, but in passive rifts uplift and melting are episodic. In addition, the presence of multiple erosional surfaces in the Elliot Lake Group indicate that many episodes of uplift occurred (Bennett 2006). Therefore, the Huronian rifting event can best be characterized as a typical passive rifting event (Tomlinson 1996). This is consistent with Jolly’s (1987) conclusion that the Thessalon Formation is a continental flood basalt sequence related to continental rifting. The upper, tholeiitic basalt flows of the Thessalon Formation (upper lava series of Tomlinson 1996; Ketchum et al. 2013) are almost uniformly greenish-grey fine- to medium-grained tholeiitic metabasalt. The essential minerals are albite, actinolite, chlorite, clinozoisite, epidote and Fe-Ti oxide. Primary clinopyroxene is present in only a few samples of basalt from the Sault Ste. Marie area. The andesitic rocks of the lower lava series are typically darker and contain stilpnomelane and biotite with green pleochroism (Fe+3 rich?) and albite and actinolite. Quartz is a minor component of the basaltic and andesitic types (Bennett 2006). 208 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 7. Stratigraphic Relationships in the Elliot Lake Group (from Bennett 2006). Figure 8. Internal stratigraphy of the Thessalon Formation, Elliot Lake Group, in the Sault Ste. Marie to Elliot Lake area (from Bennett 2006, modified from Tomlinson 1996)). 209 �Proceedings of the 68th ILSG Annual Meeting - Part 2 higher Ni and lower V contents compared to EMF1. On a Nb/Yb versus Th/Yb diagram (Pearce element diagram), the EMF-2 mafic rocks define a distinct trend separate to that of the EMF-1 and Stobie Formation mafic volcanic rocks. Comparing data from Gordon (2021) with that of Ketchum et al. (2013) on volcanic rocks in the Thessalon area (~200 km to the west), the mafic rocks of EMF-1 and the Stobie Formation are comparable to a Thessalon upper basalt-basaltic andesite unit. The rocks of EMF-2 closely resemble a stratigraphically lower basalt-andesite Thessalon unit. Based on the geochemical similarity with the Thessalon volcanic rocks, similar magmatic processes were likely responsible for the generation of the mafic volcanic rocks of the Thessalon, Elsie Mountain and Stobie Formations. Huronian Supergroup volcanic rocks of the Sudbury Area The volcanic rocks of the Sudbury area differ in terms of internal stratigraphy, overall thickness, and depositional environment from the Huronian Supergoup volcanic rocks in the Sault Ste. Marie– Elliot Lake area. The Huronian Supergroup volcanic sequence in the Sudbury area has been subdivided into the predominately mafic Elsie Mountain (1000 m thick) and Stobie Formations (1500 m thick), and the felsic Copper Cliff Formation (760 m thick). The depositional age of the Copper Cliff Formation is circa 2460 Ma (Beeker et al. 2015; Ketchum et al. 2013; Krogh et al. 1984), and it is likely that the Creighton Granite was coeval with the Copper Cliff Formation (Beeker et al. 2015). The volcanic rocks in the Sudbury area show evidence of submarine eruption from fault-controlled vents along the edge of a depositional basin into which arkosic sandstones were transported from the Archean granitic terrain to the north, with turbidites being deposited from the basin margins (Card 1978a). Sedimentary rocks associated with the Thessalon Formation Some early reports referred to the presence of quartz-pebble conglomerate in the Livingstone Creek Formation, however, this could not be confirmed by Bennett (2006). At many locations, however, a thin unit (< 1 m) of radioactive, pyritic, quartz-pebble conglomerate overlain by a few metres of coarse arkose sand was found to lie upon the Archean basement, or directly atop the Livingstone Creek Formation, where the latter is present. In Duncan Township in the Sault Ste. Marie–Thessalon area, this quartz-pebble conglomerate-arkose sequence occurs in the lower flows of the Thessalon Formation. (Hay 1963; Bennett et al. 1978; Meyer 1983) (Figures 6, 7). Based on recent mapping in the Sudbury area, a preliminary geochemical characterization of the mafic volcanic rocks of the Elsie Mountain and Stobie Formations has been presented (Gordon (2021). Mafic volcanic rocks of the Elsie Mountain Formation are divided into EMF-1 and EMF-2, which are geochemically distinct from each other. Mafic rocks of EMF-1, along with those in the Stobie Formation, are high-Fe tholeiitic basalts. EMF-1 samples in the Elsie Mountain Formation represent the basal lavas, but they are not the most primitive lavas as they have lower MgO, Ni and Cr and higher SiO2 compared mafic lavas in the overlying Stobie Formation. Mafic rocks of EMF1 and the Stobie Formation exhibit similar primitive mantle-normalized trace element profiles, characterized by LREE enrichment relative to HREE and negative Nb-Ta-Ti anomalies. EMF-2 mafic volcanic rocks are tholeiitic andesites with distinct primitive mantlenormalized REE profiles characterized by strongly depleted HREE. EMF-2 mafic rocks also have Bennett et al. (1991) proposed that the conglomerate-arkose units indicate the presence of a disconformity between the volcanic rocks of the Thessalon Formation and the Livingstone Creek Formation. The wide distribution of these units (Figure 6) suggests that they reflect an early erosional period of regional extent. The resistant nature of the mineral assemblage in the conglomerate (assuming an oxygen deficient atmosphere) points to a period of extreme weathering. Some of this quartz-rich regolith may still be visible as a quartz breccia atop the granitic basement west of Highway 639 (Optional Stop G). 210 �Proceedings of the 68th ILSG Annual Meeting - Part 2 (1989) suggested that the Matinenda Formation was deposited from shallow braided streams flowing down a south dipping paleoslope which underwent tilting to the southeast during deposition. Kimberly et al. (1980) reported that the uraniferous conglomerates contained almost no magnetite-ilmenite and had very high K/Na ratios. These are also features of the paleosols beneath the Huronian Supergroup and suggest that the sediment of the Matinenda Formation was formed by the intense weathering of a granitic source terrain, as proposed by Roscoe (1969). The Matinenda Formation, Elliot Lake Group The Matinenda Formation of the Elliot Lake Group is a sequence of arenites and intercalated quartz-pebble conglomerates which host the once strategically important uranium deposits of the Elliot Lake camp where it lies on Huronian Supergroup volcanic rocks and/or the Archean basement (Roscoe 1969; Robertson 1968, 1976). In the Thessalon, Sault Ste. Marie, and Sudbury areas, it consists predominantly of fine-to mediumgrained, subarkose to subwacke, and is probably less than 50 m thick (Bennett 1978). In Haughton Township, the Matinenda Formation lies upon grey sandstones equated with the Livingstone Creek Formation (Bennett 2006) (Figure 7). Two southeast trending ore zones were recognized since the early days of uranium mining in the Elliot Lake camp. The Nordic zone, east of the City of Elliot Lake is about 1.6 km (1 mi) wide and 5.6 km (3.6 mi) long. The Quirke Zone, in the Quirke Lake area, is about 3.2 km (2mi) wide and 9 km (6 mi) long. Basement paleotopography is thought to have had a determining influence on the position and orientation of the zones. Ore grade (approximately 0.1 % U3O8 (850 ppm U)) conglomerate occurs as persistent lenses with individual units up to 4.5 m thick. The uraniferous quartz-pebble conglomerates are commonly well developed at the base of the Matinenda Formation but also occur in the arkose up to 45 m above the base (Roscoe 1969, Robertson 1968). Total mine production from 1955 to 1990 was 160 million tonnes of ore averaging 896 g/t U3O8 for a total uranium metal production of 164,000 tonnes. In the Sudbury area, clastic units correlated with the Matinenda Formation thin rapidly eastward and are intercalated with the mainly metavolcanic rocks of the Stobie Formation and mudstones of the McKim Formation (Card 1978a). The most abundant rock type of the Matinenda Formation in the Elliot Lake area is generally described as medium- to coarse-grained subarkose, arkose and grit consisting of poorly-sorted quartz and feldspar grains set in a matrix of sericite and comminuted rock and mineral and the fragments. The ratio of potassium feldspar to plagioclase feldspar is about 8:1. Minor constituents are pyrite, calcite, chlorite, zircon and rarely, leucoxenecoated iron oxide and monazite. Varied amounts of sericite give the sandstones a green, apple green or greenish-yellow colouration. Well-sorted, quartzpebble conglomerate beds, with well-rounded pebbles and cobbles of quartz and chert, and pebbly subarkose units, are scattered throughout the coarse subarkose of the Matinenda Formation, but are more common near the base, in what has been termed the “floater-reef zone” (Robertson 1968; Pienaar 1963) (Stop 3). The quartz-pebble conglomerate consists mainly of well-rounded, pale- to dark-grey, quartz and chert pebbles in a matrix of pyrite, quartz and/or feldspar grit and sericite. Minable units contain about 15% pyrite. Radioactive minerals include uraninite, brannerite. and uranothorite (Roscoe 1969). Monazite and zircon are common heavy minerals. The sedimentological and mineralogical features of the uranium-bearing zones of the Elliot Lake camp are generally believed to support a modified paleoplacer origin of the ores as outlined by Roscoe (1969). Advocates of this hypothesis propose that prior to the accumulation of significant free oxygen Trough crossbedding, scour, and fill structures are common in the subarkose units (Robertson 1968; Roscoe 1969). Paleocurrent studies have established a northwest source area for the sediment of the Matinenda Formation (McDowell 1957; Long 1978; Figure 5). Fralick and Miall 211 �Proceedings of the 68th ILSG Annual Meeting - Part 2 in the Earth’s atmosphere, southeastward flowing streams carried quartz, pyrite and uraniferous minerals released by the extensive weathering of the Ramsey-Algoma granitoid terrain and deposited them in southeast-trending units constrained by the basement topography (Bennett 2006). McKim Formation. North of the Murray Fault the McKim Formation rarely exceeds a few hundred metres in thickness, whereas south it is at least 2400 metres thick (Debicki 1990). Card (1978) suggested that the change from laminated siltstone in the west to more wacke in the east indicated a change from more distal to proximal facies, in turn suggesting more tectonic activity and possibly a source for the McKim Formation sediments from the east. Fralick and Miall (1989) concluded that the McKim Formation in the Elliot Lake area represented a marine transgression that gradually drowned the Matinenda Formation fluvial plain. The McKim Formation The McKim Formation is the uppermost formation of the Elliot Lake Group. In diamond drill core, the contact between the Matinenda Formation and the McKim Formation is interfingering over 1 to 3 metres, consisting of clean sandstone of the Matinenda Formation and mudstone and wacke of the McKim Formation. Aweres Formation In the Sault Ste. Marie area, the Aweres Formation, a 1700 m thick sequence of conglomerate and sandstone (McConnell 1927), unconformably overlies mafic volcanic rocks of the Thessalon Formation. The internal stratigraphy and rock types of the Aweres Formation are consistent with deposition as an alluvial fan (Bennett 2006). Robertson (1968) gives a thickness of 0 to 100 m for the McKim Formation on the south limb of the Quirke Lake syncline. It is missing on the north limb. The McKim Formation is thickest in the Sudbury area, where it is up to 2400 m thick. Card et al. (1977) recognized 3 facies within the McKim Formation: The base of the formation consists almost entirely of mafic volcanic clasts whereas higher levels show a progressive increase in granitic clasts. The uppermost rocks of the Aweres Formation south of Aweres Lake are mainly arkose with thin pebble conglomerate beds. The lithological variation with stratigraphic height indicates the continual erosion of an uplifted, faultbounded, plateau of Huronian Supergoup volcanic rocks (Bennett 2006).  the “quartz sandstone” facies is equivalent to the Matinenda Formation, and represents thin interbeds of meta-quartz arenite and minor metaconglomerate in the 2 other main facies of the McKim Formation.  the “greywacke” facies of interbedded metawacke, metasiltstone and thin bedded mudstone and siltstone. Ripple marks, cross-laminations, graded beds and Bouma cycles are common. Bedding varies from a few centimetres to 50 cm thick.  the “laminated argillite facies” of thin-bedded mudstone and siltstone, with occasional beds of finegrained wacke. Bedding is commonly less than 1 cm, and rarely exceeds 10 cm. The distinct lithology of the Aweres Formation prevents its direct correlation with other Huronian Supergroup rocks. The upper surface is partly faultbounded. but is unconformably overlain by the Gowganda Formation on Highway 556. It is possible that the Aweres Formation is an erosional remnant of a more extensive alluvial fan system that extended in a more-or-less northeast direction beyond the present northern limit of the Hough Lake and Quirke Lake Groups. The Mississagi Formation may represent a distal depositional environment compared to that of the Aweres Formation (Bennett 2006). Where more highly metamorphosed, rocks of the laminated argillite facies, are best described as metapelites. The metapelites are characterized by high Al2O3 contents (20-25 weight %, Easton 2006b; Card et al. 1977), and moderate Fe/Mg ratios (~2), which is probably why metamorphic porphyroblast development is generally restricted to the laminated argillite facies. The Murray Fault appears to have exerted an important influence on the deposition of the 212 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Hough Lake Group Stratigraphic Relationships within the Elliot Lake Group, Sault Ste Marie-Elliot Lake area Introduction The stratigraphic relationship between the Matinenda, Thessalon, and Livingstone Creek Formations is revealed on a rock face near the northern boundary of Haughton Township about 30 km (18 miles) north of the town of Thessalon (Bennett 2006). Here pyritic quartz-pebble conglomerate of the Matinenda Formation directly overlies an apple-green paleosol on grey, finegrained sandstone correlated with the Livingstone Creek Formation (Bennett 2006). The Hough Lake Group (Robertson et al. 1969a; Roscoe 1969) is lowest of the 3 groups of the Huronian Supergroup that display the cyclic deposition of diamictite; mudstone-siltstone and/or carbonate; and arenite. Each cycle is generally thought to represent a sequence of glaciogenic– marine–fluvial and/or shallow marine deposition (Figure 4). Ramsay Lake Formation The Ramsay Lake Formation is the lowermost unit of the Hough Lake Group and is the oldest of 3 such conglomerate units that define the base of Hough Lake, Quirke Lake and Cobalt Groups (Roscoe 1969; Pienaar 1963) (Figure 3, 4). About 600 m northwest of the aforementioned occurrence, arkose and quartz-pebble conglomerate of the Matinenda Formation disconformably overlie a steeply dipping, eaststriking, mafic dike; the upper few metres of which is a sericite-leucoxene paleosol. The dike cuts grey sandstone and apple-green paleosol of the Livingstone Creek Formation (Bennett 2006). The Ramsay Lake Formation is a widespread, but relatively thin unit. In the Elliot Lake area, the Ramsay Lake Formation ranges from zero to just over 30 m thick (based on diamond drill logs from the Sault Ste. Marie District Geologist’s Office). The Ramsay Lake Formation is 70 to 170 m thick in the Sudbury-Manitoulin area (Card 1978a). Less than 2 km south of the above location Chandler (1976) identified a fault-bounded block of Thessalon Formation volcanic rocks with a minimum thickness of approximately 500 m. The mafic dike referred to above was a feeder for Thessalon flows, since the Thessalon Formation is the only known igneous activity at this stratigraphic level (Bennett 2006). Matrix-supported polymictic conglomerate (diamictite) is the most abundant rock type in the formation, especially near the base. Cobbles in the lowermost few metres usually reflect the underlying rock type (Robertson 1968; Parviainen 1973). Locally, minor amounts of mudstone, wacke and arenite are present. Subround to well-rounded pebbles and cobbles of grey granitic rocks and angular to rounded clasts of dark green to black volcanic rocks generally form less than 30 volume percent of the diamictite. The dark matrix consists of quartz, feldspar, chlorite, muscovite-sericiteillite and pyrite (Parvianen 1973). The above observations show that there was a period of volcanic activity, and a period of erosion, separating the Matinenda and the Livingstone Creek Formations. Since paleosols occur upon Huronian Supergroup flows in the Elliot Lake area, the sub-Matinenda unconformity seen in Haughton Township likely extends east to the Quirke Lake Syncline. In addition, the outcrop pattern of the volcanic rocks on geological maps also suggests that the volcanic rocks are erosional remnants preserved in basement depressions (Bennett 2006). The Thessalon Formation may have once extended beyond the limit suggested from its present outcrop distribution (Figure 6), especially if it were a continental flood basalt sequence. Although some writers have argued for a debris flow origin, most writers now accept the Ramsay Lake Formation as having a significant glaciogenic component (cf. Roscoe 1969; Robertson 1976). Fralick and Maill (1989) identified an ice-proximal association of pebbly sandstone and diamictite; subaqueous gravity flows and ice rainout deposits; and ice-proximal, fluvial outwash deposits. 213 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Mississagi Formation; but Long (1978) argued that the abundance of mud-grade matrix in the immature arenites, the predominance of unimodal paleocurrent directions, and the lack of quartz arenites argued against a marine environment for the Mississagi Formation. Long (1978) concluded that the Mississagi Formation was deposited from braided streams with low to intermediate sinuosity and high width to depth ratios. Pecors Formation The Ramsay Lake Formation is conformably overlain by a sequence of generally dark, bedded and laminated wacke, mudstone, siltstone and sandstone (Roscoe 1969). The Pecors Formation is 30 m thick at Quirke Lake (Robertson 1968) but is as much as 900 m thick south of the Murray Fault in the Sudbury area (Card 1978a). It was not identified in the area between Thessalon and Sault Ste. Marie (Frarey 1977). Ripple marks, graded bedding, cross-laminations parallel laminations, ball and pillow structures, clastic dikes and slumpage features have been reported in the formation. The basal part of the formation is commonly laminated, resembling varves, and in places has dropstones (Robertson 1968; Parvianen 1973). Partial Bouma sequences are common (Card 1978a; Robertson 1976). The Pecors Formation is the result of transgressive units formed in deep water by turbidity currents (Card 1978a). The presence of dropstones is evidence of a cold paleoclimate and provides supporting evidence for the glaciogenic origin of the underlying Ramsay Lake Formation. Beds are commonly about a metre thick but can range from a few centimetres to over 4 m thick. Trough cross-stratification and ripple crossstratification are common sedimentary structures (Long 1978). Cross-stratified beds may show grain-size gradation (McDowell 1957). Long (1978) measured over 2500 cross-stratified units in the Mississagi Formation (Figure 5) and recognized 2 major stream systems: a stream system flowing southeast to east from the Sault Ste. Marie area, which joined a stream system flowing southwest from the Cobalt Plain, thereby forming a southward flowing system southwest of the Sudbury area. These observations suggest that the area now occupied by the Sudbury Igneous Complex was elevated during the time of Mississagi Formation deposition (Long 1978). Mississagi Formation The Mississagi Formation is a thick sequence of predominantly grey, arenitic rocks extending most of the length of the Huronian Supergroup outcrop belt. In the Quirke Lake syncline, the Mississagi Formation is 344 to 704 m thick. South of the Murray Fault the formation is notably thicker, being more than 3000 m thick in the Sudbury area (Card 1978a; Long 1978). Quirke Lake Group Bruce Formation The Bruce Formation extends from the Garden River Indian Reserve near Sault Ste. Marie to about 70 km northeast of Sudbury. It consists mainly of matrix-supported and minor clast-supported conglomerate. Pebbly wacke, arkose, wacke and siltstone are locally present. By far the most dominant rock type in the Mississagi Formation is moderately well-sorted, medium- to coarse-grained subarkose and arkose. Small to medium quartz and/or chert pebble conglomerate is a minor component of the formation; but is more common in the western and northeastern parts of the Huronian belt. Finegrained pyrite along forsets commonly results in rusty staining of outcrops. Greenish, sericitic units form relatively thin planar-bedded units between crossbedded sandstones. Palonen (1973) provided evidence supporting a marine origin for the The Bruce Formation is from 79 to 12 m thick in the Elliot Lake area and is 26 to37 m thick under the main part of the Quirke Lake Syncline (Robertson 1968). Pebble- to boulder-sized, angular to subrounded clasts generally consist of pale-grey granitic rocks, Archean supracrustal rocks and fine-grained mafic clasts. The upper parts of the formation may contain up to 5% carbonate (Robertson 1968). 214 �Proceedings of the 68th ILSG Annual Meeting - Part 2 The Bruce Formation is generally interpreted as a tillite with minor beds and lenses of glacially derived sandstone. Dropstones have been observed in laminated units (Robertson 1968). Serpent Formation The Serpent Formation occurs throughout much of the Huronian Supergroup outcrop belt; however, it was locally removed by erosion during a period of tectonic activity preceding deposition of the Gowganda Formation of the Cobalt Group. Thickness estimates range from 150 to 1500 m (Bennett et al. 1991). According to Robertson (1968), nowhere in the Blind River–Elliot Lake area is there evidence that the total thickness of the Serpent Formation has been preserved. Casshyap (1969) concluded that the formation was deposited from terrestrial wet-base glaciers. Sims et al. (1981), however, proposed that the Bruce Formation represents an accumulation of debris flows released by normal faulting, a sudden increase in paleoslope, and a sudden increase in water depth. This is consistent with observations made in Porter and Vernon townships showing considerable down-cutting, ranging from 5 to 30 m, of the Bruce Formation into the underlying Mississagi Formation (Easton 2005). The Serpent Formation is mainly fine- to medium-grained, quartz arenite and arkose. Conglomeratic units have been noted, especially near its base. Carbonate is a significant component near the base of the formation in the Elliot Lake area (Robertson 1968). Planar and festoon crossbedding, rip-up clasts, fine-laminations, and mud cracks have been reported. Long (1976) proposed that the Serpent Formation was deposited in a distal braided stream environment with calcareous units representing a sabkha environment. Young (1982) noted that the presence of very large crossbeds and a bimodal size distribution suggest aeolian processes may have been active, at least locally. Espanola Formation The Espanola Formation is the only widespread carbonate unit of the Huronian Supergroup. It is a present from Sault Ste. Marie to the Maple Mountain area, approximately 70 km northeast of Sudbury. Its widespread distribution and distinctive lithology make it the most useful stratigraphic marker unit in the Huronian Supergroup. In the Elliot Lake area, the Espanola Formation can be subdivided into 3 members: a lower limestone member, a middle siltstone- arenite member and an upper dolomite member (Robertson 1968). The latter generally contains 3% to 4% total iron which gives it a distinct brownish hue on weathered surfaces. Contacts between members tend to be gradational. All 3 members are thinly bedded to laminated. The threefold subdivision is less well developed south of the Murray Fault (Young 1982). Cobalt Group Gowganda Formation The Gowganda Formation is a complex sequence of conglomerates, sandstones, siltstones and mudstones, and is the lowermost formation of the Cobalt Group. Its thickness ranges from 1070 m in the Sault Ste. Marie area; to 970 to 1150 m around Whitefish Falls on the north shore of Lake Huron; and from 950 to 2700 m near Sudbury. Near Dunlop Lake, in the Elliot Lake area, the Gowganda Formation is about 600 m thick. Intraformational breccias, mud cracks, ripplemarks, flame structures and ball-and-pillow structures are common sedimentary features. Hofmann et al. (1980) described stromatolites in the Espanola Formation on Quirke Lake. All these features suggest deposition in quiet shallow waters with carbonate deposition being interrupted by influx of fine-grained sediment. Matrix-supported conglomerates are common, especially in the lower parts of the formation. However, these are commonly intercalated with clast-supported conglomerates and sandstone units. Laminated mudstones and siltstone are especially prominent in the upper parts of the Gowganda Formation. Many occurrences of ice-rafted 215 �Proceedings of the 68th ILSG Annual Meeting - Part 2 dropstones have been reported in laminated mudstone-siltstone units. Individual units are generally relatively thin and discontinuous making subdivision of the Gowganda Formation difficult except in well-exposed areas. northern Ontario. It is overwhelmingly an arenite sequence, with local siltstone units present in lower parts of the formation. It is up to 2500 m thick near Sault Ste. Marie and in the LaCloche Syncline, southwest of Sudbury. It is up to 2300 m thick in the Cobalt Basin. Most granitic clasts in Gowganda Formation conglomerates have a distinctly pinkish or reddish hue, in comparison to the grey, granitic clasts in the matrix-supported conglomerates of the stratigraphically lower Ramsay Lake and Bruce Formations. Pink- and red-hued sandstones also first make their appearance in the formation. Roscoe (1969) pointed out the appearance of red coloration (i.e. ferric iron) just above the basal units of the Gowganda Formation, and argued that it represents the appearance of free oxygen in the Earth’s atmosphere, and a change from the previously reducing atmospheric conditions that allowed the accumulation of easily oxidized minerals such as pyrite and uraninite. Roscoe (1969) did, however, emphasize that glaciation is only one of several processes likely responsible for the deposition of the Gowganda Formation. In general, the lower part of the Lorrain Formation is dominated by pink, arkosic sandstone; the middle by hematite-rich subarkose and quartzarenite; and the upper part by pale grey to white mature, quartz-arenite. A distinctive jasper-pebble conglomerate found in the Sault Ste. Marie area is a popular decorative stone, known locally as “pudding stone”. Previously these jasper clasts were thought to be derived from banded iron formations from the Abitibi greenstone belt, however, Bleeker (2018) observed that the jasper clasts of the puddingstone are often angular (i.e. more or less proximal), that they suddenly become a dominant clast type (again suggesting proximal); that there are few if any real banded iron formation clasts; and that the jasper clasts are extremely fine-grained and delicately textured and they do not contain magnetite, unlike banded iron formation samples from the Abitibi greenstone belt which are noticeably more recrystallized. Thus, Bleeker (2018) concluded that the jasper clasts of the Lorrain Formation puddingstone are not of Archean derivation, but 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 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 depositional environment of the diamictites in the Gowganda Formation have been the subject of discussion since Coleman (1905) proposed a glacial origin for these matrix-supported conglomerates. Many subsequent writers including Ovenshine (1965), Casshyap (1969), Lindsay (1971) and Young and Nesbitt (1985) also have supported a glacial, glacial-marine, or glaciolacustrine, origin for the Gowganda Formation diamictites. Card (1968) concluded that, although glaciation may have supplied coarse detritus to the basin initially, debris flows and turbidity currents, related to vertical tectonic movement, may better explain the thickness variations, rock associations and distribution of units in the Gowganda Formation in the Sudbury– Manitoulin area. Lorrain Formation The presence of aluminous minerals is a characteristic feature of the uppermost quartzarenites of the Lorrain Formation. Diaspore and kaolinite are common in the Sault Ste. Marie area The Lorrain Formation is generally wellexposed throughout most of the Huronian Supergroup outcrop belt, where it commonly forms the background to some of the most scenic views in 216 �Proceedings of the 68th ILSG Annual Meeting - Part 2 and north of Elliot Lake (Wood 1973) whereas kyanite, andalusite and kaolinite occur as metamorphosed equivalents in the LaCloche Lake– Killarney area (Card 1978a). Young (1973) and Wood (1973) interpreted the presence of diaspore and kaolinite as the result of post-depositional, insitu, alteration of feldspar under hot and humid climatic conditions. Formation. He also described hematite ooliths and the abundance of grains in the 0.02 to 0.05 mm range, a relatively uncommon grain size in sedimentary rocks. Since this size is found in loess deposits, Wood (1973) proposed that the quartz silt of the Gordon Lake Formation was formed by glacial action, and then carried by the wind and deposited in a tidal flat environment. The presence of abundant detrital hematite in the Lorrain Formation and the occurrence of monazitebearing quartz-pebble conglomerate north of Elliot Lake, have been interpreted by Frarey and Roscoe (1970) as indicating an oxidizing environment. Bar River Formation The Bar River Formation is the uppermost formation of the Huronian Supergroup. It is characterized by quartz-arenite with minor ferruginous arenite and siltstone. It is approximately 300 m thick in the Flack Lake area, north of Elliot Lake. Wright and Rust (1985) concluded that the Bar River Formation was deposited in a tidal environment. Planar and trough crossbedding are common, as are ripple marks and other primary depositional structures. There is no consensus as to the depositional environment of the Lorrain Formation. Most of the sedimentary structures present can be found in either shallow marine or fluvial environments. Wood (1973), Young (1973) and Frarey (1977) favored a fluviatile setting, whereas Pettijohn (1970) supported a marine setting. Card (1976) proposed that the Lorrain Formation resulted from near-shore coastal shelf deposition during episodic marine transgression and regression. Nipissing Intrusions Sills, and minor dikes and cone sheets, of gabbro, diabase and granophyre, commonly referred in the older literature (pre-1995) as “Nipissing diabase”, are the most widespread igneous rocks associated with the Huronian Supergroup. Nipissing intrusions are widely and evenly distributed throughout the Huronian Supergroup outcrop belt but, and with few exceptions, are not recognized in the Archean Ramsey-Algoma granitoid terrane. Individual intrusions may be up to several hundred metres thick and extend over a strike-length area of 10s of kilometres. There is no current consensus on the tectonic setting for emplacement of the Nipissing intrusions. Gordon Lake Formation The Gordon Lake Formation displays a gradational contact with the underlying Lorrain Formation. It is composed predominantly of variegated mudstone, siltstone, chert and minor fine-grained sandstone. The Gordon Lake Formation in the Flack Lake area is subdivided into a lower member of reddish arenite, siltstone, and chert with anhydrite and gypsum nodules; a middle member of green siltstone and mudstone; and an upper member of reddish mudstone, siltstone and chert (Robertson 1986). Sedimentary features include small-scale crossbeds, ripple marks and desiccation cracks. Olivine-bearing hypersthene gabbro, gabbro, feldspathic pyroxenite, two-pyroxene quartz gabbro, hornblende gabbro, granophyric gabbro and granophyre have been identified in Nipissing intrusions. Many Nipissing sills are characterized by chilled margins 50 cm to 5 m wide, overlain by 10-20 m of quartz gabbro, then 100-500 m of hypersthene-poor gabbro-norite and vari-textured diabase (Lightfoot and Naldrett 1996). Some features of the Gordon Lake Formation are unique in the Huronian Supergroup. Wood (1973) noted the abundance of feldspar in marked contrast to rocks of the immediately underlying Lorrain Baddeleyite and/or zircon from Nipissing gabbro sills in the Gowganda area, the Sudbury area, the 217 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Agnew Lake area and north of Thessalon, have all given ages between 2214 and 2219 Ma (Davey et al. 2019; Noble and Lightfoot 1992, Corfu and Andrews 1986; Easton, unpublished data 2021). Buchan and Card (1985) report that paleomagnetic data suggests at least 2 periods of Nipissing intrusive activity. If so, then the 2 paleomagnetic poles formed in a short timespan of circa 5 million years.  colour variations  destruction of primary rock textures accompanied by the development of soil textures  destruction of primary minerals with formation of clay minerals or metamorphic equivalents  dikes of material from overlying sediment washed down into desiccation cracks in the soil  rip-up clasts of overlying sediments Well-preserved paleosols below the Matinenda Formation in the Elliot Lake area have been described by many workers (Roscoe 1969; Pienaar 1963; Robertson 1968; Frarey and Roscoe 1970; Gay and Grandstaff 1980; Kimberly et al. 1984; G-Farrow and Mossman 1988; Prasad and Roscoe 1991; Sutton and Maynard 1992, 1993; Easton 2013b). Lightfoot and Naldrett (1996) discuss the geochemical characteristics of the Nipissing magmas and the potential for platinum group metal deposits. They concluded that parental magmas of remarkably uniform composition underwent in-situ contamination and differentiation in the intrusions. In addition to nickel-copper-PGE mineralization (Jobin-Bevans 2014, 2016; Jobin-Bevans et al. 1998), a spatial association between Nipissing intrusions and 5-element vein-type mineralization has long been recognized, especially in the Cobalt area (cf. Fyon et al. 1992). Bennett et al. (1991) proposed that there are 3 disconformities or unconformities in the Elliot Lake Group, which all have the potential for paleosol development (Figure 7). These are in descending stratigraphic order the sub-Matinenda disconformity, the sub-Thessalon Formation disconformity, and the sub-Livingstone Creek Formation unconformity. Huronian Paleosols and Evidence for Oxygen Accumulation in the Huronian Atmosphere Paleosol Evidence The sub-Livingstone Creek Formation unconformity is the lowest unconformity and is the only entirely sub-Huronian Supergroup unconformity (Figure 3, 7). This unconformity is exposed in the Thessalon area, where the upper few metres of the Archean granitic rocks can be seen to progress from angular, slightly rotated blocks, separated by grey grit and fine-grained sandstone, upward, to more rounded boulders with a higher proportion of finer clastic material (Collins 1925). This zone may be termed a “paleo-regolith”, since there is little or no obvious development of the yellow, sericitic paleosol commonly found in the younger, sub-Matinenda paleosols. It has long been recognized that the study of paleosols (ancient soil profiles) beneath the Huronian Supergroup could provide information pertaining to the development the Earth’s atmosphere and climate during the Proterozoic. Since iron is much less soluble in the ferric state than when in the ferrous state, the behavior of iron in paleosols should provide some indication of the oxygen partial pressure of the environment. Many of the best descriptions of Precambrian paleosols have been from those associated with the Huronian Supergroup unconformity (Gall 1992). Grandstaff et al. (1986) identified 8 features of paleosols; most of which have been described in paleosols beneath the Huronian Supergroup. These features are: Prasad and Roscoe (1996) described 2 paleosols in the same diamond drill core from the Denison Mine at Elliot Lake. One was found above Huronian Supergoup volcanic rocks and another, less well-developed paleosol, was found upon Archean tonalite below a short section of quartzpebble conglomerate and grit below the 9 m thick  stratiform  relatively thin (<20 m)  transitional lower boundary-sharp upper boundary 218 �Proceedings of the 68th ILSG Annual Meeting - Part 2 volcanic unit (Prasad, personal communication, 1997 in Bennett 2006). Gay and Grandstaff (1980) concluded that the upward increase in iron content indicated the presence of free oxygen in early Huronian atmosphere, although at approximately 1% of the present level. They also suggested that the loss of iron in most Huronian paleosols could be due to local reducing environments. Some writers have concluded that the increase in potash (as sericite) in Huronian paleosols is due largely to diagenetic and metamorphic processes that may mask the environmental and hydrologic conditions operative during paleosol development (cf. Gay and Grandstaff 1980; G-Farrow and Mossman 1988). The best developed, and most studied, Huronian paleosols have been found directly below the Matinenda Formation. On mafic rocks, the subMatinenda paleosols can generally be recognized by the presence of an upper, distinctly apple-green to yellowish, sericitic zone which grades downward, over a few centimetres to several metres, to a black, fine-grained, chlorite-rich eluvial zone up to several metres thick. Abundant pseudomorphs of titanium oxide after ilmenite are a feature of paleosols on mafic igneous rocks. Ripup clasts of sericitic paleosol are commonly found in the lower few metres of the overlying Matinenda Formation. Prasad and Roscoe (1996) report significant amounts of carbonate and pyrite in subMatinenda paleosols in the Elliot Lake area. The mineralogy and geochemistry of subLorrain Formation paleosols described by Rainbird et al. (1990) and Sutton and Maynard (1992, 1993) commonly show an enrichment of Fe+3 relative to Fe+2 without a significant loss of total iron. Hematite is a common mineral in the upper parts of sub-Lorrain paleosols, in contrast to the presence of pyrite in sub-Matinenda paleosols. In this respect, the sub-Lorrain paleosols resemble many postGeon 23 paleosols and are consistent with weathering in an oxidizing atmosphere (Prasad and Roscoe 1996; Rainbird et al. 1990). The uppermost sections of sub-Matinenda Formation paleosols developed on Archean granitic rocks is generally an apple-green to yellowish rock composed mainly of quartz and sericite (Robertson 1968; Gay and Grandstaff 1980; Sutton and Maynard 1992), Where the texture of the protolith is well preserved, but the original mineralogy is replaced, the paleosol may be termed a saprolith (Rainbird et al, 1990). The chlorite-rich eluvial zone of paleosols on granitic rocks is generally lacking or relatively thin. Other Evidence Since pyrite and uraninite are unstable under oxidizing conditions, the abundance of detrital pyrite and uraninite in the paleoplacer uranium ore zones in the Matinenda Formation provide evidence for an oxygen deficient atmosphere during weathering, transport and deposition of early Huronian Supergroup sediments. Sub-Matinenda paleosols commonly show the pronounced loss of sodium typical of most paleosols. Calcium and magnesium are also depleted, but there is generally a large increase in potassium content (Gay and Grandstaff, 1980). In most cases, iron and manganese are depleted in the upper parts of the paleosol. This is held to provide evidence of weathering in a reducing environment. Gay and Grandstaff (1980), however, noted an upward increase in total iron in paleosol from the Pronto Mine area. Easton (2013b) locally reported chemical compositions approaching that of a bauxite developed over a mafic substrate in diamond drill core from Elliot Lake. In contrast to the common red beds of more modern clastic sequences (post-Geon 23), sandstones and most granitic clasts below the Cobalt Group are almost all drab coloured despite the abundance of red and pink granitic rocks in the source area (Roscoe 1969, 1973). Frarey and Roscoe (1970) proposed that the drab colour of lower Huronian Supergroup clastic rocks is due to the lack of free oxygen in the atmosphere during their deposition. 219 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Red-hued, hematite-bearing rocks, which Roscoe (1969) proposed mark the presence of an oxidizing atmosphere, make an appearance with the Gowganda Formation of the Cobalt Group, and are important in parts of the Lorrain and Gordon Lake Formations. fault, it is generally interpreted as an inverted growth fault; i.e. an early listric normal fault active during sedimentation; which during later compression became a thrust or reverse fault (Card 1978a; Jackson 2001; Zolnai et al. 1984). The rocks of the Huronian Supergroup have been subjected to several deformational events (Table 1). This is particularly evident south of the Murray Fault. In the Whitefish Falls area, Young and Nesbitt (1985) concluded that some large-scale folding was related to syn-depositional and/or postdepositional deformation of unconsolidated sediment. Early syndepositional deformation is indicated also the unconformity beneath the Gowganda Formation; and the presence of ragged, slumped contacts and large slump blocks along major faults (Card 1978a; Young 1983). Not all workers, however, accept the above explanation for the preservation of uraninite and pyrite, and the observed change in colour with stratigraphic position. For example, Ohmoto (1996) has stated “the loss of total iron in paleosols of all ages is not due to a reducing atmosphere but to the reductive dissolution of ferric hydroxides under an oxic atmosphere”. Regional Tectonic Patterns and Metamorphism Major structures in the Huronian Supergroup outcrop belt follow 2 trends: 1) west-northwest trending folds and faults in the Sault Ste. MarieElliot Lake area; and 2) east to northeast striking folds and faults in the Sudbury-Manitoulin area south of the Murray fault. These 2 orientations are associated with differing fold styles, metamorphic grade and metamorphic fabric. Convincing evidence of at least one important pre-Nipissing (circa 2217 Ma) deformational event, historical assigned to the apocryphal Blezardian orogeny (Stockwell 1982), comes from the observation that Nipissing bodies in the Sudbury-Whitefish Falls area transect axial surfaces of major folds (Card 1978a). Such relationships are not observed north of the Murray Fault (Jackson 2001; Robertson 1964). North of the Murray Fault, Nipissing sills tend to occupy structures parallel to the axial plane of the Chiblow anticline, suggesting pre-Nipissing. folding. Easton (2006a), in the Porter-Vernon area north of Espanola and north of the Murray Fault, noted that at least 2 periods of folding are present, roughly orthogonal to one another — the resulting interference forms a dome-and-basin pattern (Figure 10). F1 folds Nipissing gabbro intrusions in the lowermost part of the stratigraphy (in the Hough Lake and Quirke Lake Groups), whereas Nipissing gabbro appears to be emplaced along fractures related to F2 axial planes (all groups). This suggests either multiple periods of gabbro emplacement, or more likely, that gabbro emplacement occurred syn-folding. In either case, folding cannot be significantly younger than the emplacement age of the Nipissing intrusions. In the Sault Ste Marie–Elliot Lake area, fault and fold structures generally trend west-northwest to northwest. Folds are generally upright, and open, with gentle, variably-plunging hinges. There is only weak development of minor tectonic structures; metamorphic grade is subgreenschist (Figure 9, Card 1978b). The major structural features of the Elliot Lake area include a gently south-dipping homocline south of the Flack Lake fault, the open fold of the Quirke Lake Syncline, and the Chiblow Anticline to the south of the Quirke Lake Syncline. In the Elliot Lake area, neither Jackson (2001) nor Easton (2009, 2013a) found any evidence of a detachment at, or near, the basement-cover interface. Notable changes occur across the northeasttrending faults of the Murray Fault system, the most significant structural feature of the Huronian Supergroup outcrop belt. Because many formations show a significant increase in thickness south of the 220 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 9. Metamorphism of the Huronian Supergroup. Figure from Card (1978b). Jackson (2001) also noted evidence of preNipissing faults north of the Murray Fault zone. Easton (2006a), again in the Porter-Vernon area, made the same observation, and recognized at least 5 major fault sets, 4 of which are post-folding.  The earliest faults are north-trending and juxtapose Archean granitic basement against Huronian Supergroup strata. These faults appear to have been fluid conduits, as indicated by the presence of large quartz vein systems and microbrecciation in the Archean basement, and hydrothermal annealing of quartz in sedimentary rocks adjacent to the faults. obscured by subsequent vertical movement, and the fact that these faults are the loci for the development of extensive zones of Sudbury breccia. The localization of Sudbury breccia along this fault set suggests that it may have developed at circa 1850 Ma due to the Sudbury impact.  Finally, significant vertical displacement, occurs along a major set of closely spaced northwesttrending faults. Some of these faults are the loci of Sudbury swarm diabase dikes (circa 1240 Ma), which are undeformed and unmetamorphosed, suggesting that this fault set formed between 1850 and 1240 Ma.  East-northeast faults also juxtapose Huronian Supergroup strata against basement rocks, but are post- F1 folding, with both vertical and lateral movement. They may be associated with a set of north to northeast, dominantly normal faults, which may have an older thrust component.  Most significant in terms of map pattern, at least in the southern part of the Porter-Vernon area nearest the Murray fault system, are east to east-northeast normal faults across which major changes in stratigraphic level occur. There may be a thrust component to these faults, but if so, it has been Following emplacement of the Nipissing intrusions, but prior to the emplacement of the Sudbury Igneous Complex (1850 Ma), there was further deformation and regional metamorphism. Rb/Sr isotopic studies of Huronian Supergroup metasedimentary rocks indicate that metamorphic resetting occurred at 1900-1850 Ma (Fairbairn et al. 1969). This age range is correlative with the Penokean Orogeny of Michigan and Minnesota (Sims et al. 1981), which has a peak metamorphic age of circa 1835 Ma (Holm et al. 2001). 221 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 10. Simplified geological map of the northeast shore of Agnew Lake, showing the distribution of fold styles within Porter and southern Vernon townships. The contact between the Mississagi and Bruce formations has been highlighted to illustrate the fold pattern, and units stratigraphically above the Bruce Formation are shown by a pattern. Between the Cameron Creek and Midport faults, the area is dominated by a dome and basin geometry, indicating the presence of 2 fold generations, with approximately perpendicular axial planes. North of the Midport fault, the early, north-oriented fold style (F1) dominates. Figure from Easton (2005). 222 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Recently, increasing evidence suggests, that even in its type area, the Penokean orogeny is not as significant a regional event as had been previously thought. In fact, there is increasing evidence that the Yavapai orogeny (Geon 17) may be responsible for much of the deformation previously attributed to the Penokean (cf. Zi et al. 2022; Holm et al., 2018; Schulz and Bjornerud 2018; Raharimahefa et al. 2014). Clearly additional work on the timing and extent of deformation affecting the Huronian Supergroup is needed in Ontario, especially away from Sudbury where there has been considerable resetting of isotopic systems by the Sudbury impact event. it have deformed Huronian Supergroup sedimentary rocks that were not yet deposited. Jackson (2001) also proposed that the “inverted growth-fault” model, as applied by Zolnai et al. (1984) to structural-stratigraphic relationships in the Huronian Supergroup outcrop belt may, in some cases, be interpreted as thrust faults with flats following depositional boundaries, and ramps that cut up through the stratigraphic section. Given the data available, neither model could be rejected for major northwest-trending faults in the Sault Ste. Marie area (Jackson 2001). Jackson (1994) points out that the curvature of the Flack Lake fault is in the opposite direction to that expected if it is a thrust fault, as proposed by Zolnai et al. (1984). After the emplacement of the Sudbury Igneous Complex (SIC) at 1850 Ma (Davis 2008), and the deposition of the Whitewater Group, there is evidence of further deformation and low-grade metamorphism of the Huronian Supergroup, followed by intrusion of granite plutons at circa 1740 Ma and circa 1450 Ma, predominately in the Killarney area and in what is now the Grenville Province in the Sudbury area. The intensity of postSIC deformation and retrograde meta-morphism increase south of the Murray Fault, especially in the area between the SIC and the Grenville Front. The Murray Fault system separates moderately deformed, low grade metamorphic rocks to the north from multi-deformed, higher-grade rocks of the Sudbury–Manitoulin area to the south. The Sudbury–Manitoulin area is characterized by open to sub-isoclinal, flattened buckle folds with upright to northward overturned axial surfaces. Elongate domes and basin are formed by reversals in plunge. Penetrative axial place cleavage and steeply plunging rodding and/or mineral lineations are well developed. More than one age of major and minor structures can be discerned south of the Murray Fault (Jackson 2001). A study of magnetic fabrics, strain patterns, and microstructures in granitoid rocks of the Creighton and Murray granites and their Huronian Supergroup host rocks (Riller 1996) lent credence to the concept of a pre-2220 Ma “Blezardian orogeny” (Stockwell 1982). Riller (1996) concluded that major folding and amphibolite facies regional metamorphism in the Sudbury area was coeval with the emplacement of the Creighton the Murray granites, which at the time yielded discordant upper intercept ages of 2333+33/-22 Ma (Frarey et al. 1982) and 2388+20/-13 Ma (Krogh et al. 1984). Subsequent work on the Creighton granite by Bleeker et al. (2015), Kenny et al. (2017) and Smith (2002), and on the Murray granite by Krogh et al. (1996), have yielded ages of 2460±20 Ma and 2477±9 Ma, respectively, suggesting that the Blezardian orogeny was not widespread, nor could Metamorphism south of the Murray Fault ranges from lower greenschist to lower amphibolite facies (Figure 9). Rocks of higher metamorphic grade occur in 2 zones or nodes, one along the Murray Fault system itself and another northwest of the Grenville Front. Both zones coincide with major anticlinoria, although in detail, metamorphic isograds transect fold axes (Jackson 2001). Highergrade metamorphic nodes do not coincide with the few granitic intrusions that intrude the Huronian Supergroup rocks south of the Murray Fault. The inferred 1900 to 1850 Ma age of metamorphism is much younger than the age of the Creighton and Murray granites (circa 2460 Ma) yet older than the circa 1740 Ma and 1450 Ma Cutler and Chief Lake granites. 223 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Jackson (2001) considered the origin of the high-grade staurolite-biotite assemblages of the McKim Formation in the hanging wall of the Murray Fault as one of the most enigmatic aspects of the tectonic history of the Southern Province. He concluded that geobarometry indicates a relatively low- pressure metamorphism (2-3 kbar, bathozone 2 of Carmichael 1978) at high temperature (Figure 11). These conditions differ significantly from the 6-7 kbar pressures (bathozone 5 of Carmichael 1978) estimated for the Penokean metamorphism in Minnesota as determined by Holm and Selverstone (1990). Jackson (2001) concluded that the high-temperature metamorphism was at or below pressure corresponding to the thickness of the Huronian Supergroup rock column, thereby precluding crustal thickening as the origin of the metamorphism. Jackson (2001) concluded that a high heat flow regime, such as that developed in areas of crustal extension and related mantle upwelling, was the cause. Such a model is compatible with Card’s (1964) view that the highgrade metamorphism may be the result of rapid, focused heat flow. Figure 11. Pressure-temperature (P–T) grid showing the location of major mineral assemblages in the system KFMASH, after Spear (1993). Bathozones from Carmichael (1978). Also indicated are possible P– T paths for different parts of the Southern Province. Abbreviations: and = andalusite, as = aluminosilicate, bt = biotite, chl = chlorite, cld = chloritoid; crd = cordierite, grt = garnet, kfs = potassium feldspar, ky = kyanite, ms = muscovite, prl = pyrophyllite, qtz = quartz, sil = sillimanite, st = staurolite; B1 = Baldwin Township initial conditions, B2 = Baldwin Township peak conditions, B3 = Baldwin Township retrograde path; DK = diaspore to kyanite path. Figure from Easton (2006b). 224 �Proceedings of the 68th ILSG Annual Meeting - Part 2 In contrast, at the Great Bend of the Spanish River, interpreted as a high-grade metamorphic node (Card 1978b), and which has been the subject of detailed metamorphic studies (Card 1964; Fox 1971), the co-existing assemblage staurolitechloritoid in metapelitic rocks of the McKim Formation, along with the local presence of andalusite, places the area in bathozone 3 of Carmichael (1978) (Figure 11). Furthermore, work on McKim Formation metapelites in Baldwin Township that contain relict kyanite, north of the Murray Fault near Espanola, Easton (2006b) concluded that minimum metamorphic conditions corresponded to bathozone 4 (>4 kbar, ≥500°C; Carmichael 1978). Upon reaching peak temperatures, the metapelites cooled quickly to lower grade, likely in a fluid-rich environment, which retrogressed most minerals except for kyanite (Figure 11). Minerals such as chloritoid, andalusite and staurolite, were especially susceptible to retrograde alteration, as they would have already started to break down during the period of increasing temperature (Easton 2006b). the observations in Baldwin Township and the proposed tectonic thickening model (Easton 2006b), presents a considerably more complex metamorphic history for the south-central Southern Province than has been previously envisaged (e.g., Card 1978b; Bennett et al. 1991; Bennett 2006). If nothing else, it emphasizes the need for additional, modern, metamorphic studies of the Huronian Supergroup throughout the Sault Ste. Marie to Sudbury area. What detrital zircon studies tell us about the source region for the Huronian Supergroup The publication of the first detrital zircon analyses from the Huronian Supergroup (Rainbird and Davis 2006) took place at the same time as this field trip was last run back in May 2006. Since then, detrital zircon work has been completed on more than 25 samples of the Huronian Supergroup (Table 2), and 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. 2017; Long et al. 2011; Ménard 2017, 2019; 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 2, with age ranges and averages based on grains that are < 5% discordant, a lower cutoff than used in many studies. Key observations are: Easton (2006b) argued that tectonic thickening is the most common explanation used to account for the transition from andalusite to kyanite and provides an explanation for the syn-kinematic character of the metamorphic porphyroblasts. It also can account for the differences between Baldwin Township and the Great Bend area. The model of Easton (2006b) explains other metamorphic mineralogical anomalies in the Southern Province. For example, at low temperatures, but similar pressures, the reaction of diaspore to kyanite occurs, which would account for the presence of reported occurrences of diaspore and kyanite (Card 1978a, 1978b; Church 1967; Chandler et al. 1969). It accounts for the folding of metamorphic isograds, as reported by Jackson (2001) in the May Township area. It also provides an alternate explanation for the highgrade metamorphic nodes in the Southern Province other than the presence of focussed heat and fluid zones proposed by Card (1978b). The resulting tectonic history of the Southern Province, based on  Zircons between 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 formations.  Samples from the lower Huronian Supergroup (Elliot Lake and Hough Lake Groups) are dominated by Geon 26 detritus (see Table 2), consistent with provenance dominated by local sources characteristic of the Ramsay-Algoma granitoid complex. Where detailed stratigraphic sampling has occurred, the lowermost units have unimodal populations, 225 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Paleoproterozoic rocks of the Thessalon Formation and the East Bull Lake intrusive suite have ΕNdT values ranging from 2.58 to -2.28 (Easton 2012; Prevec 1993), suggesting derivation from a primary magma originating from a depleted mantle source, which locally was affected by minor amounts of crustal contamination. 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 Geon 27 basement in that area.  Above the Mississagi Formation, Geon 27 populations are dominant, but Geon 28, 29 and Geon 30 grains are also commonplace (see Table 2). 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. In contrast, the Matinenda Formation sandstone samples have negative ΕNdT, ranging from -0.52 to -9.21 (Easton 2012). Samples with the highest negative ΕNdT were also enriched in Th, most likely due to the presence of monazite. The magnitude of the negative ΕNdT values in the Matinenda Formation indicates a negligible contribution from the volcanic and intrusive rocks of the Whiskey Lake greenstone belt, all of which have positive ΕNdT. The Matinenda Formation data can be explained if the sandstones contain a significant component derived from a suite of radiogenic granites located 30 to 40 km north and northwest of Elliot Lake which have ΕNdT of -6.19 (Easton 2012). These radiogenic granites typically contain 8 to 33 ppm U and 30 to 50 ppm Th (Easton 2010), thus they are also a potential source of uranium. In contrast, uranium contents of other Archean felsic intrusions in the Elliot Lake area are 1 to 4 ppm (Easton 2010). It may be no coincidence that the most negative of the Matinenda Formation samples was collected only a few metres below the mineralized Main Conglomerate Bed.  The uppermost Huronian Supergroup units have ages of circa 2310 Ma (Hill et al. 2018; Rasmussen et al. 2013), meaning deposition of the entire supergroup occurred between 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 2 groups. These grains have no known local source, and as suggested by Bleeker (personal communication, 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) and non-glaciogenic sandstone units. In the Elliot Lake area, the Ramsay Lake Formation has a zircon population consisting only of Geon 26 and Geon 27 grains, similar to the population present in the underlying Matinenda Formation (Easton and Heaman 2011; Ménard 2019).  Grains >3000 Ma occur sporadically throughout the Huronian 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, 3000-3600 Ma, 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). In summary, the neodymium data, and the detrital zircon and geochemical data reported by Easton and Heaman (2011), are all consistent with a local source region that included radiogenic granites found north and northwest of Elliot Lake. The absence of similar radiogenic granites north of the Huronian Supergroup between Elliot Lake and Sault Ste. Marie may explain why the Matinenda Formation west of Elliot Lake contains no significant uranium occurrences. Nd isotope data reported by Easton (2012) supports the conclusions based on the detrital zircon studies. Archean felsic volcanic and granodiorite samples from the Whiskey Lake greenstone belt, have positive ΕNdT values close to the depleted mantle evolution curve. 226 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Table 2. 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 the Cobalt Basin are in italics. Abbreviations: cong, conglomerate; EL, Elliot Lake area; MCB, main conglomerate bed; S, Sudbury area; TH, Thessalon area. Table updated from Easton (2019). Number Range (Ma) Bar River mudstone Bar River EL Gordon Lake EL Gordon Lake EL Lorrain EL Gowganda Formation n=16 n=62 n=57 n=30 n=172 2279-2745 2523-3074 2284-2840 3 sites 2684-2890 2520-3614 Serpent EL Serpent EL-S Mississagi EL Mississagi EL Mississagi EL-S Mississagi (upper) S Mississagi S Mississagi S Ramsay Lake EL Ramsay Lake EL-S Ramsay Lake S Ramsay Lake S cong McKim S Mississagi cong Matinenda S n=46 n=10 n=19 n=125 n=63 n=22 n=130 n=117 n=72 n=84 n=65 n=25 n=37 n=36 n=210 n=39 2549-3576 2531-3317 2531-3317 2420-3499 2443-3617 2591-2832 2388-3286 2414-2978 2544-2949 2658-2781 2656-2887 2526-2719 2607-2821 2533-2752 2366-2906 2505-3774 Matinenda EL-S n=27 2451-2714 Matinenda EL Matinenda (upper) EL Matinenda EL Matinenda above MCB EL Matinenda below MCB EL Livingstone Creek TH n=30 n=47 2650-2742 2620-2897 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%) 2450, 2679 (26>27) 2466, 2692 (26>27) 2663 (26>>27) 2477, 2697 (26≈27) 2490, 2560, 2689 2683 (26>>27) 2678 (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) Main Peak (Ma) n=36 n=5 n=15 n=28 2617-2776 2634-2651 2621-2684 2546-2838 2649 (26>>>27) 2641 (5%) 2643 (10%) 2641 (26>>>27) n=37 2507-2890 2698 (26≈27) 227 24 25 26 27 28 29 >3.0 1 18 20 2 30 22 2 1 4 3 2 5 5 3 2 6 34 18 51 6 38 18 29 9 4 8 45 24 18 57 47 44 54 35 19 27 26 78 22 22 4 8 34 13 1 57 39 22 30 28 3 8 7 122 5 11 1 3 1 1 7 13 9 20 4 25 47 5 3 3 3 33 5 15 24 1 17 16 19 4 2 3 1 1 2 3 5 2 2 10 4 3 1 3 3 11 3 10 4 1 9 11 2 2 7 2 2 4 1 1 1 1 3 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Tectonic Models for the Development of the Huronian Basin More recently, Roscoe and Card (1992), noting the close stratigraphic correlation between the Paleoproterozoic sequences of the Wyoming craton and the Huronian Supergroup, proposed that the Superior and Wyoming cratons are rifted portions of what was once a single continental land mass. They suggest the direction of the Matachewan-Hearst dike swarm (2475-2460 Ma) indicates an east-west tensional regime, which resulted in, a Huronian basin elongated in a northsouth direction. On this larger craton, sediment was deposited in a southward-deepening intracratonic basin. Roscoe and Card (1992) proposed that it was during the Nipissing igneous event (2217 Ma) that successful rifting of the Superior Province took place with the eventual drifting of part of the missing Superior Province to its present location as Wyoming craton. They attribute pre-Nipissing folding to the Blezardian orogeny of Stockwell (1982) and the later, more important, deformation to be coeval with the Penokean orogeny of Michigan, Wisconsin and Minnesota (Roscoe and Card 1992). Various tectonic models have been proposed for the early development and later deformation of the Huronian basin. Many reconstructions are essentially modifications of the model put forth by Dietz and Holden (1966), which stated that the Huronian Supergroup represents a rift and passive margin sequence that was compressed, partly tectonically buried, and metamorphosed during a collision with the Superior craton and another mass which overrode its southern edge. Zolnai et al. (1984) and Bennett et al. (1991) accepted the essential aspects of the Dietz and Holden (1966) model. The model proposes that rifting and continental break-up was coeval with Huronian Supergroup volcanism (2475-2450 Ma) and that the much later regional deformation occurred coincident with the Penokean Orogeny (1860-1835 Ma). This model does not attempt to account for the multiple deformation events affecting Huronian Supergroup rocks or the origin of the Nipissing intrusions. Jackson (2001) supported the model of Roscoe and Card (1992) since the high heat flow, which he considers necessary to give the observed features of the high-grade metamorphic rocks, would be a necessary effect of mantle upwelling during continental break-up. He also interpreted some of the early high-strain deformation as being consistent with a Nipissing-age break-up of the Superior craton. Young (1982) proposed that the Huronian Supergroup was deposited in an aulocogen, an easterly-trending fault-bounded trough, which opened toward an ocean in the area now occupied by the Grenville Province. Sims et al. (1981) concluded that the Huronian Supergroup, the Marquette Range Supergroup, and Animikie Group rocks were deposited as intra-continental, fault-controlled basins along a major, Neoarchean structure, the Great Lakes Tectonic Zone. 228 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Acknowledgements The senior author wishes to thank Mike Hailstone, former Sault Ste. Marie Resident Geologist (after Gerry Bennett) for introducing him to this trip back in 2008. This was followed by a mapping program in the Elliot Lake to Pecors Lake area between 2009 and 2011. Figure 12. Index map for included geological maps and some areas mentioned in the text. Figure modified from Bennett (2006). 229 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figure 13. Legend for Figure 14. Figure modified from Bennett (2006) 230 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Figures 14. Upper. Geological map of the Quirke Lake syncline with stop locations. Lower. Geological map of the Flack Like area with stop locations. Both modified from Bennett (2006). 231 �Proceedings of the 68th ILSG Annual Meeting - Part 2 FIELD TRIP DETAILS 0.0 km. Bridge across Depot Lake (approximately 15.4 km from visitor centre). Enter into the Whiskey Lake greenstone belt (Archean). Set odometer to zero. Geological Maps Geological compilation maps covering all or parts of the area of the field trip include Giblin and Leahy (1979); Johns, McIlraith and Muir (2003); and Easton (2013a). The southern part of the field trip area, including optional Stops A-C and Stop 1, as well as the Whiskey Lake greenstone belt and the eastern part of the Quirke Lake syncline, are depicted on Easton (2013a). 1.0 km. Approximately a kilometre north of the bridge, pull over on the right shoulder on the passing lane up the hill. Examine exposures on the large rock face on the right side of the road. Highway 108. This is a new stop. Optional STOP A: Archean metasedimentary rocks of the Whiskey Lake greenstone belt ROAD LOG UTM co-ordinates 381165E 5133640N Note: Caution should be taken when parking vehicles on the shoulder of the highway and when examining outcrops located along Highways108 and 639 and on other roads along the field trip route. All UTM co-ordinates are given in NAD 83 datum, zone 17, which is essentially equivalent to NAD WGS84. Here we see thin- to medium-bedded turbidites of the Archean Whiskey Lake greenstone belt. Partial Bouma sequences are present in the thicker turbidite beds. The turbidites likely have a high felsic volcaniclastic component, as a sample from this outcrop for detrital zircon study yielded no zircons or titanite (Photo1). Note: This guidebook describes a total of 25 stops (17 stops and 8 optional stops). If one is starting in Elliot Lake, it is possible to do all of the stops in one day. If coming from Sudbury, then some stops will need to be omitted (mainly the optional stops). All 16 stops from Bennett (2006) are included in the road log, along with an additional 9 stops added by the senior author. Leave from Science North at the junction of Paris Street and Ramsey Lake Road in Sudbury. Head to Highway 17 and proceed west toward Sault. Ste. Marie. At the junction of Highway 17 and 108, turn right onto Highway 108 and head north to Elliot Lake (29 km (18 mi) from, Highway 17). Set odometer to zero at this point. Photo 1. Thick volcaniclastic bed in turbidites at Optional Stop A. Hammer handle is 33 cm long. From the junction with Highway 17 to the bridge at Depot Lake, the highway passes through the Ramsey-Algoma granitoid complex. The rocks along the highway have not been studied in detail. The granitoid rocks include xenoliths of mafic rock and are cut by numerous dikes of the Matachewan dike swarm (circa 2460 Ma), which are mediumgrained, medium-green, and locally plagioclase porphyritic. Also present are fine-grained, flinty, mafic dikes that may represent feeders to Huronian Supergroup volcanic rocks. Although not observed in this outcrop, as one heads along the highway and up stratigraphy, the turbidites pass into thinly bedded mudstones and then into a magnetite facies iron formation. These metasedimentary rocks lie atop the older (circa 2740 Ma) of the 2 volcanic sequences that comprise the Whiskey Lake greenstone belt (Easton 2013a). Return to vehicles, continue on Highway 108 232 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Archean Whiskey Lake greenstone belt (Photo 2). A U/Pb zircon sample from this outcrop yielded a CA-TIMS age of 2724.9±1.4 Ma (Hamilton in Easton 2010) indicating that these rocks are among the oldest in the greenstone belt. 2.3 km (1.44 mi). Turnoff to Elliot Lake airport on the left. Continue straight. 3.6 km (2.25 mi). Pull over on the shoulder of the road. Examine long roadcut on the right side of the highway. This is a new stop. The northwestern outcrop consists of extremely flattened tuff-breccia (Photo 2). The higher strain in this outcrop may reflect its location in the contract strain aureole of the large granodiorite body that we will examine at Stop 1. Return to vehicles, continue straight (westward) on Highway 108. Optional STOP B: Stone Ridge intrusion UTM co-ordinates 379693E 5135505N The roadcut exposes part of an east-trending metagabbro intrusion, termed the Stone Ridge intrusion. The intrusion is 700 to 1000 m wide, with a minimum strike length of 15 km. It lies 1 to 2 km south of, and roughly parallels, the Archean– Proterozoic unconformity (Easton 2009, 2013a). As seen in this roadcut, large parts of the intrusion contain preserved primary mineralogy. The predominant rock type is a weakly metamorphosed, grey to light grey weathering, medium-grained, leuconorite to leucogabbronorite. Where recrystallized, orthopyroxene is altered to amphibole, and the rock takes on a greener colour. Texturally, the body is remarkably uniform, but coarse-grained to pegmatitic patches of gabbro occur along the northern margin of the intrusion. Matachewan dikes (circa 2460 Ma), which were observed to intrude the body (Easton 2009), would preclude the Stone Ridge Intrusion being part of the Nipissing intrusive suite, which was not emplaced until circa 2217 Ma, suggesting that the Stone Ridge intrusion is more likely part of the East Bull Lake intrusive suite. 6.1km (3.81 mi). Turnoff to the left takes you to the Discovery Site lookout. Several large boulders representative of the uraniumbearing “Main Conglomerate Bed” are present in the parking area of the lookout. Looking northeast from the highway and/or the lookout, you can see a large ridge of greenish sandstone of the Matinenda Formation. The first mine in the Elliot Lake Camp, the Buckles Mine, was located at the base of this ridge opposite the turnoff. 6.8 km (4.25 mi). Pull over and park near the middle of a large roadcut on the north side of the highway This is a new stop. STOP 1. Thessalon Formation feeder dike and Archean granodiorite UTM co-ordinates 377092E 5136767N The roadcut contains a 15 m wide, near-vertical mafic dike which is vesicular (Photo 3). The vesicular nature of the dike is best observed on the roadcut on the south side of the Highway (Photo 4). The mafic dike has the composition of a tholeiitic andesite on a Jensen discrimination diagram (54.9 wt.% SiO2, 2.38 wt.% K2O, magnesium number of 36 (data in Easton (2013b)) and is a high-K basaltic andesite in the IUGS total-alkali silica classification. Return to vehicles, continue northward of Highway 108 to the junction with Nordic Road and the golf course. 5.3 km (3.31 mi.) Junction with Nordic Road. Park safely and examine outcrops on the northeast and northwest sides of the intersection. This is a new stop. Optional STOP C: Archean metavolcanic rocks of the Whiskey Lake greenstone belt UTM co-ordinates 378483E 5136436N geochronology site; 378486E 5136443 flattened tuff The northeastern outcrop consists of reversely graded felsic tuffs and felsic tuff-breccias of the 233 �Proceedings of the 68th ILSG Annual Meeting - Part 2 Photo 3. Roadcut on north side of Highway 108 at Stop 1. The dark, 10 m wide, vertical, steep-walled mafic dike is intruded Archean granodiorite, which has a U/Pb zircon age of 2674.8±0.8 Ma (Easton 2013a). This dike likely feed Thessalon Formation flows. Photo 4 Close-up of mafic dike rock on the south side of Highway 108 at Stop 1 with irregular, white vesicles in the dike. Pen is 13.5 cm long. The medium-grained granodiorite that forms the bulk of the roadcut is typical of the younger intrusive bodies within the Ramsey-Algoma granitoid belt. A sample collected from the roadcut on the south side of the road yielded a CA-TIMS age of 2674.8±0.8 Ma (Easton 2013a). The youngest age reported so far from metavolcanics of the Whiskey Lake is 2685.5±1.1 Ma (Easton 2013a). Photo 2. Felsic volcanic rocks at Optional Stop C. Upper. Lapilli tuff in the northeastern outcrop which was sampled for geochronology. Knife is 9 cm long. Middle. Moderately flattened tuff breccia from the northwestern outcrop. Lower. Strongly flattened tuff breccia from the northwestern outcrop. Hammer handle is 33 cm long. Return to vehicles, continue straight (westward) on Highway 108. 234 �Proceedings of the 68th ILSG Annual Meeting - Part 2 7.1 km (4.44 mi). Pull over and park on the shoulder of the highway. Cross over to the former route of the highway and walk west toward the intersection of the new and the old highways. This is a new stop. Optional STOP D: Paleoweathering of Archean granodiorite UTM co-ordinates 376567E 5136603N Proceeding west along the outcrops present on the north side of the former highway, you can observe the progression from grey granodiorite to reddishweathering granodiorite to greenish granodiorite, with the change in coloration reflecting increasing paleoweathering of the granodiorite (Photo 5). Key elemental changes between the 3 phases are summarized in Table 3, and include increasing iron, manganese, magnesium, potassium and aluminum contents, with decreasing silica content. Table 3. Element changes related to paleoweathering in granodiorite at Optional Stop D. Data from Easton (2013b). Major elements are in weight percent. Abbreviations. LOI, loss on ignition; CIA, chemical index of alteration. Element Photo 5. Paleoweathering in Archean granodiorite at Optional Stop D. Upper. Grey, unweathered granodiorite. Middle. red-green weathering granodiorite. Lower. Green weathering granodiorite. See Table 3 for chemistry on each type. Pen is 9 cm long. Photo 5 Grey, not weathered upper Red weathered middle Green weathered lower SiO2 69.1 60.1 52.7 Al2O3 15.7 19.4 19.6 Fe2O3total 3.2 5.0 9.9 MnO 0.02 0.06 0.12 MgO 1.9 2.7 5.5 CaO 1.0 0.5 0.7 Na2O 6.8 7.9 6.9 K2O 1.0 2.5 0.8 LOI 1.3 2.0 3.5 CIA 89 87 70 Th (ppm) 7 12 13 TiO2, P2O5, U, Zr nearly constant in all 3 samples Return to vehicles, continue straight (west) on Highway 108. Continue straight on the highway past Hillside Drive South. 235 �Proceedings of the 68th ILSG Annual Meeting - Part 2 9.2 km (5.75 mi). Roadcut to the right of the highway just west of the junction with Esten Drive north consists of coarse sandstone and pebblestone of the Matinenda Formation similar to what we will see at Stop 3. The pink colouration in this roadcut is related to its proximity to a fault located approximately along the highway route. UTM co-ordinates 370462E, 5137991N The low outcrops on north side of Spine Road are grey, buff and dark-grey sandstone and radioactive, pyritic, quartz-pebble conglomerate of the Matinenda Formation. Pebble units, located near the top of the outcrop, are rusty-weathering, about 20-30 cm thick, and dip about 10 degrees to the north (Photo 6). Pebbles in this outcrop are 1-2 cm across and are generally much smaller than the typical pebbles in the ore zones of the Elliot Lake mines. Scintillometer readings from the pebble beds, collected by R.M. Easton between 2009 and 2022, range from 15,000 to 18,000 counts per second, with 600 to 800 ppm U and 600 to 800 ppm Th, with the U content being typical of the ore grades from the Elliot Lake camp. Continue straight on the highway and past the first few stoplights to Hillside Drive North. Turn left (west) onto Hillside Drive North. If going directly to Stop 2, once on Hillside Drive North, continue west for approximately 1 km to Spine Road. Turn right (west) onto Spine Road and drive past the hospital to Lawrence Avenue at the far west end of Spine Road (~2.1 km). Park in the turn-around at the end of Spine Road for Stop 2. This is Stop 15 of Bennett et al. (1997) and Stop 1.2 of Bennett (2006). If going to Optional Stop E, continue west on Hillside Drive North for about 400 to Spruce Avenue. Turn right on Spruce Avenue and continue approximately 100 m to Valley Crescent, turn right onto Valley Crescent and follow it for approximately 350 m to Balsam Place. Turn right on to Balsam Place and stop at the end of the cull-de-sac. This is Stop 1.1 in Bennett (2006). Optional STOP E: McKim Formation and Nipissing Diabase Photo 6. Matinenda Formation at Stop 2. Hammer head is resting on radioactive, rusty-weathering pebble beds that are the focus of this stop. UTM co-ordinates 372751E, 5138695N Outcrops on the east side of the cul-de-sac are mudstone and grey sandstone of the Mckim Formation. Note the deflection of the axial plane cleavage in the mudstone units. The movement of adjacent beds inferred from the deflection of the cleavage indicates the south limb of a syncline. A gabbroic dike is separated from the sedimentary rocks by a zone of sheared and fractured rocks. The McKim Formation is missing on the north limb of the syncline. Ruzicka and LeCheminant, (1984) reported the radioactive conglomerate contains “rare-earthelement-bearing uranothorite (?), large zircons, a Ti-U-Si-Fe phase (brannerite?), chalcopyrite and chromite. The distribution of radioactive minerals in the conglomerate displays layering thus indicating a detrital origin of these grains”. STOP 2. Radioactive quartz-pebble conglomerate, Matinenda Formation 236 �Proceedings of the 68th ILSG Annual Meeting - Part 2 STOP 3. Coarse sandstone, floater reef zone, Matinenda Formation, Elliot Lake Group The dark-grey areas in the radioactive beds are due to the presence of minor amounts radioactive carbon generally known in the Elliot Lake area as “thucolite” also referred to as a hydrocarbon kerogen. Ruzicka and LeCheminant, (1984) noted that several generations of carbon occur in the conglomerates of the Matinenda Formation. The earliest generation occurs as layers concordant with the bedding or as a component of the matrix and appears to have been deposited in areas of quiescent sedimentation during the last phase of an upward-fining sedimentary cycle. UTM co-ordinates 373385E 5137830N Created in 2014, this rock face exposes greenish coarse sandstone and pebblestone of the Matinenda Formation. Compared to Stop 2, we are now only 25 m higher in the stratigraphy above the “Main Conglomerate Bed” in what is locally referred to as the “Floater Reef Zone”. The name is derived from the fact that many of the pebblestone beds are variably radioactive (typically 2,000 to 5,000 counts per second). Scintillometer data collected by the senior author from this Stop give 20 to 25 ppm U and 130 to 150 ppm Th. Later generations of thucolite are probably remobilized phases of the first generation. The carbonaceous matter in the Elliot Lake ores is comparable in occurrence and composition to hydrocarbon in the Witwatersrand gold reefs; interestingly Ruzicka and LeCheminant (1984) report elevated gold content (1000-2000 ppb) in the carbonaceous matter of the Elliot Lake ore beds. The radioactive carbon at this site is reported to be auriferous, although the gold content is not available. At the north-end of the rock face is a 5 m wide medium-grained mafic dike cutting the Matinenda Formation. The affinity of the dike is unknown, and no geochemical data are available for this dike. It cannot be a Matachewan dike as it cuts the Matinenda Formation, and it does not have the scintillometer characteristics of a Nipissing gabbro (potassium is too low). It could be related to a suite of east-trending dikes in the Elliot Lake area, such as the one at Stop 14. In 1955 Rio Algom Mines Limited completed a diamond drill hole about 30 m south of this location. The drill log shows that the radioactive beds exposed here are about 35 metres above the Archean basement rocks. This drilling indicated that there are no ore-grade units in this area. Return to vehicles. Exit the south end of the mall parking lot, turn left onto Hillside Drive South, continue east on Hillside to the traffic lights (Highway 108), approximately 300 m. Grab samples collected by G. Bennett in 1982 and reported in Bennett (2006) returned up to 0.80 lbs U3O8/ton and 0.78 lbs ThO2/ton (340 ppm U, 340 ppm Th). A continuous chip sample returned 0.31 lbs U3O8/ton and 0.53 lbs ThO2/ton (130 ppm U, 232 ppm Th). Turn left onto Highway 108 and head north. Reset odometer to zero Retrace route on Spine Road to Hillside Drive. At the traffic light, turn right onto Ontario Drive heading south. Vacant lot on the right is the site of the former Algo Mall, which had a catastrophic collapse on June 23, 2012. Continue for approximately 450 m on Ontario Avenue and where the road bends, continue straight into the retail mall parking lot. Park and examine the large rock face on the west side of the mall parking lot. This is a new stop. 2.1 km (1.31 mi), Westview Park on Elliot Lake on the left. 1.1 km (0.69 mi). Miners Memorial Park and Horne Lake on the right. Cliff on the east side of Horne Lake consists of Mississagi Formation sandstone (Photo 7). 2.4 km (1.5 mi). Stanleigh Road on the right, 237 �Proceedings of the 68th ILSG Annual Meeting - Part 2 (R.M. Easton, unpublished data). The broad age range of zircons in this sample contrasts greatly with the limited age range (2658-2781 Ma) present in a conglomerate of the underlying Ramsay Lake Formation, which was sampled by Ménard and Easton in 2017, 13 km to the east-southeast of the field trip stop (Ménard 2019). The population in the Ramsay Lake Formation sample is almost identical to the 4 Matinenda Formation samples studied by Easton and Heaman (2011), only a few hundred metres to the south of the conglomerate. Photo 7. View from the Miners Memorial across Horne Lake showing cliff of Mississagi Formation sandstone. 3.5 km (2.19 mi). Pull-over where snowmobileATV trail intersects Highway 108. Cross Highway to outcrop on the west side. This is Stop 2.3 of Young (1991), Stop 16 of Bennett et al. (1997) and Stop 1.3 of Bennett (2006). STOP 4. Mississagi Formation, Hough Lake Group UTM co-ordinates 371717E, 5140426N One-metre-thick beds of grey sandstone of the Mississagi Formation on the west side of the highway display the rusty staining on the face of the outcrop reflecting the minor pyrite content along the foreset beds of trough cross-beds (Photo 8). The paleocurrent direction (from the west) can be best observed on the upper surface of the outcrop. — Please exercise caution when walking on smooth, wet rock surfaces — The grey colour of these sandstones and the presence of apparent detrital pyrite are held by most geoscientists to indicate the very low partial pressure of free oxygen of the atmosphere during the deposition of the Mississagi Formation. Photo 8. Crossbedding in Mississagi Formation sandstone at Stop 4. Return to vehicles, continue north on Highway 108 for 1.4 km. 4.1 km (2.56 mi). Pull over on the right shoulder of the passing lane part way up the hill. This is Stop 2.4 of Young (1991); Stop 17 of Bennett et al. (1997) and Stop 1.4 of Bennett (2006). STOP 5. Nipissing gabbro, altered Mississagi Formation, Bruce Formation UTM co-ordinates 371652E, 5140948N J.A. Ménard and R.M. Easton collected a sample for detrital zircon geochronology from this stop in 2017 (Photo 8). The zircon population ranged in age from 2420 to 3499 Ma, with notable peaks at 2450 Ma (13 grains) and 2679 Ma (45 grains), and with 10 Geon 28 grains and 13 grains ≥3000 Ma (Ménard 2019). The abundance of circa 2450 Ma grains compared to other Mississagi Formation samples (see Table 2) is not unexpected given the abundance of Elliot Lake Group metavolcanic rocks to the north of this site near Dunlop Lake We will start by examining the low outcrops on the west side of Highway 108. Please exercise caution when crossing the highway. Low, rounded outcrops at the south end are Mississagi Formation sandstones similar to those we saw at Stop 4, but slightly pink in colour, likely due to the formation of albite by hydrothermal fluids from the adjacent Nipissing gabbro intrusion (east roadcut). After an outcrop gap, conglomerate 238 �Proceedings of the 68th ILSG Annual Meeting - Part 2 of the Bruce Formation is well-exposed. In particular note a clast with a gneissic fabric, as well as numerous, subrounded to subangular, white, granitoid clasts typical of the formation (Photo 9). The abundant matrix is dark-grey to black. Large, glassy, quartz grains are abundant on fresh surfaces of the conglomerate, another feature typical of the formation. The glassy, black, appearance results from the dark matrix behind the clear quartz (Bennett 2006). has been interpreted as the result of erosion by high-pressure, waterborne sediment, presumably by the melting of an adjacent Pleistocene ice sheet (Bennett 2006). North of gabbro sill, the upper portion of the Mississagi Formation is exposed along the east side of the highway and is also pinkish. A few tens of metres northward, the Mississagi Formation is overlain by diamictite of the Bruce Formation (UTM 371668E, 5141245N). The dispersed megaclasts in the Bruce Formation are predominantly grey granitic rocks with smaller mafic clasts, predominantly Thessalon Formation volcanic rocks. At this locale, there is no evidence of a significant disconformity at the base of the Bruce Formation. Return to vehicles and proceed approximately 300 m to near the top of the hill and park on the right shoulder. 4.4 km (2.75 mi). Roadcuts are present on both sides of the highway. This is Stop 18 of Bennett et al. (1997) and Stop 1.5 of Bennett (2006). STOP 6. Espanola Formation, Quirke Lake Group and Nipissing gabbro sills The base of the Espanola Formation is a green, laminated unit about a metre or so thick. Laminated silty limestones and minor thin, chert beds of the limestone member of the Espanola Formation (Photo 10), overlie the green unit. At this location, the proximity of Nipissing gabbro sills (Photo 11) has led to the development of calc-silicate minerals including: grossular garnet, diopside, idocrase (vesuvianite), and wollastonite typical of a skarn (Robertson 1968; Bennett 2006). Wollastonite (identified by X-ray diffraction) is found just below the north-dipping gabbro sill near the north end of the exposure, where it occurs as sub-parallel groups of pale grey to white prismatic crystals about 1 mm wide and up to a cm long. The pink coating on joint surfaces is apophyllite (KFCa4[Si8O20]8H20) an uncommon mineral (identified by X-ray diffraction), sometimes found in amygdules in basalts, but which is also Photo 9. Ramsay Lake Formation conglomerate at Stop 4. Note a gneissic clast just below the rusty spot at the centre of the photo. Knife is 9 cm long. A sill-like body of Nipissing gabbro is exposed at the south end of the roadcut on the east side of the highway. Rhythmic, compositional layering is visible on the vertical face of the roadcut. Additional evidence of hydrothermal activity along the intrusion contact is seen by dark-green to black chlorite deposited along fractures in the gabbro (Bennett 2006). Near the north end of the Nipissing outcrop face, note the relatively planar, striated, surface is truncated by more irregular surface that 239 �Proceedings of the 68th ILSG Annual Meeting - Part 2 associated with calc-silicates. Young (1991) states that the small scale thrust faults and folds in the limestone on the west side of the highway are probably the result of slumping during early tectonic activity. The upper, ferruginous dolostone-bearing member of the Espanola Formation and the overlying Serpent Formation are not present at this location but are well represented on the north limb of the syncline. Young (1991) and Bennett (2006) both suggested that the missing Serpent Formation and ferruginous dolostone member were removed during a period of pre-Gowganda Formation erosion, which are visible in roadcuts along the next stretch of the highway. Alternatively, a fault could be present in the linear valley at the end of the outcrop, which has truncated stratigraphy. Photo 11. Nipissing sill (dark) in contact with Espanola Formation marbles (white) at Stop 5. Note discoloration of the marbles near the sill contact. View west across Highway 108. 4.5-5.0 km (2.81-3.13 mi). Diamictite and minor sandstone of the Gowganda Formation are exposed in a near-continuous roadcut along Highway 108. In these exposures, megaclasts of pink granite, grey granite and granitic gneiss and mafic rocks are widely distributed in a dark green matrix. A typical example can be seen at 4.8 km. Most geologists now consider at least some of the diamictites in the Gowganda Formation to be tillites, although a debris-flow origin, either glaciogenic or as submarine debris flows, is a more reasonable interpretation at specific localities. Roscoe (1969) places the appearance of free oxygen in the atmosphere (“oxyatmoversion”) as coinciding with the appearance of the reddish hue of hematite just above the base of the Gowganda Formation. Return to vehicles and continue north on Highway 108. 6.5 km (4.06 mi). The stop is at a large roadcut at the top of a hill, near a communication tower. Park at the south end of the roadcut on the east side of the highway. This is Stop 2.5 of Young (1991), Stop 19 of Bennett et al. (1997) and Stop 1.6 of Bennett (2006). Photo 10. Espanola Formation at Stop 5. Note white recessive weathering limestone beds and thinly laminated, darker, calc-silicate and mudstone beds. Hammer handle is 33 cm long. This is an impressive exposure though a stratified sequence of diamictites, clastsupported conglomerates and sandstones of the Gowganda Formation. 240 �Proceedings of the 68th ILSG Annual Meeting - Part 2 STOP 7. Stratified Gowganda Formation, Cobalt Group UTM co-ordinates 371841E, 5143180N At the south end of the roadcut on the west side of the highway, massive diamictite is exposed. It is overlain by about 50 cm of laminated mudstone with dropstones, in turn overlain by a thick succession of lenticular beds of coarse pink and pink-grey arkosic sandstone interbedded with distinct beds of diamictite (Photo 12), pebbly sandstone and clast-supported polymictic conglomerate (Photo 13). Some conglomerate units display normal and reverse grading suggestive of debris flows. Some sandstones contain large clasts. Some clasts in the diamictite show striations, suggestive of a glacial origin. Clasts in the conglomerate are mainly wellrounded fragments, but some rip-up clasts of sandstone are also present. The rocks displayed here may be interpreted as debris and mass flows, possibly formed in an ice-proximal setting by resedimentation of glacial debris at a retreating glacial margin (Young 1991; Bennett 2006). Photo 12. Gowganda Formation at Stop 7 showing pink-grey arkosic sandstone (lower) overlain by matrix- to clast-supported conglomerate. Hammer handle is 33 cm long. Note the predominance of red and pink granitic clasts, in marked contrast to the pale grey clasts of the Bruce Formation seen earlier at Stop 5. There is also a significant proportion of black pebble to cobble-sized clasts. The mineral assemblage and metamorphic grade of a few mafic clasts examined by G. Bennett many years ago indicated that the clasts were probably from Thessalon Formation basaltic flows (Bennett 2006). Return to vehicles and continue north on Highway 108. 7.4 km (4.63 mi). Pink sandstone and diamictite of the Gowganda Formation. 9.4 km (5.88 mi). Diamictite with large boulder, Gowganda Formation. Photo 13. Gowganda Formation matrix-supported conglomerate (lower) overlain by clast-supported conglomerate at Stop 7. Hammer handle is 33 cm long. 12.0 km (7.5 mi). Stanrock Road. Reset odometer. Turn east onto Stanrock Road. 241 �Proceedings of the 68th ILSG Annual Meeting - Part 2 2.0 km (1.25 mi). park on the south side (right) of the road opposite a roadcut on the north side. This is Stop 2.6 of Young (1991) and Stop 1.7 of Bennett (2006). STOP 8. Laminated varvite? Gowganda Formation, Cobalt Group UTM co-ordinates 374911E, 5147007N Laminated siltstone/mudstone of the Gowganda Formation on the north side of the road. This unit has been interpreted as being similar to the varves found as deposits in Pleistocene glacial lakes. Continue east on Stanrock Road. 9.5 km (5.94 mi). Turn right onto Popeye Lake Road. 9.8 km (6.13 mi). Park and examine outcrops on the west side of the road. STOP 9. Gowganda—Serpent Formation disconformity UTM co-ordinates 381485E, 5144900N The Serpent Formation consists of fine, wellsorted sandstone and siltstone, and is typically light grey. The exposure at this Stop shows typical sandstones of the formation (Photo 14), just below the disconformity with the overlying Gowganda Formation. The contact can be seen partway up the hill above the road level exposures of the Serpent Formation. The contact is sharp but irregular, and the Gowganda Formation consists of polymictic, matrix-supported, conglomerate Photo 14. Serpent Formation. Upper. Crossbedding in sandstone at Stop 9. Lower. Indistinct, medium bedding in fine sandstone at Stop 9. Hammer handle is 33 cm long. Optional STOP F. Gowganda—Serpent Formation disconformity UTM co-ordinates 375102E, 5150499N The Serpent Formation is not present in the south limb of the Quirke Lake Syncline and in the Blind River—Sault Ste. Marie area where it was probably removed during a period of preGowganda Formation erosion. At this location, on the south side of the road, well-sorted sandstone of the Serpent Formation is overlain by polymictic conglomerate of the Gowganda Formation. The contact is sharp but irregular. Evidence of a subGowganda Formation disconformity at this location is based on the presence of pebble and cobbles of the Serpent Formation near the base of the overlying Gowganda Formation. Return to vehicles, retrace route to Highway 108. Reset odometer to zero at the junction. Turn right onto Highway 108 and continue north. 3 km (1.9 mi). Denison Mine Road - Turn east. Reset odometer to 0. 1.5 km (0.95 mi). This is Stop 2.7 of Young (1991), Stop 20 of Bennett et al. (1997) and Stop 1.8 of Bennett (2006). Return to Highway 108. Reset odometer at Highway 108 and Denison Mine Road. 242 �Proceedings of the 68th ILSG Annual Meeting - Part 2 0.5 km (0.3 mi). Disseminated carbonate in sandstone of the Serpent Formation on the east side of Highway 108. The detrital zircon sample of the Serpent Formation reported by Rainbird and Davis (2006) and Craddock et al. (2013) came from this roadcut. The population ranged from 2549 to 3576 Ma, with the dominant population at Geon 27, but with 11 Geon 28 grains (see Table 2). 1.1 km (0.68 mi). Road to Quirke Lake and the former Panel Mine. Reset odometer to zero. Turn east onto Panel Mine Road. 0.6 km (0.38 mi) and 1.5 km (0.9 mi). Two outcrop areas are exposed on the west side of the road. The southern of the 2 areas is currently betterexposed. The second area to the north is Stop 2.8 of Young (1991), Stop 21 of Bennett et al. (1997) and Stop 1.9 of Bennett (2006). Photo 15. Dropstone in laminated dolostone of the Espanola Formation at Stop 10. Pen is 13 cm long. STOP 11. Ramsay Lake Formation overlain by Pecors Formation STOP 10. Upper member of the Espanola Formation, Quirke Lake Group UTM co-ordinates 377379E, 5152019N Diamictites of the Ramsay Lake Formation contain cobbles of grey granitic rocks, mafic clasts of Huronian Supergroup metavolcanic rocks and Archean felsic metavolcanic clasts in an abundant dark-grey to black sandy matrix. The Ramsay Lake Formation is overlain by dark laminated siltstone and mudstone of the Pecors Formation (Photo 16). The latter contains a few dropstones (Photo 17). Note: the Matinenda Formation of the Elliot Lake Group, expected between the basement and the Ramsay Lake Formation, is truncated by the Ramsay Lake Formation in this area. The Matinenda Formation does occur in the mine workings down-dip from this location. UTM co-ordinates area 1, 374350E, 5151383N area 2, 375258E, 51511331N Both areas exposure ferruginous dolomite and siltstone of the upper member of the Espanola Formation. At the first stop, ripple marks are visible on some bedding surfaces. The upper member is the uppermost of the 3 members of the Espanola Formation recognized in the Elliot Lake area (Robertson 1968). It is characterized by intercalated siltstone and reddishbrown weathering, ferruginous dolostone beds containing 3-4% FeO. Intraformational breccia, ripple marks, small-scale crossbedding, and a variety of soft sediment features are present, but only faintly visible on the south outcrop (Photo 15). Near the east end of the second outcrop a grey clastic dike crosses stratification at a high angle. If you continue east to the end of the Panel Mine Road, you enter the rehabilitated area of the former Panel Mine. There is little evidence of the uranium mine and mill complex that was on this site until 1993. Return to vehicles and continue east on Panel Mine Road. 4.1 km (2.5 mi). This is Stop 2.9 of Young (1991), Stop 22 of Bennett et al. (1997) and Stop 1.10 of Bennett (2006). 243 �Proceedings of the 68th ILSG Annual Meeting - Part 2 1.9 km (1.19 mi). Near the top of the hill a fresh roadcut on the west side of Highway 639 exposes greenish, coarse-grained sandstone of the Matinenda Formation (UTM 372210E 5152225N. We are near the top of the floaterreef zone, so slightly higher stratigraphically than at Stop 3. 2.5 km (1.6 mi). This is Stop 23 of Bennett et al. (1997) and Stop 1.11 of Bennett (2006). We will not visit this stop due to logistical and accessibility reasons. Optional STOP G. Huronian Supergroup volcanic rocks of the Thessalon Formation Photo 16. Laminated mudstone of the Pecors Formation, Stop 11. Hammer handle is 33 cm long. UTM co-ordinates 371508E, 5152302N A gated road leads west from Highway 639 to one of the Quirke Mine tailings dams. Park near the gate and walk a short distance along a rough road from the gate to the base of the tailings dam. Note the very dark green to black, flattened, chlorite amygdules characteristic of the Huronian Supergroup mafic volcanic rocks between Sault Ste. Marie and Elliot Lake. Cross the stream and proceed northward a short distance along a rough road to the crest to the low hill. The Huronian Supergroup volcanic rock at this location (UTM 371428E, 5152342N) include hawaiite and mugearite (Bennett 2006). The eastward-trending, south dipping unconformity between the Archean granitic basement rocks and Huronian Supergroup volcanic rocks is visible near the crest of the hill. There appears to be no paleosol development at this location. Near the west end of the outcrop, a thin, quartz-pebble conglomerate or breccia unit, consisting mainly of angular, quartz-clasts, overlies the granitic rocks at the base of the volcanic unit. Scattered, isolated, mainly cobblesized clasts of quartz are also present along the unconformity. Photo 17. Laminated mudstone of the Pecors Formation containing a dropstone at Stop 11. Knife is 9 cm long. Return to vehicles, retrace route back to Highway 108. Reset odometer to zero at highway. Turn right and continue north on Highway 108. Tailings dam of the Quirke Mine is visible west of the Highway. 0.8 km (0.5 mi). Highway 108 ends and Highway 639 begins. 1.0 km (0.6 mi). Diamictite of the Bruce Formation is exposed on the west side of the highway. 1.5 km (0.9 mi). Outcrops of Mississagi Formation are exposed along Highway 639. Note the yellowish colour characteristic of the Mississagi Formation where it lies directly on the Archean granitic basement (Robertson 1968). Some visitors to this site have proposed that the contact between the Archean and Huronian Supergroup volcanic rocks is not an unconformity, but a fault contact. During a visit to the Stanleigh Mine in 1990, however, G. Bennett observed identical scattered, quartz pebbles along the contact of the Huronian Supergroup volcanic rocks 244 �Proceedings of the 68th ILSG Annual Meeting - Part 2 where they overlie Archean mafic volcanic rocks in a haulage drift on the south limb of the Quirke Lake Syncline (Bennett 2006). Bennett (2006) proposed that these quartz cobbles are lag deposits left behind when the finer sediment was washed off the surface. A few kilometres west of this stop, occurrences of this conglomerate unit contain more rounded quartz grains, and locally are overlain by a thin arkosic sandstone. Return to vehicles and continue north on Highway 639. 2.6 to 10 km (1.63-6.25 mi). Highway 639 passes through Archean granitoid rocks cut by Matachewan and other diabase dikes before entering into dominantly mafic volcanic rocks of the Ompa Lake greenstone belt which has similar ages to the Whisky Lake greenstone belt, namely circa 2685 Ma (see summary in Easton 2010). Photo 18. Pillow structures in Archean mafic metavolcanic rocks at Optional Stop H. Scale card is 10 cm long. Photo from Bennett (2006, p.40). 9.3 km (5.81 mi). This is Stop 24 of Bennett et al. (1997) and Stop 1.12 of Bennett (2006). 11.3 km (7.06 mi). Provincial Park. Optional STOP H. Pillowed Archean metavolcanic rocks 11.6 km (7.25 mi). Large roadcut on the east side of the road. Park on the shoulder. This is an added stop. UTM co-ordinates 368530E, 515773N Entrance to Mississagi STOP 12. Bar River Formation, Cobalt Group Archean mafic metavolcanic rocks with welldeveloped pillow structures are exposed, on a north-sloping outcrop, on the east side of the highway. The pillows are deformed, however, facing directions can easily be determined. Small amygdules are concentrated near the upper surface of many pillows. Lichen growth since 2006 has rendered this stop less spectacular than as indicated in Photo 18. UTM coordinates 367555E, 5159855N Thin to medium bedded, pale grey sandstone of the Bar River Formation with herringbone crossbedding. Return to vehicles and continue north on Highway 639. 12.3 km (7.69 mi). Jim Christ Lake to the northeast (formerly Christman Lake). Park on the right shoulder beside a low ridge of partially blasted roadcuts on the north side of the highway. This is Stop 25 of Bennett et al. (1997) and Stop 1.13 of Bennett (2006). This once spectacular roadcut has suffered from needless road construction damage in recent years. Return to vehicles and continue north on Highway 639. 10.2 km (6.38 mi). Flack Lake fault occupies a valley near this point. 10.7 km (6.69 mi). Outcrops of hematite-stained sandstone of the Bar River Formation. 245 �Proceedings of the 68th ILSG Annual Meeting - Part 2 STOP 13. Bar River Formation, Cobalt Group STOP 15. Gordon Lake Formation, Cobalt Group UTM co-ordinates 366742E, 5160652N UTM co-ordinates 363062E, 5163194N Sandstones and mudstones of the Bar River Formation at this stop display ripple marks, mud cracks and sinuous structures, which have been described as possible worm casts. Comparison with desiccation structures in the Gordon Lake Formation led Young (1969) to suggest that these features are the result of the transportation of consolidated desiccation fracture fillings. Siltstones and sandstones of the Gordon Lake Formation display ripple marks, desiccation cracks, cross bedding and a late cleavage. Note: the presence of pyrite in contrast to hematitic nature of the Gordon Lake Formation near the top of the formation. Return to vehicles and continue north on Highway 639. This outcrop area was sampled by Rainbird and Davis (2006) and Craddock et al. (2013) for detrital zircon geochronology. Population range was 25233074 Ma, with peaks at 2531, 2705 and 2726 Ma. 19.6 km (12.25 mi). Park on the right shoulder roughly midway in a lengthy roadcut on a south-facing hill which has almost near continuous exposures of Nipissing gabbro. Examine outcrops on the east side of the road. Return to vehicles and continue north on Highway 639. 16.1 km (10.01 mi). This is Stop 26 of Bennett et al. (1997) and Stop 1.14 of Bennett (2006). STOP 14. Red beds of the Gordon Lake Formation, Cobalt Group UTM co-ordinates 364395E, 5162758N Laminated, maroon buff or green siltstone and mudstone, and minor chert, represent the upper part of the Gordon Lake Formation. Desiccation cracks and ripple marks are present, as are reduction spots in the maroon beds. It was near this stop that Hill et al. (2018) collected a green siltstone sample which had a limited zircon population (27 grains), but with the 4 youngest grains giving an age of 2302±19 Ma, and with another cluster of 5 grains at 2364±16 Ma. These ages, as well as those of Rasmussen et al. (2016), suggest deposition occurred at circa 2300 Ma. Other populations were 6 grains at 2525±15 Ma, with older grains ranging from 2674 to 3158, but dominated by Geon 27 grains (Hill et al. 2018). Photo 19. Mafic dike at Stop 16 showing sparse, small, plagioclase phenocrysts. Hammer handle is 33 cm long. STOP 16. Nipissing gabbro and post-Nipissing dike UTM co-ordinates 361850E, 5164825N The medium-grained, slightly greenish gabbro is typical of the Nipissing intrusions in the Elliot Lake area. At the stop, the gabbro is cut by a 3 m wide, near-vertical, sharp-walled fine-grained mafic dike that is east-trending. Small plagioclase phenocrysts occur throughout the dike (Photo 19). Return to vehicles and continue north. 17.4 km (10.88 mi). Park on the shoulder of the road and examine outcrops on the east side of the road. This is Stop 27 of Bennett et al. (1997) and Stop 1.15 of Bennett (2006). 246 �Proceedings of the 68th ILSG Annual Meeting - Part 2 STOP 17. Lorrain Formation, Cobalt Group Similar dikes have been observed elsewhere in the Elliot Lake area (Easton 2013a, 2013b), but have yet to be assigned to any specific dike swarm. It is possible they could be related to the circa 1750 Ma Trap dike swarm, or the 2125-2105 Ma Marathon dike swarm. The senior author attempted to obtain an age on the dike, but no suitable phases were recovered for geochronology. Hunt and Roddick (1987) reported a K-Ar whole rock age of 1325 Ma for this dike. UTM co-ordinates 361771E, 5166967N White to pale pink quartz arenite of the upper Lorrain Formation is exposed on the east side of the highway. The detrital zircon sample reported by Rainbird and Davis (2006) and Craddock et al. (2013) came from this Stop. Here, Geon 27 zircons are twice as abundant as Geon 26 zircons, and there are also several Geon 28 zircons present (see Table 2). Return to vehicles and continue north on Highway 639. 23.5 km (15.69). Junction Highway 639 and 546 at Little White River Road. End of Field Trip. Retrace route back to Elliot Lake and return to Sudbury. 21.8 km (13.63 mi). This is Stop 28 of Bennett et al. (1997) and Stop 1.16 of Bennett (2006). End of road log. 247 �Proceedings of the 68th ILSG Annual Meeting - Part 2 References Buchan, K.L. and Card, K.D. 1985. Preliminary comparison of petrographic and paleomagnetic characteristics of Nipissing diabase intrusions in northern Ontario; in Current Research, Part A, Geological Survey of Canada, Paper 85-1A, p.131140. Ames, D.E., Davidson, A., Buckle, J.L. and Card, K.D. 2005. Geology, Sudbury bedrock compilation, Ontario; Geological Survey of Canada, Open File 4570, scale 1:50 000. 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, Proceedings, v.63, pt.1, p.9-10. Card, K.D. 1964. Metamorphism in the Agnew Lake area, District of Sudbury, Ontario, Canada; Geological Society of America, Bulletin, v.75, p.1011-1030. ——— 1976. 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Petrology of the Bruce and Gowganda formations and its bearing on Huronian sedimentation in the Espanola-Willisville area, Ontario, Canada; Paeleogeography, Paleoclimatology and Paleoecology, v.6, p.5-36. Chandler, F.W. 1973. Geology of McMahon and Morin townships, District of Algoma, Ontario Division of Mines, Geoscience Report 112, 77p. Davey, S., Bleeker, W., Kamo, S[.L.], Davis, D.[W], Easton, M.[R.] and Sutcliffe, R.H. 2019. Ni-Cu-PGE potential of the Nipissing sills as part of the ca. 2.2 Ga Ungava large igneous province; in Targeted Geoscience Initiative: 2018 report of activities; Geological Survey of Canada, Open File 8549, p.403-419. ——— 1976. Geology of the Saunders Lake area, District of Algoma, Ontario Division of Mines, Geoscience Report 155, 46p. Chandler, F.W., Young, G.M. and Wood, J. 1969. Diaspore in early Proterozoic quartzite (Lorrain Formation) of Ontario; Canadian Journal of Earth Sciences, v.6, p.337-340. Davidson, A., and van Breemen, O. 1994. U-Pb ages of granites near the Grenville Front, Ontario, in Radiogenic age and isotopic studies: report 8, Geological Survey of Canada, Current Research 1994-F, p.107–114. Chen, Y.D., Krogh, T.E. and Lumbers, S.B. 1995. Neoarchean trondhjemitic and tonalitic orthogneiss identified within the northern Grenville Province in Ontario by precise U-Pb dating and petrologic studies; Precambrian Research, v.72, p.263-281. 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.383-386. Church, W.R. 1967. The occurrence of kyanite, andalusite and kaolinite in lower Proterozoic (Huronian) rocks of Ontario; Geological Association of Canada–Mineralogical Association of Canada, Annual Meeting, Kingston, Ontario, Program with Abstracts, p.14. 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. Coleman, A.P. 1905. The Lower Huronian ice age; Journal of Geology, v.16, no. 2, p.149-158. Debicki, R.L. 1990. Stratigraphy, paleoenvironment and economic potential of the Huronian Supergroup in the southern Cobalt embayment; Ontario Geological Survey, Miscellaneous Paper 148, 154p. Collins, W.H. 1925. North shore of Lake Huron; Geological Survey of Canada, Memoir 143, 160p. Clough, C.E. and Hamilton, M.A. 2017. Matachewan LIP revisited: a revised, high-resolution U-Pb age for the East Bull Lake intrusion and associated units; Geological Association of Canada—Mineralogical Association of Canada, Abstracts, v.40, p.65. Dietz, R. S. and Holden, J.C, 1966. Miogeosynclines in space and time; Journal of Geology, v.74, p.566-583. Easton, R.M. 2003. 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The Grenville Province; in Variations in Tectonic Styles in Canada, Geological Association of Canada, Special Paper 11, p.263-334. 255 �� 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, 2022 Description An account of the resource Institute on Lake Superior Geology. Sudbury, Ontario. May 10-11, 2022. 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 2022 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/795d9a865d55c3086bcda137ccad1c7f.pdf 59d34dc03c74139c37d2d5698efda839 PDF Text Text Volume 67, Part 1* 67th ANNUAL MEETING VIRTUAL EVENTS May 11, 14, 18, 21, 2021 Hosted by: Peter Hollings, Mark Smyk, and Mark Jirsa Co-chairs Lakehead University and Minnesota Geological Survey The Sleeping Giant 66th *The Annual meeting originally planned for 2020 in Mountain Iron, Minnesota, was cancelled due to the Covid-19 pandemic, and thus, no Proceedings Volume 66 was produced. To mitigate potential impact from the ongoing pandemic, the 67th Annual meeting was held in a virtual format. As a result, there were no fieldtrips, and only this Part 1 of Proceedings Volume 67 was produced. We would like to thank the Lakehead University Technology Services Centre, particularly Blain Boyd, Shawn Hartviksen and John Bonofiglio, for providing the Zoom account and technical support for this virtual meeting. i �Proceedings Volume 67, Part 1: Program and Abstracts Table of Contents Institutes on Lake Superior Geology, 1955-2019 iii Sam Goldich and the Goldich Medal Goldich Medal Guidelines Goldich Medalists and Goldich Medal Committee Citation for Goldich Medal Award to Allan MacTavish v vii ix x Honoring the Pioneers of Lake Superior Geology-Newton Horace Winchell xii Memoriams for John Klasner, Thomas Waggoner, Ronald Seavoy, and John Heine xvi Eisenbrey Student Travel Awards xvii Joe Mancuso Student Research Awards xviii Doug Duskin Student Paper Awards and Award Committee Institute Board of Directors Report of the 65th Annual Meeting (2019) TECHNICAL PROGRAM ABSTRACTS xix xx xxi xxv 1-72 Reference to material in Part 1 should follow the example below: Authors(s), 2021, abstract title. 67th Institute on Lake Superior Geology Proceedings v. 67, Part 1Program and Abstracts, p. XX Published by the 67th 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 Digital version of the volumes are available on-line at http://www.lakesuperiorgeology.org. lSSN 1042-99 ii �Institutes on Lake Superior Geology, 1955-2019 # Year Meeting Location 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 Year 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 56 2010 57 58 59 60 61 62 63 64 2011 2012 2013 2014 2015 2016 2017 2018 65 66 67 2019 2020 2021 Meeting Location 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 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 International Falls, Minnesota M. Jirsa, P. Hollings, T. Boerboom, P. Hinz & M.Smyk Ashland, Wisconsin T. Fitz Thunder Bay, Ontario P. Hollings Houghton, Michigan T.J. Bornhorst & A. Blaske Hibbing, Minnesota J. Miller & M. Jirsa Dryden, Ontario R. Cundari & P. Hinz Duluth, Minnesota J. Miller, C. Schardt, & D. Peterson Wawa, Ontario A. Pace, A. Wilson, & T.J. Bornhorst Iron Mountain, Michigan L. Woodruff, W. Cannon, & E.K. Stewart Terrace Bay, Ontario P. Hollings & M.C. Smyk Meeting cancelled due to Covid-19 pandemic Virtual meeting P. Hollings, M. Smyk, M. Jirsa iv �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 MEDALLION �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. 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. vii �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 2019 Mark Severson 1983 Burton Boyum 2002 Ernest K. Lehmann 2021 Allan MacTavish 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 (Not awarded 2020) 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 Goldich Medal Committee Serving through the meeting year shown in parentheses. Dan England (2017-2022) Eveleth Fee Office (Committee Chair) Steve Kissin (2018-2023) Lakehead University Dorothy Campbell (2020-2024) Ontario Geological Survey Terms of the committee members have been extended 2 years due to cancelation of the 2020 meeting and the logistical difficulties of voting during the 2021 virtual meeting. Presentation of the award to Allan MacTavish was to have occurred at the 2020 meeting, but was delayed one year. 2021 GOLDICH MEDAL RECIPIENT ALLAN MACTAVISH ix �Citation for Allan MacTavish ILSG members, Goldich Medal recipients and guests, it is my privilege to present the citation for this year’s recipient of the Goldich Medal, Allan MacTavish. Allan received his Bachelor of Science (B.Sc.) degree in geology from Laurentian University in 1977. Over the past 44 years, he has spent much of his career working as an economic geologist in the Lake Superior region, both in the mineral exploration industry and with the Ontario government. Early on, Allan also somehow found the time to be a part-time graduate student at Lakehead University, where he completed a Master of Science (M.Sc.) thesis entitled “The Geology, Petrology, Sulphide and Platinum-Group Element Mineralization of the Quetico Intrusions, Northwestern Ontario” in 1992. Since completing this project, much of Allan’s work has been focussed on the evaluation of copper-nickel-platinum group element (Cu-Ni-PGE) deposits. While working in the private sector, Allan has led numerous field exploration programs in the Lake Superior region that have targeted Cu-Ni-PGE, volcanogenic massive sulphide and gold deposits. Most notably, he has served since 2007 as Exploration Manager for Magma Metals (Canada) Ltd./Panoramic PGMs (Canada) Ltd. and most recently as Vice President Project Manager with Clean Air Metals Inc., with a primary focus on the Thunder Bay North Cu-Ni-PGE project. At Thunder Bay North, Allan has led teams of geoscientists from the grassroots exploration phase, through to mineral resource delineation. The initial exploration success at Thunder Bay North, largely attributed to Allan’s leadership, sparked a staking rush in the Nipigon Embayment and consequently, a flurry of exploration activity and investment in the region. Although exploration activity in the project area has waxed and waned over the past 14 years, Allan’s persistence is now paying off, with Clean Air Metals having reported a significant mineral resource expansion during the winter of 2021, along with an exciting new intersection of highgrade massive sulphides. Furthermore, his work on the Thunder Bay North project, and his support of academic research work into the Current Lake intrusive complex, has made significant contributions to our understanding of chonolith-hosted magmatic sulphide deposits and the greater Midcontinent Rift system. One of Allan’s most notable skills is his competence as a field mapping geologist, an aspect of geology that has largely become a lost art. This skill was developed through his long list of positions, including a stint as a Field Geoscientist with the Ontario Geological Survey, during which he mapped the Montcalm greenstone belt. Even in the private sector, any project under Allan’s watch can be counted on to include solid geological mapping. This steadfast belief in the core fundamentals of geology has led to a significant advancement in our understanding of many geological districts in the Lake Superior region, including the Hemlo gold camp, the Coldwell x �complex, the Nipigon Embayment, the Atikokan-Quetico district and the Abitibi greenstone belt. Understanding the importance of a solid foundation in geological mapping as part of the training for future geoscientists, Allan has also been a long-time supporter of the Precambrian Research Centre (PRC), focussing his contributions toward the PRCs mapping school and the students therein. It was during Allan’s time working on his M.Sc. project, when I was a fellow graduate student, that I first got to know him. I quickly gained an appreciation for his knowledge and enthusiasm for the science of geology. I consider myself fortunate that, shortly after graduation, I became one of the many aspiring geoscientists who have benefitted from Allan’s mentorship when he hired me as a field assistant for a short-term project. More recently, the Lakehead University Student Chapter of the Society of Economic Geologists (SEG) has also been privy to Allan’s wisdom and knowledge. Since its inception in 2014, he has served as the Industry Representative for the SEG student chapter, assisting the group with fundraising and field trip endeavours, most notably trips to Arizona and Ireland. The Geology Department at Lakehead University has long known the benefit of having Allan as an associate, whether as a voice for industry to help direct departmental decisions, or as an employer hiring a geology student during or immediately following their studies. Allan’s contributions to the Institute on Lake Superior geology are many. He has tirelessly supported the ILSG by serving as a Board member, Goldich Award Committee member, Annual Conference planning committee member, and as an individual contributor to the ILSG Student Travel Scholarship. Dating back to 1985, Allan is credited as an author for 7 abstracts and 4 field trip guides, plus as an editor for 3 field trip guidebooks. Allan has led field trips for ILSG annual meetings, most recently the Coldwell Complex trip at the 2019 Terrace Bay meeting. He was also the co-organizer and field trip leader for the 2017 ILSG Iceland Geological Field Trip, and organized and led the 2020 Hawaii Geological Field Trip, which also happened to be the last ILSG in-person event to occur prior to the onset of the global Covid-19 pandemic (let’s all hope that we can once again get together for ILSG field trips in 2022). Anyone who has had the pleasure of knowing Allan over the years immediately becomes aware of his enthusiasm for geoscience and his deep knowledge of the rocks of the Lake Superior region. Allan’s contributions to the understanding of the geology of the Lake Superior region, as well as his unwavering support of the Institute on Lake Superior Geology, make him a worthy recipient of the 2021 Goldich Medal. Submitted by Mark Puumala, M.Sc., P.Geo. Senior Manager, Resident Geologist Program Ontario 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) 2018-2020 not presented 2021 Newton Horace Winchell (1839-1914) xii �2021 Pioneer of Lake Superior Geology Newton Horace Winchell (1839-1914) “Minnesota’s Geologist” Citation by Jim Miller, Former Associate Professor, University of Minnesota Duluth and Senior Geologist, Minnesota Geological Survey, with insights from author Sue Leaf. Newton Horace Winchell was first director of the Minnesota Geological and Natural History Survey at the University of Minnesota from 1872 to 1898. During his 26 year tenure, Winchell and his staff visited every corner of the state to document its bedrock geology, mineral deposits, quarry stone resources, glacial deposits and landforms, paleontological record, modern flora and fauna, and hydrologic framework. The survey’s findings were recorded in 24 Annual Reports, which summarized each year’s accomplishments and commonly totaled hundreds of pages, and in a six-volume final report. Pdfs of these reports are all available from the University of Minnesota Digital Conservancy (https://conservancy.umn.edu/handle/11299/708). As skillfully told in Sue Leaf’s (2020) engaging and intimate biography of Winchell’s private and professional life, Newton Horace Winchell (NHW) was born and raised as the middle child in a large family near the New York-Massachusetts-Connecticut borders. At the age of 15 after 8th grade, he began teaching in small schools in the area. In 1857, he moved to Ann Arbor, Michigan to live with his older (by 15 years) brother Alexander, who was a professor of natural history at the University of Michigan and director of the Michigan Geological Survey. Over the next 12 years, NHW worked as an elementary school teacher/principal in various southern Michigan towns, got married, had two children, and also intermittently took classes at the U of Michigan. In the summer of 1860, he was hired by the Michigan Geological Survey as a botanist and explored the coastlines of Lakes Huron and Michigan. NHW completed his BA degree in 1866, committed to geology as a profession, and began to work more regularly for the Michigan Survey. In 1869, he completed his MS degree at U of Michigan and became Assistant State Geologist under his brother. In 1870, the Michigan legislature defunded the geological survey, which prompted NHW to work the summers of 1871 and 1872 for the Ohio Geological Survey. In July 1872, University of Minnesota president, William Folwell, offered NHW the position of director of the newly created Minnesota Geological and Natural History Survey. With an initial legislative appropriation of $1000 per year, Folwell estimated that the survey would need 20 years to accomplish its task of conducting “a thorough geological and natural history survey of the State". In addition, he would be responsible for teaching two geology courses each year and organizing a natural history museum. NWH accepted the offer, and began field work that fall. The first six field seasons of the Survey (1872-1877), NWH and his small staff focused their attention on the geology, geomorphology, paleontology, and economic geology of southern Minnesota counties. In 1876, Winchell’s responsibility for the University museum was relieved by the hiring of zoologist, Clarence Herrick, as museum director. HNW’s other non-survey obligation of teaching college classes was removed just before the 1879 field season by the hiring of another zoologist, Christopher Hall, who was also appointed assistant director of the survey. Additionally significant hires in 1879 were those of glacial geologist Warren Upham and paleontologist Edward Urlich. Although NHW was well versed in both these disciplines, their contributions and other staff hires in the 1880’s, freed Winchell to devote many months of field xiii �work per year over the next decade and to accomplish his most challenging task –deciphering the Precambrian geology of northeastern Minnesota. The Paleozoic, Cretaceous and Quaternary geology of southern and western Minnesota was summarized in Volumes 1 (1884; 673 p.) and 2 (1888; 671 p.) of the final report written by NHW and Upham. The paleontology of Ordovician and Cambrian rocks (at the time interpreted to be Upper and Lower Silurian) and of Cretaceous rocks was summarized in the 1200+ page Volume 3 that was published in two parts (pt. 1 – 1895; Pt. 2 - 1897). The task of mapping and deciphering the Precambrian geology of northeastern Minnesota started in the summer of 1878 when NWH and three others paddled along the Lake Superior shoreline from Duluth to the Canadian border. In all, they collected over 300 samples. Then, from September to November of that same year, NHW and two Ojibwe guides paddled and portaged from Grand Portage, to Lake Vermilion, down the Embarrass, St. Louis and ultimately the Mississippi rivers to Minneapolis – a 500+ mile trek! Winchell devoted most of his remaining field seasons to northern Minnesota geology with the able assistance of his son (Horace), his sonin-law (Uly), and his brother (Alexander). With iron mining taking off on the Vermilion Range (1st ore shipped -1884), and then the Mesabi Range (1st ore shipped - 1892), much of NHW’s attention was focused on that geology. Also, beginning the mid-1880’s, Winchell regularly employed petrography to better understand the granite, gabbro, greenstone, and basalt that he encountered. In fact, in 1895, he took a year-long “sabbatical” to Paris to learn petrographic methods from the French petrographers Fonque and Michel-Levy. Still, as I highlighted in an ILSG talk in 2004, Winchell struggled mightily to understand the genesis of Precambrian crystalline rocks. In the introduction of Volume 4 of the Final Report (1899), wherein he summarizes his studies of northern Minnesota geology, NWH comments: "Here [among the crystalline rocks] the geologist is deprived of his usual guides and guys, and finds himself floundering in a muddy sea of innumerable conflicting currents". Some other notable aspects of Winchell’s remarkable career include his role in the development of the first scientific journal exclusively devoted to geology – The American Geologist. He served as its managing editor and publisher from its inception in 1888 to its merger with Economic Geology in 1905. Part of his motivation for creating this journal was to give a voice to state survey geologists which had the potential of being dismissed by academics and experienced overreach by the USGS. Also, as a founding fellow, NHW was instrumental in the creation of the Geological Society of America, also starting in 1888. After the Minnesota survey ended in 1899 and he got closure on Volumes 5 (1900, Uly Grant’s PhD thesis) and 6 (1901, atlas of county geologic maps), NHW focused the final decade of his life on archeological studies with the Minnesota Historical Society, particularly as they applied to native peoples who lived in the midcontinent before, during and after the last glacial episode. He published Aborigines of Minnesota in 1911 (743 p. 642 figures, 32 plates). As Sue Leaf points out, this well received book was “comprehensive in scope, methodical in approach, and meticulous in detail; a worthy companion to the voluminous Geology of Minnesota.” Winchell died on May 2, 1914 due to surgical complications at the age of 74. Newton Horace Winchell’s geological survey took the blank canvas of Minnesota’s landscape (Fig. 1) and began the process of revealing the rich diverse geology that we now recognize (Fig. 2) and continue to embellish and improve upon. As Sue Leaf’s book proclaims in its title, Newton Horace Winchell indeed deserves the designation as “Minnesota’s Geologist,” and as one of the great pioneers of Lake Superior geology. xiv �Figure 1 – Geological map of Minnesota in 1872 at the beginning of the Minnesota Geological and Natural History Survey References (Annual Report #1). Figure 2 – Geological map of Minnesota in 1900 at the end of the Minnesota Geological and Natural History Survey (Final Report, V. 6). References Leaf, Sue, 2020, Minnesota’s Geologist - The Life of Newton Horace Winchell. University of Minnesota Press, Minneapolis, MN, 261p. Miller, James D., 2004, N.H. Winchell's study of the Keweenawan Supergroup rocks of northeastern Minnesota, 1872-1900. Proceedings and Abstracts, 50th Institute of Lake Superior Geology, Duluth, MN, p. 117-118. xv �In Memoriams The ILSG lost four of its’ long-time members since our last in-person meeting in 2019. Brief descriptions of their professional lives follow, excerpted from online obituaries and tributes. They were active members throughout their professional careers and beyond. They all made significant contributions to the science and to the Institute. They will be missed. Ronald Seavoy passed away on March 25, 2020. He was born July 6, 1931 in New York City and was raised in Chicago. He received his Bachelor of Science degree in Geology in 1953 from the University of Michigan, and his Master of Arts (1963) and PhD (1969) in History from Michigan. He served in the U.S. Army from 1953-1955. He was hired by the Department of History at Bowling Green State University in 1965 where he taught U.S. Constitutional History, U.S. Business History, as well as numerous survey courses until his retirement in 1991 as Professor Emeritus. Ron is the author of ten books and nineteen articles on the subjects of American business history, famine in developing countries, political economy, and mining exploration. Before and while employed as a historian, Ron worked as an exploration geologist for Canadian Johns Manville, International Nickel Company, Alcoa, Burwest, Western Nuclear, and Cleveland Cliffs Mining. He was a member of the Organization of American Historians, the Society of Economic Geologists, and the Institute of Lake Superior Geology. Over the many years of Ron’s involvement with ILSG, he donated thousands of dollars to the Mancuso Student Research Award, which was named for Ron’s good friend and colleague. John Klasner passed away on July 15, 2020. John was born June 22, 1935 in Flint, Michigan. He attended Michigan State University and Michigan Technological University where he graduated with degrees in geology and geophysics. He worked in mineral and oil exploration in the US, Canada, Africa, and the Middle East. John then became a geology professor and directed the honors program at Western Illinois University. He also was employed by the USGS for mapping in the Upper Peninsula of Michigan and Northern Wisconsin. John was a member of several professional organizations, including the Institute on Lake Superior Geology. John Heine passed away August 17, 2020 in Duluth. He was born in Minneapolis and graduated from University of St. Thomas in St. Paul. John moved to Duluth and pursued graduate studies in the Geology Department at the University of Minnesota Duluth. During his 30+ year career with the Natural Resources Research Institute (NRRI), Minerals & Metallurgy Strategic Research Initiative, he worked on a wide variety of field- and laboratory-based geological studies to better understand the mineral characteristics and mineral potential of the state. He retired from the NRRI in 2020 after an accomplished career. John loved field work, and he never stopped learning and teaching. Thomas Waggoner of Negaunee, Michigan, died December 7, 2020. Tom acquired an MSc in geology at Michigan State University, and worked for Cleveland Cliffs Mining. He helped Cliffs make the transition from underground mining of direct shipping ore to surface mining and production of taconite 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. During his long career with Cliffs (1965-1997), Tom actively participated in the teaching function of the Institute on Lake Superior Geology. He led 45 field trips for various organizations, and made numerous poster and oral presentations at annual meetings of the Institute. After his retirement in 1997, he continued to serve the mining industry. Tom remained active in ILSG, leading a field trip during the Iron Mountain meeting in 2018 and attending the Terrace Bay meeting in 2019. We all benefited from his knowledge, enthusiasm, and support. He will be sorely missed by all who knew him. (excerpted from a tribute by Alan Strandlie, December 10, 2020) xvi �Eisenbrey Student Travel Awards Because the 2020 meeting was canceled, and the 2021 meeting is virtual, no awards are granted. The program will continue when it’s safe to travel and meet in-person. 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. xvii �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 by Ronald Seavoy, Joe’s colleague and friend. “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 2020, the ILSG Board of Directors selected one student to be granted research funding of $1000.00 from the Joe Mancuso Student Research Fund. The awardee was: Ann Marie Prue University of Minnesota-Duluth, MSc, Department of Earth and Environmental Sciences, TOPIC: Multi-method Geochemical Investigation of the Neoarchean Soudan Iron Formation, NE, Minnesota. [editor’s note: She is presenting her research at this meeting] xviii �Doug Duskin Student Paper Awards Because the virtual format of this meeting would likely complicate the process of judging student presentations, no awards are granted. The awards program will continue when it’s safe to travel and meet in-person. 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. xix �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 have been extended 2 years due to cancellation of the 2020 meeting, and the difficulties of virtual voting by the membership during the 2021 meeting. Mark Smyk (2019-2024) – Lakehead University Esther Stewart (2018-2023) – Wisconsin Geological & Natural History Survey Anthony Pace (2017-2022) – Ontario Ministry of Energy, Northern Development and Mines Pete Hollings - Secretary (2016-2024) – Lakehead University Mark Jirsa – Treasurer (2017-2022) – Minnesota Geological Survey (retired) xx �REPORT OF THE 65th ANNUAL MEETING OF THE INSTITUTE ON LAKE SUPERIOR GEOLOGY Terrace Bay, Ontario Lakehead University and the Ontario Geological Survey (OGS), hosted the 65th Annual Institute on Lake Superior Geology on May 07 – 10, 2019 at the Terrace Bay Cultural Centre, Terrace Bay, Ontario. The meeting consisted of two days of technical sessions with pre- and posttechnical session field trips. Mark Smyk (OGS) and Pete Hollings (Lakehead University) were cochairs for the 2019 meeting. Invaluable advice and logistical support was provided by Dean Main and Michelle Malashewski of the Township of Terrace Bay’s Community Development department. Access to the Terrace Bay Public Library was graciously provided by Mary Deschatelets, Library CEO. Patty Cobin and Ted Bornhorst (A.E. Seaman Mineral Museum, Michigan Technological University) handled pre-meeting registration. Ted also supplied the poster boards. Generous contributions to the ILSG general fund and in support of 2019 student travel scholarships came from: Abitibi Geophysics, Barrick Gold Corporation, Benton Resources Corp., Geological Society of Minnesota, Greenstone Gold Mines, Landore Resources Canada Inc., Noront Resources Ltd., North American Palladium Ltd., Northwestern Ontario Prospectors Association, Panoramic PGMs (Canada) Limited, Premier Gold Mines Limited, Stillwater Canada Inc., Superior Lake Resources Limited, Thunder Bay Community Economic Development Commission, Wesdome Gold Mines Ltd., and Wolfden Resources Corporation. Individual Contributors to Student Travel Scholarship were: Al MacTavish, Mary Kay Arthur, L. Gordon Medaris, Jr. and Nick Swanson-Hysell. Total meeting registration was 114, which included 15 students. Welcoming remarks were provided by Terrace Bay Mayor Jody Davis. Elder Raymond Goodchild of Pays Plat First Nation provided traditional blessings for the meeting. Proceedings Volume 65 was published in two parts. Part 1 – Program and Abstracts, compiled and edited by Mark Puumala (OGS), contains 51 published abstracts for 32 oral and 19 poster presentations. Students presented 6 oral and 3 poster presentations. Part 2 – Field Trip Guidebooks, was compiled and edited by Al MacTavish and Pete Hollings. It contains descriptions of four premeeting and four post-meeting field trips. The 65th ILSG marked the first time in the Institute’s long history that its annual meeting was held in Terrace Bay. The meeting location enabled organizers to offer trips that showcased a variety of Archean and Proterozoic rocks along the north shore of Lake Superior. Although the majority of the field trips had been offered at previous ILSG annual meetings (e.g. Marathon 1995; Nipigon 2005), they greatly benefitted from the new mapping, research, discoveries and interpretations that had taken place in the intervening years. Local exploration companies graciously provided information and access to their properties. Unfortunately, a late spring precluded visiting the Slate Islands (lake ice) and the Midcontinent Rift-related carbonatites and diatremes in the Dead Horse Road area (snow). Slate Islands trip delegates were provided with an impromptu, “eclectic” field trip along Highway 17 between Schreiber and Marathon by Mark Smyk. Those who planned to see the MCR carbonatites and diatremes switched to other available trips. xxi �A list of field trips is provided below: Pre-meeting field trips (and leaders) on Tuesday, May 07: 1) 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 and Rob Cundari, Ontario Geological Survey) (*denotes cancelled due to snow/ice conditions) Post-meeting field trips on Friday, May10: 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 and David Good, Western University) 7) Building and ornamental stone sites of the Marathon area (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). Pre- and post-meeting field trips attracted 172 registrants, once again reaffirming their integral role in the allure and success of ILSG annual meetings. The vast majority of registrants and invited guests attended the annual ILSG banquet on Wednesday night. Although a Homer Award overview presentation was given, no “recipients” were identified during the 2019 annual meeting! As always, a highlight of the post-banquet activities was presentation of the 2019 Goldich Medal. This year’s very deserving recipient was Mark Severson. Mark’s wife, Laurie, and daughter, Allison, attended the banquet and award ceremony. The Goldich Medal citation was presented by George Hudak, his colleague for many years. George described Mark’s contributions to ILSG and to the greater understanding of Minnesota’s geology over several decades during his time as a student, in his role at the Natural Resources Research Institute-University of Minnesota Duluth, in his collaborations with the Minnesota Geological Survey and later in the mining and exploration industry. Mark is indeed a worthy recipient of this prestigious award. The 65th ILSG included a radical departure from the usual post-banquet guest speaker tradition. In lieu of a guest speaker, Master of Ceremonies, Mark Smyk, moderated a trivia contest entitled “ILSG Geo-Pardy.” It tested our delegates’ knowledge of ILSG history (e.g. past host cities, Goldich Medalists), geology in pop culture and “fun facts” about Lake Superior geology. Scores were tabulated and winners were announced the following day. Donated prizes and bragging rights attended the “awards ceremony.” In 2019, the student paper committee had its usual difficult job of selecting the best among six excellent oral presentations and three poster presentations for the Doug Duskin Student Paper Awards. The 2019 committee comprised Katarina Bjorkman (Bjorkman Prospecting), George Hudak (Natural Resources Research Institute–UMD) and David Good (Western University). The committee awarded three prizes with best talk going to Sophie Kurucz for her talk on “Paleoproterozoic snowball earth? Sedimentology and geochemistry of a Huronian glacial cycle” xxii �and runner up prizes going to Kira Arnold (Geology and geochemistry of the Terrace Bay Batholith, N. Ontario) and Munira Afroz (Sulfur, Carbon, and Oxygen Isotope Geochemistry of ~2.93 Ga Mesoarchean Chemical Sedimentary rocks in the Red Lake Area, Ontario) Eisenbrey Student Travel Grants were given to nine students: Paul Bielski (Lakehead) $200, Kira Arnold (Lakehead) $200, Jaqueline Drazan (UMD) $400, Jackie Wrage (University of Michigan) $600, Brittany Ramsay (Lakehead) $200, Munira Afroz (Lakehead) $200, Thomas Bodden (MTU) $500, Sophie Kurucz (Lakehead) $200 and Chanelle Boucher (Lakehead) $200. The Institute’s Board of Directors met on Wednesday, May 8, 2019 and a brief overview of the meeting is provided below: 1. Accepted report of the Chairs for the 64th ILSG, Iron Mountain, Michigan; as printed in the 2019 Proceedings Volume, and minutes of last Board meeting, May 16, 2018 (Hollings) 2. Received, discussed, and accepted 2018-2019 ILSG Financial Summary (Jirsa). 3. Received, discussed, and accepted 2018-2019 Report of the Secretary (Hollings). 4. Approved Mark Smyk as on-going ILSG Board member 5. Approved Mountain Iron, Minnesota as the site for the 66th annual ILSG meeting, hosted by Mark Jirsa, with help from Amy Radakovich and Terry Boerboom, all of the Minnesota Geological Survey. The dates of the meeting were set for May 12-15, 2020. 6. There was discussion as to future meeting locations, with Sudbury being suggested as a possibility. In addition, Esther Stewart has been in discussion with Rob Lodge and Dyanna Czeck about hosting a meeting in Wisconsin. 7. Discussed and approved replacing Klaus Schultz as the “member from government” on Goldich Committee (end of term 2019) with Dorothy Campbell (OGS). 8. Liability Insurance for the Institute was discussed and approved by the Board as an ongoing, annual expense. The Treasurer will investigate further to ensure that insurance is in place for the 2020 meeting. 9. It was approved that the Treasurer should purchase 10 Goldich medals, as only one remains. 10. Discussed and approved renewal of Pete Hollings as Institute Secretary (end of term 2022). This was later approved by a vote of the membership. The 65th 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 and Lakehead University who were pressed into action as editors, field trip leaders and drivers. The Township of Terrace Bay was extremely hospitable in providing facilities (including the Cultural Centre, Library and Bowling Alley), suggestions and liaison services. We enlisted five different businesses to admirably handle our many catering needs in trying to “spread business around” in both Terrace Bay and neighbouring Schreiber. We heard many favourable comments from local people and businesses, not only about our patronage, but also about their interactions with ILSG delegates. They were thrilled to have us. We even received coverage in the local newspaper, provided by a reporter who tagged along on a field trip and popped by the Technical Sessions. We were happy that we were finally able to host the annual meeting in Terrace Bay. It was well worth the wait! Like most ILSG annual meetings hosted in smaller communities in Ontario, attendance was lower in 2019. However, those who did attend thoroughly enjoyed the meeting, facilities and field trips. Always flexible and understanding, our delegates took any changes in programming in stride. xxiii �We continue to encourage student attendance and new delegates in ensuring a healthy future for the Institute. We appreciate the support of our respective organizations, the Ontario Geological Survey and Lakehead University, and recognize the efforts that everyone made in travelling, attending, presenting and participating in activities that made the a 65th Annual Meeting what we feel was a great success. We encourage others to consider hosting and contributing to future meetings. It is truly a rewarding experience. Mark Smyk and Pete Hollings Co-Chairs, 65th Institute on Lake Superior Geology xxiv �TECHNICAL PROGRAM (Times are EDT) TUESDAY, MAY 11, 2021 SESSION 1 9:45-10:00 Welcome, introduction, and instructions (meeting chairs) 10:00-10:20 Paul Bielski and Philip Fralick Comparisons of Archean Iron Formations deposited at Shallow vs Deeper Depths 10:20-10:40 Bailey Drover, Mary Louise Hill, and Shannon Zurevinski The deformation and alteration of granitoid plutons in the Wabigoon subprovince 10:40-11:00 Manuel Duguet Geochemistry and Age of Archean Volcaniclastic and Mafic Intrusive Rocks, Georgia Lake Area, Quetico Subprovince, Northwestern Ontario 11:00-11:20 Seamus Magnus Chemostratigraphy of the western Schreiber-Hemlo Greenstone Belt—Results and Regional Implications 11:20-11:40 Brittany Ramsay, Philip Fralick, Stefan Lalonde, Paul Bielski, and Laureline Patry Environmental Control of Seawater Geochemistry in a Mesoarchean Peritidal System, Woman Lake, Superior Province 11:40-12:00 Michael Tamosauskas, Robert Lodge, M.A. Chong, Rasmus Haugaard, and Ross Sherlock The provenance, depositional environment and metallogenic implications of the Ament Bay Metasedimentary Assemblage, Sturgeon Lake Greenstone Belt, Northwest Ontario 12:00-12:20 Val W. Chandler V. W and Mark Jirsa Three-dimensional geologic mapping of Precambrian rocks in Minnesota: The creation of “removeable” geologic layers using gravity and magnetic data interpretation 12:20-12:40 Planavsky, Noah Negative Carbon Dioxide Emissions through Enhanced Silicate Weathering and the Lake Superior Region xxv �FRIDAY, MAY 14, 2021 SESSION 2 9:45-10:00 Introduction and instructions 10:00-10:20 Benjamin Drenth, Bill Cannon, Klaus Schulz, and Robert Ayuso Geophysical insights into Paleoproterozoic tectonics along the southern margin of the Superior Province, central Upper Peninsula, Michigan 10:20-10:40 Megan Landman and Mary Louise Hill Archean Orogenesis to Proterozoic Rifting: A structural history of Pass Lake, Thunder Bay, Ontario 10:40-11:00 Seamus Magnus Proterozoic Geology of the Schreiber–Terrace Bay Area 11:00-11:20 Ann Marie Prue and Latisha Brengman Preliminary pXRF results from Precambrian rocks of northern Minnesota 11:20-11:40 Nicholas Swanson-Hysell, Margaret Avery, Yiming Zhang, Eben Hodgin, and Terry Boerboom The paleogeography of Laurentia in its early years: new constraints from the Paleoproterozoic East-Central Minnesota batholith 11:40-12:00 Michael Easton The Critical Mineral potential of carbonatite and alkalic complexes in Ontario 12:00-12:20 Mark Puumala and Robert Cundari Critical Minerals Exploration and Development Potential in Ontario 12:20-12:40 Robert Cundari, Mark Puumala, Riku Metsaranta, and Dorothy Campbell Critical Mineral Potential on the North Shore of Lake Superior; Elements to ‘Structure’ our future POSTER SESSION 14:00-14:05 Introduction to 5-minute author summaries 14:05-14:10 Benjamin Gallagher and Marcia Bjornerud Reconstructing the hydration and carbonation history of the Presque Isle peridotite, Marquette Michigan: Insights into mechanisms of carbon sequestration in ultramafic rocks 14:10-14:15 Tara Lemke, Evan Weber, and Robert Lodge Characterization of Hydrothermal Alteration and Sulfide Ores at the Lynne ZnCu-Pb Deposit, Oneida Co. WI. xxvi �14:15-14:20 Trevor Nelson, Rory Johnson, Robert Lodge, Chong MA, Jeffery Marsh U-Pb geochronology and zircon trace-element geochemistry from granitoid plutons in the Neoarchean Sturgeon Lake greenstone belt, Ontario, Canada 14:20-14:25 Dana Peterson, P. Bedrosian and C. Finn 3-D Modeling of the Duluth Complex from geophysical data 14:25-14:30 Rebecca Price, Shannon Zurevinski, and Roger Mitchell A newly discovered orbicular occurrence within the Good Hope carbonatite, north of Marathon, Ontario 14:30-14:35 Shelby Short, Lillian Glodowski, and Robert Lodge Geochemistry and Petrography of Volcanic and Intrusive Rocks Hosting the Lynne Cu-Zn-Pb Deposit, Oneida County, WI 14:35-14:40 Esther Stewart, Billy Fitzpatrick, and Eric Stewart Geologic mapping identifies bedrock folds that may be significant for increased probability of arsenic detection in water wells 14:40-15:30 Breakout room discussions with Poster Presenters TUESDAY, MAY 18, 2021 SESSION 3 9:45-10:00 Introduction and instructions 10:00-10:20 Paul Bedrosian, Max Pace, and Katerina Zamudio Geophysical mapping of the eastern arm of the Midcontinent Rift in Upper Michigan 10:20-10:40 Sophie Mueller, J. Degraff, and D. Lizzadro-Mcpherson Structural Analysis and Interpretation of Deformation along the Keweenaw Fault System West of Lake Gratiot, Keweenaw County, Michigan 10:40-11:00 Nicholas Craik and Phil Fralick [presentation cancelled] Tracing redox pathways of the Mesoproterozoic Copper Harbour Conglomerate, Michigan 11:00-11:20 Ted Bornhorst Evolved seawater as the source of salinity for metamorphic-dominated oreforming hydrothermal fluids of the Keweenaw Peninsula native copper district, Michigan xxvii �11:20-11:40 Tien Grauch and S.J. Heller Integration of geophysical evidence indicates that anorthosite composes a significant portion of Grand Marais ridge, an inferred basement high in western Lake Superior 11:40-12:00 Laurel Woodruff and Tien Grauch Possible implications of a non-Archean Grand Marais Ridge, western Lake Superior 12:00-12:20 Eben (Blake) Hodgin, Nicholas Swanson-Hysell, Daniel Stolper, Andrew Turner, James DeGraff, Andrew Kylander-Clark, and Mark Schmitz Final inversion of the Midcontinent Rift during the Rigolet Phase of the Grenvillian orogeny FRIDAY, MAY 21 2021 SESSION 4 9:45-10:00 Introduction and instructions 10:00-10:20 Matthew Brzozowski, Peter Hollings, Allan MacTavish, Dawn EvansLamswood, Abraham Drost, and Derek Wilton Mineralizing processes in the Current Lake and Escape Lake conduit-type PGE– Cu–Ni deposits of the Thunder Bay North igneous complex, northwestern Ontario, Canada 10:20-10:40 David Good A Practical Approach to Using Incompatible Elements to Define Geochemical Correlation and a Common Mantle Source: Example Crystal Lake Gabbro and the Duluth Complex 10:40-11:00 Kyle Lachance, Jackie Kleinsasser, Adam Simon, Dean Peterson, and George Hudak Identifying the genesis of Fe-Ti oxide- and sulfide-bearing ultramafic intrusions in the Duluth Complex through sulfide geochemical analysis 11:00-11:20 Jackie Kleinsasser, Adam Simon, Dean Peterson, and George Hudak Textures and geochemistry of ilmenite and titanomagnetite in Fe-Ti oxide-bearing ultramafic intrusions of the Western Margin of the Duluth Complex, Minnesota 11:20-11:40 Amartya Kattemalavadi, Jackie Kleinsasser, Adam Simon, Dean Peterson, and George Hudak Olivine Geochemistry from Fe-Ti Oxide-Bearing Ultramafic Intrusions in the Duluth Complex, MN xxviii �11:40-12:00 Dave Dahl, Stacy Saari, and Thomas Lee Availability of Historical Airborne Geophysical Survey Data at Minnesota DNR 12:00-12:20 Allan MacTavish, Peter Hinz, Mary Kay Arthur, Robert Chataway, Jim Edberg, Tom Erickson, Steve Fox, Joan Furlong, Jim Gerlich, Lindsay Smith, and David Wilhelm Island of Hawaii Field Trip, February 11-21, 2020 12:20-12:30 Closing remarks MAY 21-AWARDS EVENING 19:00-20:30 • • • • • • Words of welcome (chairs) ILSG updates Goldich Medal presentation to Allan MacTavish 2022 ILSG meeting plans—Sudbury Ontario Honoring the Pioneers of Lake Superior Geology—Newton Horace Winchell (by Jim Miller and Sue Leaf) Closing remarks xxix �ABSTRACTS xxx �Geophysical mapping of the eastern arm of the Midcontinent Rift in Upper Michigan BEDROSIAN, Paul A, PACE, Max, and ZAMUDIO, Katrina D. Geology, Geophysics and Geochemistry Science Center, U.S. Geological Survey, Denver, CO The western arm of the Midcontinent Rift system (MRS), extending from western Lake Superior into Kansas, is somewhat understood due to exposures of MRS igneous and sedimentary rocks, scattered drillholes, and seismic imaging. By comparison, the eastern rift arm (ERA), extending from eastern Lake Superior through lower Michigan is poorly understood, with almost no outcropping volcanic rocks south of Michipicoten and Caribou Islands in Lake Superior, few boreholes that intercept MRS rocks, and the majority of the rift concealed beneath the Michigan basin. Seismic reflection sections in the eastern Lake Superior basin (Behrendt et al., 1990; Mariano and Hinze, 1994) and in Lake Michigan (Cannon et al., 1991) provide the primary constraints on the gross structure of the ERA. Unlike the western rift arm, which has undergone considerable post-rift shortening (Cannon et al., 1993), the ERA appears less deformed, with perhaps as much as 30 km of rift clastic and volcanic rocks preserved within a broadly symmetric basin (Behrendt et al., 1988) and about 5 km of anticlinal relief on the top of the volcanics (Mariano and Hinze, 1994). Under Paleozoic cover in the eastern part of Michigan’s Upper Peninsula, the ERA western margin is inferred from linear northwest-trending potential-field anomalies (Fig. 1a, b) and the occurrence of MRS basalt and rhyolite beneath ~2 km of Paleozoic and MRS clastic rocks in the St. Amour drillhole (Ojakangas and Dickas, 2002). In this examination of the ERA, we present results of a recent magnetotelluric survey over the ERA and a 3D electrical resistivity model derived from these data. The ERA is clearly imaged as a 150-km wide conductive rift basin with a sharp western margin and an irregular eastern margin that traces the distribution of mapped MRS clastics in Ontario (Fig. 1c). The lowest resistivities (10 Ω·m) correspond to the Jacobsville Ss, imaged as an east-dipping horizon extending to 2-3 km depth beneath the edge of the Michigan basin. Underlying the Jacobsville are moderately conductive rocks (100-500 Ω·m) that thicken basinward to 8-10 km depth with a nested inner basin that may be 20 km or more deep. The rift-fill rocks are attributed to a combination of Oronto Grp. clastics and MRS basalts, which cannot be separately distinguished within the resistivity model. In map view, the resistivity model at 10 km depth (Fig. 1d) shows along-rift changes in resistivity with a wavelength of ~50 km. Though data coverage is limited, both magnetic and gravity data (extending north into Lake Superior) show a similar pattern of highs and lows. A seismic reflection profile parallel to the rift axis (LS-15; Mariano and Hinze, 1994, their Fig. 7) interprets folding of the rift sedimentary and volcanic section and infers several kilometers of relief along strike. The linearity of the western rift margin suggests that both the thickest sections of rift basalt and Jacobsville are fault bounded. Together with the evidence for along-strike folding and faulting, we suggest the current structure of the ERA reflects post-rift NW-SE compression oriented along the rift axis, with the western margin marked by a transform fault (likely a reactivated normal fault of MRS age). Extrapolating farther north, we speculate that left-lateral offsets within the broad magnetic high that follows the western side of the ERA reflect a series of en echelon faults that together conspired to shorten the ERA by several tens of kilometers or more. �Figure 1. (a) magnetics, (b) gravity, (c) resistivity at 1 km depth and (d) resistivity at 10 km depth. Shaded regions show mapped or inferred MRS clastic rocks. White circles, stars, and lines indicate magnetotelluric stations, deep drillholes, and seismic reflection lines, respectively. Faults are shown in black. White dashed lines show interpreted structures. Potential field data are from Anderson and Grauch (2018). Anderson, E.D., and Grauch, V.J.S., 2018, Updated gravity stations and anomaly compilation over Lake Superior: U.S. Geological Survey data release, doi:10.5066/F7F18X8S. Behrendt, J.C., Green, A.G., Cannon, W.F., Hutchinson, D.R., Lee, M.W., Milkereit, B., Agena, W.F., and Spencer, C., 1988, Crustal structure of the Midcontinent Rift system: Results from GLIMPCE deep seismic reflection profiles: Geology, v. 16, p. 81, doi:10.1130/0091-7613(1988)016<0081:CSOTMR>2.3.CO;2. Behrendt, J., Hutchinson, D., Lee, M., Thornber, C., Trehu, A., Cannon, W., and Green, A., 1990, GLIMPCE seismic reflection evidence of deep—Crustal and upper-mantle intrusions and magmatic underplating: Tectonophysics, v. 173, p. 595–615. Cannon, W.F., Lee, M.W., Hinze, W.J., Schulz, K.J., and Green, A.G., 1991, Deep crustal structure of the Precambrian basement beneath northern Lake Michigan, midcontinent North America: Geology, v. 19, p. 207–210. 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, doi:10.1029/93TC00204. 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. 31, p. 619–628. 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: Sedimentary Geol., v. 147, p. 13–36. 2 �Comparisons of Archean Iron Formations deposited at Shallow vs Deeper Depths BIELSKI, Paul and FRALICK, Philip Department of Geology, Lakehead University, Thunder Bay, ON, Canada The occurrence of iron formation (IF) during the Archean is well documented, however the mechanisms of their genesis are poorly understood within shallow waters and even more 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, multiple shallow and deeper water iron formations were analysed by coarse and fine-scale geochemical analysis. The fine-scale method entailed using Laser Ablation Inductively Coupled Mass Spectrometry (LA-ICP-MS) alongside Scanning X-Ray Fluorescence (XRF), while the coarse-scale analysis is done with bulk rock geochemistry. The IFs used in this study were all deposited at approximately 2.7 Ga in different depositional settings. The shallow water IF is from Beardmore and Lake St. Joseph, which are both instances of deltaic sediments intermingled with sections of cherty-IF, and magnetite layers directly within deltaic distributary mouth bars. The deeper water IFs are from Temagami and Timmins in Ontario and Soudan, Minnesota. These IFs contain no discrete siliciclastic strata within the main body of IF, where chert and iron oxides are the dominant lithology. While the true depths of these deeper water formations are unknown, the lack of non-chemical sedimentary rock, excluding thin slate strata, and no storm-related sedimentary structures within layering suggests some distance from landmasses and depths from below storm wave base to abyssal. Siliciclastic turbidites from Beardmore have been included in this study to assist in identifying trends related to detrital contamination. Rare earth element (REE) concentrations, normalized with post-Archean Australian shale (PAAS), and analysed by bulk rock geochemistry, show similarities between deeper and shallow water deposits. The deeper water IFs contain pronounced HREE enrichment with prominent Eu and Y anomalies while shallow water IF do not show consistent Y anomalies but do contain relatively smaller Eu anomalies (Figure 1) despite siliciclastic contamination. Y/Ho is largest within the deep water IF with mainly supercondritic ratios from 28 to 55. Shallow water IF contains near chondritic (~28) ratios. The smaller Eu and Y anomalies are interpreted to be a combination of weaker Eu anomalies present in shallower water (cf. Planavsky et al., 2010) and contamination by siliciclastics, which have non-anomalous values. The lack of superchondritic Y/Ho in the shallow water IF samples is odd as Y anomalies are thought to be largest in shallow water during the Archean (cf. Planavasky et al., 2010) which may relate to a scavenging disequilibrium with the surrounding seawater as deeper water IF contain a consistent Y anomaly (Figure 1). Trace element analysis of redox sensitive elements U, V, and Cr have a strong relationship with Al2O3, indicating significant contributions of these elements from detrital material, especially within shallow water IF. However, V/Fe and Cr/Fe ratios in deeper water IF are comparable to shallow water ratios while having much lower levels of detrital contamination (i.e., Al2O3) (Figure 2). This indicates Cr and V were enriched in the precipitated phase, adding extra Cr and V to the sediment. These elements are more soluble in oxidizing fluids, but their solubility greatly decreases in reducing environments, e.g. the Archean deeper ocean. Thus, they required the presence of some oxygen in the weathering environment, but were reduced and precipitated in the Archean deeper ocean. Th/U ratios of IF samples may be used to identify enrichments of U, or depletions of Th, 3 �relative to the ratio found within igneous rock at a given time. At 2.7 Ga, the Th/U ratio within crustal rock was approximately 3.9, which is imprinted into the seawater ratio from the weathering and breakdown of minerals containing this ratio. Deep water IF contains low Th/U ratios in bulk rock geochemistry (Figure 2) and has even lower ratios (<1) in laser ablation data, which is less affected by detrital contamination. Shallow water IF bulk geochemistry shows values near that of average crustal rock, with variance around this value (~3.9). A more detailed understanding of differences between shallow and deeper water IF is obscured by the inevitable presence of detrital contamination. However, it is evident that Eu anomalies are larger and redox elements are enriched in deeper water IF. Shallow water IF may be more susceptible to changes in water chemistry within the deltaic depositional environment causing anomalous and/or variable REE and trace element geochemistry compared to the similar geochemistry between different deeper water IFs. Taylor-McLennan 1985-REEs 80 70 60 1 50 Y/Ho Rock/Post-Arch. Aust. Shale-PAAS 10 40 30 .1 20 10 .01 Ce La Nd Pr Eu Sm Tb Gd Y Dy Er Ho 0 .01 Yb Tm .1 1 Al 2 O 3 Lu 10 100 Figure 1. REE analysis. Note that the flatter pattern of the shallow IF is due to contamination. 5 100 4 10 3 Th/U Al 2 O 3 1 2 .1 1 0 .01 .1 1 10 .01 .01 .1 Th 1 V/Fe 2 O 3 10 100 Figure 2. Redox element analysis. Shallow water IF and turbidites correlate while deep water IF contains similar V/Fe2O3 ratios despite much lower amounts of Al2O3 (detritus). References Planavsky, N., Bekker, A., Rouxel, O. J., Kamber, B., Hofmann, A., Knudsen, A., & Lyons, T. W. (2010). Rare earth element and yttrium compositions of Archean and Paleoproterozoic Fe formations revisited: new perspectives on their significance and mechanisms of deposition. Geochimica et Cosmochimica Acta, 74(22), 6387-6405. 4 �Evolved seawater as the source of salinity for metamorphic-dominated ore-forming hydrothermal fluids of the Keweenaw Peninsula native copper district, Michigan BORNHORST, Theodore J. Department of Geological and Mining Engineering and Sciences, Michigan Technological University, Houghton, Michigan 49931 The Mesoproterozoic Midcontinent rift of Michigan’s Keweenaw Peninsula hosts the world’s largest accumulation of native copper. For more than 60 years, it has been generally accepted that burial metamorphic processes generated ore-forming hydrothermal fluids. However, Brown (2006) reasoned that metamorphic processes could not generate high enough salinity. This constraint requires an outside source of salinity. He proposed a hybrid hydrodynamic evolved meteoric water and metamorphic ore-forming fluid at depth and down-dip below the native copper deposits in the fluid source zone. Brown’s envisioned meteoric waters evolved during descent through the Oronto Group clastic sedimentary rocks and the salinity was dissolved from postulated evaporite horizons and/or scattered cement. An additional genetic constraint is that the ore-forming fluids were abnormally low in sulfur leading to the precipitation of native copper rather than copper sulfides (White, 1967). Because of very low sulfur magmas and degassing of sulfur upon eruption, the rift-filling Portage Lake Volcanics (PLV), host rocks of the native-copper deposits, and stratigraphically equivalent rocks of the source zone are sulfur-poor. Metamorphism of the source zone rocks resulted in sulfur-poor metamorphic ore-forming fluids. Brown’s hydrodynamic evolved meteoric water model lacks a mechanism to satisfy the constraint of very low sulfur fluids. Further, the reasoning presented by Bornhorst and Mathur (2017 and 2018) suggests that hydrodynamic evolved meteoric water had limited to no significant role in the ore-forming fluids. A role for seawater in generation of the native copper ore-forming fluids was proposed by Livnat (1983). However, seawater has not been widely accepted since the PLV were subaerially deposited and until recently there was only the debatable possibility of an incursion of seawater into the rift during deposition of the Nonesuch Formation. Recent evidence (Jones et al., 2019) suggests that the Nonesuch and underlying Copper Harbor Formations were deposited in a “braided fluvial-evaporitic shoreline-marine embayment triplet.” Johnson (1985) documented an extensive subaqueous emplaced volcanic layer in the Keweenaw Peninsula and Isle Royale in the upper part of the PLV basalts which are otherwise subaerially deposited. This layer may have been deposited during an incursion of seawater into the rift. Incursions of seawater are also possible during times of PLV interbedded minor clastic sediments when volcanism waned. Thus, there is likely a long history of seawater incursion into the Midcontinent rift and a significant volume of seawater could have penetrated deeply into the PLV. As the PLV was buried into the fluid source zone, the contained seawater (formation water) evolved. A plausible mechanism is needed to satisfy the constraint of very low sulfur fluids. Blättler et al. (2020) have shown that low sulfate levels in seawater occurred during the late-Mesoproterozoic which would have required less depletion of sulfur. When seawater penetrates mid-ocean basalts in vicinity of hydrothermal vents it becomes heated and evolves as it reacts with and alters the host basalt. At 130-150oC anhydrite begins to precipitate; the solubility of anhydrite is retrograde (Antonelli et al., 2017). The precipitation of anhydrite (CaSO4) depletes sulfate from seawater and if there is sufficient Ca, precipitation of anhydrite removes most of the sulfate. As the partially evolved seawater and its host basalts continue to be heated, albitization of basalt releases Ca into the seawater which results in complete removal of any remaining sulfate. Breakdown of olivine and pyroxene releases Mg into the fluid but it is bound up in secondary Mg minerals, such as chlorite, leaving the fluid both sulfur- and Mg-poor. At yet higher temperatures 5 �(> 250oC) the water-rock reactions continue to release Ca to the fluid, which accumulates, producing a Ca-rich brine (Antonelli et al., 2017; Tivey, 2007). Oceanic and continental rifts undergoing greenschist to amphibolite facies metamorphism are known for CaCl2-rich, Mg-poor brines, also enriched in leached Na, K, and Cu among other elements (Tivey, 2007; Hardie, 1983). This model for modern mid-ocean basalts is applicable to the generation of native copper ore-forming fluids. The PLV basalts are tholeiitic, similar to mid-ocean ridge basalts, and have undergone extensive albitization. The rocks in the source zone have been subjected to temperatures exceeding the temperature of precipitation of anhydrite, hence the seawater could have become depleted in sulfur and enriched in Ca before undergoing burial metamorphism. Kelly (2020) demonstrated that probable primary fluid inclusions formed during precipitation of native copper are filled by Ca-rich fluids with up to 30 equivalent weight percent CaCl2. Published light stable isotope data are permissive of a hybrid evolved seawater and metamorphic-dominated ore-forming fluid originating from the source zone. Such a hybrid cannot be isotopically distinguished from metamorphic-only fluid. The combination of low sulfate seawater and anhydrite precipitation provides a viable mechanism to generate evolved seawater that provides an outside source of salinity for the ore-forming fluids without sulfur. Lengthy incursions of seawater into the Midcontinent rift supports the hypothesis that evolved seawater may have played an important role in the generation of burial metamorphic-dominated native copper ore-forming fluids. References: Antonelli, M.A., Pester, N.J. Brown, S.T, DePaolo, D.J., 2017, Effect of paleoseawater composition on hydrothermal exchange in midocean ridges, Proceedings of the National Academy of Sciences,114: 12413-12418. Blättler, C.L., Bergmann, K.D., Kah, L.C., Gómez-Pérez, I., Higgins, J.A., 2020, Constraints on MesoNeoproterozoic seawater from ancient evaporite deposits, Earth and Planetary Science Letters, 532: 115951. Bornhorst, T.J, Mathur, R., 2017 and 2018, Copper isotope constraints on the genesis of the Keweenaw Peninsula native copper district, Michigan, USA, Minerals: 7, 185 and 8, 508 Brown, A.C., 2006, Genesis of native copper lodes in the Keweenaw Peninsula, northern Michigan: A hybrid evolved meteoric and metamorphogenic model, Economic Geology, 101: 1437–1444. Hardie, L. A., 1983, Origin of CaCl2 brines by basalt-seawater interaction; insights provided by some simple mass balance calculations, Contributions to Mineralogy and Petrology: 82, 205-213. Johnson, R.C., 1985, Documentation of a subaqueously emplaced volcanic horizon in the upper Portage Lake Volcanics, Keweenaw Peninsula, Michigan, ILSG Proceedings, 31, part 1: 38-39. Kelly, David., 2020, Fluid inclusion study of selected calcite associated with native copper, Quincy mine, Keweenaw Peninsula, Michigan, Open Access M.S. Report, Michigan Technological University, 165. Livnat, A., 1983, Metamorphism and Copper Mineralization of the Portage Lake Lava Series, Northern Michigan. Ph.D. Dissertation, University of Michigan, Ann Arbor, MI, USA; 1–292. Tivey, M.K., 2007, Generation of seafloor hydrothermal vent fluids and associated mineral deposits, Oceanography, 20: 50-65. White, W.S., 1968, The native-copper deposits of northern Michigan, In Ore Deposits of the United States, 1933–1967 (Graton Sales Volume), Ridge, J.D., Ed.; American Institute of Mining, Metallurgical, and Petroleum Engineers: New York, NY, USA, 303–325. 6 �Mineralizing processes in the Current Lake and Escape Lake conduit-type PGE–Cu–Ni deposits of the Thunder Bay North igneous complex, northwestern Ontario, Canada BRZOZOWSKI, Matthew1, HOLLINGS, Peter1, MACTAVISH, Allan2, EVANS– LAMSWOOD, Dawn2, DROST, Abraham2, WILTON, Derek3. 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1, Canada Clean Air Metals, 1004 Alloy Drive, Thunder Bay, ON, P7B 6A5, Canada 3 Earth Sciences, Memorial University, 230 Elizabeth Avenue, St. John’s, NL, A1C 5S7, Canada 2 The Thunder Bay North igneous complex (TBNIC), located approximately 50 km northeast of Thunder Bay, comprises a series of small mafic–ultramafic intrusions that occur in the vicinity of several other similar intrusions, including the Sunday Lake, Saturday Night, and Thunder intrusions to the southwest, and the Seagull, Disraeli, and Hele intrusions to the northeast. The intrusions of the TBNIC comprise the 1106.6 ± 1.6 Ma Current Lake (Bleeker et al., 2020), Escape Lake, Lone Island Lake, and 025 intrusions, all of which were emplaced proximal to the east–west trending Quetico Fault System. Of these intrusions, the Current Lake (CL) and Escape Lake (EL) chonoliths host significant orthomagmatic Ni, Cu, and platinum-group element (PGE) mineralization, and have been in active exploration since 2005. Although few theses have been completed on the Current Lake (Chaffee, 2015) and Escape Lake (D’Angelo, 2013) chonoliths, the mechanisms that generated the variable styles of PGE–Cu–Ni mineralization in these spatially associated chonoliths, and their genetic relationships have never been formally characterized. This presentation highlights some of the preliminary results from a three-year collaborative effort between Clean Air Metals, Lakehead University, and Memorial University, with the ultimate goals of developing robust mineralization models for and characterizing the genetic relationships between the CL and EL intrusions. Although both the CL and EL chonoliths have been actively explored, the most extensive work has been conducted on the CL chonolith, resulting in an indicated resource of 12 Mt at 1.48 g/t Pt, 1.4 g/t Pd, 0.28% Cu, and 0.17% Ni for CL, and 4.3 Mt at 0.92 g/t Pt, 1.18 g/t Pd, 0.52% Cu, and 0.28% Ni for EL (Kuntz and Jones, 2021). The CL chonolith contains five variably mineralized zones (from north to south) — the Current Lake, Bridge, Beaver Lake, Cloud, and 437 zones. The southernmost portion of the conduit is represented by the unmineralized Southeast Anomaly. The EL chonolith has been divided into two sections (from north to south) — the Escape and Steepledge Lake sections. Both conduits dip gently to the south, with their lowest points being adjacent to the E–W trending Escape Lake Fault. Primary base-metal sulfides (BMS) in both systems largely comprise pyrrhotite, pentlandite, and chalcopyrite; BMS occur in the CL system occur throughout the chonolith in the Current Lake and Bridge zones, and at the base and top of the chonolith in the Beaver Lake and Cloud zones, respectively. Metal tenors in the CL and EL systems are similar, being enriched in Cu and the Pt-group PGEs relative to Ni and the Ir-group PGEs, and exhibiting a distinct negative Ru anomaly, a feature not commonly observed in other Ni–Cu–PGE systems globally (e.g., Eagle and Jinchuan). The anomalously low Ru concentrations are likely the result of clinopyroxene fractionation at depth. Copper/Pd ratios in the Current Lake and Beaver Lake zones of the CL chonolith, and Escape Lake and Steepledge Lake sections of the EL chonolith, are indistinguishable and largely within the range of mantle values, suggesting that significant amounts of sulfide were not removed at depth. Although most of the samples from both systems have mantle-like Cu/Pd and Pd/Pt ratios, data 7 �from the Southeast Anomaly, Escape Lake, and Steepledge Lake zones trend towards high Cu/Pd and low Pd/Pt ratios, suggesting that some Pd may have been removed by segregation of limited sulfide liquid at depth. If the Southeast Anomaly is considered the feeder to the CL system, then this may suggest that sulfides are located deeper in the feeder system. Numerical modeling suggests that, although the CL system is characterized by slightly higher R factors than the EL system (~12,000 vs. ~7,000), there was not a significant variation in R factor along the length of the conduits, with variations in bulk-rock metal concentrations largely representing variable accumulations of sulfide. In both systems, the Pt-group and Ir-group PGEs, and Au correlate with Pd, suggesting that the physical separation of residual Cu-rich liquid from monosulfide solid solution did not play a role in generating the mineralization. Variations in Ni/Cu–Mg and Pd/Pt– Pd/Ir ratios suggest that the concentrations of Ni and Ir in the CL system appear to be controlled, in part, by the accumulation of olivine towards the Southeast Anomaly. The fact that rocks in this zone are olivine-poor suggests that olivine may be located at depth in the feeder system. Systematic inverse correlations between Fe–S–Cu–Ni–Pd–Pt (+ spikes) and Mg–Ca–Cr–Ti (– spikes) throughout the thickness of the Current Lake Zone suggest multiple influxes of metal-rich magma into the system. The abundance of these inverse correlations at the base and top of the Beaver Lake–Cloud zones suggest that the dynamics of magma flow varied along the length of the conduit, likely as a result of the variable conduit morphology. Bulk assimilation of the granitic and sedimentary country rocks is unlikely to have played a role in sulfide saturation as there are no systematic mixing trends in La/Zr–Th/Nb between the host and country rocks. Although it is possible that the large variation in S/Se values could be the result of contamination by BMS from the granitic and sedimentary country rocks, it is also possible that the variation resulted from S loss and very low R factors (< ~1,000). The importance of local S addition for, and the timing of, sulfide liquation remains unclear, but is one of the critical questions that will be assessed in the Clean Air Metals–Lakehead University–Memorial University partnership. References Bleeker, W., Smith, J., Hamilton, M., Kamo, S., Liikane, M., Hollings, P., Cundari, R., Easton, M., Davis, D., 2020. The Midcontinent Rift and its mineral systems: Overview and temporal constraints of Ni-CuPGE mineralized intrusions (No. 8722), Targeted Geoscience Initiative 5: Advances in the understanding of Canadian Ni-Cu-PGE and Cr ore systems - Examples from the Midcontinent Rift, the Circum-Superior Belt, the Archean Superior Province, and Cordilleran Alaskan-type intrusions. Chaffee, M.R., 2015. Petrographic and Geochemical Study of the Hybrid Rock Unit Associated with the Current Lake Intrusive Complex (MSc). University of Minnesota. D’Angelo, M., 2013. Igneous textures and mineralogy of the Steepledge Intrusion, Northern Ontario (BSc). Lakehead University. Kuntz, G., Jones, L., 2021. NI 43-101 Technical report and mineral resource estimate for the Thunder Bay North project, Thunder Bay, Ontario (NI 43-101). 8 �Three-dimensional geologic mapping of Precambrian rocks in Minnesota: The creation of “removeable” geologic layers using gravity and magnetic data interpretation CHANDLER V. W. (chand004@umn.edu) and JIRSA M.A.(jirsa001@umn.edu) Minnesota Geological Survey (retired), 2609 Territorial Road, St, Paul, MN 55114 -1009 As part of the National Geological and Geophysical Data Preservation Program (NGGDP) of the U. S. Geological Survey, the Minnesota Geological Survey (MGS) compiled geophysical models from past and on-going investigations to create three-dimensional representations of several Proterozoic basins, consisting of elevation contours and inferred bedrock geology of the underlying basement surfaces. Thus defined, the Proterozoic basins are represented as removeable, three-dimensional layers in the Precambrian bedrock map of the state. The NGGDP-MGS work focused on several basins including those associated with the Sioux Quartzite, the Animikie Group, the sedimentary rocks of the Midcontinent rift system, and the somewhat more complicated basin enclosing the North Shore Volcanic Group and the Duluth Complex (NSVG-DC). In this presentation we will focus on the results from the main bowl of the Animikie basin and the adjoining NSVG-DC basin. Figure 1 presents the elevations estimated for the basement of these two basins. The basement elevations estimated for the main bowl of the Animikie basin are primarily based on Euler analysis, a semi-automated depth interpretation scheme, in which a non-magnetic basin sequence is assumed to overlie an assortment of anomaly sources that suitably approximates a basin floor. Correspondingly much of the Animike sequence consists of non-magnetic slates and grauwackes, with anomaly sources restricted to iron-formation near the base of the sequence, or to sources within the underlying Archean basement. The Euler results were cross-checked in a few areas by conventional, two-dimensional modeling, and the geology of the inferred basin floor is based on extrapolating the bedrock geology of adjacent areas, as guided by gravity and magnetic anomaly signatures. The results (Figure 1) indicate that northern and western margins of the Animike basin are rimmed by a shallow (0-1.5 km elevation) shelf. The northern part of the shelf has been previously inferred from earlier investigations, but the western shelf is somewhat novel, and it appears to be significantly controlled by NW-striking structures. Several prominent sources underlying the Animikie sequence are interpreted to be strongly magnetic, consistent with ironformation-bearing horizons, and the steep northward dips interpreted for some of these sources are most consistent with Archean rocks in the region. The deeper parts of the basin are locally below -4 km elevation, including the area along the basal contact of the NSVG-DC basin. This ~5 km.thick Animikie sequence could presumably continue to the east in some fashion beneath the NSVG-DC sequence. Such a scenario has implications for the metallogenesis and upper crustal geology of the region. Euler analysis cannot resolve the base of a magnetic sequence, so the base of the NSVG-DC basin is estimated along twenty two-dimensional gravity and magnetic models that transect the complex in strategic areas. Of these, ten models have been recovered and revised from earlier investigations and ten models have been created specifically for this study. Wherever possible, modeling has been constrained by geology mapped at the bedrock surface (Jirsa and others, 2012) and by rock-property data (Chandler and Lively, 2011). The modeling results were used to compile basement elevation contours beneath the land surface, and these contours were smoothed and merged with elevation contours beneath Lake Superior, which were compiled from 3-d gravity modeling and seismic reflection interpretations (Allen and others, 1997). The two lowest parts of the NSVG-DC sequence are interpreted lie ~-15 km. below MSL, and are plausible candidates for major feeder zones. A prominent basement high separates these two lows, and it appears to 9 �ultimately connect with major basement ridges beneath Lake Superior, including Walter White ridge to the south, and with the Grand Marais ridge to the east. Estimated dips along the base of the NSVG-DC basin generally range from 25 to 60 degrees, with the steepest dips inferred along the northern and western basal contacts. A shelf-like structure inferred beneath the northwestern margin of the NSVG-DC basin lies roughly along-strike with the shelf inferred beneath the northern margin of the Animikie basin. The interpretations presented here represent our current “best guess” of subsurface structure, based on available data, and many improvements should be possible in the future. As such, the interpretations presented here should serve as a helpful starting point for future three-dimension investigations, including the gravity, magnetic, and electromagnetic studies that has been recently initiated by the U. S. Geological Survey in the region. References cited: Allen, D. J., Hinze, W. J., Dickas, A. B., Mudrey Jr., 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, p 47-72. Chandler, V.W.; Lively, R.S.. (2011). Density, Magnetic Susceptibility, and Natural Remanent Magnetization of Rocks in Minnesota: An MGS Rock Properties Database. Minnesota Geological Survey. Retrieved from the University of Minnesota Digital Conservancy, http://hdl.handle.net/11299/175580. Jirsa, M.A.; Boerboom, T.J.; Chandler, V.W.. (2012). S-22, Geologic Map of Minnesota, Precambrian Bedrock Geology. Retrieved from the University of Minnesota Digital Conservancy, http://hdl.handle.net/11299/154540. Figure 1. Estimated elevation of the basement surface below the main bowl of Animikie basin and the combined igneous sequences of the North Shore volcanic Group and the Duluth Complex. Elevations are relative to mean sea level, and range from 0 to 2,000 m. (yellow), -2,000 to -5,000 m. (light green), -5,000 to -15,000 m. (green), and <15,000 m. (blue). Elevations beneath the North Shore Volcanic Group- Duluth Complex basin are highlighted with 1000 m. contours. 10 �Tracing redox pathways of the Mesoproterozoic Copper Harbour Conglomerate, Michigan CRAIK, Nicholas1 and FRALICK, Philip1 1 Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The history of Earth’s oxygenation is key to understanding the evolution of life, the emplacement of important mineral deposits, and the cycling of oxidative weathering products through time. Many studies have been particularly focused on atmospheric and oceanic oxygenation during the Archean to Paleoproterozoic – denoting the Mesoproterozoic as part of the ‘Boring Billion’, a time of general stability or stagnation of many Earth processes. During the ‘Boring Billion’ oxygen was thought to have been stalled at low concentrations relative to today. This idea has largely been put forth by researchers utilizing deep sea black shales as a proxy for continental weathering patterns Many studies of atmospheric oxygen within the Mesoproterozoic are contradictory to one another, likely due to sampling of deep ocean settings where oxygen cannot penetrate, and lack of abundant studies on terrestrial sediment. To reach results more representative of the ancient atmosphere, choosing a depositional setting that has been well documented as subaerial, and/or being able to constrain the weathering history of the system is imperative. In this work sedimentological and geochemical evidence from the subaerial red-bed deposits and carbonates of the unmetamorphosed Copper Harbour Conglomerate were collected with the aim of better understanding and constraining the spatial and temporal fluctuations of atmospheric oxygen levels during the Mesoproterozoic. The Copper Harbour Conglomerate of the Keweenaw Peninsula, Michigan is a progradational alluvial to lacustrine sedimentary sequence representing the first continuous infilling of the Midcontinent Rift basin after the cessation of associated volcanic activity at ~1.1 Ga. Within the formation are pervasive carbonate cement (calcrete) lenses as well as a stromatolitic horizon located in its uppermost portion. The purpose of this study was to examine the oxidative weathering products (siliciclastics), groundwater-precipitated carbonates (calcrete), and pond/lacustrine-precipitated carbonates (non-biogenic precipitates and stromatolites) found within the Copper Harbour Conglomerate to determine the relative oxygenation of the atmosphere at the time of deposition (~1085 Ma). ICP-MS and ICP-AES methods, both whole rock and preferential carbonate extraction, were utilized to determine the whole rock and carbonate geochemistry of the rock types within the upper Copper Harbour Conglomerate. By analyzing the redox sensitive metals (e.g., Fe, Mn, V, and U) and rare earth elements (e.g., Ce anomaly), theoretical constructions of the hydrological pathways of these elements can be developed. This allows understanding of the redox environment during early diagenesis. 11 �Figure 1 shows the Ce and La anomalies of the samples taken in the study area. The calcrete samples have the strongest negative Ce anomalies, indicating they have the least abundance of Ce and therefore have previously had the greatest interaction with oxygen, either due to abundance of oxygen or elapsed reaction time. This makes sense because calcrete is formed by the evaporation of groundwater near surface (i.e., subaerial) and takes thousands of years to form. The stromatolite Figure 1: La and Ce anomalies (PAAS normalized) of samples show a slight negative Ce calcrete, stromatolite, and siliciclastic samples in the anomaly which indicates the fluids study area. they formed from had previously interacted with oxygen. The anomaly for the stromatolites is less than that of the calcrete which can be attributed to the lower residency time of surface water (in which the stromatolites form) compared to groundwater. The siliciclastic samples do not show a positive or negative Ce anomaly. This is likely due to abundance of primary, unweathered material contained within the samples themselves. Similar results were obtained from the redox sensitive elements Fe, Mn, V, and U. Surficial weathering of the basaltic sediment induced by significant atmospheric oxygen levels resulted in the development of Fe hydroxides and oxyhydroxides pervasive throughout the sediment of the Copper Harbour Conglomerate. In this system Mn was staying in solution for a longer time, on average, than Fe and became concentrated within the groundwater and precipitated out with the calcrete. The oxidized forms of V and U were readily weathered from the basalt and held in solution until encountering the reducing biochemical reactions of the decaying bacterial component of the stromatolites or becoming oversaturated within the groundwater and precipitating out with the calcrete. By breaking down the sub-environments of the three sample types in this study, a clear progression and hydrologic geochemical evolution of the broader depositional environment could be traced out. There was enough oxygen in the atmosphere for pervasive oxidative weathering at surface and for the studied redox sensitive elements to maintain their oxidized state through surface and groundwater transportation – only being removed through oversaturation or local reduction via organic decomposition. These results indicate that atmospheric oxygen was present in greater concentrations ~1085 Ga than previously thought. This also leads into the idea that Earth’s oxygenation may have been variable in relatively short time spans. 12 �Critical Mineral Potential on the North Shore of Lake Superior; Elements to ‘Structure’ our future CUNDARI, Robert1, PUUMALA, Mark1, METSARANTA, Riku2 and CAMPBELL, Dorothy1 1 Resident Geologist Program, Ontario Geological Survey, Thunder Bay, ON, P7E 6S7 2 Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, Sudbury, ON P3E 6B5 In March 2021, the Ontario Government released a Critical Minerals Framework Discussion Paper which outlines Ontario’s proposal for developing a critical minerals strategy as well as a draft critical mineral list. Critical minerals are defined as “….raw materials needed to produce many products and specialized technologies. The minerals that a jurisdiction deems “critical” depends on its geology, as well as its own domestic and economic priorities.” (https://www.ontario.ca/page/critical-minerals). This abstract and associated presentation provides a general synopsis of the geology and existing models for several critical mineral enriched mineralization styles (“deposits” sensu lato) related to the Proterozoic Midcontinent Rift (MCR) along the northwestern shore of Lake Superior. The specific deposit types highlighted are: orthomagmatic platinum group elements deposits in MCR-related intrusions (PGE-Ni-Cu-Cr-Co) , polymetallic hydrothermal vein deposits (Ag-rich, Pb-Zn-Cu-rich U-rich types; U-Ni-Zn-Cu-Co-Sb-Bibarite-fluorite) and hydrothermal polymetallic breccia complex related deposits (Deadhorse Creek diatreme type; U-Be-Zr-Hf-Y-Ce-Nb-Ti-V-Cr-REE). Platinum Group Elements (PGE) have long been a subject of exploration interest north of Lake Superior. Since the discovery of PGE mineralization at the Thunder Bay North (TBN) project, the region has seen relatively steady exploration efforts targeting similar chonolith-hosted magmatic sulphide deposits. Larger MCR-related mafic-ultramafic intrusions, including Sunday Lake and Seagull, have also been explored for magmatic sulphides. These PGE prospective intrusions are interpreted to have been emplaced early in MCR history through crustal-scale structures (Bleeker et al., 2020). Hart and MacDonald (2007) stated the importance of deep-seated structures, specifically the Black Sturgeon fault, as conduits for early, primitive melts and subsequent emplacement of the early mafic-ultramafic units (e.g. Hele, Seagull, Kitto, Disraeli). Furthermore, Heggie (2012) recognized that the emplacement of the TBN complex was primarily controlled by a macro-structure (i.e. Archean Quetico fault) during MCR magmatism which provided a zone of weakness for the intrusive complex to develop. Many of these mineralized intrusions are relatively removed spatially from the historically mapped extent of rocks related to the MCR highlighting the larger structural footprint of the rift and its related fault systems. The silver-bearing veins in the Thunder Bay area have been subject to considerable historical mining and more recently, renewed exploration interest. The critical mineral potential of these vein systems is not well established, although examples of similar vein systems (e.g. deposits at Cobalt) were a source of critical minerals (e.g. Kissin, 1992). The vein systems are organized in 2 groups known as the Mainland and Island belts (cf. Franklin et al., 1986). The Mainland Belt deposits (e.g., Beaver, Badger, Silver Mountain) occur in Rove Formation sedimentary rocks immediately below the contact with Logan diabase sills. The veins locally extend upwards into the sills, with silver-rich portions of the systems bound by locally silicified shale. The Island Group deposits (e.g. Silver Islet Mine) are characterized by fracturefilled vein systems oriented perpendicular to a northeast-trending swarm of gabbroic (Pigeon River) dykes. The Mainland and Island group veins contain similar mineral assemblages of acanthite and native silver associated with base-metal sulphides, fluorite, barite, quartz, and calcite. The Island group differs from the Mainland group in that they locally contain a nickel-cobalt sulpharsenide suite of minerals often termed 5elements veins (Ag-Bi-Co-Ni-As). The Island and Mainland group are present in or near crustal-scale, extensional listric faults that formed throughout the main episode of MCR extension which then became reverse faults at the termination of that event (Smyk and Franklin, 2006). Similar to the mafic-ultramafic 13 �intrusions hosting PGE mineralization, polymetallic vein systems have a wider spatial distribution than the mapped extent of MCR related rocks and thus may serve as a tool for mapping rift related structures and hydrothermal systems. Uraninite (pitchblende) was documented by Franklin (1978) and occurs in quartz veins and vein breccias in association with hematite, magnetite and/or pyrite within the Archean basement rocks of the Sibley basin (Sutcliffe, 1991). Uranium mineralization is associated with fracture zones related to north- to northwest-striking regional structures such as the Black Sturgeon fault, and has been dated at 1094 Ma (Ruzicka and LeCheminant, 1984), supporting a model similar to the genesis of lead-zinc veins in the Dorion area (Franklin and Mitchell, 1977; Smyk and Franklin 2007). The model invokes mineralizing hydrothermal fluids generated by MCR-heating inducing leaching of uranium and other metals from uraniferous basement pegmatites or basal sedimentary strata in the Sibley Group (ibid). The fluids which migrated through permeable Pass Lake formation sandstone precipitated metals in stratigraphic traps at or near the Archean unconformity (ibid). Although the timing of these veins is considered broadly coeval with the Ag-rich veins of the Mainland and Island silver-belts, the variable metal signatures of different vein systems (Ag-rich vs U-rich vs Pb-Zn-Cu-rich) remains relatively poorly understood. In addition, their trace critical mineral contents have not been investigated extensively. The Dead Horse Creek volcaniclastic breccia complex is host to a diverse group of highly altered heterolithic breccias and mafic to intermediate dikes that have been explored for rare earth elements, uranium, yttrium, cerium, zirconium and hafnium (Zurevinski et al., 2019). The Dead Horse Creek Complex has been divided into 5 subcomplexes, referred to as North, South, East, West and Central and described as altered heterolithic breccias that have undergone varying degrees of alteration and are variably radioactive (Sage 1982; Smyk et al., 1993). Exploration has mainly been focused on the mineralized zones at the West Dead Horse and the North Dead Horse subcomplexes. The West Dead Horse subcomplex has been described as “diverse, exotic, hydrothermally altered and rare metal mineralized” (ibid). The main mineralized zone at the West Dead Horse Subcomplex is small; however, the grade and style of mineralization highlights the importance of following up on larger, underexplored structures elsewhere in the Dead Horse Complex. Uranium-beryllium-zirconium mineralization was introduced via A-type granitic fluids along fault structures that crosscut the genetically unrelated breccias. Subsequently, niobiumtitanium-vanadium-chromium-bearing alkaline fluids were introduced into the same fault system. These reacted with the pre-existing mineral assemblage and created the observed exotic mineralogy (Potter and Mitchell 2005). The recognition of structures that were formed or re-activated during the Midcontinent Rift event is integral to understanding the location and timing of all the deposit types noted above. Bleeker et al, (2020) have made significant advances in the understanding of the timing of intrusions and subsequently the overall geodynamic evolution of the MCR. It has been noted that younger and more voluminous magmatism was focussed into the central rift as the lithosphere thinned and rifted apart, whereas the older intrusions (e.g. Thunder Bay North, Tamarack and others) were emplaced farther afield. This lithospheric thinning may have been an important factor in the location of structurally controlled hydrothermal deposits surrounding the Lake Superior basin as the majority of the vein systems are hosted in faults parallel to the rift axis, proximal to the rift margin. Distance from rift axis may have also had a bearing on the type of mineralization present in the vein systems as specific arrays of deposits appear to have differing metal signatures. The structural controls on the deposits, whether deposit controlling faults are primary MCR-related or reactivated Archean structures, remain important to understanding the structural evolution of the MCR and may help to further elucidate the nature of metallogeny in the region. Apart from the intrusion-hosted PGE deposits, the polymetallic vein deposits outlined here (Ag-rich, U-rich, Pb-Zn rich, Deadhorse Creek-type) are all generated by fluids derived from heat related to the MCR (Smyk and Franklin, 2007). The composition of the polymetallic veins is a product of the nature of the diverse ore-related fluids, and the geochemistry of the host rocks and their relative interactions and contributions of metals. Understanding 14 �the structures which host all the deposit types outlined here and their relative timing, along with refining tectonic/structural evolution of the MCR, will be integral to furthering our understanding on the controls on these critical deposit types. References Bleeker, W., Smith, J., Hamilton, M., Kamo, S., Liikane, D., Hollings, P., Cundari, R., Easton, M., and Davis, D., 2020. The Midcontinent Rift and its mineral systems: Overview and temporal constraints of Ni-Cu-PGE mineralized intrusions; in Targeted Geoscience Initiative 5: Advances in the understanding of Canadian Ni-CuPGE and Cr ore systems – Examples from the Midcontinent Rift, the Circum-Superior Belt, the Archean Superior Province, and Cordilleran Alaskan-type intrusions, (ed.) W. Bleeker and M.G. Houle; Geological Survey of Canada, Open File 8722, p. 7–35. https://doi.org/10.4095/326880. Franklin, J.M., Kissing, S.A., Smyk, C. and Scottt, S.D. 1986. Silver deposits associated with the Proterozoic rocks of the Thunder Bay district, Ontario. Canadian Journal of Earth Sciences 23, p. 1576- 1591. Franklin, J.M. and Mitchell, R.H. 1977. Lead-zinc-barite veins of the Dorion area, Thunder Bay District, Ontario. Canadian Journal of Earth Sciences, 14: 1963-1979. Franklin, J.M. 1978. Uranium mineralization in the Nipigon area, Thunder Bay District, Ontario; in Current Research, Part A, Geological Survey of Canada, Paper 78-1 A, p.275-282. Heggie, G., MacTavish, A., Johnson, J., Weston, R., and Ma, L., 2012. Structural control on the emplacement of the TBN-Igneous Complex in Institute on Lake Superior Geology Proceedings, 58th Annual Meeting, Thunder Bay, Ontario, Part 1 – Proceedings and Abstracts, v.65, p.37-38. 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. Kissin, S.A., 1992. Five-element (Ni-Co-As-Ag-Bi) veins. Geosci. Can. 19, 113–124. 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, v.150, p.212-229. Ruzinka, V. and LeCheminant, G.M. 1984. Uranium deposit research, 1983, in Current research, part A. Geological Survey of Canada, Paper 84-1A, p39-51. Sage, R.P. 1982. Mineralization in diatreme structures north of Lake Superior, Ontario Geological Survey, Study 27, Ontario Ministry of Natural Resources, Toronto, 79p. 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 44, p. 1041-1053. Smyk, M.C., Taylor, R.P., Jones, P.C. and Kingston, D.M. 1993. Geology and geochemistry of the West Deaf Horse Creek rare metal occurrence, northwestern Ontario; Exploration and Mining Geology, v.2, p.245- 251. 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 Volume 4 Part 1. P. 627-672. Zurevinski, S., Campbell, D. and Puumala, M. 2019. Midcontinent Rift-related carbonatites and diatremes; in Institute on Lake Superior Geology Proceedings, 65th Annual Meeting, Terrace Bay, Ontario, Part 2 - Field Trip Guidebook, v.65, part 2, p.2-10. 15 �Availability of Historical Airborne Geophysical Survey Data at Minnesota DNR DAHL, David1, SAARI, Stacy1, and LEE, Thomas1 Minnesota Department of Natural Resources – Lands and Minerals Division, 1525 3rd Ave. East, Hibbing, MN 55746 USA dave.dahl@state.mn.us stacy.saari@state.mn.us thomas.lee@state.mn.us 1 The collections of exploration and scientific research data housed at the Minnesota Department of Natural Resources include a large volume of unpublished airborne geophysical survey deliverables, including maps, flight line profiles, flight records, interpretation reports, and other materials. An effort to discover, organize and make accessible these materials has uncovered 131 non-government pre-GPS mineral exploration airborne geophysical surveys representing 156,192 line kilometers of acquired data. The effort has accomplished several goals: • From a collections perspective, segregation of these airborne survey deliverables from other exploration project records has succeeded in greatly reducing the physical volume of mineral exploration materials remaining to be handled in other curation and cataloging activities. Large maps, typically on Mylar, have been indexed and collected into vertical map cabinets. Folders of flight line profiles have been indexed at the survey level and stored in archival quality office boxes on a space efficient mobile shelving system. Contractor technical reports and irregular documents such as 35mm strip films and nine-track magnetic tapes have also been indexed, boxed and stored in the mobile shelving system. This single episode of curation activity has addressed some 30% of the volume in the collections. • The effort has introduced the geography of mineral exploration data collections to newer mineral potential geoscientists in the Division, and provided big-picture orientation to historical exploration program data that is often reused in support of geologic survey mapping, university research, and natural resource land management activities. • Scanned products of the organizing effort include 1) PDF/A archive format files of all maps, reports and similar documents, 2) color-indexed TIFF image files of all inventoried airborne geophysical survey map sets, and 3) georeferencing links files for all inventoried map images (georeferenced to UTM zone 15, NAD83). These files provide a critical backup to otherwise irreplaceable hard copy. An effort to make these and other collections products more easily accessed, viewed, searched and downloaded is currently in development. • Four index files have been developed: 1) a surveys index based on flight line extents contains acquisition footprints (multi-part polygons) and survey attributes for the 131 surveys, 2) a maps index based on the 860 maps contains map tile footprints and map tile attributes (some survey map sets have up to 18 tiles, with up to 4 map sheets per tile), 3) a flight lines index contains line traces and line attributes for the 8,778 flight lines represented on the 129 map sets, and 4) a documents index spreadsheet based on inventory of the document collection contains metadata for each of the 1,089 documents curated during the effort. Availability of these indexes serves dual purpose, first as a finding aid, and secondly as an orientation and quality control aid for the curation of much of the follow up geologic, ground geophysical, and geochemical mineral exploration data sets archived at Minnesota DNR. Mineral exploration magnetic-EM survey data comprise 89% of flight line content (138,786 line km). Magnetics-only surveys represent 10% of the data, and 1% is from a regional AFMAG (passive EM) reconnaissance survey over the Duluth Complex. Locations in Minnesota targeted by these surveys include the Archean Wawa and Wabigoon Subprovinces, Proterozoic bedrock on 16 �the Mesabi and Cuyuna iron ranges and in areas of west-central, east-central, and southwestern Minnesota. Most of the surveys are based on ¼ mile flight line spacing. For geologic mappers these surveys provide context for site surveys and drill core locations, and additional information about Minnesota’s concealed bedrock terranes. For instance, many of the surveys have flight orientations different than government sponsored magnetic surveys; and formational EM conductors mapped in areas such as the southern part of the Animikie basin may enhance geologic interpretation in areas mostly devoid of magnetic anomaly detail. EM data may also provide information on surficial material features such as gravels or conductive clays. Obviously, more work can be done to round out the inventory of non-government airborne geophysical surveys in Minnesota. GPS-based airborne survey data sets, which are digital and less susceptible to physical loss, are yet to be indexed, and indexing staff are aware of about 20 additional historical non-government airborne surveys in Minnesota that might potentially be targets for future donation (if they still exist). These additional historical airborne surveys are partially evidenced, either through coarse index maps, or clusters of named ground target sites, or staff knowledge of survey area vicinities for which no footprints are found in the collections. The flight line profiles, which hold the actual magnetic, EM, and occasional radiometric survey data are now curated at the survey level and indexed at the flight line level. But they are still without backup and vulnerable to physical loss. These flight line profiles are an obvious target for a future National Geological and Geophysical Data Preservation Program (NGGDPP) proposal to scan and convert the hard copy records to PDF/A format, for preservation and access. Figure 1: Acquisition footprints for the non-government pre-GPS airborne geophysical surveys inventoried during current effort. Bedrock geology - Minnesota Geological Survey map S-22. 17 �Geophysical insights into Paleoproterozoic tectonics along the southern margin of the Superior Province, central Upper Peninsula, Michigan DRENTH, Benjamin J.1, CANNON, William F.2, SCHULZ, Klaus J.2, and AYUSO, Robert A.2 1 U.S. Geological Survey, PO Box 25046, MS 964, Denver Federal Center, Denver, CO 80225 U.S. Geological Survey, 12201 Sunrise Valley Dr., MS 954, Reston, VA 20192 2 The southern margin of the Archean Superior Province in the central Upper Peninsula of Michigan was a nexus for key Paleoproterozoic tectonic events involved in the ~2.1 Ga rifting of proposed Archean supercraton Superia and subsequent assembly of Laurentia. Interpretations of the region’s tectonic history have historically been hampered by extensive Pleistocene glacial and Paleozoic sedimentary cover and a lack of appropriate geophysical data. These rifting and orogenic events formed geologic effects that are readily mappable with modern geophysical methods. New aeromagnetic and gravity data provide a critical means of mapping and interpreting the complex geological framework through cover, allowing development of significantly richer geographical and process-based perspectives on all these tectonic events (Drenth et al., in press). Existing interpretations of Archean and Paleoproterozoic contacts and structure (e.g., James et al., 1961; Bayley et al., 1966) are carried >30 km eastward under Paleozoic cover (Fig. 1). The progression of ~2.1 Ga rift-related sedimentation and magmatism recorded in rocks of the Dickinson Group are clearly expressed as a geographically extensive and largely concealed tectonic feature of the southern Superior Province. Geophysical interpretations provide evidence for plausible ~2.1 Ga intrusive magmatism, such as a previously unrecognized swarm of mafic northeast-striking dikes cutting Archean rocks and multiple granitic bodies. Effects of the ~1.87-1.83 Ga Penokean orogeny include the clearly imaged Niagara fault zone suture, abundant evidence for thin-skinned thrusting and folding in the Menominee iron district, and speculative emplacement of an allochthonous sedimentary sequence in the Calumet trough. Numerous east-west trending structures likely originated, or were significantly reactivated, by post-Penokean deformation. Metamorphic events at ~1.76 Ga and ~1.65 Ga may correspond to orogenies involving younger, outboard Paleoproterozoic crustal provinces recognized in southern Laurentia. For example, the previously unrecognized West Branch fault, separating the Dickinson Group from Archean rocks, is shown to be a major structure in the region and is a proposed expression of ~1.76 Ga thick-skinned deformation documented elsewhere in the region. Oblique disruptions of crudely east-west striking structures have robust geophysical expressions and are speculatively connected to transpressive deformation at ~1.65 Ga. References Bayley, R. W., Dutton, C. E., Lamey, C. A., and Treves, S. B., 1966, Geology of the Menominee ironbearing district Dickinson County, Michigan and Florence and Marinette Counties, Wisconsin: U.S. Geological Survey Professional Paper 513, 96 p. Drenth, B. J., Cannon, W. F., Schulz, K. J., and Ayuso, R. A., in press, Geophysical insights into Paleoproterozoic tectonics along the southern margin of the Superior Province, central Upper Peninsula, Michigan, USA: Precambrian Research. 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. 18 �Figure 1: Plan view of interpretations (Drenth et al., in press). All units Paleoproterozoic unless otherwise noted. 19 �The deformation and alteration of granitoid plutons in the Wabigoon subprovince DROVER, Bailey, HILL, Mary Louise, ZUREVINSKI, Shannon Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Ongoing research suggests that the brittle deformation, ductile deformation and associated metasomatism/hydrothermal alteration studied in 12 granitoid plutons across the Wabigoon subprovince provide evidence for coeval brittle-ductile deformation at the amphibolite facies of metamorphism consistant with Archean-aged oblique convergence (transpression). Approximate outcrop locations where the brittle and/or ductile deformation were studied are shown in figure 1. Four hundred and two structural measurements of fractures and faults along with approximately 95 thin sections were used to assess the degree of deformation /metamorphism and the metasomatism of various granitoid plutons across the subprovince. Figure 1. Modified map from Blackburn et al. (1991) showing the approximate outcrop locations where the brittle and/or ductile deformation of the granitoid plutons were studied. Brittle deformation features common to all plutons include chlorite ± epidote shear veins and unfilled fractures. Shear veins show a conjugate strike-slip relationship in the field and are seen with dips ranging from 15° to 90°. Average and median dips of the shear veins is roughly 70° ± 10° for each pluton studied. These shear veins are seen with sub-horizontal slickenlines that have pitches ranging from 00° to 33°. Although conjugate relationships are noted in the field, the orientation of the shear veins are quite variable across outcrops of individual plutons and the subprovince. These shear veins are also commonly seen defining the fabric in narrow shear zones present in the plutons across the Wabigoon subprovince. Many of the plutons also have quartz veins that have been subsequently ductiley deformed. These ductiley deformed quartz veins are commonly associated with narrow shear zones and are typically boudinaged with strong undulose extinction, subgrains and serrated grain boundaries. The ductile overprint of the quartz veins demonstrates coeval brittle-ductile deformation and is seen in the Sabaskong batholith, the Dryberry batholith, the Ottertail pluton, the Atikwa batholith and the Revell batholith. Although the plutons studied appear to lack a notable macroscopic deformation fabric, thin section analysis shows that dislocation creep mechanisms were active in both quartz and feldspar mineral phases during peak metamorphic conditions. Deformation microstructures commonly seen 20 �in quartz across all plutons studied include undulose extinction, serrated grain boundaries and subgrains. Feldspars have undulose extinction, subgrains, serrated grain boundaries, intragranular microfractures and deformation twins. Evidence for dislocation creep in feldspars indicates that the metamorphic grade of the plutons is consistent with the amphibolite facies of metamorphism. Alteration of the plutons is directly related to brittle fracturing. Alteration minerals commonly seen adjacent to brittle fractures include chlorite, epidote group minerals, biotite, sphene, calcite, sericite, and muscovite with lesser amounts of hematite. The degree of alteration is most notable near the margins of the plutons where strain is concentrated. REFERENCES Blackburn, C.E., John, G.W., Ayer, J., Davis, D.W., 1991, Wabigoon Subprovince. Ontario Geological Survey Special Volume 4, Pt. 1, pp. 303-383. 21 �Geochemistry and Age of Archean Volcaniclastic and Mafic Intrusive Rocks, Georgia Lake Area, Quetico Subprovince, Northwestern Ontario DUGUET, Manuel Earth Resources and Geoscience Mapping Section, Ontario Geological Survey, Sudbury, Ontario manuel.duuet@ontario.ca The metaturbidites of the Georgia Lake area in the Quetico Subprovince host 2 peculiar types of rock: volcaniclastic rocks that are mafic to intermediate in composition and mafic intrusive rocks that likely were emplaced when the turbidites were unconsolidated. The volcaniclastic rocks were described first on the shoreline of Lake Jean by Williams (1988), who interpreted them as sedimentary rocks derived from the erosion of nearby ultramafic bodies. A similar interpretation was given by Valli, Guillot and Hattori (2004), but they suggested a komatiite basalt as the source. Field work conducted during the summer 2019 showed that these rocks are very likely the result of deposition and local reworking of either ash from a volcanic plume or a distal part of a pyroclastic flow. Analyses of newly acquired and historical geochemical data were performed to, first, assess the level of contamination of these volcaniclastic rocks by the surrounding turbidites and, second, to compare it with possible source in the area such as the numerous Archean mafic sills and veins that intruded the turbidites. The least contaminated members of the volcaniclastic rocks fall in the fields of boninites and siliceous high-magnesium basalts in various discriminant diagrams such as SiO2 versus MgO, TiO2 versus MgO, SiO2 versus Cr and TiO2 versus Cr (Pearce and Reagan 2019). TiO2 content of the mafic intrusive rocks compared to that of the volcaniclastic rocks (respectively >1 wt.% and less than 0.6 wt. %) show that the former are unlikely to be the source of the volcaniclastic rocks. On primitive mantle– normalized incompatible element diagram, volcaniclastic rock compositions show moderate to strong enrichment in light rare earth elements (LREE) and depletion in heavy rare earth elements (HREE), with negative anomalies in niobium, tantalum and titanium and positive anomalies in zirconium and hafnium, as summarized in Duguet (2020). Although being more enriched in REE, the mafic intrusive rocks display similar fractionation pattern than those of the volcaniclastic rocks but with negative anomalies in zirconium and hafnium. LA-ICP-MS U-Pb dating on zircon of a volcaniclastic rock yielded a mean age of 2702 ± 3 Ma that is interpreted as its depositional age (Figure 1). These volcaniclastic rocks came either from the Wabigoon Subprovince to the north or the Wawa Subprovince to the south. Such rocks remain to be identified in these areas and the closest candidates may be the andesitic volcanic rocks of the Shebandowan greenstone belt to the south. Different process can generate boninitic magmas, but based on the geological setting and previous studies, an arc/fore-arc environment is favoured. The mafic intrusive rocks are also very interesting because they were intruded at the time when the sediments were deposited. This would rule out an accretionary prism environment for the Quetico Subprovince because active volcanism coeval with turbidite sedimentation is more common in back-arc and/or rifted arc settings. However, the calc-alkaline chemistry, which also display similarities with sanukitoid plutons in the area, contradicts this interpretation. The geodynamical implications of this remain to be investigated and a more complex scenario involving either rifted fore-arc/arc or subduction polarity flips should be investigated. 22 �Figure 1. Concordia plot for a volcaniclastic rock from the Georgia Lake area showing the main zircon population at 2702±3 Ma (n=28) reflecting the age of deposition. Geon 28, 29 and 30 grains are also present. Data from Sutcliffe (2020). Duguet, M. 2020. Geochemistry of Archean volcaniclastic and mafic intrusive rocks, Georgia Lake area, Quetico subprovince, northwestern Ontario; in Summary of Field Work and Other Activities, 2020; Ontario Geological Survey, Open File Report 6370: 8-1 to 8-10. Pearce, J.A. and Reagan, M.K. 2019. Identification, classification, and interpretation of boninites from Anthropocene to Eoarchean using Si-Mg-Ti systematics; Geosphere, 15: 1008-1037. doi.org/10.1130/GES01661.1 Sutcliffe, C.N. 2020. U-Pb geochronology by LA-ICP-MS in samples from northern Ontario; internal report prepared for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, Ontario, 128p. Valli, F., Guillot, S. and Hattori, K.H. 2004. Source and tectono-metamorphic evolution of mafic and pelitic metasedimentary rocks from the central Quetico metasedimentary belt, Archean Superior Province of Canada. Precambrian Research, 132: 155-177. Williams, H.R. 1988. Geological studies in the Wawa, Quetico, and Wabigoon subprovinces, with emphasis on structure and tectonic development; in Summary of Field Work and Other Activities, 1988, Ontario Geological Survey, Miscellaneous Paper, 141: 169-172. 23 �The Critical Mineral potential of carbonatite and alkalic complexes in Ontario EASTON, Robert Michael Earth Resources and Geoscience Mapping Section, Ontario Geological Ontario P3E 6B5 mike.easton@ontario.ca Survey, Sudbury, There are more than 50 carbonatite and/or alkalic complexes in Ontario ranging in age from Neoarchean to Jurassic (Figure 1). Although there are numerous reports and compendia on them (Sage 1991; Woolley and Kjarsgaard 2008), from a mineral exploration standpoint, there is little guidance available with respect to determining which complexes might have higher mineral potential than the others, especially for niobium and the rare earth elements (REE). Recent academic studies on alkaline and carbonatitic magmas by Nabyl et al. (2020) and Ballouard et al. (2020) indicate that calcio-carbonatites associated with coeval feldspathoidal-dominated silicate rocks (e.g., ijolite, nepheline syenite) are more likely to host niobium and/or REE mineralization than are calcio-carbonatites found with other igneous rock associations. This presentation examines this suggestion with respect to carbonatite and alkalic complexes in Ontario; more details are found in (Easton 2020). It should be noted that many of the complexes in Ontario, especially the carbonatite complexes, are typically poorly exposed, known only through limited diamonddrill hole data, and only major element geochemical data are available for most. Neoarchean alkaline intrusions in Ontario do not have phonolite to phono-trachyitic rock compositions and thus are not considered as having high REE potential. This also applies to the circa 1090-1065 Ma Kensington-Skootamatta intrusive suite in the Grenville Province. In contrast, mafic end-members of the Frontenac intrusive suite (circa 1180-1155 Ma) in the Grenville Province are characterized by high TiO2, P2O5 and alkali (Ba, K, Rb, Sr) contents, have bulk rock compositions of basaltic trachyandesite to trachyandesite, and have the potential to host iron-oxide (P-rich) mineralization (Easton 2019). Using the suggestion that phonolite to phono-trachyitic rock compositions may be a predictor or favourable rare element mineralization in carbonatite and alkalic complexes reduces the number of potential targets from more than 40 to ~10. The more favourable complexes are associated with feldspathoidal-bearing intrusive rocks and known pyrochlore and/or significant apatite occurrences. Most occur in 3 main areas of the province associated with areas of crustal extension and/or mantle upwelling. Specifically, 1) the Paleoproterozoic Argor Complex is proximal to the Circum-Superior Orogen; 2) the Manitou Island, Iron Island and Burritt Island complexes, and the Lavergne fenite, are all associated with the Ottawa–Bonnechere graben where it impinges on the former trace of the Sudbury dike swarm mantle plume at circa 1238 Ma; and, 3) Although proximal to the Midcontinent Rift (MCR), the Firesand River, Good Hope and Prairie Lake complexes were emplaced between 1163 and 1143 Ma, suggesting that they may be a far-field effect of the Grenville Orogen rather than the MCR. The Mesoproterozoic Lackner, Nemegosenda and Seabrook complexes along the Kapuskasing Structural Zone also have exploration potential. Acquisition of modern trace element geochemistry from these 10 complexes may help to further refine their potential to host rare element mineralization. Ballouard et al. 2020; Earth Science Reviews, v.203, article 103115, 31p. Easton, R.M. 2019. in Ontario Geological Survey, Miscellaneous Release—Data 351 Easton, R.M. 2020. in Ontario Geological Survey, Open File Report 6370, p.9-1 to 9-10. Nabyl, Z. et al. 2020. Geochimica et Cosmochimica Acta, v.282, p.297-323. Sage, R.P. 1991. Chapter 18, in Ontario Geological Survey, Special Volume 4, Part 1, p.682-709. Woolley, A.R. and Kjarsgaard, B.A. 2008. Geological Survey of Canada, Open File 5796. 24 �Figure 1. Locations of carbonatite complexes, alkalic complexes, aeromagnetically inferred alkalic and/or carbonatite complexes and areas of fenite and carbonatite dikes in Ontario. Also indicated are the traces of the southern limit of the Circum-Superior Orogen, the Kapuskasing Structural Zone, the Midcontinent Rift, and the Ottawa–Bonnechere graben. Complexes with higher estimated exploration potential are indicated in red. Figure modified from Sage (1991). 25 �Reconstructing the hydration and carbonation history of the Presque Isle peridotite, Marquette Michigan: Insights into mechanisms of carbon sequestration in ultramafic rocks GALLAGHER, Benjamin and BJORNERUD, Marcia Geosciences Department, Lawrence University, 711 East Boldt Way, Appleton WI 54911 USA The Presque Isle peridotite is a serpentinized lherzolite that occupies about 500,000 m2 on the northern outskirts of Marquette, Michigan. The age of the Presque Isle peridotite is not known, although it and a similar body, the Deer Lake Peridotite, exposed about 25 km to the WSW, have been assumed to be Neoarchean (Sims, 1991; Sasso, 2016). Both ultramafic masses lie within an area of Neoarchean (2.7 Ga) rocks -- the tonalitic to granodioritic Compeau Creek Gneiss -- near the southern edge of the Superior Craton, about 10 km north of the Great Lakes Tectonic Zone. If the peridotites are Neoarchean, they may represent cumulate bodies related to metabasalt/ greenstone units mapped as the Mona and Kitchi Schists in the Marquette region. However, there is another ultramafic body in the area, the mineralized Yellow Dog peridotite found 40 km northwest of Marquette, that is related to the ca. 1.0 Ga Midcontinent Rift – part of a dike swarm emplaced into Paleoproterozoic metasedimentary rocks of the Baraga Group. In the absence of direct geochronologic constraints, a Mesoproterozoic age for the Presque Isle and Deer Lake peridotites cannot be ruled out, although their irregular, non-tabular shapes are quite different from the fracture-controlled form of the Yellow Dog intrusions. The Presque Isle peridotite is overlain nonconformably by a unit that has been mapped as the ca. 1.0 Ga Jacobsville Sandstone (a late rift-filling unit). However, the sedimentary material just above the unconformity is a rubbly conglomerate about 2 m thick with a greenish matrix and irregularly shaped redorange clasts of chalcedony -- very different in character from the typically well sorted, quartzose Jacobsville Fm. In places along the contact, it is in fact difficult to distinguish the highly altered peridotite from the overlying sedimentary rocks. Despite the uncertainties about the igneous age and origin of the Presque Isle peridotite and the time at which it was exposed to weathering in the geologic past, extensive unvegetated exposures of the peridotite at the northern tip of Presque Isle Park make it possible to reconstruct a detailed relative chronology of fluid flow, deformation and weathering processes that altered the rock over time. Based on field, thin section, XRD and stable isotope analyses, we infer the following sequence of events, recording progressive cooling and exhumation of the peridotite from depth. 1) Percolative infiltration by water-rich fluids The least altered parts of the Presque Isle peridotite still have recognizable igneous textures and record an early period of pervasive, percolative flow by aqueous fluids that altered the original olivine crystals to lizardite along grain boundaries and transgranular cracks. Phase relations for ultramafic rocks (Philpotts and Ague, 2009) suggest this happened at temperatures below 500-600°C, depending on depth. XRD analysis also indicates the presence of minor talc in these rocks, which could indicate metamorphism at slightly higher temperatures (up to ca. 650°). Serpentine also occurs in centimeter-scale curviplanar seams within the Presque Isle peridotite. These seams are typically meters apart but variably oriented and intersecting, and thus do not appear to have been influenced by a consistent external stress regime. Although the seams represent a morphologically distinct style of serpentinization, they may be temporally linked with the grain-scale alteration of the rocks. The serpentine seams exhibit small-scale transverse fractures whose spacing is comparable to their width; such features have been interpreted in other peridotites to reflect the significant volume increase (close to 40%; 26 �Klein and Le Roux, 2020) that occurs with serpentinization. This early serpentinization event does not seem to have been associated with significant deformation. 2) Tectonically-influenced hydrofracturing and reaction-induced fracturing by CO2-rich fluids Carbonate material in different forms cuts across serpentinized areas of the peridotite. The simplest carbonate occurrences are NW-striking, vertical dolomite-serpentine veins with fibrous crack-seal textures indicating incremental, dilational opening. These are spatially associated with vertical WNW-trending dextral faults with dolomite slickenfibers. The orientation of the veins relative to the faults is consistent with their formation as tensile fractures in a dextral shear zone. The crack-seal texture of the vein fillings suggests repeated, perhaps seismically induced, fluid flow and hydrofracturing events at depths shallow enough that the rocks deformed brittlely. In a few samples, carbonate crystals have terminated faces, suggesting they grew into open space or a dense vapor phase; this too implies shallow depths of formation. More geometrically complex carbonate masses are 5-10 cm thick ‘rinds’ around pods of less altered peridotite. These have crude ‘onion-skin’ layers and transverse veins that create a mesh-like appearance, and are similar to features attributed to reaction-induced fracturing by Jamtveit et al (2008). Like the crackseal veins, the carbonate meshes consist mainly of dolomite. Oxygen isotope values for both the vein and mesh carbonates are consistently negative (18O VPDB between -6.57 and -11.89) and suggest that the mineralizing fluids came from meteoric water. These fluids must have carried enough dissolved CO2 to cause reactions such as: Mg2SiO4 (ol) + CaMgSi2O6 (cpx) + CO2 + H2O → Mg3Si2O5(OH)4 (serp) + 2CaMgCO3 3) Deep near-surface weathering under high CO2 conditions Beginning about 5 m below the unconformity, the peridotite has a strikingly different color and composition. The rock mass consists of at least 30% carbonate (dolomite, magnesite, rhodochrosite), some of it massive, with intervening areas of reddish material that represents extremely altered olivine and serpentine. As one approaches the unconformity, even the serpentine has been altered to a residuum of fine-grained quartz or chalcedony and opaque minerals, likely Fe and Mg oxides. This deep weathering may reflect long-term leaching by surface fluids in late Proterozoic time under higher atmospheric CO2 levels than at present. The large amount of carbonate near the unconformity indicates that the rock continued to be reactive even after the original olivine had already been altered to serpentine. The complex hydration and carbonation history of the Presque Isle peridotite provides a natural analog for physical and chemical processes that might be emulated for carbon sequestration in ultramafic rocks if carbon capture becomes technologically and economically feasible. References cited Jamtveit, B., Putnis, C., and Malthe-Sorenson, A., 2008. Reaction induced fracturing during replacement process. Contributions to Mineralogy and Petrology, 157: 127-133. Klein F. and Le Roux, V., 2020. Quantifying the volume increase and chemical exchange during serpentinization. Geology, 48: 552-556. Philpotts, A. and Ague, J., 2009. Principles of Igneous and Metamorphic Petrology. Cambridge Univ Press. Sasso, A., 2016. Geochemical and Petrological Characterizations of Peridotite and Related Rocks in Marquette County, Michigan. Master’s Thesis, Western Michigan University. 100 pp. Sims, P.K., 1991, Great Lakes Tectonic Zone in Marquette Area, Michigan Implications for Archean Tectonics in the North-Central United States. US Geological Survey Rep. 1904-E. 27 �A Practical Approach to Using Incompatible Elements to Define Geochemical Correlation and a Common Mantle Source: Example Crystal Lake Gabbro and the Duluth Complex GOOD, David Department of Earth Sciences, Western University, London, ON N6A 5B7 Canada This presentation focuses on fundamental principles of igneous geochemistry to provide a practical approach for identifying geochemical correlations between mafic rock units in the MCR. A common approach will aid in development of a rift-wide classification strategy and knowledge of the geographic extent of defined mafic rock series. This level of information is key to understanding the nature of mantle sources across the rift. The ultimate goal is to provide useful criteria to guide ongoing exploration for Cu-Ni-PGE deposits in the MCR. There is a seemingly infinite array of methods to combine geochemical criteria and trace element data to derive useful petrogenetic models to explain mafic rocks, but most are not necessary for a regional scale study such as this. By defining a clear hypothesis, it is possible to simplify the number of elements used to a manageable and reliable few. In this case, a widely known concept that igneous rock sequences which have not been contaminated by crustal material, and that have similar incompatible trace element patterns were derived from a similar mantle source. With this objective in mind, it is not necessary to consider elements that are susceptible to hydrothermal alteration or are associated with plagioclase. Thus, U, Th, Ta, Nb, Zr and the REE provide a definitive range of elements for making comparisons and drawing conclusions regarding correlation (but do not necessarily show rocks are co-genetic). In those rare instances where there is evidence for mantle metasomatism, K, Ba and Sr abundances should be considered. A practical approach applies 4 steps: 1) elimination of samples considered to have undergone crustal contamination by evaluation Th/Nb (>0.2), 2) cleaning the data set of unrelated or outlier samples using REE patterns or the λ1 vs λ2 plot, 3) testing for evidence of clinopyroxene fractionation (La vs. Sm), and 4) inspection and comparison of twice normalized (primitive mantle and Lu) extended trace element diagrams. Two examples are presented: 1) the Babbitt intrusion of the Duluth Complex is compared to the Crystal Lake gabbro, and 2) the Current Lake intrusion is compared to the Lower Suite (Simpson Island traverse) of the Osler Volcanic Group. Sources: Current Lake: Cundari, R.M., Puumala, M.A., Smyk, M.C. and Hollings, P.N. 2021. OGS MR Data 308 Osler Volcanic Group: Barnes et al., 2021, Magmas Through Time, J. Petrology, in press. Crystal Lake Gabbro: O’Brien, S. 2018, MSc Thesis, Lakehead University. Duluth Complex: Ripley et al., 1998, Geochimica et Cosmochimica Acta 62, 3349-3365. λ-λ plot: O’Neill, H. 2016, J. Pet. 57, 1463-1508 28 �Integration of geophysical evidence indicates that anorthosite composes a significant portion of Grand Marais ridge, an inferred basement high in western Lake Superior GRAUCH, V.J.S. 1 and HELLER, S.J. 2 1 U.S. Geological Survey, MS 973, Federal Center, Denver, CO 80225 U.S. Geological Survey, MS 939, Federal Center, Denver, CO 80225 2 The Midcontinent Rift System (MRS) is expressed geophysically by a semi-linear, regional gravity high that trends across the Midcontinent and Great Lakes region of North America. The gravity high is interrupted by two prominent, semicircular gravity lows, which have been interpreted from modeling and seismic-reflection sections as basement highs of Archean granite (Allen et al., 1997). One is centered southwest of Isle Royale in western Lake Superior (Grand Marais ridge) and the other over Bayfield Peninsula (White’s Ridge). Allen et al. (1997) suggest that the Archean granite highs were pre-rift features that remained high while lava basins of the MRS subsided adjacent to them. Hart et al. (1994) questioned the presence of granitic rocks underlying Grand Marais ridge (GMR) because heat flow measurements there are much lower than is typical for Archean granitic upper crust. They argued that the region must instead be underlain by rocks of low radiogenic heat production, such as gabbro, extending to at least 15 km depth. However, gabbro has high densities and would not produce the observed gravity low. Thus, the geophysical observations appear contradictory. Figure 1. Physical properties of rock types compared to the ones inferred for Grand Marais ridge. (A) All velocitydensity point data are from Christensen and Mooney (1995), representing means at 5, 10, 15, 20, 25 km depths, except for two additional sites for anorthosite (orange), which are from Birch (1961), measured at 2 kbars. The regression line representing Keweenawan basalt and diabase is from Halls (1969). Typical radiogenic heat production by rock type (in brackets) in µW/m3 are from Hasterok et al. (2018) except for anorthosite, which is from Roy et al. (2021). (B) Statistics for Duluth Complex and Animikie Basin densities were analyzed from the database of Chandler and Lively (2011). Representative densities for the NE Grenville and Egersund massifs are from Keary and Thomas (1979) and Smithson and Ramberg (1979), respectively. N is the number of samples measured. Bouguer gravity, seismic-reflection, seismic-refraction, and heat-flow data suggest that the rock types composing GMR to a depth of about 15 km have densities of 2,650-2,720 kg/m3 (Hutchinson et al., 1990; Allen et al., 1997), compressional-wave velocities of about 6.5 km/s (Luetgert and Meyer, 1982; Hutchinson et al., 1990; our own velocity analyses), and radiogenic heat production 29 �values of no more than 0.7 µW/m3 (Hart et al., 1994). These inferred GMR physical properties are compared to the same properties compiled globally and measured locally for rock types likely to occupy the subsurface of western Lake Superior (Fig. 1). As already noted, the granitic rocks have densities and velocities closest to the GMR suite of properties, but the heat production is too high. The mafic and intermediate rocks (Fig. 1A) and metasedimentary rocks of the Animikie Basin (Fig. 1B) are too dense. The latter also are likely to have heat production that is too high (noted next to MGW on Fig. 1A). If the values compiled for anorthosite (Fig. 1A) are representative for the MRS, the low heat production fits but velocities are too high for anorthosite to be the sole rock type underlying GMR. However, if added in significant volume to felsic rock types, for example, the velocity and the heat generation could be combined to result in the desired physical properties for the whole rock column. Alternatively, the lower densities for anorthosite xenoliths and massifs elsewhere (Fig. 1B) hint that the compiled values are not representative; the xenolith source may have velocities that match those of GMR. In any case, anorthosite likely constitutes a significant volume of GMR. 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: Geological Society of America Special Paper 312, p. 47-72. doi: 10.1130/0-8137-23124.47 Birch, F., 1961, The velocity of compressional waves in rocks to 10 kilobars, Part 2: Journal of Geophysical Research, v. 66, no. 7, p. 2199-2224. Chandler, V. W., R. S. and Lively, 2011, Density, magnetic susceptibility, and natural remanent magnetization of rocks in Minnesota: Minnesota Geological Survey online services, http://dx.doi.org/10.13020/D63S3D, accessed 22 January 2018. Christensen, N.I., and Mooney, W.D., 1995, Seismic velocity structure and composition of the continental crust: A global view: Journal of Geophysical Research, v. 100, no. B6, p. 9761-9788. doi:10.1029/95JB00259. 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. Hart, S.R., Steinhart, J.S., and Smith, T.J., 1994, Terrestrial heat flow in Lake Superior: Canadian Journal of Earth Sciences, v. 31, p. 698-708. Hasterok, D., Gard, M., and Webb, J., 2018, On the radiogenic heat production of metamorphic, igneous, and sedimentary rocks: Geoscience Frontiers, v. 9, p. 1777-1794. Hutchinson, D.R., White, R.S., Cannon, W.R., 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, no. B7, p. 10,869-10,884. Keary, P., and Thomas, M.D., 1979, Interpretation of the gravity field of the Lac Fournier and Romaine River anorthosite massifs, eastern Grenville Province: Significance to the origin of anorthosite: Journal of the Geological Society of London, v. 136, p. 725-736. Luetgert, J. H., and Meyer, R. P., 1982, Structure of the western basin of Lake Superior from cross structure refraction profiles, in Wold, R. J., and Hinze, W. J., eds., Geology and Tectonics of the Lake Superior Basin, Geological Society of America Memoir 156, p. 245-255. Roy, D.J., Merriman, J.D., Whittington, A.G., Hofmeister, A.M., 2021, Thermal properties of carbonatite and anorthosite from the Superior Province, Ontario, and implications for non-magmatic local thermal effects of these intrusions: International Journal of Earth Sciences, published online 29 March 2021, doi: 10.1007/s00531-02102032-w. Smithson, S.B. and Ramberg, I.B., 1979, Gravity interpretation of the Egersund anorthosite complex, Norway: Its petrological and geothermal significance: Geological Society of America Bulletin, v. 90, p. 199-204. 30 �Final inversion of the Midcontinent Rift during the Rigolet Phase of the Grenvillian orogeny HODGIN, Eben B., SWANSON-HYSELL, Nicholas L., STOLPER, Daniel A., TURNER, Andrew C. Department of Earth and Planetary Science, University of California, Berkeley, CA, USA DeGRAFF, James M., Department of Geological and Mining Engineering and Sciences, Michigan Technical University, Houghton, MI, USA KYLANDER-CLARK, Andrew Department of Earth and Planetary Science, University of California, Santa Barbara, CA, USA SCHMITZ, Mark D. Department of Geosciences, Boise State University, Boise, ID, USA The Keweenaw Fault is a major structure associated with inversion of the Midcontinent Rift; it thrusts ca. 1093 to <1086 Ma rift volcanic and sedimentary rocks atop the post-rift Jacobsville Sandstone in northern Michigan. Given the position of folded Jacobsville Sandstone in the footwall of the fault, its depositional age provides a maximum age constraint on faulting. Previous detrital zircon U-Pb geochronology used laser ablation-inductively coupled plasma mass spectrometry (LA-ICPMS) that led to an interpretation of a ~950 Ma maximum depositional age for the Jacobsville Sandstone and to argue that motion on the Keweenaw Fault must postdate the ca. 1080 to 980 Grenvillian orogeny (Craddock et al., 2013; Malone et al., 2016). In this study, we reproduced similar low-precision maximum age constraints on detrital zircons using LA-ICPMS, then we dated the youngest detrital grains at high-precision by chemical abrasion-isotope dilutionthermal ionization mass spectrometry (CA-ID-TIMS). Our revised maximum depositional age is based on the youngest detrital zircon dated by CA-ID-TIMS and contemporaneous with the 1000– 980 Ma Rigolet Phase of the Grenvillian orogeny. This revised maximum age constraint overlaps with U-Pb LA-ICPMS dates on syn- to post-kinematic calcite veins within the brecciated fault zone of the Keweenaw Fault that yield lower intercept dates that are also contemporaneous with the Rigolet Phase of the Grenvillian orogeny. From these age constraints, we infer that the Jacobsville Sandstone was deposited during the final inversion of the Midcontinent Rift during the Rigolet Phase of the Grenvillian orogeny, just before reverse motion on the Keweenaw Fault. This deposition and subsequent reverse faulting is contemporaneous with development of the Grenville Front during the Rigolet phase of the Grenvillian orogeny. These data demonstrate the propagation of Rigolet-aged orogenesis into the continental interior ca. 580 km from the Grenville Front. Craddock, J.P., Konstantinou, A., Vervoort, J.D., Wirth, K.R., Davidson, C., Finley-Blasi, L., Juda, N.A., & Walker, E. (2013). Detrital zircon provenance of the Mesoproterozoic Midcontinent Rift, Lake Superior region, USA. The Journal of Geology, 121(1), 57–73. doi:10.1086/668635 Malone, D.H., Stein, C.A., Craddock, J.P., Kley, J., Stein, S., & Malone, J.E. (2016). Maximum depositional age of the Neoproterozoic Jacobsville Sandstone, Michigan: Implications for the evolution of the Midcontinent Rift. Geosphere, 12(4), 1271–1282. 31 �Olivine Geochemistry from Fe-Ti Oxide-Bearing Ultramafic Intrusions in the Duluth Complex, MN KATTEMALAVADI, Amartya , KLEINSASSER, Jackie , SIMON, Adam , PETERSON, Dean , and HUDAK, George 1 2 1 2 1 1 2 Department of Earth & Environmental Sciences, University of Michigan Natural Resources Research Institute, University of Minnesota–Duluth The Duluth Complex is a series of mafic intrusions in northeastern Minnesota which formed 1.1 billion years ago as part of the Midcontinent Rift System. On the southern side of the Complex, specifically on the western margin, it hosts 14 different ultramafic intrusions which bear Fe-Ti oxides, known as oxide bearing ultramafic intrusions, or OUIs. There are currently three competing hypotheses about how these OUIs formed, but none of them have been rigorously tested. However, there are many minerals present in these OUIs which can be used as proxies to Figure 1: Depth vs Forsterite content in Titac and Longnose. Both discover more about their the surrounding host troctolites and the OUI rocks themselves were genesis, such as olivine. analyzed, and results can be seen here. Since olivine is a solid solution mineral and is typically one of the first minerals to crystalize from a melt phase, we can gain a lot of information about the way the magma formed from its composition alone. Two OUIs in particular, Titac and Longnose, were studied in thin section using a variety of tools, including a petrographic microscope, a scanning electron microscope (SEM), and an electron probe microanalyzer (EPMA). The petrographic microscope and SEM showed the different textures in which olivine formed. In Titac, olivine appears to be tubular in shape and fractured. In Longnose, olivine appears to form in a ropy shape, and is heavily striated. There is evidence of fracturing, but not as much as in Titac. In addition, there is evidence of deformation, shown by grain boundary bulging of Fe-Ti oxides and plagioclase present. In both intrusions, the olivine is heavily serpentinized. Under the EPMA, compositions of olivine were measured, of olivine in OUIs as well as olivine in surrounding host rocks (troctolites). The results of this analysis can be seen in Figure 1. In the OUI rocks, it was found that forsterite content was overall higher than the host rock in both intrusions (~Fo60 to ~Fo70), with Longnose having a higher content than Titac. It also appears to be very consistent throughout the drill core for both. In addition, the host rocks have much lower amounts of forsterite (~Fo60 to ~Fo35), with the Longnose (both OUI rock and host rock) having a higher forsterite content overall. The higher forsterite content in the OUI rocks can likely be attributed to the abundance of Fe-Ti oxide minerals, which leach out iron from the area, leaving the olivine with a higher forsterite content. The higher forsterite content in Longnose 32 �indicates that the magma which formed it is relatively more primitive than the magma which formed Titac. The forsterite content is consistent with the findings of Miner (1995) in her thesis about the Longnose body1. We also compare the forsterite content to the content of nickel in the olivine (trace). We can see that there is overall more nickel in the Longnose OUI, and this is most likely due to the lower modal abundance of sulfide minerals compared to Titac. 1 Miner, G. C., 1995. Aspects of the Petrogenesis of the Longnose Fe-Ti-Oxide-Rich Ultramafic Body, Duluth Complex, MN. Washington University Department of Earth and Planetary Sciences. 33 �Textures and geochemistry of ilmenite and titanomagnetite in Fe-Ti oxide-bearing ultramafic intrusions of the Western Margin of the Duluth Complex, Minnesota KLEINSASSER, Jackie1, SIMON, Adam1, PETERSON, Dean2, and HUDAK, George2 1 Department of Earth & Environmental Sciences, University of Michigan, Ann Arbor, MI, USA Natural Resources Research Institute, University of Minnesota–Duluth, Duluth, MN, USA 2 The Duluth Complex hosts numerous undeveloped mineral deposit types, including CuNi, PGE, Mn, Fe-Ti±V, and others. Significant Fe-Ti±V mineralization has remained particularly unstudied over the past 20 years and represents one of the United States’ most promising domestic resources of these energy- and infrastructure-critical metals. Fe-Ti±V mineralization is hosted in ~14 Fe-Ti oxide-bearing ultramafic intrusions (known in the literature as OUIs) along the Western Margin. Although most are broadly similar in mineralogy, three contrasting genetic models have been proposed to explain the source of Fe and how the deposits formed. Because there exists incredibly scarce geochemical evidence to (a) support or refute any of the three models, ilm and the main objective of this project is to unravel the genesis of Fe-Ti±V oxidebearing ultramafic intrusions. pln ilm Focusing on the Longnose and Titac (Section 34) OUIs, two well-drilled intrusions ~40 km apart hosted in the ttm Partridge River and Western Margin intrusions, respectively, we used detailed optical and microbeam methods to compare the textures and compositions of 100µm Fe-Ti oxides (ilmenite, titanomagnetite, and magnetite) to gain insight into their (b) origin and cooling history. Both intrusions show similar complex textures despite ttm differing proportions of ilmenite and titanomagnetite, with Longnose typically having higher amounts of ilmenite. ‘dog tooth’ Titanomagnetite contains extensive pln ilmenite and pleonaste ((Fe,Mg)Al2O4). ilm Ilmenite is present as primary grains and as ‘oxy-exsolution’ lamellae in titanomagnetite (e.g., Buddington and Lindsley, 1964) as thin trellis-type 100µm lamellae, thicker bands of sandwich-type lamellae, and titanomagnetite-hosted Figure 1 – BSE images of Fe-Ti oxides from Longnose (a) and granular exsolutions (Fig. 1). Pleonaste is Titac (b). Multiple types of exsolution and oxy-exsolution are present in ilmenite at or near the contacts present. Ilmenite – ilm; titanomagnetite – ttm; pleonaste – between titanomagnetite and ilmenite as pln. large, granular blebs and as ‘dog tooth’ symplectites that are triangular intergrowths with ilmenite (Fig. 1b). In titanomagnetite, pleonaste exsolution is present as lamellae and cubic granules and within and around ilmenite sandwich-type 34 �lamellae as both vermicular symplectites (Fig. 2a) and blebs. Hematite lamellae are present in ilmenite from both OUIs, which indicates the original ilmenite-hematitess decomposed during cooling below 750–675℃ to form the hematite exsolution rods (Lindsley 1991). These textures indicate that both magnetite-ulvöspinelss and ilmenite-hematitess were present and that cooling and oxidation-reduction reactions induced the various exsolution textures observed, starting at >900℃ (Turnock and Eugster 1962). (a) (b) ilm pln hem pln ttm ilm 10µ m 30µ m Figure 2 – BSE images of finer exsolution textures from Titac. In (a), ilmenite oxyexsolution bands contain vermicular pleonaste symplectites while titanomagnetite contains pleonaste exsolution lamellae. In (b), the presence of hematite lamellae that exsolved during cooling through decomposition of an original ilmenite-hematitess below 750–675℃. Ilmenite – ilm; titanomagnetite – ttm; pleonaste – pln, hematite – hem. Major and minor element compositions of titanomagnetite and ilmenite are broadly similar between Longnose and Titac, with the possibility of remobilization and distribution of V and Co during alteration in Titac. The initial results of this study point towards a similar complex cooling history for Fe-Ti oxide liquids in both the Longnose and Titac OUIs, despite being emplaced in different rock types and hosted in different intrusions. Future work will involve measuring stable Fe and Ti isotopes of the ilmenite and titanomagnetite to fingerprint the metal sources and directly assess whether assimilation of the Biwabik Iron Formation is necessary to form the OUIs. Another avenue for future work involves using stable O isotopes to test whether volatiles released during the assimilation of Virginia Formation footwall were responsible for transporting metals and the serpentinization of olivine. Buddington, A., and Lindsley, D., 1964. Iron-Titanium oxide minerals and synthetic equivalents. Journal of Petrology, 5: 698–719. Lindsley, D., 1963. Equilibrium relations of coexisting pairs of Fe-Ti oxides. Carnegie Institution of Washington Year Book, 62: 60–66. Turnock, A. and Eugster, H., 1962. Fe-Al oxides: Phase relationships below 1000. Journal of Petrology, 3: 533–565. 35 �Identifying the genesis of Fe-Ti oxide- and sulfide-bearing ultramafic intrusions in the Duluth Complex through sulfide geochemical analysis LACHANCE, Kyle1, KLEINSASSER, Jackie1, SIMON, Adam1, PETERSON, Dean2, and HUDAK, George2 1 Department of Earth & Environmental Sciences, University of Michigan Natural Resources Research Institute, University of Minnesota–Duluth 2 During the formation of the 1.1 Ga Midcontinent Rift System, a series of massive mafic intrusions were emplaced in northeast Minnesota and are referred to as the Duluth Complex. Rock types include a series of gabbroic, troctolitic and anorthositic intrusions that protruded into the Archean and Paleoproterozoic footwall, which comprises banded iron formation and other metasedimentary rocks. The western margin of the complex is home to 14 different Fe-Ti oxidebearing bodies containing ilmenite, magnetite and titanomagnetite. While these oxide-bearing ultramafic intrusions (OUIs) have been identified Figure 1: (Femol+ Nimol)/Smol vs. Smol ratio plot demonstrating the distribution of electron probe in field studies, their exact genesis is unknown. microanalysis data for both deposits Three hypotheses have emerged that attempt to explain the origin of the OUIs, none of which has been rigorously tested. To investigate the origin of the OUIs, we studied the sulfide geochemistry of two spatially distant OUIs: Longnose, which is located in the Partridge River Intrusion, and Titac (sec. 34), which is located in the Western Margin Intrusion. Published compositional data for sulfides from these OUIs is sparse, with most studies reporting a very brief overview of the sulfide abundances, mineralogies and stable sulfur isotopes. One study, for instance, reports modal abundances of sulfides within Longnose ranging from 0.2-5%, 99% of which being chalcopyrite (Linscheid 1991). Our petrographic study of samples from Titac and Longnose revealed a greater abundance and variety of sulfides in samples from both deposits. As the first systematic geochemical comparison of sulfides between OUIs, our research began with scanning electron microscopy (SEM) to image the sulfide grains, followed by energy dispersive spectroscopy (EDS) to identify a range of different sulfides present within the OUI, and finally electron probe microanalysis (EPMA) to quantitatively determine the elemental makeup of the sulfides. Initial analysis of these data demonstrates that chalcopyrite is the modally dominant sulfide in both Titac and Longnose, with trace bornite, sphalerite and millerite found in both deposits as well. Pyrite was found only in the Titac samples, whereas chalcocite was found only at Longnose. Using the electron probe data, metal/sulfur ratio plots can be produced, such as in figure 1, to visualize the full spectrum of data including non-stoichiometric points that do not distribute into the expected mineral phases. As seen in this figure, which takes the ratio of moles of Fe and Ni to moles of S, the data indicates areas of points consistent with the Fe+Ni values for sulfides of interest, such as chalcopyrite around the value of 0.5 and bornite around 0.25. Many points were present that did not match any known mineral elemental compositions, suggesting points of exsolution, mixing or inclusions. Some Co-rich data points were also present that we were unable to characterize, so further analysis will be required to identify, if any, Co-sulfides within the deposits. 36 �With similar bulk sulfide content across the two deposits, these initial results suggest a consequently similar genesis across both of the OUIs. Future work regarding the stable sulfur isotopes will be necessary to identify the origin of the sulfur content to determine whether it is magmatic or through assimilation from the nearby Virginia Formation. Due to the heavier footprint of the Virginia Formation sulfur, an increase in the sulfur isotope ratio in the deposits would indicate some form of assimilation rather than solely a magmatic origin. Furthermore, a complete study on trace elements, such as the Pt group elements, will be necessary to further this project. Reference Linscheid, E., 1991. The Petrography of the Longnose Peridotite and Its Relationship to the Duluth Complex: MSc thesis, University of Minnesota Duluth, 121p. 37 �Archean Orogenesis to Proterozoic Rifting: A structural history of Pass Lake, Thunder Bay, Ontario LANDMAN, Megan and HILL, Mary Louise Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON, P7B 5E1 Canada New roadside outcrops along the Trans-Canada Highway 11/17 near Pass Lake, ON, expose the basal unconformity between Archean basement rock and the Proterozoic Gunflint Formation. Shear fractures, joints, and the Blende Lake Fault damage zone seen in these two outcrops record the brittle deformation history of the area before and after the Gunflint Formation was deposited. The unconformity represents a temporal gap in the stratigraphic sequence of at least 800 million years. The Archean rock belongs to the Wawa subprovince of the Superior province, in close proximity to the Wawa-Quetico subprovince boundary. Structural measurements, stereographic projections, and qualitative observations have allowed deeper insight and analysis of how structural controls have changed over an approximately 2.7 billion year-old geological history. The Archean basement unit underneath the unconformity is a coarse-grained amphibolite that contains accessory epidote and biotite, and is homogeneous throughout the length of both high-standing outcrops. This is interpreted to be a mafic pluton that has undergone amphibolitefacies metamorphism. The amphibolite records a scatter in orientation of joints and shear fractures, but some trends align well with data from Mackenzie River granite plutons, including an overall east-northeast and west-southwest strike. Later post-Proterozoic features, including the Blende Lake fault, have a common strike of east-northeast, which closely align with the orientation of the 1.1 Ga Mid-Continent Rift in Thunder Bay. This similarity is further reflected by the Blende Lake fault being oriented subparallel to the Mid-Continent Rift related silver veins. Similarities between trends in the amphibolite and Gunflint Formation suggest that the Mid-Continent Rift in Thunder Bay may have reactivated some Archean-aged orogenic-related faults and shear fractures. Minor folding in the Gunflint Formation directly adjacent to the Blende Lake Fault may be evidence of compression during the later stage of the Mid-Continent Rift. 38 �Characterization of Hydrothermal Alteration and Sulfide Ores at the Lynne Zn-Cu-Pb Deposit, Oneida Co. WI. LEMKE, Tara C., WEBER, Evan M., and LODGE, Robert W.D. Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI The volcanogenic massive sulfide (VMS) Lynne Zn-Cu-Pb deposit was discovered by Noranda Exploration in 1990 within Oneida County, WI (Dematties, 1994; Adams, 1996). The Lynne deposit is one of many of VMS deposits across northern Wisconsin found in the Pembine-Wausau terrane. The Pembine-Wausau terrane is an accreted volcanic arc that formed during the Paleoproterozoic Penokean Orogeny. VMS deposits form from metallic-rich fluid exhaling on or near the ocean floor during extensional submarine volcanism (Galley et al., 2007). Interestingly, the hydrothermal alteration at the Lynne deposit contains talc-carbonate and calc-silicate mineral assemblages (Adams, 1996) and is unique compared to other deposits in Wisconsin (DeMatties, 1994). In addition, the sulfide ores contain more Pb than other deposits (DeMatties, 1994). The purpose of this study is to better characterize these unique hydrothermal rocks to improve our knowledge of the ore-forming environment at the Lynne deposit. The drill core from the Lynne deposit was obtained from the Natural Resources Research Institute in Duluth, Minnesota, and brought to Eau Claire where it was re-logged and sampled. Samples taken from the hydrothermal alteration and mineralized zones were cut into thin sections for petrographic analyses using transmitted- and reflected-light microscopes and scanning electron microscope (SEM) to determine the paragenetic sequence of ore minerals. Major and trace element geochemistry of altered rocks were determined using X-Ray Fluorescence (XRF) at UW-Eau Claire to characterize the element mobility during alteration and to determine the nature of the protoliths. The main sulfide mineralization within the Lynne deposit with abundance in decreasing order includes sphalerite, pyrrhotite, pyrite, galena, and chalcopyrite. Gold, silver, and copper are also found but in more localized sections within the Lynne deposit (Adams, 1996). The mineralization can be further divided into units A, B, and C, based on their stratigraphic position and compositional differences (Adams, 1996). Sphalerite is formed early on and is replaced by other minerals over time. In the deeper ore zones, pyrite tends to form relatively early and is replaced over time but in the ore zones closer to the surface, pyrite forms later. Chalcopyrite forms later in the ore zones replacing other minerals as it grows. Galena, pyrrhotite, and magnetite form in the middle replacing some minerals such as sphalerite and getting replaced by others such as pyrrhotite over time. The main alteration types found within the Lynne contain are carbonate and talc alteration. The unusually high amounts of carbonate material - unique to VMS deposits - found within the alteration zones are credited to CO2-rich hydrothermal vent fluid (Adams, 1996). The talc alteration is highly variable throughout the deposit but is largely found within two major zones. There is also unique skarn-type alteration is noted by chloritized diopside and garnet. A brecciated alteration texture is also notable as it shows progressive replacement of volcanic rocks with alteration minerals. Major chemistry indicates Mg-metasomatism with MgO values up to 30 wt% and five times that of the hosting volcanic rocks. Trace element chemistry reveals that these unique alteration assemblages had felsic protoliths with similar trace element abundances as the host rhyolites. 39 �Figure 1: A cross section of the Lynne deposit shows the distinct sulfide units. A) massive sphalerite and pyrite, B) disseminated sphalerite in carbonate host rock, C) massive sphalerite with semi-massive pyrite and disseminated galena, D) talc alteration, E) carbonate alteration, and F) skarn alteration (Adams, 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, p. 161-179. DeMatties, T.A., 1994, Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Economic Geology, 89: 1122-1151. Galley, A., et al., 2007, Volcanogenic massive sulfide deposits in Goodfellow W.D., ed., Mineral Deposits of Canada, Special Publication 5, p. 141-161. 40 �Island of Hawaii Field Trip, February 11-21, 2020 MACTAVISH(1), Allan, HINZ(2), Peter, ARTHUR(3), Mary Kay, CHATAWAY(4), Robert, EDBERG(3), Jim, ERICKSON(3), Tom, FOX(3), Steve, FURLONG(3), Joan, GERLICH(3), Jim, SMITH(5), Lindsay, WILHELM(3), David 1 Clean Air Metals Inc., Thunder Bay, ON; 2Ontario Ministry of Energy, Northern Development and Mines, Thunder Bay, ON; 3Minnesota Geological Society, Minneapolis, MN; 4Consulting Geologist, Thunder Bay, ON; 5Big Rock Exploration, Vancouver, BC Between February 11 and 21, 2020 the Institute on Lake Superior Geology facilitated a field trip to the Island of Hawaiʻi. This 10-day trip was led by Allan MacTavish and Peter Hinz, and consisted of 11 participants that included 4 professional geologists and 7 members of the Minnesota Geological Society. The trip circumnavigated the entirety of the Big Island of Hawaiʻi (Figures 1 and 2) and visited all 5 exposed volcanoes on the island. The aims of the field trip were to observe the characteristics of modern volcanoes in an intra-plate, non-rift environment (see) and to provide contrast and comparison to modern and ancient rift environments (i.e. the MCR and Iceland; Figure 3). The Big Island is made up of five distinct volcanoes: Kohala, Hualālai, Mauna Kea, Mauna Loa and Kīlauea. A sixth volcano associated with the Big Island, Loʻihi, is an active undersea volcano located approximately 20 km (12mi) south of Kīlauea. Stops were made at all five (5) extinct, dormant, or active volcanoes with the intent to focus on volcanology, igneous petrology, and geomorphology as well as a few stops at Hawaiʻian cultural and historical locations The 9 main islands comprising the Hawaiʻian islands are part of a chain of over 100 individual volcanoes in various states or formation and erosion (Lockwood and Hazlett, 2010). The Hawaiʻian-Emperor Chain is composed of eighteen islands, atolls, and seamounts stretching for greater than 6000km from the Island of Hawaiʻi and Loihi Seamount, in the east, northwestward and then north-northwestward across the Pacific Ocean to the Aleutian Trench (Lockwood and Hazlett, 2010; Best, 2003). It is postulated that this chain of islands was formed over a greater than 65 million year period as the Pacific Ocean plate migrated over a mantle “hotspot” initially in a north-northwest direction (Emperor Chain) and later in a northwesterly direction (Hawaiʻian Chain) (Barker, 1983). The volcanism creating the islands was, and is, primarily basaltic in chemistry with the occurrence of both effusive (dominant) and explosive eruptions (Lockwood and Hazlett, 2010). Explosive activity is primarily due to the interaction between basaltic magma and ground or sea water. This presentation will summarize the highlights of our trip and provide comparisons of Hawaiʻian geology, the Midcontinent rift, and Iceland (?). Featured highlights will include historical eruptive features from the eruptions of Kohala (100,000BP). Hualālai (1801 To 1802), Mauna Kea, and the 2018 Kīlauea Lower East Rift Zone eruption (Fissures 8 and 9). The presentation will also display local features such as the Koaʻe fault zone, Hōlei Pali (cliff), Mauna Ulu, the Pu‘u ‘Ō‘ō eruption lava field, Kīlauea Iki, fossil footprints in 18th century Kīlauea ash, the 2018 Kīlauea caldera collapse (Halemaʻumaʻu), and the Papakōlea Green Sand Beach. 41 �Kaua‘i Ni‘ihau Oahu Ka‘ula Molokaʻ i Maui Lanaʻi Kahoʻolaw e Hawai‘i Figure 1: Topographic relief and bathymetric map of the nine Hawai‘ian Islands. Modified from Eakins et al., 2003. Figure 2: Daily ILSG 2020 Field Trip routes, Island of Hawai‘i. Modified from Hazlett and Hyndman, 1996. 42 �Iceland MCR Figure 3: Location of the Hawai‘ian Islands with respect to world tectonic plates, the MCR, and Iceland. Modified after Tilling et al., 2010. REFERENCES Barker, D.S. 1983. Igneous Rocks; Prentis-Hall Inc., 417p. Best, Myron, G. 2003. Igneous and Metamorphic Petrology; Blackwell Science Ltd., 729p. Eakins, Barry W.; Robinson, Joel E.; Kanamatsu, Toshiya; Naka, Jiro; Smith, John R.; Takahashi, Eiichi; Clague, David A. 2003. Hawaii's volcanoes revealed: U.S. Geological Survey Geologic Investigations Series Map I-2809, 1 plate, https://pubs.usgs.gov/imap/2809/. Hazlett, R.W. and Hyndman, D.W. 1996. Roadside Geology of Hawaiʻi; Mountain Press Publishing Company, MT, 308p. Lockwood, Jock P. and Hazlett, Richard W. 2010. Volcanoes: Global Perspectives; Wiley-Blackwell, 536p. 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. 43 �Chemostratigraphy of the western Schreiber-Hemlo Greenstone Belt, Results and Regional Implications MAGNUS, Seamus Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5 Canada seamus.magnus@ontario.ca The Neoarchean Schreiber–Hemlo greenstone belt is located within the Wawa–Abitibi terrane of the Superior Province. The western (Schreiber) and eastern (Hemlo) parts of the belt are separated by the cross-cutting Mesoproterozoic Coldwell Alkalic Complex. Recent mapping by the Ontario Geological Survey in the western part of the belt was conducted to produce an updated genetic model for the western part of the greenstone belt and to better understand its relationship with the rocks in the Hemlo area and surrounding greenstone belts (Hastie and Magnus, 2021a; Magnus 2021a, 2019a, 2017). New geochemical and geochronological data (Hastie and Magnus, 2021b; Magnus 2021b, 2019b, 2018) have changed the model presented by the author in a field trip guide for the 2019 ILSG conference (Magnus, 2019c). The supracrustal rocks in the western Schreiber–Hemlo greenstone belt are arranged in 4 depositional packages composed of metavolcanic and metasedimentary rocks with distinct petrographic and geochemical characteristics. Chemical and clastic metasedimentary rocks mark major disconformities between the depositional packages. For the sake of brevity, the clastic and chemical metasedimentary rocks of the greenstone belt will not be discussed in this presentation. The oldest recognized supracrustal rocks are exposed south of the Town of Schreiber, west of the Terrace Bay Pluton. In this package, predominantly tholeiitic mafic flows are interbedded with plagioclase porphyritic felsic flows and volcaniclastic breccias, wackes and mudstones. One mafic flow has major and trace element concentrations that are transitional between tholeiitic and calc-alkalic mafic volcanic rocks. Felsic volcaniclastic rocks near the top of this package are crosscut by a quartz porphyritic felsic intrusive rock west of Worthington Bay, dated at 2722.6 ±1.1 Ma by zircon (Kamo, 2019), which constrains this depositional package to >2720 Ma. Quartz porphyritic felsic volcanic and volcaniclastic rocks with U-Pb zircon ages of circa 2720 Ma occur along the entire length of the greenstone belt and along the shores of Lake Superior in Tuuri and Walsh townships (Hastie and Magnus, 2021; Magnus, 2021a, 2019a, 2017). These felsic rocks are closely associated with mafic flows that commonly contain an abundance of both equant plagioclase phenocrysts and coarse amygdules. These flows contain major and trace element concentrations that are transitional between tholeiitic and calc-alkalic volcanic rocks. Unlike the single flow of “transitional” mafic rock in the older package, the transitional mafic rocks in this package are the overwhelming majority. No contacts between the >2720 Ma and the circa 2720 Ma depositional packages have been observed. Both packages are disconformably overlain by an extensive package of tholeiitic mafic flows with distinct variolitic textures and trace element geochemistry consistent with a back-arcbasin volcanic environment. These rocks are cross-cut by quartz and feldspar porphyritic felsic dikes dated 2698.1±4.5 Ma (Sutcliffe and Davis, 2019), constraining their deposition to between circa 2720 Ma and circa 2700 Ma. The disconformity between the younger tholeiitic rocks and the older transitional mafic rocks is not easily distinguishable by geophysics or by field observations, especially where the rocks are deformed. The distribution of geochemical samples in this study area has helped trace this disconformity and locate cryptic folds along that contact that otherwise would not have been identified. This geochemical mapping technique may be helpful for delineating the stratigraphy in the nearby greenstone belts, it could be critical for reconciling the stratigraphy between the east 44 �and west parts of the Schreiber–Hemlo greenstone belt, and it may be applicable to other nearby greenstone belts in the western Wawa–Abitibi Terrane. In addition to the major and trace element geochemical analyses, 27 samples of mafic, intermediate and felsic metavolcanic rocks from the previously described packages and 12 samples of the varied granitoid plutons that cross-cut and surround the greenstone belt were submitted for Sm-Nd isotope analysis (Magnus, 2021b). εNd values for all of the volcanic and intrusive rocks cluster around the value of the depleted mantle at circa 2700 Ma, which supports the model that the volcanic rocks of the Wawa-Abitibi Terrane were generated in a juvenile arc magmatic system. Figure 1. Simplified geological map of the western Schreiber–Hemlo greenstone belt, highlighting the three metavolcanic packages described in this abstract References Hastie, E.C.G. and Magnus, S.J. 2021. Ontario Geological Survey, Preliminary Map P.3846, scale 1:20 000. Hastie, E.C.G. and Magnus, S.J. 2021b. Ontario Geological Survey, Miscellaneous Release—Data 382. Kamo, S.L. 2019. Internal report for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 4p. Magnus, S.J. 2017. Ontario Geological Survey, Preliminary Map P.3812, scale 1:20 000. Magnus, S.J. 2018. Ontario Geological Survey, Miscellaneous Release—Data 361. Magnus, S.J. 2019a. Ontario Geological Survey, Preliminary Map P.3826, scale 1:20 000. Magnus, S.J. 2019b. Ontario Geological Survey, Open File Report 6357, 41p. Magnus, S.J. 2019c. Ontario Geological Survey, Open File Report 6357, 41p. Magnus, S.J. 2021a. Ontario Geological Survey, Preliminary Map P.3845, scale 1:20 000. Magnus, S.J. 2021b. Ontario Geological Survey, Miscellaneous Release—Data 381. Sutcliffe, C.N. and Davis, D.W. 2019. Internal report for the Ontario Geological Survey; Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 146p. 45 �Proterozoic Geology of the Schreiber–Terrace Bay Area MAGNUS, Seamus Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, 933 Ramsey Lake Road, Sudbury, ON P3E 6B5 Canada seamus.magnus@ontario.ca The Mesoproterozoic Keweenawan Midcontinent Rift event emplaced a multitude of intrusive and volcanic igneous rocks around present-day Lake Superior. Previous work on these rocks along the north shore of Lake Superior has mainly focused on the Thunder Bay, Lake Nipigon and Marathon areas. Recent mapping by the Ontario Geological Survey along the north shore in the Schreiber–Terrace Bay area was conducted, in part, to bridge this gap (Hastie and Magnus, 2021a; Magnus, 2017, 2019a, 2021a). Dikes with orientations and geochemical and geophysical signatures consistent with known Paleoproterozoic dike swarms, including the Matachewan, Biscotasing and Marathon dike swarms, have been identified in the Schreiber–Terrace Bay area. Similarly, dikes with other orientations and geochemical and geophysical signatures, assumed to be Mesoproterozoic and belong to the Midcontinent Rift event. A total of 156 samples of Mesoproterozoic diabase and lamprophyre dikes were submitted for whole rock geochemical analysis throughout this project (Hastie and Magnus, 2021b; Magnus, 2018, 2021b). The results of mapping and geochemical analysis suggest potential correlation between several smaller populations of dikes and known Keweenawan units. East-northeast trending dikes in the study area are parallel to, along strike with (separated by 200 km), and share similar wholerock geochemistry with the Pigeon River dikes of the Thunder Bay area (Cundari et al. 2021). Dikes of varied orientation share similar whole-rock geochemistry with Osler Group I and Group II volcanic rocks (Hollings et al., 2007). A newly recognized, circular intrusion north of the town of Terrace Bay, dated 1107.9±1.4 Ma by baddeleyite (Kamo, 2019), also shares similar geochemistry with Osler Group II volcanic rocks. The results of mapping and geochemical analysis have revealed one new and significant population of diabase dikes. These west-northwest trending dikes with reversely polarized aeromagnetic signature cross-cut the entire study area. In the east, they were observed crosscutting, and being cross-cut by, syenite at the western edge of the circa 1108 Ma Coldwell Alkalic Complex, which constrains them to emplacement at around the same time. The dikes consist of an ophitic to sub-ophitic textured olivine gabbro composed of augite, plagioclase and generally fayalitic olivine (average composition Fo30, as low as Fo10, up to Fo70), which is an expected composition of olivine for an alkalic rock. The dikes have alkalic major element concentrations and enriched, “ocean island basalt”-like trace element concentrations and chondritic to slightly depleted Sm-Nd and Sr-Rb isotopic values. There are two small subsets of these dikes – one that is plagioclase megacrystic and is geochemically indistinguishable from the main population; and one that is heavily depleted in calcium, strontium, barium and phosphorous, consistent with apatite fractionation, and which displays a “trachytic” texture with abundant plagioclase laths. South of the town of Schreiber a parallel set of alkalic dikes with normally-polarized aeromagnetic signature cross-cut the peninsula between the Schreiber Channel and Worthington Bay. Geophysical data, air photos and previous maps indicate that these dikes continue along strike westward, cross-cutting the islands in Schreiber Channel and south of Nipigon Bay. A targeted 46 �sampling program in these islands would be helpful to study the extent of these normally-polarized alkalic dikes. The petrographic and geochemical characteristics of the alkalic diabase are similar to the Wolfcamp Lake alkalic basalts that overlie the Coldwell Alkalic Complex (Davis et al., 2017), the Geordie Lake gabbro, which cross-cuts the complex (Meghji, 2016), and a subset of dikes that crop out south of the complex in Pukaswka National Park (Cundari et al., 2021). Altogether, the mafic alkalic magmatism in the northeast part of Lake Superior represents an emergent and potentially important part of the Keweenawan geological history (Good et al., in review). References Cundari, R.M., Puumala, M.A., Smyk, M.C. and Hollings, P. 2021. New and compiled whole-rock geochemical and isotope data of Midcontinent Rift–related rocks, Thunder Bay area; Ontario Geological Survey, Miscellaneous Release—Data 308 – Revised. 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. Hastie, E.C.G. and Magnus, S.J. 2021a. Precambrian geology of Strey Township, northwestern Ontario; Ontario Geological Survey, Preliminary Map P.3846, scale 1:20 000. Hastie, E.C.G. and Magnus, S.J. 2021b. Geological, geochemical and petrographic data from Strey Township, western Schreiber–Hemlo greenstone belt, Wawa–Abitibi terrane, Superior Province, northwestern Ontario; Ontario Geological Survey, Miscellaneous Release—Data 382. Hollings, P., Fralick, P and Cousens, B. 2007. Early history of the Midcontinent Rift inferred from geochemistry and sedimentology of the Mesoproterozoic Osler Group, northwestern Ontario; Canadian Journal of Earth Sciences, v.44, p.389-412. Good, D., Hollings, P., Dunning, G., Epstein, R., McBride, J., Jedemann, A., Magnus, S., Bohay, T. and Shore, G. in review. A new model for the Coldwell Complex and associated dykes in the Midcontinent Rift, Canada; Journal of Petrology, 43p. Kamo, S.L. 2019. Part A: Report on U-Pb ID-TIMS geochronology for the Ontario Geological Survey: Bedrock Mapping Projects, Ontario, Year 4: 2018-2019, internal report for the Ontario Geological Survey, Jack Satterly Geochronology Laboratory, University of Toronto, Toronto, Ontario, 26p. Magnus, S.J. 2017. Precambrian geology of Tuuri and Walsh townships, northwestern Ontario; Ontario Geological Survey, Preliminary Map P.3812, scale 1:20 000. Magnus, S.J. 2018. Geological, geochemical, geophysical and petrographic data from Tuuri and Walsh townships, Schreiber–Hemlo greenstone belt, Wawa–Abitibi terrane, Superior Province; Ontario Geological Survey, Miscellaneous Release—Data 361. Magnus, S.J. 2019a. Precambrian geology, Syine Township; Ontario Geological Survey, Preliminary Map P.3826, scale 1:20 000. Magnus, S.J. 2019b. Geological, geochemical and petrographic data from Syine township, Western Schreiber–Hemlo greenstone belt, Wawa–Abitibi terrane, Superior Province, Northwestern Ontario; Ontario Geological Survey, Miscellaneous Release—Data 375. Magnus, S.J. 2021a. Precambrian geology of Priske township; Ontario Geological Survey, Preliminary Map P.3845, scale 1:20 000. Magnus, S.J. 2021b. Geological, geochemical and petrographic data from Priske Township and Nd, Sm and Sr isotopic data from Priske, Strey, Syine, Tuuri and Walsh townships, Western Schreiber–Hemlo Greenstone Belt, Wawa–Abitibi Terrane, Superior Province, Northwestern Ontario; Ontario Geological Survey, Miscellaneous Release—Data 381. Meghji, I. 2016. The character and distribution of Cu-PGE mineralization at the Geordie Lake Deposit within the Coldwell Complex, Ontario. Unpublished M.Sc. thesis, University of Western, Ontario, London, Canada, 336 p. 47 �Structural Analysis and Interpretation of Deformation along the Keweenaw Fault System West of Lake Gratiot, Keweenaw County, Michigan MUELLER, S.A., DEGRAFF, J.M., and LIZZADRO-MCPHERSON D.J. Michigan Technological University, Houghton, MI 49931 The Keweenaw fault extends along the southern margin of the Midcontinent Rift System from northwest Wisconsin to near Keweenaw Point in Michigan, making it longest known fault associated with the rift. Reverse movement on the NW-dipping fault has thrust Portage Lake Volcanics (PLV, 1.1 Ga) over younger, mostly flat-lying Jacobsville Sandstone (JS) (Fig. 1), regionally tilting PLV strata northwestward on the southeast limb of the Lake Superior syncline (1). Published geologic maps from the 1950s (2-4) depict the fault as a single sinuous trace parallel to strike of hanging-wall PLV strata that are juxtaposed against JS nearly everywhere. Published cross-sections generally show hanging-wall strata facing northwest with a simple listric geometry and footwall strata tilted steeply southeast near the fault. The maps and crosssections show a few complications along the fault trace (anomalous sinuosity) and in its hanging-wall (lesser faults, folds), hints of deformation complexity that may reveal the nature and cause of faulting. Midway along the Keweenaw Peninsula, published maps show a prominent reentrant of JS into PLV along the fault and hint at a major splay in the Keweenaw fault’s hanging wall (3-4). New mapping near Lake Gratiot and Lac La Belle, combined with drill core logs, suggests that the postulated splay in the Keweenaw fault’s hanging wall is instead the main fault of an orderly fault system (Fig. 2). The Keweenaw fault system here consists of: 1) segments striking east-northeast with steep northerly dip (main trend), 2) segments striking east-southeast also dipping steeply north, and 3) segments striking north-northeast with moderate-to-shallow westerly dip. Members of these fault sets define a multistranded fault system and several, large, PLV blocks southeast of the main fault. Set 2 faults have a left-stepping arrangement similar to analogous faults mapped in a previous project to the east along Bête Grise Bay (5). Mapping also has revealed folding in hanging-wall Portage Lake Volcanics and has resolved fold geometry in footwall Jacobsville Sandstone. Fold axes are generally subparallel to adjacent faults and therefore are probably related to fault movement. A Set 3 fault crossing Bruneau Creek is associated with a faulted anticline in hanging-wall PLV and multiple anticlines and synclines in the footwall JS, indicating significant shortening across a west-dipping thrust fault. A Set 2 fault crossing a Snake Creek tributary has an anticline-syncline pair in hanging-wall PLV and a single asymmetric syncline in footwall JS, resulting from mostly right-lateral strike slip and lesser north-side-up reverse slip on a steeply dipping fault. Shortening across this ESE-trending fault is relatively minor because of its inferred steep dip and dominant strike slip motion. The pattern and relationships of faults and folds suggest a fault system dominated by dextral shear rather than by reverse movement as in the conventional model. Indicators of slip direction and sense measured on 55 small faults demonstrate that the fault system here is in fact dominated by strike slip, with a 2.5:1 ratio of strike-to-dip slip. Inversion of fault-slip data yields a maximum shortening direction of 285°105° during faulting, which implies overall dextral shear on the fault system and differences in slip kinematics between fault sets. Set 1 faults defining the overall trend of the Keweenaw fault system in this area have strike slip > dip slip; Set 2 faults with ESE-trend have strike slip >> dip slip; Set 3 faults with N to NE trends have dip slip >> strike slip. New mapping and structural analyses in this area have revealed a multistranded Keweenaw fault system that is transpressional in nature, dominated by dextral strike slip, and has lesser reverse slip with north side up. Fault-bounded blocks with ENE-oriented long dimensions generally moved eastward along set 1 and 2 faults, as thrusting of Portage Lake Volcanics over Jacobsville Sandstone occurred on set 3 faults. The estimated WNW-ESE maximum shortening direction associated with fault movement strongly suggests that the Grenville Orogeny was primarily responsible for movement of the Keweenaw fault system, with possible reactivation occurring during the Appalachian Orogeny. 48 �Acknowledgements: We appreciate primary funding by the USGS EDMAP program, sponsorship by the Michigan Geological Survey, a Michigan Space Grant Consortium award to the lead author, and a generous donation by the Keweenaw Community Forest Company. Figure 1 (left): Keweenaw Peninsula where Portage Lake Volcanics are thrust over Jacobsville Sandstone. Green rectangle near middle of peninsula marks area of Figure 2. (adapted from 1). Figure 2 (below): Study area along Keweenaw fault between Mohawk and Bête Grise Bay. Main units: PLV mafic = greens; PLV felsic = reds; JS = yellow. Red line marks newly interpreted main trace of the Keweenaw fault system. Other major faults as black solid and dashed lines. 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 Delaware Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-51, scale 1:24,000. Davidson, E.S., Espenshade, G.H., White, W.S. and Wright, J.C., 1955, Bedrock Geology of the Mohawk Quadrangle, Michigan: U.S. Geol. Survey, Washington, D.C., Geologic Quadrangle Map GQ-54, scale 1:24,000. Wright, J.C. and Cornwall, H.R., 1954, Bedrock Geology of the Bruneau Creek Quadrangle, Michigan: U.S. Geological Survey, Washington, D.C., Geologic Quadrangle Map GQ-35, scale 1:24,000. DeGraff, J.M., Tyrrell, C.W., and Hubbell, G.E., 2018, Keweenaw Fault Geometry, Secondary Structures, and Slip Kinematics along the Bête Grise Bay Shoreline: U.S. Geological Survey, Final Technical Report, 21 p. 49 �U-Pb geochronology and zircon trace-element geochemistry from granitoid plutons in the Neoarchean Sturgeon Lake greenstone belt, Ontario, Canada NELSON, Trevor J.1, JOHNSON, Rory M.1, LODGE, Robert W.D.1, MA, Chong2, MARSH, Jeffery H.2 1 Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI 54701 USA Metal Earth, Mineral Exploration Research Centre, Harquail School of Earth Sciences, Laurentian University, Sudbury, Ontario, Canada P3E 2C6 2 The Sturgeon Lake granite-greenstone belt is in the eastern portion of the Western Wabigoon subprovince of the Superior province and is bordered to the north, east, and south by the Winnipeg River terrane. The belt is comprised of felsic to mafic volcanic successions, synorogenic sedimentary assemblages, and calc-alkaline to alkalic plutons (Sanborn-Barrie and Skulski, 2005; Lodge et al., 2019). The emplacement of granitoid plutons in the Sturgeon greenstone belt can provide a more complete magmatic and deformational history than volcanic assemblages and sedimentary rocks. The origin and timing of their melts, mantle-crust interaction, and subsequent magmatic-hydrothermal activity record important tectonic events throughout the petrogenetic history of Sturgeon Lake greenstone belt and the western Superior Province. The primary objective of this project is to better constrain the magmatic and tectonic history of the Sturgeon Lake greenstone belt through U-Pb geochronology and trace element geochemistry of zircons from the granitoid plutons. Previous mapping in the Sturgeon Lake greenstone belt (Sanborn-Barrie and Skulski, 2005) constrained spatial and temporal relationships between plutonic rock and supracrustal assemblages largely by describing the geometry of plutons through aeromagnetic surveys and their relationship with regional foliations in the surrounding rocks. However, geochronologic data was largely constrained to the volcanic and sedimentary assemblages and the timing of plutonic events were mostly estimated. Plutons were broadly characterized into three categories based on these relationships: synvolcanic, early- to syntectonic, and late- to post-tectonic (Sanborn-Barrie and Skulski, 2005). This study obtained representative samples from the different plutons along the Metal Earth Sturgeon transect. After petrographic and whole-rock geochemical characterization (Nelson et al., 2020), zircons were separated for further analyses. Zircon grains were analyzed using Laser Ablation Split Stream Inductively Coupled Plasma Mass Spectrometry (LASS-ICP-MS) at Laurentian University for simultaneous collection of U-Pb isotopic abundances and trace-element geochemistry. The interpreted U-Pb ages of the sampled plutons allowed for refinement of the prior tectonic classifications. Some plutons were re-assigned to refined classifications based on the interpreted age: those that formed at the same time as the greenstone belt volcanic assemblages (syn-volcanic), those that formed during major deformation events (syn-tectonic), and those that post-date most deformation events (post-tectonic) (Figure 1). Trace-element geochemistry of the zircons not only supports the updated plutonic classifications, but also reveals a temporal evolution in the petrogenesis of granitoid plutons from anhydrous melting of mantle-derived sources to hydrous, oxidized melts that have significant interaction with the evolved crust. (Figure 1, inset). 50 �Figure 1: Regional geology of Sturgeon Lake granite-greenstone belt modified from Sanborn-Barrie and Skulski (2005) with new U-Pb ages of granitoid plutons and their timing relative to major tectonic events. Inset diagram displays U/Yb versus 207Pb/206Pb age displaying increasing role of evolved crust in the genesis of felsic melts over time. REFERENCES Lodge, RWD., Ma, C, Etienne, M, Nelson, TJ, Chandler, MD, Brock, NM (2019). Geologic Overview and petrogenetic history of the Sturgeon Lake Transect, Western Wabigoon. Summary of Field Work and Other Activities, Ontario Geological Survey Open File Report 6260. Nelson, TJ, Lodge, RWD, Ma, C (2020). Geologic overview of Sturgeon Lake greenstone belt granitoid plutons, Western Wabigoon Subprovince. Geological Society of America Abstracts with Programs, 52(6), paper 97-7. Sanborn-Barrie, M, and Skulski, T (2005) Geology, Sturgeon Lake greenstone belt, western Superior Province, Ontario; Geological Survey of Canada, Open File 1763, scale 1:100 000. 51 �3-D Modeling of the Duluth Complex from geophysical data PETERSON, D.E., BEDROSIAN, P.A. and FINN, C.A. U.S. Geological Survey, MS 973, Federal Center, Denver, CO 80225 The Mesoproterozoic Duluth Complex in northeastern Minnesota 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. The Duluth Complex is composed of a series of individual mafic and felsic intrusions emplaced 1110-1098 Ma within Paleoproterozoic sedimentary rocks of the Animikie basin and volcanic flows of the Midcontinent Rift. Prior work has included 2-D modeling and qualitative geologic interpretations of gravity and magnetic data (e.g., Chandler, 1990; Chandler and Ferderer, 1989), much of which is still preliminary (V. Chandler, written commun., 2020). Three-dimensional modeling has been limited, with only one 3-D model created using Bouguer gravity data constrained by seismic-reflection interpretations as part of a PhD thesis (Allen, 1994). Given the complex geology of the area, 3-D modeling is useful for providing a complete picture of the variable densities, susceptibilities, and electrical resistivities throughout the Duluth Complex and associated volcanic rocks as well as their depth extent beneath sedimentary cover. Models of these geophysical properties at depth enable more accurate geologic mapping in the subsurface which can lead to an improved understanding of the formation history of the Duluth Complex. In this study, we use aeromagnetic data acquired between 1979-1991 (Chandler, 2007), Bouguer gravity data collected since 1950 (Chandler and Lively, 2019), and magnetotelluric data collected in 2019 to create new 2-D and 3-D geophysical models of the Duluth Complex constrained by seismic reflection, geologic, and rock property data. An inversion of the Bouguer gravity data for thickness of the Duluth Complex using constant densities of 3110 kg/m3 and 2670 kg/m3 for the Duluth Complex and surrounding crustal rocks, respectively, results in thicknesses ranging from ~3-28 km for the Duluth Complex and related intrusions and volcanic rocks (Figure 1A). A 3-D model of the magnetotelluric data reveals low resistivity anomalies at ~5-10 km depth below the northern margin of the Duluth Complex and below the Greenwood Lake intrusion (Figure 1B). We expect to encounter low resistivities at depth associated with the Paleoproterozoic Animikie basin, which makes up the floor of the Duluth Complex, and therefore interpret these anomalies as either the base of the complex or as fragments of Animikie sediments interfingered with igneous intrusive rocks. Finally, 3-D voxel models of density and susceptibility illuminate the subsurface distribution of rock properties below the Duluth Complex which, in combination with resistivity and thickness models, can be used to create a 3-D geologic map of this area. 52 �Figure 1. Geophysical models of the Duluth Complex. A) Modeled thickness of the Duluth Complex and associated igneous rocks (dashed black outline in A and B) resulting from constant density inversion of Bouguer gravity data. B) Depth slice through 3-D resistivity model at 5 km below the surface. GL=Greenwood Lake intrusion. 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 pp. Chandler, V. W., 1990, Geologic interpretation of gravity and magnetic data over the central part of the Duluth Complex, northeastern Minnesota: Economic Geology, v. 85, p. 816-829. Chandler, V. W., 2007, Upgrade of aeromagnetic databases and processing: Minnesota Geological Survey Open File Report OFR07_06. Chandler, V. W., and Ferderer, R. J., 1989, Copper-nickel mineralization of the Duluth Complex, Minnesota; a gravity and magnetic perspective: Economic Geology, v. 84, p. 1690-1696. Chandler, V. W., and Lively, R. S., 2019, Upgrade of the gravity database at the Minnesota Geological Survey: Retrieved from the University of Minnesota Digital Conservancy, 28 August 2020. 53 �Negative Carbon Dioxide Emissions through Enhanced Silicate Weathering and the Lake Superior Region PLANAVSKY, Noah Department of Earth and Planetary Sciences, Yale University, 210 Whitney Avenue, New Haven, CT, USA Carbon capture is now widely recognized as necessary to stabilize atmospheric carbon dioxide to keep global warming below 1.5˚C, a temperature threshold mandated by the 2015 Paris Agreement. The Intergovernmental Panel on Climate Change (IPCC) projects significant degradation of livelihoods, food security, and water supply, if this threshold is exceeded. There are multiple proposed mitigation approaches that could curb climate catastrophe. Although no one carbon capture approach is perfect, all are on the table by necessity. Each has specific capture potential in terms of uptake rate, storage capacity, and storage longevity. I will give an overview of work determining whether adding basalt to agricultural lands is an effective, low-risk carbon capture strategy—and make a case that Lake Superior region is great place to implement this means of climate change mitigation. The idea behind this carbon dioxide removal (CDR) strategy is simple. Carbon dioxide in the atmosphere chemically reacts with silicates minerals—and reacts especially rapidly with mafic minerals. The byproducts of this reaction lead to marine carbonate precipitation, which transfers carbon from the atmosphere into the rock record. This is, on geologic timescales, how Earth has sequestered almost all the carbon continuously sourced to the atmosphere from Earth’s interior (e.g., volcanic carbon dioxide outgassing). In principle, carbon capture through mineral weathering involves enhancing the rate of a process that the Earth does naturally. In practice, the overriding controls on this process are complex. Carbon capture rates are highly variable and dependent on local conditions and implementation strategy. Nonetheless it is clear that capture through weathering can be greatly accelerated by milling of mafic minerals to increase its surface exposure. This approach holds great potential and is certain to be carbon negative. IPCC estimates are that the agricultural sector contributes about 25 percent of total greenhouse gas emissions worldwide. I will present modeling work and empirical studies that suggests addition of mafic rocks to agricultural lands has the potential to simultaneously offset a significant component of this anthropogenic greenhouse gas source and increase agricultural yields and pest resistance. If subsidized, this process could provide a stable source of income for farmers increasingly vulnerable to climate extremes and promote regional economic growth. The Lake Superior region is the ideal area to carry out this carbon capture process and widespread implementation could revitalize the economy of Lake Superior region towns that are (or were formerly) heavily dependent on mining. Compilations of the extensive knowledge of Lake Superior Region geology are essential to support that enhanced silicate weathering is a viable means climate change mitigation, sway policy makers, and foster investment in case studies of this process. 54 �A newly discovered orbicular occurrence within the Good Hope carbonatite, north of Marathon, Ontario PRICE, Rebecca, ZUREVINSKI, Shannon, and MITCHELL, Roger H. Department of Geology, Lakehead University, 955 Oliver Road, Thunder Bay, ON P7B 5E1 Canada The Good Hope carbonatite is emplaced within Archean gneisses of the Wawa Subprovince. This occurrence is located along the northwest margin of the Prairie Lake carbonatite complex (1160 Ma, Wu et al. 2017), although their genetic relationship remains speculative (Mitchell et al. 2020). The Good Hope carbonatite is a niobium occurrence with the niobium primarily occurring within pyrochlore-group minerals, minor niobium in ferrocolumbite, and consists of multiple generations of carbonatite that include calcite-, dolomite-, ferrodolomite-, and quartz fluorite-carbonatite (Mitchell et al. 2020). The purpose of this study is to characterize the orbicular occurrence and suggest a possible petrogenesis for this rare textural occurrence. The orbicular occurrence was first identified from the drill core, where it occurs as narrow bands or possible clasts within the carbonatite groundmass. These orbicules are mm-scale, oblate to spherical in shape with variable degrees of monomineralic shell development ranging from welldeveloped shells to spherical aggregates of the minerals which define the orbs (Figure 1). The orbicules consist of biotite, Ti-bearing aegerine, and magnetite with ilmenite lamellae within an allotriomorphic calcite carbonatite groundmass that contains some Fe-dolomite. The orbicules commonly consist of a ‘core’ of magnetite + biotite followed by a shell of aegirine and one or more outer shells of monomineralic biotite and/or magnetite. Within these orbicules there are variable amounts of calcite and Fe-dolomite, which can occur both as monomineralic shells, commonly associated with the outer shells, as well as entrained within the ‘core’ of the orbicules. This type of texture has been observed within other carbonatites (eg. Haggerty and Fung, 2006), but is more commonly observed within granitic and dioritic compositions (eg. Smillie and Turnbull, 2013). It is largely accepted that the orbicular textures are of magmatic origin, however a variety of genetic models exist, due to the unique nature of each occurrence. In general, these models broadly fit into three groups with orbicular formation controlled by (1) liquid immiscibility (eg. Ai et al. 2020) (2) superheating of the magma that lowers the nucleation rate and increases the growth rate (eg. Lindh and Näsström, 2006) or (3) quenched recrystallization following the injection of magma (eg. Zurevinski and Mitchell, 2015). Orbicular textures within carbonatites are commonly modelled by liquid immiscibility, given this model lends itself to the orbicules having a different composition than the groundmass. The second group has variations in the compositions of the shells and groundmass that lead to variability between the models in the source of the ‘orbicular magma’, the cause of the increased temperature within the magma, and the process of crystallization of the orbicule shells. The third group also has variability, with the orbicule and groundmass being either the same composition (Zurevinski and Mitchell, 2015) or different compositions (Zhang and Lee, 2020) due to the composition of the initial magma and the latter injected magma. The Prairie Lake orbicular occurrence mimics ‘contact metamorphism-type’ recrystallization of the orbicules within a compositionally identical groundmass (Zurevinski and Mitchell, 2015). The Good Hope occurrence differs significantly from the Prairie Lake occurrence; the composition of the orbicules and the groundmass are distinctly different, the orbicules are variably oblate indicating soft-shell deformation, and they do not exhibit quenched or recrystallization textures. 55 �Figure 1: Scan of the standard polished petrographic thin section of the orbicular occurrence of the Good Hope carbonatite, highlighting the predominantly oblate, mm-sized orbicules. References Ai, J., Lu, X., Li, Z., and Wu, Y. 2020. Genesis of the graphite orbicules in the Huangyangshan graphite deposit, Xinjiang, China: Evidence from geochemical, isotopic and fluid inclusion data; in Ore Geology Reviews, 122: 103505 Haggerty, S. E., and Fung, A. 2006. Orbicular oxides in carbonatitic kimberlites; in American Mineralogist, 91: 1461-1472 Lindh, A., and Näsström, H. 2006. Crystallization of orbicular rocks exemplified by the Slättemossa occurrence, southeastern Sweden; in Geological Magazine, 143: 713-722 Mitchell, R.H., Wahl, R., and Cohen, A. 2020. Mineralogy and genesis of pyrochlore apatitite from the Good Hope Carbonatite, Ontario: A potential niobium deposit; in Mineralogical magazine, 84: 8191 Smillie, R. W., and Turnbull, R. E. 2013. Field and petrographical insight into the formation of orbicular granitoids from the Bonney Pluton, southern Victoria Land, Antarctica; in Geological Magazine, 151 (3): 534-549 Wu, F., Mitchell, R.H., Li, Q., Zhang, C., and Yang, Y. 2017. Emplacement age and isotopic composition of the Prairie Lake carbonatite complex, Northwestern Ontario, Canada; in Geological Magazine, 154: 217-236 Zhang, J., and Lee, C. A. 2020. Disequilibrium crystallization and rapid crystal growth: acase study of orbicular granitoids of magmatic origin; in International Geology Review 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; in Lithos, 239: 234-244 56 �Preliminary pXRF results from Precambrian rocks of northern Minnesota PRUE, Ann Marie1 and BRENGMAN, Latisha1 1 Department of Earth and Environmental Sciences, University of Minnesota-Duluth, 229 Heller Hall, 1114 Kirby Drive, Duluth, MN 55812 Precambrian iron-rich chemical sedimentary rocks are an important archive of early seawater chemistry (Konhauser et al., 2017). Interpreting the origin of these iron-rich rocks is complicated because they can form at the intersection of sedimentary and hydrothermal activity. One tool that can be used to identify signatures of past seawater and hydrothermal fluids are major-, trace-, and rare-earth-elements. Drawbacks to laboratory-based geochemical approaches include extensive sample preparation and destruction of the material. Advances in XRF technology have led to the development of portable x-ray fluorescence (pXRF) instruments capable of producing laboratory-quality element analyses in the field when used with a custombuilt, rock-type-specific calibration (Steiner et al., 2017; Al-Musawi and Kaczmarek, 2020). With the ability to create a custom calibration, it is possible to measure major- and trace- elements in the field or in drill core. One goal of this project is to create a customized, rock-type specific calibration for the Bruker 5g Tracer pXRF instrument to analyze iron-rich rocks that outcrop near Soudan, MN. Field and drill-core geochemical analysis can help identify units of unknown origin, while aiding in the identification of key samples for additional destructive analyses. To create the custom iron formation calibration for the pXRF instrument, we first need an internal reference material. We analyzed select iron formation samples from the Soudan iron formation for a suite of 46 major, trace-, and rare-earth-elements using inductively coupled plasma optical emission spectrometry (ICPOES) and inductively coupled plasma mass spectrometry (ICP-MS) methods following the procedures outlined in Brengman et al. (2020). We can directly compare elements analyzed using ICP-OES and ICPMS to the pXRF data (Figure 1A). The pXRF instrument was set up in tabletop working mode, and powdered samples were analyzed in standard XRF mount cups. To test the effect of pressing the sample into the mount cup, we performed a comparison test on one sample (PL-19-3c; Figure 1B) using the internal mudrock calibration. First, the un-pressed powder was run, then the pressed powder was run 1.5 hr. later, with five minutes between each analysis to allow for cooling. Measurements for the element silicon (Si) in the three un-pressed powers ranged from 23.075 wt. % (±0.083) to 23.197 (±0.083) wt. % (Figure 1B). The same three powders were pressed, and measurements were repeated using the same analytical setup. Measurements for the element silicon (Si) in the three samples (now pressed powders) ranged from 23.733 (± 0.084) to 23.891 (±0.0847) wt. %. We observe a difference of ~0.7 wt. % between Si measured in the un-pressed and pressed powders and recommend pressing powders to ensure uniform measurements. Previous studies identified a “warm-up” period in which the pXRF instrument took roughly 1 hour to stabilize. To test whether the Bruker Tracer 5g instrument also required this “warm-up” period, we conducted continuous acquisition of one sample (V-19-05A) for 21 consecutive runs, with 5 minute-spacing between analyses. At this point, a second “split” of the same initial powder (V-19-05B) was analyzed in the same way to determine if there was a measurable difference between the two splits of the same sample (V-19-05A and V-19-05B). Data for Si (wt. %) and Ca (wt. %) from sample V-19-05A and B are reported in Figure 1C, and D. We identified a steady increase (Figure 1C, D) until the instrument stabilized (run #110 for Si and #102 for Ca). We determined the instrument requires a significant “warm-up” period to ensure consistency between analytical runs. Future work will focus on comparison of ICP-OES and ICPMS data to pXRF data to build the custom-calibration for powdered, and non-powdered sample materials. 57 �Figure 1: Preliminary geochemical data from iron-rich rocks near Soudan, MN. (A) Rare-earth-element data for iron formation (V-19-05) and regional greywacke (PL-19-3c). (B) Initial test run of pressed versus un-pressed powders using the pXRF instrument. (C, D) Consecutive pXRF analyses of iron formation sample V-19-05 to determine the duration of the “warm-up” period for the instrument. REFERENCES Al-Musawi, M., Kaczmarek, S., 2020. A new carbonate-specific quantification procedure for determining elemental concentrations from portable energy-dispersive X-ray fluorescence (PXRF) data. Applied Geochemistry 113. Brengman, L.A., Fedo, C.M., Whitehouse, M.J., Jabeen, I., Banerjee, N.R. 2020. Textural, geochemical, and isotopic data from silicified rocks and associated chemical sedimentary rocks in the ~ 2.7 Ga Abitibi greenstone belt, Canada: Insight into the role of silicification, Precambrian Research, 351. Konhauser, K.O., Planavsky, N.J., Hardisty, D.S., Robbins, L.J., Warchola, T.J., Haugaard, R., Lalonde, S. V., Partin, C.A., Oonk, P.B.H., Tsikos, H., Lyons, T.W., Bekker, A., Johnson, C.M., 2017. Iron formations: A global record of Neoarchaean to Palaeoproterozoic environmental history. Earth-Science Reviews 172, 140–177. Steiner, A. E., Conrey, R. M., Wolff, J. A. 2017. PXRF calibrations for volcanic rocks and the application of in-field analysis to the geosciences. Chemical Geology, 453: 35-54. 58 �Critical Minerals Exploration and Development Potential in Ontario PUUMALA, Mark and CUNDARI, Robert Ontario Geological Survey, Ministry of Energy, Northern Development and Mines, Resident Geologist Program, Suite B002, 435 James Street South, Thunder Bay, Ontario, P7E 6S7 On March 10, 2021, Ontario released a discussion paper that outlines the province’s proposal to develop a critical minerals strategy (Ministry of Energy, Northern Development and Mines 2021). The discussion paper is focused on: 1. Supporting partnership opportunities with Indigenous peoples; 2. Finalizing an Ontario critical minerals list; 3. Enhancing investment in mineral exploration and development; 4. Regulatory and policy reform; and, 5. Supply chain and manufacturing opportunities. Public consultation is underway, with the strategy scheduled for release by the end of 2021. While there is no universal definition of critical minerals and various jurisdictions define them differently, the term generally applies to minerals that have specific industrial, technological and strategic applications for which there are few viable substitutions. These minerals are also at higher supply risk due to geopolitical considerations and market demand. Ontario is uniquely positioned to meet rising global demand for critical minerals and can attract potential investment opportunities by creating a provincial critical minerals list to guide mineral exploration and development. An Ontario critical minerals list will be of interest to jurisdictions that are seeking to secure a reliable supply of raw materials for their own domestic markets, such as the United States, the European Union, Japan and South Korea. Ontario has created a draft list of critical minerals that includes commodities that are currently being produced commercially at existing mines, commodities that have a reasonable prospect of being developed in the near term (i.e., advanced projects that could produce within 5 years), commodities that have demonstrated exploration potential, and commodities that were not originally mined in Ontario, but are being smelted or refined here. The entire list is provided in Table 1, with the critical minerals that are in common with the United States draft list (Fortier et al. 2018) indicated in bold type. Ontario’s current critical mineral production is derived from magmatic copper-nickelcobalt-platinum group element (Cu-Ni-Co-PGE) deposits (8 Sudbury mines and Lac des Iles) and volcanogenic massive sulphide zinc-copper (Kidd Creek) deposits. The Sudbury mines also produce selenium and tellurium as by-products, while indium is a by-product at Kidd Creek. Due to current high prices for platinum group elements (most notably palladium) and battery metals (which include nickel and cobalt), current exploration interest in magmatic Cu-Ni-Co-PGE deposits is strong in Ontario, with numerous active projects. These include several in the Midcontinent Rift area that are primarily targeting PGEs, including the Clean Air Metals Inc. Thunder Bay North project, where the Escape Lake deposit is currently being delineated, and Generation Mining Limited’s Marathon project, which is undergoing an Environmental Assessment to support future mine development. Another notable advanced-stage magmatic sulphide project is Canada Nickel Company’s Crawford Nickel-Cobalt project near Timmins, where a high-tonnage low-grade resource has been defined in an Archean ultramafic intrusion. Strong interest in battery metals has also fuelled significant exploration interest in lithium and cobalt in recent years. Numerous areas in the Superior Province of northwestern Ontario are known to host lithium-bearing rare element pegmatites. These deposits are associated with fertile peraluminous granites (Breaks et al. 2003), with notable examples currently moving toward 59 �development in the Pakeagama Lake (Frontier Lithium Inc.), Separation Rapids (Avalon Advanced Materials Inc.) and Georgia Lake (Rock Tech Lithium Inc.) areas. Recent cobalt exploration has largely been focussed on the Proterozoic 5-element vein systems of the Cobalt Embayment, where First Cobalt Corp. is developing a refinery to produce battery-grade cobalt sulphate. Some cobalt exploration has also occurred in the Thunder Bay area, where Honey Badger Exploration Inc. recently discovered a new style of disseminated cobalt mineralization in close proximity to silver-bearing veins at the Beaver Mine, near the contact between the Rove Formation and a diabase sill (Puumala et al. 2019). Another previously unrecognized style of cobalt-coppernickel mineralization that has been described as a possible skarn or IOCG occurrence is hosted in altered calcareous sedimentary rocks of the Rossport Formation north of Thunder Bay, near Disraeli Lake (Ontario Geological Survey 2019). Both deposit types present new critical mineral research and exploration opportunities in Ontario. Table 1. Ontario’s draft list of critical minerals. Asterisk indicates commodity also processed in Ontario. Critical minerals listed in bold type are also found on the United States draft list (Fortier et al. 2018). Producing Cobalt * Copper * Indium Nickel * Platinum Group Elements * Selenium * Tellurium * Zinc Advanced Projects Barite Chromite Graphite Lithium Magnesium Niobium Exploration Potential Antimony Beryllium Bismuth Cesium Fluorspar Manganese Molybdenum Phosphate Rare Earth Elements Tantalum Tin Titanium Tungsten Vanadium Zirconium Processing Only Uranium References 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. Fortier, S.M, Nasser, N.T., Lederer, G.W., Brainard, J., Gambogi, J. and McCullough, E.A. 2018. Draft critical mineral list – summary of methodology and background information; United States Geological Survey, Open File Report 2018-1021, 15p. Ministry of Energy, Northern Development and Mines 2021. Critical minerals framework discussion paper; Queen’s Printer for Ontario, 32p. Ontario Geological Survey 2019. Caro Lake; Mineral Deposit Inventory Record No. MDI000000002293. Puumala, M.A., Campbell, D.A., Tuomi, R.D., Fudge, S.P., Pettigrew, T.K. and Hinz, S.L.K. 2019. Report of Activities 2018, Resident Geologist Program, Thunder Bay South Regional Resident Geologist Report: Thunder Bay South District; Ontario Geological Survey, Open File Report 6353, 109p. 60 �Environmental Control of Seawater Geochemistry in a Mesoarchean Peritidal System, Woman Lake, Superior Province 1 RAMSAY, Brittany, 1FRALICK, Philip, 2LALONDE, Stefan, 1BIELSKI, Paul, and 2PATRY, Laureline 1Department of Geology, Lakehead University, 955 Oliver Road Thunder Bay, ON P7B 5E1 Canada 2Ocean Geosciences Laboratory, European Institute for Marine Studies, Technopôle Brest, Place Nicolas Copernic, 29280 Plouzané, France The 2.857±5 Ga (this study) carbonate platform at Woman Lake, Ontario, Canada, presents a unique opportunity to fill a 130 million year knowledge gap on early carbonate sedimentology and ocean chemistry between similar platform occurrences at Steep Rock Lake (2.80Ga)(Fralick & Riding, 2015) and Red Lake (2.93Ga)(MacIntyre & Fralick (2017). Woman Lake carbonates are among the few very early and thick carbonate platforms to develop in the Mesoarchean. Field, petrographic, and geochemical investigations were performed on the limestone sequence to better understand the paleoenvironmental context of this understudied, 90-meter-thick succession. At the base of the carbonate platform, lying atop felsic subaerial Archean tuff, are stratiform stromatolites interbedded with thin beds of massive carbonate grainstone, followed by laterally linked low domal stromatolites, which gradually become larger domes, then bioherms with walled pseudocolumnar stromatolites. They are overlain by cross-stratified and parallel laminated carbonate grainstones and more pseudocolumnar stromatolites. A variety of fenestral microbialites overly this unit, including thrombolites, stromatactis-bearing low domal stromatolites, and narrow isolated columnar stromatolites. This is followed by a cyclic succession of low domal stromatolites alternating with microbial carbonate and carbonate grainstone. Three main stromatolitic morphologies exist and represent a range of low to moderate current energies from upper intertidal to subtidal environments. They are: 1) low relief stratiform to undulating stromatolites 2) laterally linked low domal and pseudocolumnar stromatolites, and 3) isolated to locally isolated domes and narrow columnar stromatolites. Evidence here supports mainly peritidal environments on a carbonate platform with fluctuating sea-level and water energies in an overall deepening succession. The diverse carbonate facies are comprised of geochemical features reminiscent of both Archean and modern signatures in shale normalized REE patterns. Trace elements indicate that the carbonates precipitated from a mixture of two different fluids: anoxic seawater that carried a positive Eu anomaly, and oxygenated waters that imparted significant negative Ce anomalies. On a microscopic scale, using LA-ICP-MS, there is less compositional contrast between carbonate phases, which indicates that dissolution and precipitation on a small spatial scale homogenized localized areas, but did not affect changes on a metric scale. Geochemical trends paired with stratigraphic depth show decameter cycles of gradual declines in Mg, Fe, Mn, Ba and Sr substitution into the calcite lattice followed by sharp increases throughout the platform’s deposition, possibly reflecting changing accommodation space effecting precipitation rate (Fig. 1). Typical Archean values for δ13C ranging from -3.83‰ to 1.30‰, with an average of 0.53‰ (±0.59, n=31) occur with Y/Ho ratios ranging from 27 to 117 and 87Sr/86Sr isotopic values from 0.700346 to 0.711313 (±0.00098 (1σ)). The observed trends suggest that the precipitating carbonates were able to record and retain the effects of an evolving water column had in the local environment. Importantly, the Woman Lake carbonate platform provides context for, and evidence of, free oxygen approximately 500 million years before the Great Oxygenation Event, during a relatively undocumented period in time. 61 �Fig. 1. Five meter moving averages of molar weight ratios normalized to Ca are plotted against stratigraphic depth. Ratios are of common carbonate group elements that substitute into various carbonate minerals. Three cycles of increasing and decreasing concentrations are evident. Decreasing trends represent times of higher substitution rates and an increase represents lower substitution rates relative to Ca (numerator). Fralick, P., & Riding, R. (2015). Steep Rock Lake: Sedimentology and geochemistry of an Archean carbonate platform. Earth-Science Reviews, 151, 132–175. https://doi.org/10.1016/j.earscirev.2015.10.006 MacIntyre, T., & Fralick, P. (2017). Sedimentology and Geochemistry of the 2930 Ma Red LakeWallace Lake Carbonate Platform, Western Superior Province, Canada. The Depositional Record, 3(2), 258–287. https://doi.org/10.1002/dep2.36 62 �Geochemistry and Petrography of Volcanic and Intrusive Rocks Hosting the Lynne Cu-ZnPb Deposit, Oneida County, WI SHORT, Shelby R., GLODOWSKI, Lillian N., and LODGE, Robert W.D. Department of Geology, University of Wisconsin-Eau Claire, Eau Claire, WI The Lynne Zn-Cu-Pb deposit was first discovered in 1990 by Noranda Exploration in Oneida County, WI and is one many volcanogenic massive sulfide (VMS) deposits in northern Wisconsin (DeMatties, 1994). These deposits formed 1.8 to 1.9 Ga when the Pembine-Wausau Terrane collided and accreted onto the Superior Craton during the Paleoproterozoic Penokean Orogeny. VMS deposits are created in extensional submarine environments that are preserved in the rock record in a variety of tectonic settings. The VMS-forming environment in Wisconsin has previously been interpreted to be back-arc rifting in a continental setting during the collision of the Archean Marshfield Terrane onto the Pembine-Wausau Terrane and the rest of the Superior Craton (Schulz and Cannon, 2007). However, studies at other deposits throughout Wisconsin have highlighted variable volcanic and tectonic settings (e.g. Jackson et al, 2016; Jacobson and Lodge, 2018). The purpose of this study is to better constrain the volcanic and tectonic setting at the Lynne deposit and improve tectonic and metallogenic models across the Penokean Orogen. Research on Penokean VMS systems depends on drill core as outcrop exposure in the region is limited due to a thick cover of glacial deposits and Paleozoic Sedimentary strata (Dematties, 1994; Adams, 1996). The Noranda drill cores from the Lynne deposit were obtained from the archives at the Natural Resources Research Institute of Duluth, Minnesota and re-logged and sampled at an off-campus lab associated with UW-Eau Claire. Core samples from the leastaltered felsic host rocks and various intrusive rocks were then processed for petrographic and geochemical analyses. Major and trace elements were analyzed using X-ray Fluorescence at the Materials Science Center at UW-Eau Claire. Interpretations from this new geochemical data yielded an improved volcanic stratigraphy and tectonic model that describes the volcanic system forming the Lynne VMS deposit. The stratigraphy of the Lynne deposit was described by Adams (1996) based on drilling during the exploration program. The stratigraphy consists of the Upper Rhyolite, Upper Dacite, Upper Volcaniclastic (VCS), Lower Rhyolite, Lower Dacite, and Lower Volcaniclastic based on their relative position to the ore horizon (Adams, 1996). After a geochemical analysis of the felsic volcanic units was conducted, it revealed that there are no chemical distinction between the upper and lower horizons. Therefore, unit descriptions are constrained by composition and not stratigraphy. Rhyolites are crystal rich lapilli tuff to ash with plagioclase being the most abundant crystal. The VCS unit is greywackes and siltstones with some bedding and localized alteration. The Dacite unit is a crystal to crystal-lithic lapilli tuff with plagioclase crystals. The Lynne drill core also intersects felsic and mafic dykes, along with a large granophyre/granodiorite pluton at the bottom of the deposit which is geochemically indistinguishable from intersecting felsic dykes. The felsic dykes are fine grained and light to medium grey in color and contain plagioclase and quartz veins. The mafic dykes appear fine grained, dark grey in color, and contain plagioclase phenocrysts. The granodiorite/granophyre contains medium- to coarse-grained quartz and feldspars which vary between the phenocrystic granophyre near the contact to equigranular away from the contact region. The geochemical signature of the intrusive units of the Lynne deposit indicate that they are derived from two distinct 63 �sources that formed in a volcanic arc setting much like the extrusive volcanic stratigraphy. Therefore, it is likely that they are genetically related. REFERENCES Figure 1: Generalized stratigraphic column of the Lynne Deposit (from Quigley, 2016) with core pictures from representative volcanic and intrusive units. A. Upper Dacite B. Upper Rhyolite C. Upper VCS D. Felsic Dyke E. Lower VCS F. Lower Rhyolite 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, p. 161-179. DeMatties, T.A., 1994, Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview. Economic Geology, 89: 1122-1151. Jackson, N.R., Moura Merss, B.H., and Lodge, R.W.D., 2016, Lithostratigraphy and Ore Petrology of the Eisenbrey Zn-Cu-Pb Deposit, Rusk County, Wisconsin: Institute on Lake Superior Geology Proceedings, 62 nd Annual Meeting, Duluth, MN. Jacobson, R.E. and Lodge, R.W.D., 2018, Reconstructing Paleoproterozoic volcanism in northwestern Wisconsin: Geochemistry of the Flambeau Cu-Zn-Au Mine: Institute on Lake Superior Geology Proceedings, 64 th Annual Meeting, Iron Mountain, MI. Schulz, K.J. and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region. Precambrian Research 157: 4-25. Quigley, P.O., 2016, The Spectrum of Ore Deposit Types, their Alteration and Volcanic Setting in the Penokean Volcanic Belt, Great Lakes Region, USA, Colorado School of Mines, Master’s Thesis, 2016 64 �Geologic mapping identifies bedrock folds that may be significant for increased probability of arsenic detection in water wells STEWART, Esther K.1, FITZPATRICK, Billy1, and STEWART, Eric D.1 1 Wisconsin Geological & Natural History Survey, 3817 Mineral Point Road, Madison, WI 53705 Recently completed bedrock mapping of Dodge County, southeast Wisconsin, provides an improved geologic framework that informs the region’s geologic history and adds to our understanding of groundwater resources (Stewart, E.D. et al., 2021; Stewart, E.K. 2021). Dodge County is underlain by glacial sediments, Precambrian Quartzite, and Cambrian through Silurian siliciclastics and carbonates and is situated proximal to the WI Arch, a basement high that influenced Paleozoic basin development. We identify bedrock folds that likely increase the probability of arsenic (As) detection in groundwater wells. New petrographic observations constrain the As host and suggest processes that resulted in As mineralization. Mapping indicates episodic movement along the WI Arch influenced deposition of Paleozoic strata. Most significant deformation during early Ordovician time resulted in localized uplift and erosion of the dolomitic Prairie du Chien Group prior to deposition of the overlying siliciclastic Ancell Group. Continued deformation is associated with folding of the overlying Sinnipee Group, fracture development, and mineralization. Marcasite and carbonate-filled fractures in one locality along an anticlinal axis show well developed cataclasis evidencing deformation during and after formation of trace MVT occurrences. Most elevated hydraulic conductivity (K) estimates from the Ancell Group are within 2 km of mapped folds, indicating fractures associated with folds locally increases groundwater flow (Stewart, E.D. et al., 2021). Water wells completed in the Ancell Group in Dodge County have increased likelihood of detecting As close to folds, suggesting fold induced fractures promoted a focus or change in As host minerals that increased the probability of arsenic release. No As was detected in SEM-EDS analysis of oxide phases, while two sulfide As hosts were identified in mineralized core samples: (1) Small (10-20 µm), cubic, Ni-Co-Fe sulfides with up to 3 wt.% As are scattered within marcasite that infills localized fractures and sulfide cement in Cambrian Elk Mound Group sandstone . These crystals are similar in morphology, size, and chemistry to Ni-Co-Fe sulfides identified as VaesiteBravoite from trace MVT-mineralized samples in Brown and Oconto Counties reported by Luczaj et al., 2016. (2) Tiny (3-5 µm), circular, rimmed Fe-sulfides with 1.4-2.2 wt.% As are only present within thin clay rims on quartz grains in the Ancell Group sandstone , in contrast to all other sulfides observed from the same sample, which lack detectable As. Rimmed grains may be the mineralized remains of specialized microbacterial colonies that metabolically reduced dissolved aresenate and sulfate to produce As-bearing Fe-sulfide (e.g. Stolz et al., 2006). Observation of Fesulfide frambroids supports a biologic control on sulfide mineralization. High-resolution SEMBSE imaging reveals common As-poor, Fe-sulfides present as hexagonal, small to large, euhedral marcasite plates and subordinate framboidal clusters of tiny Fe-sulfide crystals . Framboids are common as isolated clusters encased within larger euhedral sulfide or as nucleation points for single grains of euhedral sulfide. Framboidal sulfide clusters are commonly interpreted as mineralized remains of sulfate-reducing microbacterial colonies (e.g. Folk, 2005), and sulfatereducing bacterial colonies are known to seed adjacent areas with reduced sulfur allowing for inorganic precipitation of larger volumes of sulfide cement (Schieber, 2002). Bacterial reduction 65 �of sulfate sourced from regionally migrating hydrothermal brines coming out of the Michigan Basin is proposed as a broader mechanism that drove precipitation of Fe-sulfides forming the bulk of mineralization within the trace MVT occurrences of Dodge County. References: Folk, R.L., 2005, Nanobacteria and the formation of framboidal pyrite: Textual evidence: Earth System Science, v. 114, no. 3, p. 369-374. Luczaj, J.A., McIntire, M.J., and Olson-Hunt, M.J., 2016, Geochemical characterization of trace MVT mineralization in Paleozoic sedimentary rocks of northeastern Wisconsin, USA: Geosciences, v. 6, no. 29. Schieber, J., 2002, The role of an organic slime matrix in the formation of pyritized burrow trails and pyrite concretions: PALAIOS, v. 17, p. 104-109. Stewart, E.D., Stewart, E.K., Bradbury, K.R., and Fitzpatrick, W., 2021, Correlating bedrock folds to higher rates of arsenic detection in groundwater, SE Wisconsin, USA, Groundwater. Stewart, E.K., 2021, Bedrock geology of Dodge County, 1:100,000-scale. Wisconsin Geological and Natural History Survey map series, plate 1. Stolz, J.F., Basu, P., Santini, J.M., and Oreland, R.S., 2006, Aresenic and selenium in microbial metabolism: Annual Reviews in Microbiology, v. 60, p.107-130. Figure 1: (A) Electron backscatter image of marcasite-cemented vertical zone within Elk Mound group. Frame is centered on an unusual cubic Fe-Ni-Co-As sulfide grain. (B) Electron backscatter image of unusual rimmed, circular As-bearing Fe-sulfides from Tonti Member of the St. Peter Formation, Ancell Group. Smaller, unrimmed grains outboard of main rimmed sulfides contain detectable As but at reduced concentration relative to main rimmed grain (<.5 wt%). Platy medium gray matrix enclosing sulfides is a thin rim of kaolinitic clays. Light gray at bottom right of image is edge of detrital quartz grain. (C) Electron backscatter image of single framboidal cluster at high magnification illustrating tiny individual grains making up structure. (D) Electron backscatter image of cluster of sulfide grains within Tonti Member of St. Peter. Note how framboidal clusters appear to be serving as nucleation points for larger, euhedral hexagonal sulfide grains. 66 �The paleogeography of Laurentia in its early years: new constraints from the Paleoproterozoic East-Central Minnesota batholith SWANSON-HYSELL, Nicholas L., AVERY, Margaret S., ZHANG, Yiming, HODGIN, Eben B. Department of Earth and Planetary Science, University of California, Berkeley BOERBOOM, Terrence, J. Minnesota Geological Survey, St. Paul, MN, USA The ca. 1.83 Ga Trans-Hudson orogeny resulted from collision of an upper plate consisting of the Hearne, Rae, and Slave provinces with a lower plate consisting of the Superior province. While the geologic record of ca. 1.83 Ga peak metamorphism within the orogen suggests that these provinces were a single amalgamated craton from this time onward, a lack of paleomagnetic poles from the Superior province following Trans-Hudson orogenesis has made this coherency difficult to test. Figure 2: Map of Laurentia showing the location of Archean provinces and younger Proterozoic crust (simplified from Whitmeyer and Karlstrom 2007)). The localities of paleomagnetic poles that constrain Laurentia’s position just after its amalgamation are shown with stars including the new pole from this study developed from the EastCentral Minnesota Batholith (ECMB). This right map shows interpreted Precambrian geology for the state of Minnesota (simplified from Jirsa et al., 2012)) including in regions covered by Phanerozoic sedimentary rocks where the bedrock is inferred from geophysical data and drill cores. We develop a high-quality paleomagnetic pole for northeast-trending diabase dikes of the post-Penokean orogen East-Central Minnesota Batholith (Holm et al., whose age we constrain to be 1779.1 ± 2.3 Ma (95% CI) with new U-Pb dates. Demagnetization and low-temperature magnetometry experiments establish dike remanence be held by low-Ti titanomagnetite. Thermochronology data constrain the intrusions to have cooled below magnetite blocking temperatures upon initial emplacement with a mild subsequent thermal history within the stable craton. The similarity of this new Superior province pole with poles from the Slave and Rae provinces establishes the coherency of Laurentia following Trans-Hudson orogenesis. This consistency supports interpretations that older discrepant 2.22 to 1.87 Ga pole positions between the provinces are the result of differential motion through mobile-lid plate tectonics. The new pole supports the NENA connection between the Laurentia and Fennoscandia cratons. The pole can be used to jointly reconstruct these cratons ca. 1780 Ma thereby strengthening the paleogeographic position of these major constituents of the hypothesized late Paleoproterozoic supercontinent Nuna. 67 �Figure 2: Paleogeographic reconstructions at five different times in the Paleoproterozoic and the position of the provinces at present. Paleomagnetic poles within 20 Myr of the given time (10 Myr for 1888 and 1868 Ma) are shown from the compilation of Evans et al. (2021) as well as the new ECMB pole. These data illustrate differential plate motion between the Superior and Slave Provinces that is required by the data leading up to the closure of the Manikewan Ocean and the assembly of Laurentia during the Trans-Hudson orogeny. The ECMB pole is consistent with an assembled Laurentia following the Trans-Hudson orogeny which contrasts with the disparate orientations and paleolatitudes between Laurentia’s constituent provinces prior to the orogeny. Evans, D. A. D., Pesonen, L. J., Eglington, B. M., Elming, S.-Å., Gong, Z., Li, Z.-X., … Zhang, S. (2021). An expanding list of reliable paleomagnetic poles for Precambrian tectonic reconstructions. In Ancient supercontinents and the paleogeography of the Earth. Holm, D. K., Van Schmus, W. R., MacNeill, L. C., Boerboom, T. J., Schweitzer, D., & Schneider, D. (2005). U-Pb zircon geochronology of Paleoproterozoic plutons from the northern midcontinent, USA: Evidence for subduction flip and continued convergence after geon 18 Penokean orogenesis. Geological Society of America Bulletin, 117(3), 259–275. doi: 10.1130/b25395.1 Jirsa, M., Boerboom, T., & Chandler, V. (2012). S-22, Geologic Map of Minnesota, Precambrian Bedrock Geology (Tech. Rep.). Minnesota Geological Survey. Whitmeyer, S., & Karlstrom, K. (2007). Tectonic model for the Proterozoic growth of North America. Geosphere, 3(4), 220–259. doi: 10.1130/GES00055.1 68 �The provenance, depositional environment and metallogenic implications of the Ament Bay Metasedimentary Assemblage, Sturgeon Lake Greenstone Belt, Northwest Ontario TAMOSAUSKAS, Michael1, LODGE, Robert2, MA, Chong1, HAUGAARD, Rasmus, SHERLOCK, Ross1 1 Mineral Exploration Research Centre (MERC), Harquail School of Earth Sciences, Goodman School of Mines, Laurentian University, 935 Ramsey Lake Rd., Sudbury, ON, P3E 2C6, Canada 2 Department of Geology, University of Wisconsin-Eau Claire, 105 Garfield Ave, Eau Claire, WI 54701, United States The Ament Bay Metasedimentary Assemblage (ABMA) is the youngest supracrustal assemblage in the Savant Lake-Sturgeon Lake greenstone belt (SSGB), and has previously been interpreted to have alluvial-fluvial origin (Sanborn-Barrie and Skulski, 2005). However, its timing, tectonic setting, and metallogenic significance is unclear. The SSGB makes up the easternmost part of the Western Wabigoon terrane of the Superior Province. The supracrustal assemblages in this belt form a regional-scale syncline, with the ABMA exposed along the hinge, and is intersected by the Sturgeon Lake fault zone (fig. 1). The ABMA is surrounded by felsic and mafic volcanic rocks that compose the Central and South Sturgeon Assemblages, however its contact relationships are poorly understood (Sanborn-Barrie and Skulski, 2005). Figure 1. Regional geology of the Sturgeon Lake greenstone belt. Modified after Sanborn-Barrie and Skulski (2005). Inset map showing close-up view of central portion of ABMA (outlined). 69 The goal of this study is to determine if the ABMA represents a Timiskaming-type basin. Timiskamingtype basins are found in various Neoarchean greenstone belts throughout the Superior. These are synorogenic basins which consist of clastic rocks with fluvial and marine �origins representing molasse basin-fill successions (Hyde, 1980; Thurston and Chivers, 1990; Mueller and Corcoran, 1998). These basins are commonly coeval with alkalic magmatism, as alkalic rocks can be both interbedded and intruded into the sediments, however they are also incorporated as clasts. Timiskaming-type basins are of interest because of their association with gold mineralization, as numerous prolific gold occurrences in the Abitibi greenstone belt are hosted within Timiskaming Group sediments (Bleeker, 2012). The ABMA consists of many features consistent with Timiskaming-type basins, as it is interpreted as a molasse basin-fill succession and Figure 2. Close-up view of central portion of the ABMA and its interpreted lithofaices. is spatially associated with alkalic magmatism. The ABMA consists of mostly weakly foliated, massive- to- planar bedded polymictic conglomerates and arkosic- to- lithic-rich arenites and greywackes. Based on petrographic analysis and lithology distribution, three lithofacies have been attributed to the ABMA: (i) arkosic greywacke lithofacies; (ii) polymictic conglomerate-sandstone lithofacies; (iii) lithic greywackemafic-rich conglomerate lithofacies (fig. 2). These are comparable to some lithofacies which make up other molasse basin-fill successions, some of which are Timiskaming-type (Corcoran and Mueller, 2007). Despite having similar characteristics with Timiskaming-type basins, the ABMA is poorly endowed relative to some of the gold-rich Timiskaming Group sediments in the Abitibi greenstone belt. This study explores possible factors which hindered gold mineralization in the Sturgeon Lake greenstone belt. References Bleeker, W., 2012. Lode gold deposits in Deformed and metamorphosed terranes: The role of Extension in the Formation of Timiskaming Basins and Large Gold Deposits, Abitibi Greenstone Belt–A Discussion, in Parker, J., ed., Summary of Field Work and Other Activities 2012: Ontario Geological Survey, Open File Report 6280, p. 47-41 to 47-12 Corcoran, P., & and Mueller, W., 2007. Time-Transgressive Archean Unconformities Underlying Molasse Basin-Fill Successions of Dissected Oceanic Arcs, Superior Province, Canada. Journal of Geology, volume 115, p. 655–674. Hyde, R.S., 1980. Sedimentary facies in the Archean Timiskaming Group and their tectonic implications, Abitibi greenstone belt, northeastern Ontario, Canada; Precambrian Research, v.12, p.161-195. Mueller, W.U. and Corcoran, P.L., 1998. Late-orogenic basins in the Archean Superior Province, Canada: Characteristics and inferences; Sedimentary Geology, v.120, p.177-203. Sanborn-Barrie, M. and Skulski, T. 2005. Geology, Sturgeon Lake greenstone belt, western Superior Province, Ontario; Geological Survey of Canada, Open File 1763, scale 1:100 000. Thurston, P.C., Chivers, K.M., 1990. Secular variation in greenstone sequence development emphasizing Superior Province, Canada. Precambrian Res. 46, 21 58. 70 �Possible implications of a non-Archean Grand Marais Ridge, western Lake Superior WOODRUFF, L.G. 1, and GRAUCH, V.J.S. 2 1 U.S. Geological Survey, 2280 Woodale Drive, St. Paul, MN, 55112 U.S. Geological Survey, MS 973, Federal Center, Denver, CO, 80225 2 The Grand Marias Ridge (GMR) is a gravity low but a structural high hidden beneath the waters of western Lake Superior. It has been interpreted as a preserved basement block of low-density Archean granite that remained a positive topographic high as adjacent volcanic basins subsided (Allen et al., 1997). Recent geophysical analysis (Grauch and Heller, 2021, this volume), however, suggests that the GMR likely comprises a significant volume of anorthosite, possibly mixed with felsic rocks. This part of the Midcontinent Rift System (MRS) in western Lake Superior is an area where many unusual features occur, including extraordinarily large concentrations of rapidly intruded mafic intrusions of the Duluth and Beaver Bay Complexes (DC and BBC), and a high percentage of felsic rocks in the North Shore Volcanic Group (NSVG) compared to other MRS volcanic sections. The thick basalt and overlying sedimentary packages in western Lake Superior also occur in a series of sag basins rather than the graben-like basins that characterize the MRS in the nearby St. Croix horst and the western arm extending away from Lake Superior. Here we address implications of recent modeling results and speculate on the potential significance of a GMR composed of anorthosite and felsic rocks, as well as the evolution of the MRS in this unique part of the Lake Superior region. Perhaps the most challenging aspect of this new compositional interpretation is how the GMR came to be a topographic high if it is not a block of remnant Archean basement. 1)What is the origin of anorthosite in the GMR? Option 1. Anorthosite represents pre-MRS basement rocks (e.g., Allen et al., 1997). This age is possible but unlikely as the regional Archean basement is granite-greenstone with little reported major units of anorthosite, and thus does not match the observed geophysical characteristics. Option 2. Anorthosite is part of the large complex suite of plagioclase-rich cumulate intrusions of the DC. Anorthosite intrusions are widespread throughout the DC, from the exposed western margin to where it disappears under the younger BBC and the NSVG. It is plausible that anorthosite in the GMR is an eastern extension of the DC, although it is challenging to determine how an older, deeper DC anorthosite pluton would now represent a high structural block. Option 3. The striking presence of abundant and, in places, very large anorthosite inclusions in the BBC Beaver River gabbro has been attributed to incorporation at depth of older DC anorthosite (Miller and Chandler, 1997). A recent paleomagnetic intensity study by Zhang (2020) found that anorthosite inclusions and host Beaver River diabase have similar paleomagnetic poles, which suggests that inclusions could be related to the BBC rather than the slightly older DC. Thus, anorthosite in the GMR could be products of BBC igneous events. In either case, anorthosite inclusions appear to be related to the MRS rather than pre-MRS basement and may be evidence of significant anorthosite accumulation at some depth near the GMR. 2)What is the origin of the felsic component in the GMR? Option 1. Archean granitic rocks that were intruded by plagioclase-rich magma during MRS time. Option 2. Partial melting of older basement rocks caused by staging of mafic magmas at depth during MRS time would lead to large accumulations of felsic magma (Vervoort and Green, 1997). Felsic magma could have been incorporated into an anorthosite cumulate melt and intruded higher in the crust. Option 3. The GMR is located just offshore of two of the largest rhyolite lava or ash flows in the NSVG (Green and Fitz, 1993), all temporally related to the DC. Nicholson et al. (1997) suggested that the proximity of large felsic flows on land and the GMR was not coincidental. It is possible that the GMR was a felsic volcanic center during DC time that erupted these large lava flows along with other regional felsic volcanics. 71 �3)What dynamic processes could have created the GMR anorthosite/felsic structural high? Option 1. Anorthosite and layered troctolite-gabbro plutons were intruded as inclined sheets between a footwall of Paleoproterozoic Animikie Group sedimentary rocks and Archean granites and a hangingwall of the NSVG (Miller et al., 2002). Thus, an Archean GMR would have to have been a topographic block within the Animikie Basin, as suggested by Allen et al. (1997). This basement structure could have been intruded by one of the many anorthositic igneous bodies generated during DC magmatic events as well as by felsic magmas derived from partial melting of some basement rocks. Option 2. Recent high precision 206Pb/238U zircon dates show that multiple intrusions of DC were emplaced together around 1096 Ma within less than 1 million years (Swanson-Hysell et al., 2020), suggesting that magma supply from deep crustal chambers for both intrusions and overlying volcanics was effectively continuous. Rapid, repetitive magmatic pulses would have elevated crustal temperatures around DC feeder zones and crustal magma chambers. Vervoort and Green (1997) suggested that the large rhyolite flows of the NSVG were products of crustal melting produced by the large thermal flux of DC magmatism. Perhaps remnants of a felsic volcanic center are located above a now crystallized plagioclase-rich magma chamber in the GMR, although the relative geometry of a middle-crust anorthosite intrusion, an upper crust felsic center, and a resulting GMR topographic high is perplexing. Option 3. A large volcanic sag basin just offshore from the BBC is bounded on the northeast by the GMR and on the southwest by the Archean (probable) structural block of White’s Ridge. The ~14 km-thick basalt section that fills this bowl-shaped basin, which is not fault-bounded as shown by Grant-Norpac legacy seismic profiles, was likely erupted from the curvilinear Beaver River diabase dike and related sheets that acted as major conduits for igneous activity (Miller and Chandler, 1997). Loading by a rapid succession of relatively dense basalt flows into an initially thermally subsiding basin could create an unstable environment (e.g., Mukherjee et al., 2020) where less dense underlying plagioclase-rich cumulates could be forced upward into a zone of weakness created by earlier high temperature felsic volcanism. This scenario, however, would require ductile behavior relatively high in the crust, but could explain how MRS anorthosite and felsic rocks became a topographic high as adjacent volcanic basins subsided. Perhaps localization of magmatic events over a short duration in this area created this seemingly implausible scenario. Allen, D. J., Hinze, W. J., Dickas, A. B., 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: Geol. Soc. Amer. Special Paper 312, 47-72. Grauch, V.J.S., Heller, S., 2021, Integration of geophysical evidence indicates that anorthosite composes a significant portion of Grand Marais ridge, an inferred basement high in western Lake Superior: this volume. Green, J.C., Fitz, T.J., III, 1993, Extensive felsic lava and rheoignimbrites in the Keweenawan Midcontinent Rift plateau volcanics, Minnesota: petrographic and field recognition: Jour. Volcan. Geotherm. Res., 54, 177-196. Miller, J.D., Jr., Chandler, V.W., 1997, Geology, petrology, and tectonic significance of the Beaver Bay Complex, northeastern Minnesota: Geol. Soc. Amer. Special Paper 312, 73-96. Miller, J.D., Jr., Green, J.C., Severson, M.J., Chandler, V.W., Hauck, S.A., Peterson, D.M., Wahl, T.E., 2002, Geology and mineral potential of the Duluth Complex and related rocks of northeastern Minnesota: Minn. Geol. Sur. Report Invest. 58, 207 p. Mukherjee, A.B., Das, S., Sen, D., Bhattacharya, B., 2020, Buoyant rise of anorthosite from a layered basic complex triggered by Rayleigh-Taylor instability: Insights from a numerical modeling study: Amer. Min., 105, 437-446. Nicholson, S.W., Shirey, S.B., Schulz, K.J., Green, J.C., 1997, Rift-wide correlation of 1.1 Ga Midcontinent rift system basalts: implications for multiple mantle sources during rift development: Can. Jour. Earth Sci., 34, 504520. Swanson-Hysell, N.L., Hoaglund, S.A., Crowley, J.L., Schmitz, M.D., Zhang, Y., Miller, J.D., Jr., 2020, Rapid emplacement of massive Duluth Complex intrusions within the North American Midcontinent Rift: Geol., 49, 185-189. Vervoort, J.D., Green, J.C., 1997, Origin of evolved magmas in the Midcontinent rift system, northeastern Minnesota: Nd-isotope evidence for melting of the Archean crust: Can. Jour. Earth Sci., 34, 521-535. Zhang, Y., 2020, Pairing paleointensity results with coercivity spectra: providing support for selection criteria: Inst. Rock Mag. Quar., 30, 2-4. 72 � 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 2021 Description An account of the resource Institute on Lake Superior Geology. Virtual meeting. May 11, 14, 18, 21, 2021. 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 2021 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. 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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. 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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. 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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/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. 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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. 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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/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. 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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