SETTING OF VOLCANOGENIC MASSIVE SULFIDE DEPOSITS IN THE PENOKEAN VOLCANIC BELT, GREAT LAKES REGION, USA
by Ashley Kaye Quigley
A thesis submitted to the Faculty and Board of Trustees of the Colorado School of Mines in partial fulfillment of the requirements for the degree of Master of Science (Geology).
Golden, Colorado
Date ______
Signed: ______Ashley Kaye Quigley
Signed: ______Dr. Thomas Monecke Thesis Advisor
Golden, Colorado
Date ______
Signed: ______Dr. Paul Santi Professor and Head Department of Geology and Geological Engineering
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ABSTRACT
The Paleoproterozoic (ca. 1875 Ma) Penokean volcanic belt represents one of the world’s most significant orogens hosting volcanogenic massive sulfide (VMS) deposits. Sporadic exploration from 1970-1995 has identified a large number of VMS deposits and prospects throughout the belt, including the world-class Crandon deposit that comprises an estimated 61 million tonnes of polymetallic massive sulfide ore. Despite successful exploration and the significant economic potential in the Penokean volcanic belt, only limited academic research has been conducted focusing on constraining the tectonic, structural, and volcanic setting of the VMS deposits. Many key aspects of the regional geology are not well understood, which is in part due to extensive glacial cover of the Paleoproterozoic bedrocks.
As part of the present study, aeromagnetic and gravity data were used to study the geological make-up of the Penokean volcanic belt. The data were integrated with existing mapping to constrain the bedrock geology under the glacial cover. Interpretation of the geophysical data showed that several distinct geological domains could be distinguished. Whole- rock major and trace element geochemical data were used to identify magmatic affinities of volcanic rocks within the geophysically defined domains. The results of the geochemical analyses reveal subtle differences between the geophysical domains. The majority of the volcanic rocks sampled have a tholeiitic affinity with fewer calc alkaline and transitional rocks. The geochemical evidence suggests that volcanism was arc-related, with the massive sulfide deposits being presumably formed in zones of intra-arc or back-arc extension.
High-precision chemical abrasion ID-TIMS U-Pb dating was performed on zircon grains separated from felsic volcanic samples collected from the host rock successions of some of the major VMS deposits and prospects within the Penokean volcanic belt. Results of the geochronological investigations showed that four of the deposits, namely Bend, Horseshoe, Lynne, and Pelican River, were all formed at about 1874 Ma. This suggests that several of the deposits of the Penokean volcanic belt formed during a major, but short-lived, period of rapid extension.
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The Back Forty massive sulfide deposit, located at the east end of the Penokean volcanic belt, is hosted by rhyolite that yielded an apparent age of about 1833 Ma. This is approximately 50 million years younger than the host rock successions of the other deposits of the belt. There are two possible explanations for this apparent age. The first explanation is that this represents a crystallization age and the host rocks to the Back Forty are part of a distinctly younger volcanic succession. Alternatively, a thermal event at 1833 Ma may have reset the U-Pb isotopic system. A crystallization age at around 1833 Ma is potentially consistent with an age recorded for a felsic volcanic rock from the Mountain area to the southwest.
The rhyolite sampled from the host rock successions of the Lynne and Back Forty deposits were found to contain rare Archean-aged zircon grains. These zircon grains yielded U-Pb ages of approximately 2700 Ma and are presumably inherited from an Archean basement. The existence of inherited ages supports the model that volcanism of the Pembine-Wausau terrane occurred, at least in part, on older Archean basement.
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TABLE OF CONTENTS
ABSTRACT ...... iii LIST OF FIGURES ...... vii LIST OF TABLES ...... viii ACKNOWLEDGEMENTS ...... ix CHAPTER 1: INTRODUCTION ...... 1 1.1. The Penokean Volcanic Belt ...... 1 1.2. Previous Research ...... 4 1.3. Research Approach and Methods ...... 5 1.3.1. Compilation of Existing Data ...... 7 1.3.2. Acquisition of New Data ...... 7 1.4. Thesis Organization ...... 8 1.5. References ...... 10 CHAPTER 2: VOLCANOGENIC MASSIVE SULFIDE DEPOSITS IN THE 1.8 GA PENOKEAN VOLCANIC BELT, MICHIGAN AND WISCONSIN: GEOLOGICAL FRAMEWORK AND PALEOTECTONIC SETTING ...... 13 2.1. Introduction ...... 13 2.2. Geological Background ...... 14 2.3. Magnetic and Gravity Data ...... 18 2.3.1. Methodology...... 18 2.3.2. Geophysical Domains ...... 23 2.3.2. Distribution of Volcanic and Plutonic Rocks ...... 26 2.3.3. Lineaments ...... 28 2.3.4. VMS Deposits ...... 30 2.4. Geochemical Data ...... 30 2.4.1. Methodology...... 31 2.4.2. Geochemistry of Volcanic Rocks ...... 32 2.5. Discussion and Conclusions ...... 38 2.6. Acknowledgements ...... 40 2.7. References ...... 42
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CHAPTER 3: ID-TIMS U-PB GEOCHRONOLOGY OF THE PALEOPROTEROZOIC PENOKEAN VOLCANIC BELT, MICHIGAN AND WISCONSIN: TIMING OF VOLCANOGENIC MASSIVE SULFIDE FORMATION AND EXPLORATION IMPLICATIONS ...... 45 3.1. Introduction ...... 45 3.2. Geological Setting ...... 47 3.3. Previous Geochronology ...... 52 3.4. Materials and Methods ...... 53 3.5. Results ...... 57 3.6. Discussion and Conclusions ...... 66 3.7. Acknowledgments ...... 74 3.8. References ...... 75 CHAPTER 4: CONCLUSIONS AND RECOMMENDATIONS FOR FUTURE WORK ...... 81 APPENDIX A: SUPPLEMENTAL ELECTRONIC FILES ...... 85
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LIST OF FIGURES
FIGURE 1-1. Simplified Geology map of the Penokean volcanic belt ...... 2 FIGURE 1-2. Location of the Penokean orogeny ...... 3 FIGURE 1-3. Schematic cross-sections of the tectonic evolution of the Penokean orogeny ...... 6 FIGURE 2-1. Simplified Geology map of the Penokean volcanic belt ...... 15 FIGURE 2-2. Magnetic TMI-RTP and Gravity maps of the Penokean volcanic belt ...... 24 FIGURE 2-3. Detailed Geology map of the Penokean volcanic belt ...... 25 FIGURE 2-4. Geology and geophysics illustrating the geophysical signature of volcanic and intrusive rocks ...... 29 FIGURE 2-5. Rock Classification diagrams for volcanic rocks ...... 33 FIGURE 2-6. Magmatic affinity diagram for volcanic rocks ...... 34 FIGURE 2-7. REE plots for mafic and felsic volcanic rocks from each geological domain ...... 37
FIGURE 2-8. Harker diagrams of SiO2 vs. P2O5, TiO2, and Zr ...... 38 FIGURE 2-9. Relative timeline of volcanism, plutonism and orogenic events ...... 41 FIGURE 3-1. Location of the Penokean orogeny ...... 46 FIGURE 3-2. Detailed geology of the Penokean volcanic belt with U-Pb date locations ...... 48 FIGURE 3-3. Concordia plots of CA-TIMS U-Pb dates from zircon grains ...... 63 FIGURE 3-4. Plots of 207Pb/206Pb dates from zircon analyzed by the CA-TIMS method ...... 64 FIGURE 3-5. Cathodoluminescence and backscattered electron images of zircon grains from host rocks to the Mountain prospect ...... 65 FIGURE 3-6: U-Pb geochronology of Paleoproterozoic intrusive and volcanic rocks of the Penokean volcanic belt ...... 67 FIGURE 3-7: Inherited zircon grains from PVB2014-001and PVB2014-004 compared to Archean aged intrusions in the Penokean volcanic belt ...... 71
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LIST OF TABLES
TABLE 2-1. VMS deposits of the Penokean volcanic belt ...... 17 TABLE 2-2. Magnetic susceptibility measurements on volcanic and intrusive rocks from the Penokean volcanic belt ...... 20 TABLE 3-1. Existing U-Pb geochronology of intrusive and volcanic rocks of the Penokean volcanic belt ...... 54 TABLE 3-2. Isotopic composition of zircon grains analyzed by CA-TIMS ...... 60
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ACKNOWLEDGEMENTS
I thank my advisor, Thomas Monecke, for taking me on as a student and for providing constant support. I have learned so much from Thomas and I am grateful to be a part of this ongoing research effort in the Penokean volcanic belt. My research has also greatly benefited from the many discussions with my committee, Nigel Kelly and Murray Hitzman, as well as, Tom Quigley, Jim Franklin, Klaus Schulz, Ted DeMatties, and Randy Van Schmus. Thank you to Eric Anderson for collaborating with me on the geophysics and geology of the Penokean volcanic belt and to Jim Crowley and Mark Schmitz for their help and expertise provided during the laboratory research at Boise State University.
Logistical support in the field was provided by Aquila Resources Inc. Thank you to Aquila Resources for allowing me access to the Back Forty property and drill core from the Bend and Horseshoe deposits.
And thank you to my husband, Patrick Quigley, for being my field partner, sounding board, and support system over the past three years.
This research was funded with the help of generous grants from the Society of Economic Geologists and the Geological Society of America. Additional funding was provided by the Barrick Gold Endowed Scholarship Fund at the Colorado School of Mines.
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CHAPTER 1 INTRODUCTION
This chapter outlines the objectives of this study, provides a brief project background, describes previous research, and details the organization of the thesis.
1.1. The Penokean Volcanic Belt
The Paleoproterozoic (ca. 1875 Ma) Penokean volcanic belt among the world’s most significant orogens containing volcanogenic massive sulfide (VMS) deposits (Franklin et al., 2005). The belt covers an area of about 200 by 300 km in north-central Wisconsin and the south portion of the Upper Peninsula of Michigan (Fig. 1-1). Sporadic exploration from 19701995 has identified a large number of VMS deposits and prospects throughout the belt, including the Crandon deposit that comprises an estimated 61 Mt of polymetallic massive sulfide ore (Lambe and Rowe, 1987). However, due to the public opposition to mining in Wisconsin, the VMS deposits of the Penokean volcanic belt have not been exploited. So far, only the oxide zone of the Flambeau deposit has been mined. Between 1993 and 1997, the deposit yielded 1.71 Mt of ore grading 10.3% Cu, 3.63 ppm Au, and 57.4 ppm Ag (May and Dinkowitz, 1996).
Despite successful exploration and the significant economic potential, only limited academic research has been conducted in the Penokean volcanic belt constraining the tectonic, structural, and volcanic setting of the VMS deposits. The current understanding of the regional geology of the Penokean volcanic belt is largely based on mapping conducted several decades ago (compiled by Nicholson et al., 2007). Many key aspects of the regional geology are still not well understood, which is mostly due to extensive glacial cover of the Paleoproterozoic bedrock. Knowledge of the geological framework of the VMS deposits has fallen significantly behind other well-studied orogens containing VMS deposits as many of the field and analytical techniques used today have only become available during the few past decades.
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Figure 1-1: Geology map of the Penokean volcanic belt (modified from DeMatties, 1994). NFZ: Niagara fault zone; EP: Eau Pleine shear zone; JR: Jump River shear zone; MT: Mountain shear zone; AT: Athens shear zone.
2 A better understanding of the geologic framework of the Penokean volcanic belt and the setting and timing of VMS formation could aid in future exploration efforts and potentially change the way companies explore for these deposits in Wisconsin and the Upper Peninsula of Michigan. Renewed research on the Penokean volcanic belt is also important from a scientific perspective as the belt represents key step in the assembly of the North American continent but is one of the most understudied Paleoproterozoic orogens in North America (Fig. 1-2).
Figure 1-2: North American continent and the location of the Penokean orogeny (modified from Hoffman, 1989).
3 1.2. Previous Research
The Penokean volcanic belt was first identified as a greenstone belt in the late-1960s, which sparked interest in the economic potential of central Wisconsin (Weis et al., 1969). The earliest scientific studies in the Penokean volcanic belt were primarily focused on the U-Pb geochronology of the volcanic and intrusive rocks (Goldich et al., 1966a,b; Banks and Cain, 1969) and geologic mapping by Dutton and Bradley (1968).
Following Kennecott’s discovery of the Flambeau deposit in 1968 (May and Dinkowitz, 1996), a number of companies conducted exploration in the Penokean volcanic belt, including Exxon, Noranda, and U.S. Steel. The early years of exploration were highly competitive and little was published except for the summary papers by Van Schmus (1976) and Sims (1976). Most of the research on the Penokean volcanic belt was conducted during the 1980s. LaBerge and Mudrey (1979), Sims (1980), and Van Schmus and Bickford (1981) were among the first to publish ideas about the stratigraphic and tectonic framework of the Penokean volcanic belt. A detailed map (1:100,000 scale) of the bedrock geology of Marathon County, Wisconsin, was published by LaBerge and Myers (1984). The most current compilation of the regional geology was published as three 1:250,000 scale maps by Sims (1989a, 1990a,b). Sims (1989b) also published a 1:24,000 scale map of the Waupee Volcanics near Mountain in Oconto County, Wisconsin. More recently, LaBerge and Klasner (2001) published a 1:24,000 scale map of the Monico area, northern Wisconsin.
Research on the regional metallogeny, geology, and tectonic framework of the Penokean volcanic belt was conducted by Sims et al. (1987, 1989), and DeMatties (1989, 1994). Van Wyck and Johnson (1997) used Nd and Pb isotope data to present an interpretation of the tectonic framework of the belt that challenged conventional knowledge about the how far south the Archean basement of the Superior craton extends. Using mostly existing data, several more recent papers have made new interpretations of the tectonic setting of the belt and how the Penokean volcanic belt fits into the assembly of North America (Schulz and Cannon, 2007; Holm et al., 2007). The most recent tectonic interpretation of the Penokean volcanic belt was proposed by Schulz and Cannon (2007). The basic concept is a southward subducting
4 intraoceanic arc that is accreted to the southern margin of the Superior craton. This collision caused a subduction flip, switching the direction of subduction to the north. The northward subduction resulted in the accretion of the Marshfield terrane and subsequent compression (Fig. 1-3).
The first publication that discussed and summarized individual deposits and their regional context was by DeMatties (1994). The most recent and most comprehensive compilation of the VMS deposits of the Penokean volcanic belt was published in 1996 as commemorative volume of the Institute of Lake Superior Geology. The volume describes the geology of the Bend (DeMatties and Rowell, 1996), Crandon (Erickson and Cote, 1996), Eisenbrey (May, 1996), Flambeau (May and Dinkowitz, 1996), and Lynne deposits (Adams, 1996).
1.3. Research Approach and Methods
The primary objective of the present study is to provide a better understanding of the geologic framework and tectonic setting of the Penokean volcanic belt. Specifically, the research aimed to construct one of the first comprehensive geological maps of the entire Penokean volcanic belt through integration of previous mapping results and aeromagnetic data. The research involved the compilation of available geochemical analyses of volcanic rocks and the acquisition of new data to constrain the magmatic affinity of the volcanism and to draw conclusions about the magmatic evolution of the belt. In addition, new high-precision dating was conducted on selected felsic volcanic rocks from the host rock successions of major VMS deposits and prospects to complement existing U-Pb zircon age data. The geochronological research was conducted to determine whether the different deposits and prospects in the belt formed as part of a single magmatic event or were associated with volcanic centers that developed at different times. The new age constraints have significant implications for the understanding of the tectonic setting of the Penokean volcanic belt.
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Figure 1-3: Schematic cross-sections illustrating the tectonic evolution of the Penokean orogeny as proposed by Schulz and Cannon (2007).
6 1.3.1. Compilation of Existing Data
During the initial phase of the research, existing map data was compiled from the literature as well as the U.S. Geological Survey and the Wisconsin Geological Survey databases. Known outcrop locations and available geological maps were integrated in a common GIS environment to produce a geological map of the belt. In addition, whole-rock geochemical data was compiled following quality screening to ensure that only high-quality XRF and ICP-MS analyses were used. To provide as many age constraints as possible, existing U-Pb zircon dates for extrusive and intrusive volcanic rocks from the Penokean volcanic belt were compiled from the literature. The data were screened to exclude dates having large analytical errors. Dates of volcanic rocks having a precision of more than 10 Ma were excluded from the database while dates of intrusive rocks having a precision of more than 15 Ma were not considered in the present study. The error cutoffs of 10 Ma and 15 Ma were chosen to ensure that the database is both inclusive and meaningful.
To supplement existing map data and to better define the geology in areas of the belt with extensive glacial cover, available aeromagnetic data over the Penokean volcanic belt was processed utilizing various transformations including reduction-to-pole, tilt derivative, first vertical derivative, and upward continuation methods. The processed aeromagnetic and gravity data were added to the GIS environment to correlate with the compiled geological information.
1.3.2. Acquisition of New Data
Fieldwork was conducted over the course of 10 weeks during the summer of 2014. The fieldwork focused primarily on sampling volcanic rocks from the host rock successions of significant VMS deposits and prospects for whole-rock geochemical investigations and geochronological research. To constrain the geological context of the sampled volcanic rocks, core logging and limited surface mapping was conducted. In addition to the sampling, a total of 63 magnetic susceptibility measurements of volcanic and intrusive rocks were collected in the field to help confirm the magnetic signature of these units, facilitating interpretation of the compiled aeromagnetic data.
7 During the field season, 35 volcanic rocks considered least altered based on context were sampled from across the Penokean volcanic belt. The samples were analyzed to determine the magmatic affinity of the volcanic rocks and to further constrain the tectonic setting of the VMS deposits. Initially, the samples were prepared for geochemical analysis by removing any obvious vein material or amygdules. The samples were then crushed and milled using the facilities at the Department of Geology and Geological Engineering, Colorado School of Mines. The pulps were submitted for geochemical analysis to Actlabs in Ancaster, Ontario. The major element geochemistry of the rocks was determined by XRF, with the loss on ignition being measured through gravimetry following sample combustion. Trace element analysis was performed by ICP-MS following four acid digestion.
To obtain the absolute ages of the volcanic successions hosting VMS deposits in the Penokean volcanic belt, seven samples were collected for high resolution U-Pb zircon geochronology. Four samples were collected from drill core of felsic volcanic units forming part of the host rock successions of some of the most important VMS deposits. The other three samples were obtained from outcrops located in proximity to known VMS deposits and prospects. The geochronology samples were crushed and milled. Zircon grains were separated by processing the milled sample using standard techniques, including heavy mineral separation on a Wilfley table, treatment with heavy liquids, and magnetic separation. Sample preparation was in part conducted at Colorado School of Mines and Boise State University. Following hand- picking, selected zircon grains were analyzed by ID-TIMS at Boise State University.
1.4. Thesis Organization
This thesis is composed of four chapters. Chapter 1 provides a brief introduction to the geology of the Penokean volcanic belt and its VMS endowment. This chapter also discusses the importance of the new research and the methodology used to answer the scientific questions defined. Following this introductory chapter, the thesis contains two main chapters, to be submitted as separate papers to Economic Geology.
8 Chapter 2 presents the results of the geophysical investigations and describes how the aeromagnetic data and the compilation of previous mapping results were used to derive the first comprehensive regional geologic map of the Penokean volcanic belt. The mapping data are combined with whole-rock geochemical data to provide a better geological framework for the Penokean volcanic belt and to constrain the tectonic setting in which volcanism has occurred. Geophysically defined geologic domains were found to have differences in geochemical characteristics. This chapter will be submitted to Economic Geology with co-authors, Ashley Quigley, Thomas Monecke, and Eric Anderson. Ashley Quigley and Thomas Monecke are the graduate student and academic advisor, respectively, and Eric Anderson of the U.S. Geological Survey was heavily involved in the interpretation of the geophysical data.
As part of the present study, a geochronological study (Chapter 3) was conducted to provide new constraints on the timing of volcanism in the Penokean volcanic belt and to test whether known VMS deposits across the belt formed during the same period of volcanism. Chapter 3 presents and discusses the results of the ID-TIMS U-Pb zircon geochronology study conducted on samples collected from the host rock successions of some of the most important VMS deposits and prospects in the belt. The host rocks to deposits within the central part of the belt (Bend, Horseshoe, Pelican River, and Lynne) have, within error, identical ages of about 1874 Ma. However, rhyolites hosting the Back Forty deposit and the Mountain prospect have U- Pb zircons dates of c. 1833 and c. 1843 Ma, respectively. The implications of these dates are discussed that chapter; these two samples are from the east end of the belt and could represent their own volcanic) events. This chapter will be submitted to Economic Geology with co-authors, Ashley Quigley, Thomas Monecke, Nigel Kelly, Jim Crowley, and Patrick Quigley. Nigel Kelly of the University of Colorado-Boulder and Jim Crowley of Boise State University were heavily involved in the interpretation of the geochronological data. Patrick Quigley also contributed to the results of this study.
The results of the research are summarized in Chapter 4. This chapter provides key conclusions regarding the geologic framework and tectonic setting of the Penokean volcanic belt and discusses exploration implications. In addition, recommendations for future work in the Penokean volcanic belt are made.
9 1.5. 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: Cable, Wisconsin, Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, v. 42, part 2, p. 161–179. Banks, P.O., and Cain, J.A., 1969, Zircon ages of Precambrian granitic rocks, northeastern Wisconsin: The Journal of Geology, v. 77, p. 208–220. 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. DeMatties, T.A., 1994, Early Proterozoic volcanogenic massive sulfide deposits in Wisconsin: An overview: Economic Geology, v. 89, p. 1122–1151. DeMatties, T.A., and Rowell, W.F., 1996, The Bend deposit: An Early Proterozoic copper-gold VMS deposit, in LaBerge, G.L., ed., Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume: Cable, Wisconsin, Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, v. 42, part 2, p. 143–160. Dutton, C.E., and Bradley, R.E., 1968, Map of northern half of Wisconsin showing geologic data of Precambrian area: U.S. Geological Survey Open-File Report 68-91, 4 maps.
Erickson, A.J., and Cote, R., 1996, Geological summary – Crandon deposit, in LaBerge, G.L., ed., Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume: Cable, Wisconsin, Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, v. 42, part 2, p. 129–142. Franklin, J.M., Gibson, H.L., Jonasson, I.R., and Galley, A.G., 2005, Volcanogenic massive sulphide deposits: Economic Geology 100th Anniversary Volume, p. 523–560. Goldich, S.S., Muehlberger, W.R., Lidiak, E.G., and Hedge, C.E., 1966a, Geochronology of the Midcontinent Region, United States 1. Scope, Methods, and Principles: Journal of Geophysical Research, v. 71, p. 5375–5388. Goldich, S.S., Lidiak, E.G., Hedge, C.E., and Walthall, F.G., 1966b, Geochronology of the Midcontinent Region, United States 2. Northern Area: Journal of Geophysical Research, v. 71, p. 5389–5408. Hoffman, P.F., 1989, Precambrian geology and tectonic history of North America, in Bally, A.W., and Palmer, A.R., eds., The geology of North America – An overview: Geological Society of America, Boulder, Colorado, p. 447512. 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 Paleoproterzoic accretionary boundaries of the north-central United States based on a new aeromagnetic-geologic compilation: Precambrian Research, v. 157, p. 71–79.
10 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 Geologic Investigations Series Map, I-2739, scale 1:24,000. LaBerge, G.L., and Mudrey, M.G., Jr., 1979, Stratigraphic Framework of the Wisconsin middle Precambrian: Wisconsin Geological and Natural History Survey Miscellaneous Paper, v. 79-1, p. 22. 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. Lambe, R.N., and Rowe, R.G., 1987, Volcanic history, mineralization, and alteration of the Crandon massive sulfide deposit, Wisconsin: Economic Geology, v. 82, p. 1204–1238. May, E.R., 1996, Eisenbrey: A structurally complex Proterozoic copper-zinc massive sulfide deposit, Rusk County, Wisconsin, in LaBerge, G.L., ed., Volcanogenic massive sulfide deposits of northern Wisconsin: A commemorative volume: Cable, Wisconsin, Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, v. 42, part 2, p. 107–128. 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 deposits of northern Wisconsin: A commemorative volume: Cable, Wisconsin, Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, v. 42, part 2, p. 67–94. Nicholson, S.W., Dicken, C.L., Foose, M.P., and Mueller, J.A.L., 2007, Preliminary integrated geologic map databases for the United States: Minnesota, Wisconsin, Michigan, Illinois, and Indiana: U.S. Geological Survey Open-File Report 2004–1355. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 4–25. Sims, P.K., 1976, Precambrian tectonic and mineral deposits, Lake Superior region: Economic Geology, v. 71, p. 1092–1118. Sims, P.K., 1980, Boundary between Archean greenstone and gneiss terranes in northern Wisconsin and Michigan: Geological Society of America Special Paper, v. 182, p. 113–124. Sims, P.K., 1989a, Geologic map of Precambrian rocks of Rice Lake 1 x 2 degree quadrangles, northern Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map, I- 1924, scale 1: 250,000. Sims, P.K., 1989b, Geologic map of Proterozoic rocks near Mountain, Oconto County, Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map I-1903, scale 1:24,000. Sims, P.K., 1990a, Geologic map of Precambrian rocks of Eau Claire and Green Bay 1 degree x 2 degrees quadrangles, central Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map, I-1925, scale 1:250,000
11 Sims, P.K., 1990b, Geologic map of Precambrian rocks of Iron Mountain and Escanaba 1 degree x 2 degrees quadrangles, northeastern Wisconsin and northwestern Michigan: U.S. Geological Survey Miscellaneous Investigations Series Map, I-2056, scale 1: 250,000. Sims, P.K., Kisvarsanyi, E.B., and Morey, G.B., 1987, Geology and metallogeny of Archean and Proterozoic basement terranes in the Northern Midcontinent, U.S.A. – An overview: U.S. Geological Survey Bulletin, v. 1815, p. 1–51. 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 Science, v. 26, p. 2145–2158. Van Schmus, W.R., 1976, Early and middle Proterozoic history of the Great Lakes area, North America: Philosophical Transactions of the Royal Society of London, Series A, v. 280, p. 605–628. Van Schmus, W.R., and Bickford, M.E., 1981, Proterozoic chronology and evolution of the Midcontinent region, North America, in Kroner, A., ed., Precambrian plate tectonics: Amsterdam, Elsevier, p. 261–296. 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. Wies, L.W., LaBerge, G.L., and Dutton, C.E., 1969, Central Wisconsin Volcanic Belt [Fieldtrip Guide]: Institute of Lake Superior Geology Annual Meeting, 15th, Wausau, Wisconsin, 1969.
12 CHAPTER 2 VOLCANOGENIC MASSIVE SULFIDE DEPOSITS IN THE 1.8 GA PENOKEAN VOLCANIC BELT, MICHIGAN AND WISCONSIN: GEOLOGICAL FRAMEWORK AND PALEOTECTONIC SETTING
A paper to be submitted to Economic Geology. Ashley Quigley, Thomas Monecke, Eric Anderson
2.1. Introduction
The Penokean volcanic belt of northern Michigan and Wisconsin represents one of the most important Paleoproterozoic (ca. 1875 Ma) orogens hosting volcanogenic massive sulfide (VMS) districts worldwide (Franklin et al., 2005). The greenschist to amphibolite facies, bimodal mafic volcanic belt is host to more than 13 known deposits and prospects that contain in excess of 100 million metric tons (Mt) of massive sulfides (DeMatties, 1994; Schulz et al., 2008). The world-class Crandon deposit, discovered by Exxon Minerals Company in the 1975, is the largest deposit in the belt and estimated to contain 61.1 Mt of ore grading 1.04% Cu, 5.56% Zn, 0.48% Pb, 1.2 g/t Au, and 43 g/t Ag (Erickson and Cȏté, 1996; Lambe and Rowe, 1987). However, due to public opposition to mining in Wisconsin, the deposits in the belt have not been developed or mined. The only exception is the supergene-enriched portion of the Flambeau deposit, which was mined from 1993 to 1997 (May and Dinkowitz, 1996).
Over the past three decades, relatively little research has been conducted in the Penokean volcanic belt. Many key aspects of the geology of the belt and the setting of the VMS deposits are not well understood. In particular, the tectonic and structural setting of VMS deposit formation is not well constrained as no mapping across this prolific belt is limited by poor exposure.
The present contribution aims to contribute to a better understanding of the regional geology of the Penokean volcanic belt. Existing aeromagnetic data were compiled, reprocessed,
13 and integrated with new field magnetic susceptibility measurements to map major lithological units and lineaments across the belt. An updated geological map for the Penokean volcanic belt was generated based on interpretation of geophysical and geochemical data and reconciliation with previous geological mapping in areas of bedrock exposure. Whole-rock geochemical analyses were compiled from the literature to identify magmatic affinities of the volcanic successions across the belt. In addition, new sampling complemented the compilation of geochemical data to ensure adequate sample coverage of the immediate host rock succession of known massive sulfide deposits. The results of the new research support models that view the Penokean volcanic belt as an assemblage of different volcanic complexes formed at different times along the southern margin of the Archean Superior craton. These volcanic terrains were tectonically juxtaposed during the Penokean orogeny.
2.2. Geological Background
The Penokean volcanic belt is an east-west trending assemblage of volcanic and intrusive rocks located in north-central Wisconsin and the southern portion of the Upper Peninsula of Michigan (Fig. 2-1). The belt spans an area that is about 300 km from east to west and about 200 km from north to south. It is interpreted to extend under cover to the west into the state of Minnesota (Southwick and Morey, 1991). The volcanic rocks are Paleoproterozoic in age with previous dates ranging from 19001830 Ma (Sims et al., 1989) and are overlain by Paleozoic sedimentary rocks to the south and west. The 1100 Ma mid-continent rift transects and obscures the western most part of the belt (Holm et al., 2007).
Regional geology
The Penokean volcanic belt is located south of the Superior craton and comprises the Pembine-Wausau and the Marshfield terranes (Fig. 2-1). The Pembine-Wausau terrane is thought to be comprised of largely juvenile Paleoproterozoic volcanic and intrusive rocks that formed in a suprasubduction environment. Sims et al. (1989) and Schulz and Cannon (2007) suggested that volcanism in the Pembine-Wausau terrane occurred in an intraoceanic arc setting, whereby
14
Figure 2-1: Geology map of the Penokean volcanic belt (modified from DeMatties, 1994). NFZ: Niagara fault zone; EP: Eau Pleine shear zone; JR: Jump River shear zone; MT: Mountain shear zone; and AT: Athens shear zone.
15 oceanic arc crust would have occurred on both sides of the arc. In contrast, Van Wyck and Johnson (1997) favored a continental margin arc setting as isotopic evidence suggests the existence of an Archean basement below the Paleoproterozoic volcanic successions.
Within the Pembine-Wausau terrane, two distinct volcanic complexes have been identified. The Pembine volcanic complex, also known as the Ladysmith-Rhinelander volcanic complex, in the northern part of the terrane was formed at approximately 18801870 Ma (DeMatties, 1994). The Wausau volcanic complex in the southern portion of the terrane presumably formed between
18451835 Ma (Sims et al., 1989). Within the Marshfield terrane, Paleoproterozoic volcanic rocks overly Archean schists and gneisses (DeMatties, 1994). The Archean rocks are thought to represent a micro-continent that was docked against the arc during the Penokean orogeny.
Within the Penokean volcanic belt, there are several documented east-west and southwest- northeast trending fault zones. The Niagara fault zone to the north of the Penokean volcanic belt separates the Superior craton from the Pembine-Wausau terrane. The Eau Pleine shear zone marks the suture between the Pembine-Wausau terrane and the Marshfield terrane (LaBerge et al., 2003; Sims et al, 2005; Schulz and Cannon, 2007). Additional major structures have been mapped throughout the belt, including the Athens shear zone and the Jump River shear zone (Fig. 2-1).
Various mafic and felsic intrusions are found across the east-west trending Penokean volcanic belt. Many of the mafic intrusions are mapped within a poorly exposed portion of the Pembine-Wausau terrane (Nicholson et al., 2007). Although some of these mafic intrusions are exposed at surface, many of them have been mapped as a result of distinct magnetic and gravity anomalies that are sometimes accompanied by drill hole data. The age of these intrusions is unknown. By contrast, the felsic intrusions have been observed in outcrop and several have been dated via U-Pb geochronology of zircon grains (Sims et al., 1989). A line of felsic intrusions occur just south of and trending parallel to the Niagara fault zone on the northern margin of the Pembine-Wausau terrane. These intrusions are formed at around 1760 Ma (Sims et al., 1989) and are thought to be the result of continued plate tectonic activity along the southern margin of the
16 Superior craton (Holm et al., 2007). These intrusions are likely related to the East-Central Minnesota Batholith, which is characterized by granitic rocks of the same age (Jirsa et al., 2011).
Massive sulfide deposits
The Penokean volcanic belt is host to a number of polymetallic VMS deposits and prospects, which primarily occur along an east-west trend in the south-central part of the belt (Fig. 2-1; Table 2-1). The VMS deposits in the belt occur in the Ladysmith, Somo, and Crandon districts (DeMatties, 1994). Although these districts are defined based on geographic location and not geological criteria such as stratigraphic controls, some general trends can be observed. The most significant Cu-rich deposits, Flambeau and Bend, occur in the Ladysmith district in the western part of the belt. The Zn-Cu and Zn-Pb-Cu deposits are located in the central Somo district and the eastern Crandon district. The Au/Ag ratios appear to generally decrease from west to east along the belt (DeMatties, 1994). The recently discovered Back Forty deposit is the most eastern deposit in the belt, just east of the Wisconsin-Michigan border. The Back Forty deposit does not fit the Au/Ag trend as discussed by DeMatties (1994), as it is a gold-rich
Table 2-1: VMS deposits of the Penokean volcanic belt (P. Quigley, 2016).
Deposit Total Indicated and Cu Zn (%) Pb Au Ag (g/t) Resource measured (%) (%) (g/t) (Mt) reserves (Mt)
Back Forty 17.5 15.1 0.33 2.94 0.23 2.04 24.75 Bend 3.9 3.4 1.87 - - 3.14 13.26 Catwillow 2.6 1.50 2.60 - 0.69 15.43 Crandon 65.8 61.1 1.04 5.56 0.48 1.20 42.86 Flambeau ~5.9 1.7 4.10 1.00 - 2.91 30.17 Hawk 1.4 0.80 2.70 - - - Horse Shoe 0.7 2.45 5.35 0.90 2.06 36.00 Lynne 5.1 0.47 9.27 1.71 0.72 81.60 Pelican River 2.1 1.00 4.50 - - 17.49 Ritchie Creek 0.8 0.5 2.11 0.37 - 0.34 - Thornapple 2.7 1.50 ~3.50 - - -
17 deposit. The deposit is located in a poorly explored felsic volcanic succession east of the Wolf River batholith.
2.3. Magnetic and Gravity Data
As much of the Penokean volcanic belt is obscured by glacial cover, the use of geophysical data is critical in understanding the regional geology. Prior to the present study, the Wisconsin aeromagnetic and gravity datasets have been used to extrapolate the geology in areas with little to no outcrop (Sims, 1989) and to interpret the tectonic setting and the accretionary boundaries of the region (Holm et al., 2007). In the present study, additional detailed information was extracted from the existing magnetic and gravity datasets with the help of modern filtering techniques (Fig. 2-2). The resulting geologic map is shown in Figure 2-3.
2.3.1. Methodology
As part of the present study, the best publicly available aeromagnetic datasets acquired by the U.S. Geological Survey over the Penokean volcanic belt were compiled and reprocessed. Most of the data were collected along north-south flight lines spaced 800 meters at a nominal height of 150 meters. Individual survey data were processed using industry standard techniques (Luyendyk, 1997) and merged to produce a uniform anomaly map with grid spacing of 250 meters.
Several data transforms were applied to highlight different components of the anomalous magnetic field, including reduction-to-pole, upward continuation, tilt derivative, and analytic signal. The results, supplemented with ground gravity data, were used to systematically verify and modify existing geology units from a GIS compilation of current U.S. Geological Survey bedrock geology maps (Nicholson et al., 2007). Magnetic susceptibility measurements were taken using an SM30 instrument from every outcrop visited during the field season to ground truth the magnetic signature in key areas of the belt (Table 2-2). This data was used to define geology in areas with extensive glacial cover, identify structures, and locate potential intrusions associated with VMS deposits.
18 Total magnetic intensity
The total magnetic intensity (TMI) data are the resultant magnetic field after diurnal and regional variations are removed from the measured magnetic field. Transformations are applied to the TMI data to enhance various components of the anomalous field that may represent geologic units or structures (Milligan and Gunn, 1997).
Reduced-to pole
The reduced-to-pole (RTP) transform is useful for data collected from a latitude where the magnetic field is inclined. The transform allows data projection as if the magnetic field is vertical and, therefore, simplifies anomaly interpretation by better aligning them over their causative sources (Baranov and Naudy, 1964). The transform was calculated using a field strength of
59,489 nT, an inclination of 74.3, and a declination of 1.8. The other transforms of magnetic data were applied to the RTP data. The RTP transform is particularly helpful to differentiate volcanic and plutonic rocks. Figure 2-2a shows the RTP aeromagnetic map of the Penokean volcanic belt.
Tilt derivative
The tilt derivative (TDR) was applied to the RTP dataset. TDR helps to normalize the large dynamic range of the magnetic data and tends to be positive over the source, crosses through zero at or near the edges of a vertical contact, and is negative outside the source region. The tilt derivative transform is used to highlight magnetic contacts and lineaments (Miller and Singh, 1994). The TDR was primarily used to identify structures within the Penokean volcanic belt and was also helpful for identifying the locations of banded iron formations north of the Niagara fault zone.
19 Table 2-2: Magnetic susceptibility measurements on volcanic and intrusive rocks from the Penokean volcanic belt.
Mean No. Easting Northing Rock Comment n (10-3 SI)
MS-001 435358 5034119 Volcanic Quartz-phyric rhyolite 0.004 9 MS-002 435286 5033730 Volcanic 0.020 7 MS-003 434812 5032976 Volcanic 0.000 6 MS-004 434749 5032955 Volcanic 0.000 4 MS-005 431227 5032061 Intrusive Granite 0.000 4 MS-006 430514 5033122 Intrusive Granite 0.000 4 MS-007 431150 5032162 Intrusive Granite 0.000 1 MS-008 431550 5032966 Intrusive Heterogeneous rock 0.046 5 MS-009 437256 5039699 Volcanic Quartz-feldspar porphyry 0.000 5 MS-010 434479 5039137 Intrusive Granite 0.000 5 MS-011 433244 5042179 Intrusive Granite 0.000 5 MS-012 434639 5047477 Volcanic Porphyritic mafic rock 0.917 6 MS-013 433289 5047654 Volcanic Porphyritic mafic rock 29.960 5 MS-014 432887 5045982 Volcanic Felsic volcanic rock 0.020 4 MS-015 431251 5036777 Volcanic Quartz-feldspar porphyry 0.132 5 MS-016 431667 5036591 Intrusive Granite 0.000 5 MS-017 429925 5036028 Intrusive Mafic intrusive 0.400 5 MS-018 427901 5036079 Intrusive Felsic intrusive rock 0.020 5 MS-019 431559 5035989 Intrusive Granite 0.144 5 MS-020 433254 5046008 Volcanic Volcaniclastic rock 0.000 5 MS-021 433254 5046008 Volcanic Intermediate volcanic rock 0.090 4 MS-022 416655 5021014 Volcanic Intermediate volcanic rock 0.000 5 MS-023 426107 5035120 Volcanic Felsic dike 0.008 4 MS-024 426107 5035120 Volcanic Felsic volcanic rock 0.138 4 MS-025 388683 5002940 Volcanic Felsic volcanic rock 0.300 10 MS-026 388799 5002852 Volcanic Felsic volcanic rock 0.271 8 MS-027 387466 5002893 Volcanic Felsic volcanic rock 0.258 5 MS-028 391352 5006347 Volcanic Mafic volcanic rock 0.554 5 MS-029 215120 5034698 Intrusive Granite-diorite 0.075 8 MS-030 215252 5034397 Intrusive Granite-diorite 0.095 4 MS-031 199644 5007619 Volcanic Mafic volcanic rock 0.946 4 MS-032 203163 5029146 Intrusive Granite 0.093 9 MS-033 203173 5029158 Volcanic Porphyritic rock 0.058 5 MS-034 322569 5049819 Volcanic Felsic volcanic rock 0.097 8
20 Table 2-2 (continued).
Mean No. Easting Northing Rock Comment (10-3 SI) n
MS-036 330604 5051103 Volcanic Quartz porphyry 0.051 8 MS-039 327796 5045318 Volcanic Felsic volcanic rock 2.897 11 MS-040 327565 5045497 Volcanic Felsic volcanic rock 0.102 9 MS-041 331052 5045240 Intrusive Granite 0.015 6 MS-042 335804 5045483 Volcanic Altered felsic volcanic rock 0.174 8 MS-043 332673 5049067 Volcanic Mafic volcanic rock 0.237 7 MS-044 295500 4983153 Volcanic Brecciated rhyolite 2.811 8 MS-045 295541 4983181 Volcanic Brecciated rhyolite 0.021 6 MS-046 290135 4982480 Intrusive Syenite-quartz syenite 2.753 10 MS-047 280371 4991958 Volcanic Massive basalt 0.139 9 MS-048 288406 4962976 Intrusive Granite 0.005 5 MS-049 277077 4978708 Volcanic Mafic volcanic rock 0.165 6 MS-050 280667 4980907 Intrusive Wolf River batholith 2.845 10 MS-050A 280587 4981009 Volcanic Mafic dike 0.388 4 MS-051 313952 4984676 Metamorphic Foliated mafic 0.436 6 MS-052 300502 4980012 Intrusive Fine-grained intrusive 7.973 6 MS-053 289857 4989232 Volcanic Mafic volcanic dike 7.514 7 MS-054 281367 5020605 Metamorphic Quartz-mica gneiss 0.053 8 MS-055 281594 5021113 Intrusive Granite 0.000 4 MS-056 286298 5027278 Intrusive Granite 0.098 7 MS-057 335191 5045759 Volcanic Felsic volcanic rock 0.459 7 MS-058 338482 5042461 Volcanic Felsic volcanic rock 0.000 4 MS-059 334461 5041681 Volcanic Felsic volcanic rock 0.042 7 MS-060 335084 5044195 Volcanic Mafic volcanic rock 0.257 10 MS-061 280349 4991975 Volcanic Massive mafic volcanic 0.283 10 MS-062 276200 4993095 Volcanic Altered volcanic rock 0.085 7 MS-063 300609 4994531 Intrusive Felsic intrusive 0.308 7 n = number of measurements.
21 Analytic signal
The analytic signal (AS) was applied to the RTP dataset. This transform displays maxima over magnetization contrasts, such that edges of broad magnetic sources, as well as, small, discrete bodies are highlighted (Roest et al., 1992). In this study, the AS was useful for mapping the edges of mafic intrusive bodies within the Penokean volcanic belt and helped to identify potential intrusions associated with VMS deposits.
Upward continuation
The upward continuation transform calculates the magnetic field at an elevation higher than it was originally measured which filters high-frequency anomalies relative to low-frequency anomalies. The effect of the transform suppresses near-surface effects and accentuates the anomalies from deeper magnetic sources (Blakely, 1995; Milligan and Gunn, 1997). The data for this study were continued to multiple heights between 500 and 15,000 m. This transform was used to highlight deep intrusive bodies and to map geological domains.
Gravity data
The ground gravity data of Wisconsin and Michigan used for the present study represent a compilation of several gravity surveys by the U.S. Geological Survey that were based on over 60,000 gravity stations throughout the two states. The coverage over the Penokean volcanic belt is excellent and falls within at least five large surveys (Daniels and Snyder, 2002; Daniels et al., 2009). The U.S. Geological Survey reduced the observed gravity values to the Bouguer anomaly using the 1967 gravity formula and reduction density of 2.67 g/cc. The data was then converted to a 1-km grid using minimum curvature techniques (Daniels and Snyder, 2002).
For the present study, the complete Bouguer anomaly was transformed to the isostatic anomaly to correct for low-density topography roots. This was followed by the transformation to total horizontal gradient of the isostatic anomaly. These gradients were used to define the
22 geological domains detailed in this study. Figure 2-2b shows the complete Bouguer gravity anomaly map of the Penokean volcanic belt.
2.3.2. Geophysical Domains
The various data transformation techniques allowed us to identify five domains. The domains are defined by their distinct geophysical signatures and are interpreted to potentially represent different volcanic settings. The domains are highlighted over the RTP magnetic map and the complete Bouguer gravity anomaly map in Figure 2-2.
Domain 1 (Continental Margin Assemblage)
Domain 1 is characterized by long wavelength gravity and magnetic anomaly highs expressed in the upward continued transforms. The anomalies trend east-west and occur mostly north of the Niagara fault zone. The rocks consist of Archean granites and gneisses that are overlain by Paleoproterozoic continental margin rocks that include banded iron formations, mafic to felsic volcanic rocks, and sedimentary rocks.
Domain 2 (Northern Pembine domain)
Domain 2 is characterized by long wavelength magnetic anomaly highs. These anomalies trend both northeast and east-west. The gravity data show long wavelength highs and lows that are subparallel to the long wavelength magnetic anomalies. The rocks in this domain consist of Paleoproterozoic aged mafic to felsic volcanic rocks, volcanic breccias and ultramafic to felsic intrusive rocks.
23
Figure 2-2: Geophysical maps of the Penokean volcanic belt. a. Magnetic RTP map. b. Gravity (complete Bouguer) anomaly map. Domain inlay: 1. Continental margin assemblage; 2. Northern Pembine domain; 3. Southern Pembine domain; 4. Wausau domain; 5. Marshfield terrane.
24
Figure 2-3: Geological map of the Penokean volcanic belt. The map was modified from Nicholson et al. (2007).
25 Domain 3 (Southern Pembine domain)
Domain 3 is characterized by a long wavelength magnetic anomaly low. The rocks consist of Paleoproterozoic aged ultramafic to felsic volcanic flows and volcaniclastic rocks, syn-volcanic mafic and felsic intrusive rocks, and post volcanic felsic intrusions.
Domain 4 (Wausau domain)
Domain 4 is characterized by a continuous, long wavelength magnetic anomaly high that trends northeast. Several ovoid moderate wavelength magnetic anomaly highs are also evident. The rocks consist of Paleoproterozoic aged mafic to felsic volcanic rocks, syn-volcanic mafic and felsic intrusive rocks, and post-volcanic syenitic and granitic intrusions.
Domain 5 (Marshfield terrane)
Domain 5 exhibits a complex geophysical signature. Magnetic anomaly highs and lows occur throughout. Z-shaped anomalies that have been interpreted to represent Archean rocks in the Marshfield terrane (Holm et al., 2005) are evident within this domain and help to define its extent. The gravity data are complicated by the regional extent of the apparently low density Wolf River batholith. Nevertheless, moderate wavelength gravity anomaly highs are imaged throughout the domain. The rocks consist of Archean gneisses, migmatites, and amphibolites overlain by Paleoproterozoic aged mafic to felsic volcanic flows and volcaniclastic rocks.
2.3.3. Distribution of Volcanic and Plutonic Rocks
Using the different data transformation techniques, the distribution of volcanic and plutonic rocks was mapped utilizing aeromagnetic and gravity data (see Appendix A). The distribution of the volcanic and plutonic rocks as well as the overall geophysical
26 patterns were used to differentiate geophysically and geologically distinct domains within the Penokean volcanic belt.
Volcanic rocks
The distribution of the volcanic rocks of the Penokean volcanic belt has been established through geologic mapping, drilling, and geophysical surveys. Volcanic rocks in the north of the Pembine volcanic complex are dominated by magnetic highs associated with the mafic-ultramafic volcanic and intrusive rocks. Within the southern half of the Pembine volcanic complex, the RTP data show a broad magnetic low with isolated, short wavelength, moderate amplitude highs. This east-west striking area that is a magnetic low has the highest VMS potential, but the lowest proportion of outcrops. Drill hole data from this area show that the volcanic rocks are bimodal in composition.
Plutonic rocks
Felsic intrusive rocks within the Penokean volcanic belt are typically well exposed and mappable. These rocks exhibit lows in both isostatic anomaly and RTP maps (Fig. 2- 4). In contrast, mafic intrusive rocks are typically not exposed, but differ geophysically from the bimodal volcanic rocks and felsic intrusions because of their strong magnetic signature. The northern half of the Pembine volcanic complex is dominated by large mafic and felsic intrusive bodies of varying ages, including the suite of felsic intrusions emplaced 1760 Ma ago. The southern half of the Pembine volcanic complex is dominated by volcanic rocks but contains small mafic and felsic intrusive bodies. Previous U-Pb geochronological research on felsic intrusive rocks has helped to identify the spatial pattern of syn-volcanic and post-volcanic intrusions (Sims et al, 1989). Syn-volcanic intrusive rocks are generally located in the core of the Pembine volcanic complex while post-volcanic intrusive rocks are generally found within the Marshfield terrane and the Wausau volcanic complex. The syn-volcanic intrusive rocks have both positive and negative magnetic anomalies, whereas the post-volcanic intrusive rocks show negative magnetic anomalies.
27 Syn-volcanic intrusions, which may have acted as heat sources for VMS formation, were identified and imaged as TMI, RTP, and AS highs. These intrusions are assumed to be magnetite-series plutonic rocks (Ishihara, 1981) because they produce magnetic anomalies within the broad magnetically low bimodal volcanic rocks. A minimum of three flight lines must have crossed an intrusion for it to be recognized confidently as a magnetic anomaly.
2.3.4. Lineaments
The tilt derivative transform applied to the RTP data shows that the belt is dominated by east-northeast trending lineaments. These lineaments have been generally classified into three categories, namely banded iron formations (BIFs), rift-shear lineaments, and Z-shaped lineaments.
The TDR lineaments near or north of the Niagara fault zone are associated with mapped BIFs. These lineaments are interpreted to be the magnetite-rich iron oxide facies within the BIFs. As ultramafic volcanic units south of the Niagara fault zone have a similar magnetic signature, differentiation between the two rock types had to be based on available map data. The analytic signal also maps the magnetite-rich zone within the BIF units. However, the TDR lineaments appear more continuous and therefore are preferred for mapping the extent of the BIF north of the Niagara fault zone.
Rift-shear lineaments are thought to map step offsets in magnetic material such as graben boundaries that juxtapose magnetic and nonmagnetic material. The structures may have been exploited by mafic intrusions that are magnetic. These lineaments are generally less than 2 km wide and more than 5 km long. They tend to trend parallel to mapped faults, foliation, and strong EM conductors attributed to graphitic and sulfide- bearing mudstone (DeMatties, 1994).
28
Figure 2-4: Geology and geophysical maps of a representative area within the Penokean volcanic belt illustrating the geophysical signature of volcanic and intrusive rocks.
29 Discrete, irregular Z-shaped anomalies have been mapped within the Marshfield terrane and the Wolf River batholith. Although not as apparent, additional Z-shaped anomalies occur east of the batholith. Previous workers interpreted the Z-shaped anomalies to be associated with the Archean aged gneissic basement within the Marshfield terrane (Holm et al., 2007). Thus, a similar basement may continue to the east and underlie the Wolf River batholith. At a larger scale, the lineaments are truncated to the south by the Spirit Lake tectonic zone, the distinct terrane boundary between the rocks of the Penokean orogen and the rocks of the adjacent Yavapai orogen (Holm et al., 2007).
The two most prominent structures known within the Penokean volcanic belt are the Niagara fault zone on the north side of the belt and the Eau Pleine shear zone on the south side of the belt. Other major structures include the Monico, Owens, and Spring Lake faults as well as the Athens and Jump River shear zones (Fig. 2-3). These faults trend in an east-northeast direction and are associated with prominent geophysical lineaments that are best highlighted in the TDR transform. The TDR data show similarly trending, but generally shorter, lineaments associated with massive sulfide deposits.
2.3.4. VMS Deposits
The VMS deposits in the Penokean volcanic belt are located along an east-west trending domain of bimodal volcanic rocks showing a magnetic low. As there is very little to no outcrop where these deposits are located, discovery of the deposits much relied on aeromagnetic and electromagnetic data. In some cases, the deposits are spatially associated with magnetic highs and second-order lineaments, which are thought to represent syn-volcanic intrusions and extensional faults, respectively.
2.4. Geochemical Data
The volcanic rocks of the Penokean volcanic belt have been overprinted by regional metamorphism ranging from lower greenschist to amphibolite facies (DeMatties, 1994).
30 Seafloor alteration and metamorphism are known to potentially affect the major element compositions of ancient volcanic rocks due to the mobility of alkali elements, the alkaline earth elements, or silica (Floyd and Winchester, 1975, 1978; Winchester and Floyd, 1977). Because of potential major element mobilization, immobile trace elements were used in the present study to constrain rock classification and magmatic affinity. In addition, it was tested whether volcanic rocks from the different volcanic complexes within the Penokean belt could be confidently distinguished based on their trace element abundances.
2.4.1. Methodology
New geochemical data was collected from 31 least-altered, coherent volcanic rock samples (see Appendix A). The samples were only collected from the Pembine-Wausau domain which encompasses domains 2, 3, and 4 (Fig. 2-2). The rock samples were taken from outcrop and drill core, as appropriate, to provide reasonable coverage across the Pembine-Wausau domain. Initially, outcrop locations were identified using maps of the U.S. Geological Survey showing where structural measurements were collected in the past. These outcrops were described and sampled. Typically, a magnetic susceptibility reading was taken at the same location.
The 2501,000 g samples were individually examined and any amygdales or vein material were removed using a tile saw. The samples were the crushed in a jaw crusher and then pulverized in a ring mill using a mild steel milling head. Sample preparation was conducted at the sample preparation facilities at Colorado School of Mines. Approximately 40 g of pulp from each sample were sent to ACTLABS in Ancaster, Ontario, for major and trace element analysis. Whole rock and trace element data were determined by fusion X-ray fluorescence (XRF) and inductively-coupled plasma-mass spectrometry (ICP-MS), respectively.
In addition to new geochemical sampling, legacy geochemical data available through the U.S. Geological Survey were used for the present study (see Appendix A).
31 The legacy data set was filtered to ensure that only high-quality data were used. Only whole-rock data were used that were obtained by combining XRF and ICP-MS analysis. The loss-on-ignition values were used to filter out samples that showed evidence for hydrothermal alteration (i.e. LOI>5 wt.%).
2.4.2. Geochemistry of Volcanic Rocks
Two types of immobile element discrimination diagrams were used to constrain rock type. Firstly, the geochemical data were plotted in the Zr/Ti vs. Nb/Y diagram (Fig. 2-5). A spread of compositions from sub-alkaline basalt to rhyolite was obtained. The range of Nb/Y values is restricted while the Zr/Ti values show a broad range.
The Ti/Zr vs. SiO2 diagram was also used for rock classification (Fig. 2-5). The value for Ti/Zr is helpful for classification in rocks that have experienced alteration, metamorphism, or weathering which prevents the use of major element geochemistry. The elements tend to diverge in concentration during magmatic differentiation and are generally immobile except in cases of extreme alteration and metamorphism (Pearce and Cann, 1973; Winchester and Floyd, 1977; Berry et al., 1992). The geochemistry samples were plotted on the diagram from Berry et al. (1992) and the fields were adjusted to fit most of the data. This diagram shows a strong negative correlation between the Ti/Zr ratio and the whole-rock SiO2 content. The outliers are likely rocks that have undergone either silicification or a silica removal. Based on Figure 2-5, the approximate boundaries for Ti/Zr values between the major rock types in the Penokean volcanic belt are rhyolite- dacite (Ti/Zr = 15), dacite-andesite (Ti/Zr = 35), and andesite-basalt (Ti/Zr = 100).
Northern Pembine domain
Based on the Ti/Zr values, volcanic rocks from the northern domain of the Pembine volcanic complex range from basalt to rhyodacite in composition. Figure 2-6 shows a plot of the Zr vs. Y values for the mafic volcanic rocks of the northern Pembine domain, establishing that these rocks have a tholeiitic affinity. Basaltic and andesitic volcanic
32
Figure 2-5: Top: Classification diagram from Winchester and Floyd (1977); Bottom: Classification diagram from Berry et al. (1992). Black symbols = northern Pembine domain (domain 2); red symbols = southern Pembine domain (domain 3); green symbols = Wausau domain (domain 4); circles = mafic volcanic rocks; triangles = felsic volcanic rocks; large symbols: 2014 data; small symbols: = legacy geochemistry data from the U.S. Geological Survey.
33
Figure 2-6: Plot of Zr vs. Y from Pearce and Cann (1973) showing the magmatic affinity of least-altered volcanic rocks from the Penokean volcanic belt. Black symbols = northern Pembine domain (domain 2); red symbols = southern Pembine domain (domain 3); green symbols = Wausau domain (domain 4); circles = mafic volcanic rocks; triangles = felsic volcanic rocks.
rocks show a subduction-related geochemical signature with low TiO2 (mostly <1 wt%), low incompatible element concentrations, and low Nb/Y values (Fig. 2-5). The low to moderate values for Ni, Cr, Ti, Zr, Hf, P, Y, and Nb indicate that the magma is derived from a primitive source (Stern et al., 1995).
The rare earth element plots show that the basalt and andesite are enriched in light rare earth elements (LREEs) relative to the heavy rare earth elements (HREEs). They have relatively steep slopes compared to basalt and andesite from the other geophysically defined domains (Fig. 2-7). All of the mafic rocks have parallel REE patterns, but the andesite samples have a higher total REE content than the basalt and basaltic andesite
34 samples. The shape of the REE patterns is indicative of crustal contamination. Harker diagrams show that a cluster of mafic rocks just east of the Pelican River deposit plot at distinctly higher TiO2 and P2O5 concentrations than other mafic rocks of the belt (Fig. 2- 8).
Dacitic and rhyolitic volcanic rocks of the northern Pembine domain have a transitional to calc-alkaline affinity (Fig. 2-6). The normalized REE patterns of these rocks have a similar slope as the mafic rocks with slightly higher total REE values and subtle negative Eu anomalies (Fig. 2-7). A few of the samples show a small relative enrichment of Er, Yb, and Lu.
Southern Pembine domain
The southern domain of the Pembine volcanic complex is very poorly exposed, limiting sampling for whole-rock geochemical analysis. Most of the samples collected are rhyolite, rhyodacite, and dacite, with only few being of basaltic composition. The lack of andesite in the southern Pembine domain seems to suggest that the volcanic rocks are bimodal, but more samples would be needed to strengthen this claim.
The mafic samples show a tholeiitic to transitional affinity based on their Zr vs. Y values (Fig. 2-6). The mafic rocks also show an enrichment of LREEs relative to the HREEs. Thus, the samples have relatively steep negative slopes, two of which are much steeper than the rest. This is indicative of crustal contamination of the mafic magma. The samples with higher total REE contents have a subtle negative Eu anomaly, while samples with lower total REEs have a subtle positive Eu anomaly (Fig. 2-7).
Dacitic and rhyolitic volcanic rocks show scattered values in Figure 2-6, which suggests that they are of tholeiitic, transitional, and calc-alkaline affinities. The more calc-alkaline volcanic rocks cluster in areas at the western and eastern ends of the southern Pembine domain, while the more tholeiitic volcanic rocks are in the center of the domain. The felsic rocks are more enriched in LREEs and have higher total REE
35 contents than the mafic volcanic rocks of this domain, but the slope is roughly parallel. This steep negative slope is indicative of strong magmatic fractionation. One group of samples has a steeper REE pattern with or without a subtle Eu anomaly. The other group of samples shows steeper and more pronounced negative Eu anomalies.
The volcanic rocks with the pronounced negative Eu anomaly are generally correlated with a tholeiitic affinity while samples without the Eu anomaly more typically are of calc-alkaline affinity.
Wausau domain
Based on the Ti/Zr vs. SiO2 classification diagram, volcanic rocks of the Wausau domain of the Wausau volcanic complex range from basalt to rhyolite in composition. As only few outcrops of mafic volcanic outcrops could be identified, this group of rocks is underrepresented in the present study.
The mafic volcanic rocks have very low Nb/Y values that do not plot in Figure 2-5. Figure 2-6 shows that the mafic rocks have a tholeiitic affinity. They are slightly more depleted in the LREEs relative to the HREEs, which is a distinct contrast from the other mafic rocks in this study. Figure 2-7 shows a shallow positive slope for basalt and andesite samples. This indicates that the melt came from a primitive mantle source that was able to ascend rapidly and experience little to no crustal contamination. This is supported by low Ni (1150 ppm), Cr (8300 ppm), TiO2 (0.30.6 wt %), Zr (1030 ppm), and Nb (0.751.0 ppm) values (Stern et al., 1995).
The silicic volcanic rocks of the Wausau domain are rhyolitic to dacitic in composition and have a calc alkaline affinity (Fig. 2-6). One exception is a felsic sample (PVB2014-008) that has a higher Nb/Y value that plotted in the trachyandesite field (Fig. 2-5). The felsic volcanic rocks are enriched in LREEs relative to the HREEs
36
Figure 2-7: Rare earth element plots for least-altered volcanic rocks from the Penokean volcanic belt. a. Basalt and andesite from the northern Pembine domain (domain 2). b. Dacite and rhyolite from the northern Pembine domain (domain 2); c. Basalt and andesite from the southern Pembine domain (domain 3). d. Dacite and rhyolite from the southern Pembine domain (domain 3). e. Basalt and andesite from the Wausau domain (domain 4). f. Dacite and rhyolite from the Wausau domain (domain 4).
37
Figure 2-8: Harker diagrams of P2O5 and TiO2 vs. SiO2. Black symbols = northern Pembine domain (domain 2); red symbols = southern Pembine domain (domain 3); green symbols = Wausau domain (domain 4); circles = mafic volcanic rocks; triangles = felsic volcanic rocks. Large symbols = data from this study; small symbols = legacy geochemistry data from the U.S. Geological Survey. and show a subtle negative Eu anomaly. The steep negative slope in Figure 2-7 indicates strong fractionation of the melt and/or a crustal source. The mafic and felsic trends appear to have an inverse relationship but have similar HREE abundances.
2.5. Discussion and Conclusions
The present study shows that different geologic domains can be identified within the Penokean volcanic belt using aeromagnetic and gravity data. The northern domain of the Pembine volcanic complex is magnetically distinct and punctuated by numerous mafic and felsic intrusions. In contrast, the southern domain of the Pembine volcanic complex is defined by its broad, magnetically low signature and the occurrence of relatively few mafic intrusions. It is also distinct in that it is host to the majority of the VMS deposits. The Wausau domain contrasts the southern Pembine domain in that it has good outcrop exposure, a high magnetic signature, and is intruded by numerous mafic and felsic plutons. Previous age dating revealed that volcanic rocks within the Wausau volcanic complex are significantly younger than volcanic rocks in the Pembine volcanic complex (Sims et al., 1989), implying that volcanism has occurred in a geologically distinct environment.
38 The geochemical study conducted as part of the present study shows that volcanic rocks collected from across the Penokean volcanic belt form a coherent group with a restricted range of Nb/Y values and a broad range of Zr/Ti values (Fig. 2-5). This is similar to those of calc-alkaline lavas from modern intraoceanic and continental margin- arc settings (Ewart, 1979; Gill, 1981). The geochemical signature of the volcanic rocks indicates that volcanism was subduction-related (Stolz, 1995). Most of the volcanic rocks are tholeiites, but there are also volcanic rocks with transitional and calc-alkaline affinities.
The whole-rock major and trace element geochemistry of the least-altered volcanic rocks revealed some subtle differences between the geophysically defined domains. The mafic volcanic rocks of the Wausau volcanic complex have a primitive mantle source and experienced little crustal contamination. This contrasts with the mafic volcanic rocks sampled in the Pembine volcanic complex where there is strong evidence of contamination of mantle-derived melts. However, although the northern and southern domains of the Pembine volcanic complex are geophysically distinct, there are no obvious geochemical differences.
Within the northern and southern domains of the Pembine volcanic complex, there are subgroups of rocks that are geochemically distinct. The present study reports for the first time on the occurrence of a high TiO2 and P2O5 volcanic suite in the northern Pembine domain. These samples appear to cluster in two areas in the center and east of the domain. The southern Pembine domain has a group of rocks with pronounced negative Eu anomalies. These rocks are the most felsic in composition and appear to be of tholeiitic affinity. These felsic rocks occur primarily in the center and eastern part of the domain. Although a few of these volcanic rocks are associated with VMS deposits, none of them have an LOI value over 2% which suggests that the Eu anomaly is not the result of hydrothermal alteration, implying that the negative Eu anomaly can be related to the fractionation of plagioclase during the melt evolution.
39 Figure 2-9 shows a synoptic diagram establishing the relationships between the volcanic stratigraphy, intrusive history, and the regional orogenic events. The figure is centered on the Penokean orogeny, which is distinctly older than the Yavapai orogeny to the south. The rocks of the Pembine-Wausau and Marshfield terranes are similar in age and directly related to the Penokean orogeny, which this study suggests was an island-arc or continental margin setting. The volcanic rocks are associated with syn-volcanic intrusions as imaged using aeromagnetic data. The Wausau volcanic complex is distinctly younger and has its own syn-volcanic intrusions. The post-volcanic 1760 Ma suite of intrusions are similar in age to the Yavapai orogeny.
2.6. Acknowledgments
This study has benefited from discussions with M. Hitzman, J. Franklin, N. Kelly, P. Quigley, T. Quigley, and K. Schulz on the geology of the Penokean volcanic belt. Logistical support in the field was provided by Aquila Resources. We also thank Aquila Resources for giving us access to the drill core from the Bend and Horseshoe deposits and G. Hudak for allowing us to sample drill core from the Lynne deposit. This research was supported by the Society of Economic Geologists. Additional funding was provided by the Barrick Gold Endowed Scholarship Fund at the Colorado School of Mines.
40
Figure 2-9: Synoptic diagram showing the relationships between volcanic stratigraphy, plutonic history, and orogenic events.
41 2.7. References
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42 Reinterpretation of Paleoproterozoic accretionary boundaries of the north-central United States based on a new aeromagnetic-geologic compilation: Precambrian Research, v. 157, p. 7179. Ishihara, S., 1981, The granitoid series and mineralization: Economic Geology 75th Anniversary Volume, p. 459484. Jirsa, M.A., Boerboom, T.J., Chandler, V.W., Mossler, J.H., Runkel, A.C., and Setterholm, D.R., 2011, Geologic map of Minnesota - Bedrock geology: Minnesota Geological Survey, State Map Series, S-21, scale 1:500,000. 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: Annual Meeting of the Institute of Lake Superior th Geology, 49 , Iron Mountain, Michigan, 2003, Field Trip Guidebook, p. 132. Lambe, R.N., and Rowe, R.G., 1987, Volcanic history, mineralization, and alteration of the Crandon massive sulfide deposit, Wisconsin: Economic Geology, v. 82, p. 12041238. Luyendyk, A.P.J., 1997, Processing of airborne magnetic data: AGSO Journal of Australian Geology and Geophysics, v. 17, p. 3138. 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 deposits of northern Wisconsin: A commemorative volume: Cable, Wisconsin, Institute on Lake Superior Geology Proceedings, 42nd Annual Meeting, v. 42, part 2, p. 6794.
Miller, H.G., and Singh, V., 1994, Potential field tilt A new concept for location of potential field sources: Journal of Applied Geophysics, v. 32, p. 213217. Milligan, P.R., and Gunn, P.J., 1997, Enhancement and presentation of airborne geophysical data: AGSO Journal of Australian Geology and Geophysics, v. 17, p. 6375. Nicholson, S.W., Dicken, C.L., Foose, M.P., and Mueller, J.A.L., 2007, Preliminary integrated geologic map databases for the United States: Minnesota, Wisconsin, Michigan, Illinois, and Indiana: U.S. Geological Survey Open-File Report 2004– 1355. Pearce, J.A., and Cann, J.R., 1973, Tectonic setting of basic volcanic rocks determined using trace element analyses: Earth and Planetary Science Letters, v. 19, p. 290300. Quigley, P., 2016, Spectrum of ore deposit types, their alteration and volcanic setting in the Penokean volcanic belt, Great Lakes Region, USA: Unpublished M.Sc. thesis, Golden, Colorado, USA, Colorado School of Mines, 10 p.
43 Roest, W.R., Verhoef, J., and Pilkington, M., 1992, Magnetic interpretation using the 3D analytic signal: Geophysics, v. 57, p. 116125. Schulz, K.J., and Cannon, W.F., 2007, The Penokean orogeny in the Lake Superior region: Precambrian Research, v. 157, p. 425. Schulz, K.J., Nicholson, S.W., and Van Schmus, W.R., 2008, Penokean massive sulfide deposits: Age, geochemistry, and paleotectonic setting [ext. abs.]: Institute of Lake Superior Geology Annual Meeting, 54th, Marquette, Michigan, 2008, Program and Abstracts, p. 7677. Sims, P.K., 1989, Geologic map of Precambrian rocks of Rice Lake 1 x 2 degree quadrangles, northern Wisconsin: U.S. Geological Survey Miscellaneous Investigations Series Map, I-1924, scale 1: 250,000. 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 Science, v. 26, p. 21452158. Sims, P.K., Saltus, R.W., and Anderson, E.D., 2005, Preliminary Precambrian basement structure map of the continental United States – An interpretation of geologic and aeromagnetic data: U.S. Geological Survey Open-File Report 2005-1029, 31 p. 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, p. 117. Stern, R.A., Syme, E.C., Bailes, A.H., and Lucas, S.B., 1995, Paleoproterozoic (1.90– 1.86 Ga) arc volcanism in the Flin Flon belt, Trans-Hudson orogen, Canada: Contributions to Mineralogy and Petrology, v. 119, p. 117141. Stolz, A.J., 1995, Geochemistry of the Mount Windsor Volcanics: Implications for the tectonic setting of Cambro-Ordovician volcanic-hosted massive sulfide mineralization in northeastern Australia: Economic Geology, v. 90, p. 10801097. 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. 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. 325343.
44 CHAPTER 3 ID-TIMS U-PB GEOCHRONOLOGY OF THE PALEOPROTEROZOIC PENOKEAN VOLCANIC BELT, MICHIGAN AND WISCONSIN: TIMING OF VOLCANOGENIC MASSIVE SULFIDE FORMATION AND EXPLORATION IMPLICATIONS
A paper to be submitted to Economic Geology. Ashley Quigley, Thomas Monecke, Nigel Kelly, Jim Crowley, and Patrick Quigley.
3.1. Introduction
The Penokean orogenic belt represents the oldest Paleoproterozoic orogen at the southern margin of Laurentia (Fig. 3-1), stretching from Minnesota east to the Grenville orogen in the Lake Huron region (Sims, 1996) and southwest to the Central Plains orogen (Sims and Peterman, 1986; Hoffman, 1988). The continental margin assembly (external domain) of the Penokean orogen consists of northward-thrusted continental margin- foreland basin successions overlying the Archean Superior craton (Cannon and Gair, 1970; Morey, 1973), while the internal domain of this orogeny south of the Niagara fault encompasses an assemblage of accreted arc terranes (Sims et al., 1989; Sims and Schulz, 1996). Two of these accreted terranes form the Penokean volcanic belt. The belt is exposed in an area that is approximately 300 by 200 km in size, stretching from north- central Wisconsin to the Upper Peninsula of Michigan (Fig. 3-2).
The Penokean volcanic belt represents one of the most important Paleoproterozoic hosts of volcanogenic massive sulfide (VMS) deposits (Lambe and Rowe, 1987; DeMatties, 1990; DeMatties, 1994; Adams, 1996; DeMatties and Rowell, 1996; Erickson and Cȏté, 1996; May and Dinkowitz, 1996). At present, over 13 significant VMS deposits and prospects have been recognized, totaling over 100 million metric tons (Mt) of polymetallic massive sulfide ore (Quigley, 2016). This includes the world-class Crandon deposit discovered in 1975, which is composed of a total resource of about 61 million metric tons of massive sulfides grading approximately 1.1% Cu, 5.6% Zn, 0.5%
45
Figure 3-1: North American continent and the location of the Penokean orogeny (modified from Hoffman, 1989).
Pb, 1.0 g/t Au, and 37 g/t Ag (Erickson and Cȏté, 1996; Lambe and Rowe, 1987). Exploration from 2002 to present resulted in the discovery and subsequent delineation of the Back Forty deposit, which at present is composed of a measured and indicated mineral resource of 15.1 Mt of ore grading 0.33 wt.% Cu, 3.06 wt.% Zn, 0.22 wt.% Pb, 2.03 g/t Au, and 24.48 g/t Ag. The Flambeau deposit, discovered in 1968, was the only VMS deposit in the Penokean volcanic belt to reach commercial production (May and Dinkowitz, 1996). Only the supergene enriched portion of the deposit was mined which accounts for less than 2% of the known mineral reserves by tonnage.
Despite of the overall economic significance, only limited research has been conducted in the Penokean volcanic belt, which has gone some way to constrain the tectonic controls on the formation of the major VMS deposits. This paper presents the results of a study utilizing isotope dilution thermal ionization mass spectrometry (ID-
46 TIMS) to determine the precise chronostratigraphic framework for VMS deposit formation in the Penokean volcanic belt. The results of this geochronological study demonstrate that the major deposits and prospects across the belt likely formed during a short-lived phase of rapid extension that was less than a few million years in duration. Identification of volcanic centers of favorable age within the Penokean volcanic belt has important implications for future exploration.
3.2. Geological Setting
The Paleoproterozoic Penokean volcanic belt forms an east-west trending assemblage of volcanic rocks that is intruded by syn- and post-orogenic intrusive rocks. The boundary between the Penokean volcanic belt and the Archean Superior craton to the north is marked by the steeply dipping Niagara fault zone (Larue and Ueng, 1985).
Traditionally, the Penokean volcanic belt has been divided into different volcanic complexes that likely formed in distinct geologic settings as suggested by previous geochemical and geochronological data (Sims et al., 1989). This includes the Pembine and the Wausau volcanic complexes as well as the smaller Waupee volcanic complex at the eastern end of the belt. The Marshfield volcanic complex, which unconformably overlies the Archean basement of the Marshfield terrane, is inferred to be similar in age to the Pembine volcanic complex.
Pembine volcanic complex
The Pembine volcanic complex forms an east-west trending group of volcanic rocks in the northern part of the Pembine-Wausau terrane (Sims et al., 1989; DeMatties, 1994). It is truncated to the north by the Niagara fault zone and to the south by the Eau Pleine and Athens shear zones (Fig. 3-2). The Pembine volcanic complex can be further divided based on differences in magnetic signature in the northern and southern domains (see Chapter 2).
47
Figure 3-2: Geological map of the Penokean volcanic belt showing the locations of previous published U-Pb dates. Ages with associated errors greater than ±10 Ma for volcanic rocks and ±15 Ma for intrusive rocks not plotted. See Table 3-1 for a complete list of dates and associated errors. The map was modified from Nicholson et al. (2007).
48 The northern domain of the Pembine volcanic complex is dominated by tholeiitic basalt and basaltic andesite flows and volcaniclastic deposits, which are intruded by syn- volcanic mafic to ultramafic intrusions (DeMatties, 1994). The domain hosts some of the main VMS deposits and prospects of the Penokean volcanic belt and includes Eisenbrey, Flambeau (May and Dinkowitz, 1996), Lynne (Adams, 1996), and Pelican River (DeMatties, 1994). All of these VMS deposits and prospects are located near the boundary with the southern domain.
The southern domain of the Pembine volcanic complex is composed of a mafic volcanic succession with several major felsic centers of dacitic to rhyolitic composition (DeMatties, 1994). The felsic volcanic centers of the southern domain of the Pembine volcanic complex host some of the most economic VMS deposits and prospects, such as Bend (DeMatties and Rowell, 1996), Ritchie Creek (DeMatties, 1990), Horseshoe (DeMatties, 1994), and Crandon (Erickson and Cȏté, 1996). The host rocks of the Back Forty deposit may also represent the easternmost extent of the southern domain of the Pembine volcanic complex.
Age controls on volcanism in the Pembine volcanic complex are scarce. A rhyolite porphyry from the northern Pembine volcanic complex has a U-Pb zircon date of 1869 ± 6 Ma (Sims et al., 1989). Schulz et al. (2008) report a U-Pb zircon date of 1874 ± 4 Ma for the host rock succession of the Back Forty deposit. Intrusions in the Pembine volcanic complex range from 1890 ± 15 Ma (Banks and Cain, 1969) to 1739 ± 8 Ma in age (Sims et al., 1989).
Wausau volcanic complex
The Wausau volcanic complex is located in the central part of the Penokean volcanic belt, south of the Spring Lake fault and southeast of the Athens shear zone. The Eau Pleine shear zone delimits the Wausau volcanic complex to the south. To the east, the volcanic rocks are truncated by the 1484 ± 2 to 1468 ± 4 Ma, (Dewane and Van Schmus, 2007) anorogenic Wolf River batholith (Fig. 3-2).
49 The rocks of the Wausau volcanic complex unconformably overly intensely deformed late Archean to early Proterozoic amphibolite facies quartzofeldspathic gneisses, amphibolites, and locally occurring migmatitic gneisses (LaBerge and Myers, 1983, 1984). The volcanic rocks have been overprinted at greenschist facies conditions and generally show well-preserved volcanic textures. They vary from mafic to felsic composition and have a calc-alkaline affinity. Mafic volcanic rocks occur as pillowed or massive flows, although mafic volcaniclastic deposits also occur. Felsic volcaniclastic deposits are common in addition to massive and flow-banded coherent rhyolite (LaBerge and Myers, 1983, 1984). Age constraints include a U-Pb zircon age of a rhyolite micro- porphyry from Wausau (1836 ± 11: Van Schmus, 1980; Sims et al., 1989) and dacitic dike from Pittsville (1837 ± 9 Ma: Sims et al., 1989). Although there are no massive sulfide deposits reported from the Wausau volcanic complex, a low-grade disseminated Au-Cu deposit (Reef) is hosted by the volcanic rocks of the complex (DeMatties, 1994).
The volcanic rocks of the Wausau volcanic complex have been intruded extensively by quartz monzonite and syenite. With a quartz monzonite sampled near Kalinke and a quartz diorite near Mosinee having dates between 1847 ± 16 and 1836 ± 13 Ma (Van Schmus, 1980).
Waupee volcanic complex
The Waupee volcanic complex of Oconto County in Wisconsin represents a small outcrop area of volcanic rocks east of the Wolf River batholith. In the northwestern portion of the outcrop area, the Waupee volcanic complex is overprinted by the northeast- trending Mountain shear zone, which forms a 2 km wide and over 12 km long zone of intense ductile deformation (Sims et al., 1990). Within this shear zone, volcanic rocks having a felsic precursor composition are represented by interlayered biotite and amphibole schists, while rocks of mafic precursor composition occur as amphibole schists (Sims, 1989; Day et al., 1991). The metamorphosed rocks in the shear zone are intruded by the 1812.7 ± 3.6 Ma (U-Pb zircon age; Sims et al., 1990) Hines quartz
50 diorite, which provides a minimum age for the volcanic rocks and the overprinting ductile deformation.
Outside of the Mountain shear zone, the volcanic rocks of the Waupee volcanic complex are texturally well-preserved. Felsic rocks predominately form massive to thinly bedded volcaniclastic deposits, while mafic rocks occur as massive and pillow lava flows with locally occurring intercalated mafic volcaniclastic deposits (Sims, 1989). The felsic volcanic rocks of the Waupee volcanic complex are host to the Mountain VMS prospect. Regional metamorphism of the volcanic rocks was at upper greenschist facies conditions (Sims et al., 1990).
At present, the tectonic relationships between the Waupee volcanic rocks and the main Penokean volcanic belt are unresolved. The Waupee volcanic complex could represent either an extension of the Wausau volcanic complex, an extension of the southern domain of the Pembine volcanic complex, or form part of a poorly defined volcanic complex of different age.
Marshfield volcanic complex
The Marshfield terrane, located to the south of the up to 7 km wide Eau Pleine shear zone (Fig. 3-2), is composed of Archean (Van Schmus and Anderson, 1977; Myers et al., 1980; Sims et al., 1989; Van Wyck and Johnson, 1997) gneisses and small erosional remnants of stratigraphically overlying Early Paleoproterozoic volcanic rocks of the Marshfield volcanic complex (Sims et al., 1989).
The volcanic rocks of the Marshfield volcanic complex range from mafic to felsic in composition and host subordinate porphyritic intrusions (Sims et al., 1989). Volcanism is inferred to have occurred 18701860 Ma ago. A felsic volcanic rock in the Eau Claire River valley yielded a U-Pb zircon age of 1858 ± 5 Ma and a possibly syn-volcanic tonalite in the Eau Pleine shear zone yielded an age of 1860 ± 7 Ma (Sims et al., 1989). The basement rocks and the overlying volcanic rocks have been intensively deformed
51 18601835 Ma ago (Sims et al., 1989). Regional metamorphism occurred at upper greenschist facies conditions (Sims et al., 1989).
The supracrustal rocks of the Marshfield terrane have been intruded by syn- and post-orogenic tonalite to granite bodies (Sims et al., 1989). A regional Nd isotope study revealed that intrusions immediately to the north of the Eau Pleine shear zone show little contamination by old crust, whereas intrusions within the Marshfield volcanic complex have a significant isotopic contribution from the Archean basement rocks. The pronounced change in the nature of the basement across the Eau Pleine shear zone may suggest that this shear zone represents a suture (Van Wyck and Johnson, 1997)
3.3. Previous Geochronology
Over the past decades, several U-Pb geochronological studies have been conducted in the Penokean volcanic belt, providing first-order information on the timing of the Penokean orogeny and the assembly of Laurentia (Banks and Cain, 1969; Van Schmus et al., 1975; Maass and Van Schmus, 1980; Van Schmus, 1980; Sims, 1990a,b; Sims et al., 1989; Holm et al., 2005; Medaris et al., 2007; Schulz et al., 2008). A summary of U-Pb ages from the Penokean volcanic belt obtained prior to the present study is presented in Table 3-1 and the locations of the samples are highlighted in Figure 3-2.
The previous geochronological research showed that Mesoarchean rocks dating between 3100 and 2700 Ma (W.R. Van Schmus, pers. comm., 2014; Sims et al., 1989; Maass and Van Schmus, 1980) as well as Neoarchean/Paleoproterozoic rocks dating between 2600 and 2400 Ma (Maass and Van Schmus, 1980; Sims et al., 1989; Sims, 1990a) occur only to the north of the Niagara fault zone and south of the Eau Pleine shear zone (Fig. 3-2). Volcanism in the Penokean volcanic belt took place 18901860 Ma ago (Sims et al., 1989), with most felsic volcanic centers hosting VMS deposits probably being formed around 18801870 Ma (Schulz et al., 2008). Several workers have attempted to establish the age of VMS deposits in the Penokean volcanic belt through the analyses of zircons contained in felsic volcanic rocks. However, issues were encountered
52 as many volcanic samples contain few or no zircon grains (W.R. Van Schmus, pers. comm. to A. Afifi, 1979; Z.E. Peterman, pers. comm. to A. Afifi, 1979; Afifi et al., 1984).
As dating of the volcanic rocks has proven difficult, much of the geochronological research has focused on unraveling the age of syn- to post-orogenic magmatism in the Penokean volcanic belt (Table 3-1). Syn-orogenic intrusions mostly date between about 1860 and 1830 Ma, while post-orogenic intrusions throughout the belt have ages of approximately 1770 to 1740 Ma. The multiple phases of the anorogenic Wolf River batholith, formed from 1484 ± 2 to 1468 ± 4 Ma (Dewane and Van Schmus, 2007), represents the youngest dated intrusions in the Penokean volcanic belt.
3.4. Materials and Methods
The present study focused on U-Pb zircon dating by ID-TIMS of seven felsic volcanic units sampled from surface exposures and diamond drill core. The samples originated from the host rock successions of significant VMS deposits and prospects within the Penokean volcanic belt (Table 3-2). Geochemical analyses of the samples in the field using a hand-held Thermo Scientific Niton X-ray fluorescence spectrometer suggested that the volcanic rocks generally have low Zr contents ranging from approximately 50 to 150 ppm. As zircon recovery was predicted to be difficult, large samples (13–15 kg) were collected at each site.
Zircon separation was conducted at the mineral separation facility at the Boise State University Isotope Geology Laboratory. Heavy mineral recovery was performed using standard crushing and grinding of the volcanic rocks, followed by separation on a Wilfley table and by heavy liquids. Sorting of the separates was performed by magnetic separation using a Frantz isodynamic separator and hand-picking under a binocular microscope. A relatively small number of zircon grains were recovered in most samples
(Table 3-2). The zircon grains recovered were typically below 100 m in size and
53 Table 3-1: Compilation of previously reported U-Pb ages of intrusive and volcanic rocks of the Penokean volanic belt. All coordinates are reported in NAD 83 Zone 15.
No. Age (Ma) Rock type Sample Easting Northing Reference
1 2984±51 Granitic gneiss VS79-87 702855 5088037 1 2 2870±13 Granitic gneiss VS86-3 749935 4916716 8 3 2815±20 Gneiss 14844 680909 4919734 9 4 2780±26 Migmatitic gneiss VS75-6 767161 4931565 8 5 2752±16 Gneiss VS77-212 900676 5089226 1 6 2741±36 Granitic gneiss VS75-10 722851 4923443 8 7 2535±10 Gneiss 14845 690830 4934686 9 8 2530±7 Augen gneiss VS76-25 691099 4933409 8 9 2522±22 Gneiss S-1-86 636220 4989630 8 10 2516±5 Gneiss VS80-2 684191 4922959 8 11 2454±22 Granitic gneiss VS81-65 765166 4951067 6 12 1905±30 Rhyolite 1885 885176 5057296 8 13 1890±15 Granite 1882 880624 5075136 11 14 1889±7 Tonalite VS73-40 657226 5045509 8 15 1880±15 Gneiss 1884 873766 5070136 11 16 1874±4 Rhyolite 108408/108415 902751 5044931 2 17 1871±5 Granite VS76-18 687758 4938567 8 18 1870±7 Granite VS82-103 758105 5028745 8 19 1869±6 Rhyolite porphyry VS79-57 804956 5049916 8 20 1862±5 Gneiss no name 874135 5069736 8 21 1860±7 Tonalite VS84-1 760728 4954780 8 22 1858±5 Metarhyolite VS74-10 659150 4954918 8 23 1857±19 Granite W402Z 736609 4979559 8 24 1852±15 Gneiss VS79-61 812391 5064197 8 25 1852±13 Granophyre VS79-51 761130 4952682 8 26 1852±6 Biotite Schist RL-2 706631 5094261 8 27 1847±16 Tonalite VS73-18 763061 4965837 8 28 1846±11 Granodiorite VS79-84 776941 5026013 8 29 1842±10 Tonalite VS73-25 635734 4963273 10 30 1841±10 Tonalite VS73-21 755548 4924746 8 31 1838±6 Tonalite W314 752558 5022661 8 32 1837±9 Dacite VS74- 2 722851 4923443 8 33 1837±10 Tonalite VS86-4 725014 4969982 8 34 1836±15 Granite 1880 887697 5029444 11, 7 35 1836±11 Rhyolite VS73-17 767268 4986317 8 36 1836±13 Quartz Monzonite VS73-16 784366 4992621 8 37 1835±10 Gneiss VS78-6 634946 4964368 8 38 1835±6 Granite VS73-22 671493 4907437 8 39 1835±6 Granite no name 893246 5075211 8 40 1833±5 Tonalite ATH-12 730293 4991328 1 41 1833±4 Granite VS74- 68 722279 4934346 8 42 1831±7 Tonalite VS75-7 767161 4931565 6 43 1831±7 Tonalite VS80-1 689115 4959631 8 44 1812.7±3.6 Quartz Diorite 147-85,150-85 856858 5012288 8 45 1773±8 Granite VS77-251 693677 5063722 5 46 1768±11 Granite VS79- 85 691993 5060223 5 47 1765±6 Granite VS73-37 638411 5068918 5
54 Table 3-1 (continued).
No. Age (Ma) Rock type Sample Easting Northing Reference
48 1763±7 Granite porphyry VS81-61 763329 4894137 4 49 1760±9 Tonalite VS75-8 772615 4932571 10 50 1752±8 Quartz Monzonite VS73-8 891783 5053762 5 51 1749±8 Gneissic granite VS79-81 825319 5087034 1 52 1739±8 Granite VS73-11 799669 5050188 8 53 1477.3±3.2 Granite TD02-01 812865 4925917 3 54 1475.1±2.2 Wiborgite VS70-91 813129 4962958 3 55 1468.1±5 Granite TD00-03 855168 5012009 3 References: 1. W.R. Van Schmus, pers. comm. (2014); 2. Schulz et al. (2008); 3. Dewane and Van Schmus (2007); 4. Medaris et al. (2007); 5. Holm et al. (2005); 6. Sims (1990a); 7. Sims (1990b); 8. Sims et al. (1989); 9. Maass and Van Schmus (1980); 10. Van Schmus (1980); and 11. Banks and Cain (1969). commonly cracked. Most of the zircon grains were too small and sparse to be successfully put into epoxy mounts and polished so very few cathodoluminescence images were collected. The grains that were large enough for imaging showed faint oscillatory zoning. The samples were dated at the Boise State University Isotope Geology Laboratory by chemical abrasion isotope dilution thermal ionization mass spectrometry (CA-TIMS) using a modified version of the method of Mattinson (2005).
Zircon was placed in a muffle furnace at 900°C for 60 hours in quartz beakers. Single grains were then transferred to 3 ml Teflon PFA beakers and loaded into 300 μl Teflon PFA microcapsules. Fifteen microcapsules were placed in a large-capacity Parr vessel and the grains partially dissolved in 120 μl of 29 M HF for 12 hours at 140° to 180°C, lower than the normal 190°C because the grains would have dissolved entirely at that temperature due to substantial metamictization. The contents of the microcapsules were returned to 3 ml Teflon PFA beakers, HF removed, and the residual grains immersed in 3.5 M HNO3, ultrasonically cleaned for an hour, and fluxed on a hotplate at
80°C for an hour. The HNO3 was removed and grains were rinsed twice in ultrapure H2O before being reloaded into the 300 μl Teflon PFA microcapsules (rinsed and fluxed in 6 M HCl during sonication and washing of the grains) and spiked with the Boise State University mixed 233U-235U-205Pb tracer solution. Zircon was dissolved in Parr vessels in
120 μl of 29 M HF with a trace of 3.5 M HNO3 at 220°C for 48 hours, dried to fluorides,
55 and re-dissolved in 6 M HCl at 180°C overnight. U and Pb were separated from the zircon matrix using an HCl-based anion-exchange chromatographic procedure (Krogh,
1973), eluted together and dried with 2 µl of 0.05 N H3PO4.
Pb and U were loaded on a single outgassed Re filament in 5 µl of a silica- gel/phosphoric acid mixture (Gerstenberger and Haase, 1997), and U and Pb isotopic measurements made on a GV Isoprobe-T multicollector thermal ionization mass spectrometer equipped with an ion-counting Daly detector. Pb isotopes were measured by peak-jumping all isotopes on the Daly detector for 100 to 160 cycles, and corrected for 0.16 ± 0.03%/a.m.u (1 sigma error) mass fractionation. Transitory isobaric interferences due to high-molecular weight organics, particularly on 204Pb and 207Pb, disappeared within approximately 30 cycles, while ionization efficiency averaged 104 cps/pg of each Pb isotope. Linearity (to ≥1.4 x 106 cps) and the associated dead time correction of the Daly detector were monitored by repeated analyses of NBS982, and have been constant + 12 since installation. Uranium was analyzed as UO2 ions in static Faraday mode on 10 ohm resistors for 300 cycles, and corrected for isobaric interference of 233U18O16O on 235U16O16O with an 18O/16O of 0.00206. Ionization efficiency averaged 20 mV/ng of each U isotope. U mass fractionation was corrected using the known 233U/235U ratio of the Boise State University tracer solution.
U-Pb dates and uncertainties were calculated using the algorithms of Schmitz and Schoene (2007), 235U/205Pb of 77.93 and 233U/235U of 1.007066 for the Boise State University tracer solution, and U decay constants recommended by Jaffey et al. (1971). 206Pb/238U ratios and dates were corrected for initial 230Th disequilibrium using a Th/U[magma] = 3.0 ± 0.3 using the algorithms of Crowley et al. (2007). All common Pb in analyses was attributed to laboratory blank and subtracted based on the measured laboratory Pb isotopic composition and associated uncertainty. U blanks are difficult to precisely measure, but are estimated at 0.075 pg.
56 Four to nine grains were analyzed from each of six samples. Concordia plots are given in Figure 3-3. Ranked 207Pb/206Pb date plots are given in Figure 3-4. Data are given in Table 3-2. Isoplot 3.0 (Ludwig, 2003) is used to calculate weighted mean 207Pb/206Pb dates from four to seven equivalent dates per sample. The weighted mean dates are used to interpret igneous crystallization ages. The 207Pb/206Pb dates are used rather than 206Pb/238U dates due to some 206Pb/238U dates being slightly younger, presumably due to Pb loss. In all samples except Back Forty, the amount of discordance in analyses used in weighted mean calculations is minimal (-0.25 to 0.12%). In Back Forty, there is substantial discordance (-0.06 to 4.06%). Five dates are older than the weighted mean dates and are interpreted as being from grains with inherited components. Two dates are younger than the weighted mean dates and are interpreted as being from grains that suffered ancient Pb loss. Errors on the weighted mean dates are given at 2σ and are the internal errors based on analytical uncertainties only, including counting statistics, subtraction of tracer solution, and blank and initial common Pb subtraction. These errors are ± 0.5-0.7 Ma. Including the U decay constant uncertainties propagated in quadrature increases the errors to ± 2.4 Ma.
3.5. Results
The results of U-Pb isotope analysis are presented in Table 3-2. Weighted mean 207Pb/206Pb ages were calculated using Isoplot 3.0 (Ludwig, 2003), and are presented at the 2σ confidence interval. Consideration of blank and common lead, tracer solution, and counting statistics are incorporated into calculation of the internal error of weighted average ages and are based on these analytical uncertainties and decay constant errors.
Sample PVB2014-001, Back Forty VMS deposit (NAD 83 Zone 15, 904437 m. E., 5046418 m. N.): The sample was taken from a surface outcrop of a quartz-phyric rhyolite that forms part of the host rock succession of the Back Forty deposit. The main ore lenses of the deposit occur stratigraphically below and above this rhyolite unit. Eight individual zircon grains were analyzed. Analyses of two of the grains (z6 and z8) are concordant and have 207Pb/206Pb ages of 1833.59 ± 1.36 and 1831.57 ± 1.77 Ma, respectively (Fig. 3-
57 3a, Table 3-2). Analyses of five grains (z2, z3, z5, z7, and z9) are discordant. A regression through these data give an upper intercept of 1832.58 ± 0.95 Ma and a lower intercept close to zero, suggesting they were affected by modern Pb-loss. By polling all dates, including the five discordant analyses, results in a weighted mean 207Pb/206Pb age of 1832.98 ± 0.52 Ma (Fig. 3-4). The zircon grain that was not included in the calculation (z1) had a 207Pb/206Pb age of 2725.56 ± 1.64 Ma which and interpreted to be an inherited grain.
Sample PVB2014-002, Bend VMS deposit (NAD 83 Zone 15, 688371 m. E., 5018605 m. N.): The sample is a sericite-altered, quartz-phyric, coherent rhyolite occurring in the hanging wall of the Bend deposit, approximately 150 m above the massive sulfide lens. This sample was taken from diamond drill core B92-14 at a down- hole interval of 288.7 to 294.2 m (947–965 ft). About 0.5% very fine-grained disseminated pyrite is present in the sample, which complicated mineral separation. Pyrrhotite is present within rare dark colored veinlets. Five zircon grains were analyzed. All analyses are concordant and overlapping within error (Fig. 3-3b, Table 3-2). The weighted mean of all 207Pb/206Pb dates is 1872 ± 0.61 Ma (Fig. 3-4). This is interpreted to be the crystallization age of the sampled rhyolite.
Sample PVB2014-003, Horseshoe VMS deposit (NAD 83 Zone 15, 764061 m. E., 5037562 m. N.): The sample from this deposit originated from a coherent rhyolite unit proximal to the ore horizon of the Horseshoe deposit. The sample was collected from exploration drill core HS94-4 at a down-hole interval of 208.2 to 213.4 m (683–700 ft). The rhyolite has been strongly silicified and shows evidence of minor to moderate sericite alteration and patchy pink alteration. Disseminated pyrite is present in trace amounts. Four individual zircon grains were analyzed. All analyses are concordant and overlap within error (Fig. 3-3c, Table 3-2). The weighted mean of all four 207Pb/206Pb dates is 1874.52 ± 0.66 Ma (Fig. 3-4), which is interpreted to be the crystallization age of the rhyolite.
58 Sample PVB2014-004 Lynne VMS deposit (NAD 83 Zone 15, 736108 m. E., 5063657 m. N.): The Lynne sample, taken from diamond drill core Lyn90-16 at a down- hole interval of 81.7 to 87.8 m (268–288 ft), is a quartz-rich, feldspar porphyry that intruded the host rock succession of the Lynne deposit. This unit was chosen because no coherent rhyolite could be identified in the drill core available. However, the porphyry unit is traceable across several drill holes within a section and appears to be nearly conformable to the volcanic stratigraphy, implying that this unit was emplaced as a volcanic sill. Seven individual zircon grains were analyzed. Analyses of five grains (z1, z2, z3, z5, and z6) are concordant and are overlapping within error (Fig. 3-3d). The weighted mean of the five 207Pb/206Pb dates is 1874.4 ± 4.2 Ma (MSWD = 4.2). The zircon grain, z3, has a slightly younger age than the z1, z2, z5, and z6. When the zircon grain (z3) is left out, the weighted mean of the remaining four 207Pb/206Pb dates is 1874.99 ± 0.68 Ma (Fig. 3-4), which is interpreted to be the crystallization age of the syn- volcanic porphyry unit. The two zircon grains that were not included in the calculations were significantly older than main zircon population. The grain z4 had a 207Pb/206Pb age of 1899.24 ± 1.23 Ma and the grain z7 had a 207Pb/206Pb age of 2697.40 ± 0.97 Ma. Both grains are interpreted to be inherited.
Sample PVB2014-005, Mountain VMS prospect (NAD 83 Zone 15, 860272 m. E., 5012130 m. N.): The sample was collected from a volcanic unit exposed about 10 meters away from oxidized pyrrhotite outcrops of the Mountain VMS prospect. Zircon grains from the sample were comparably large and therefore, cathodoluminescence and backscattered electron images were taken (Fig. 3-5). Judging from the grain shapes and oscillatory zoning patterns observed, the zircon grains interpreted to be of igneous origin. Eight individual zircon grains were analyzed, yielding a concordant spread in ages from about 1849 to 1838 Ma (Fig. 3-3e). The analyses cluster into discrete groups. Two analyses (z6 and z8) overlap within uncertainty and yield a weighted mean 207Pb/206Pb date of 1848 ± 1.0 Ma (Table 3-2), while a second group of five analyses (z1, z2, z3, z7, and z9) overlaps within uncertainty and give a weighted mean 207Pb/206Pb date of 1843.6 ± 0.6 Ma which is interpreted to most closely represent the crystallization age of the
59 Table 3-2: U-Pb isotopic data of zircon grains analyzed by ID-TIMS.
Compositional Parameters Radiogenic Isotope Ratios
206Pb* Th mol % Pb* Pb 206Pb 208Pb 207Pb % 207Pb x10-13 c U 206Pb* Pb (pg) 204Pb 206Pb 206Pb err 235U mol c (a) (b) (c) (c) (c) (c) (d) (e) (e) (f) (e)
PVB2014-001 (Back Forty) z1 2.303 0.4459 99.09% 50 0.34 1990 0.644 0.188093 0.100 12.44926 z2 0.281 5.5235 99.88% 252 0.54 15279 0.082 0.112078 0.064 5.033462 z3 0.232 1.9003 99.15% 33 1.42 1927 0.067 0.112069 0.085 5.041965 z5 0.235 1.1918 99.74% 115 0.25 7066 0.069 0.112089 0.071 4.850472 z6 0.145 1.0404 99.79% 138 0.18 8725 0.042 0.112091 0.075 5.074391 z7 0.302 0.7956 99.75% 119 0.16 7243 0.088 0.111958 0.079 5.035193 z8 0.220 0.5225 99.69% 95 0.13 5897 0.064 0.111966 0.098 5.076660 z9 0.345 0.6547 99.69% 96 0.17 5766 0.101 0.112073 0.088 4.847025 PVB2014-002 (Bend) z2 0.432 0.3492 96.97% 10 0.94 556 0.125 0.114503 0.226 5.335062 z4 0.497 1.6895 99.67% 94 0.47 5447 0.144 0.114576 0.073 5.327193 z6 0.468 1.3197 99.83% 185 0.18 10740 0.136 0.114518 0.072 5.321311 z7 0.524 1.0773 99.78% 141 0.20 8100 0.152 0.114584 0.078 5.327695 z9 0.511 0.6523 99.44% 56 0.30 3230 0.148 0.114566 0.093 5.319602 PVB2014-003 (Horseshoe) z1 0.328 3.8927 99.91% 333 0.29 19994 0.095 0.114601 0.065 5.333839 z3 0.327 2.5027 99.84% 190 0.33 11437 0.095 0.114689 0.068 5.336991 z4 0.357 1.2875 99.69% 98 0.33 5831 0.103 0.114674 0.083 5.334535 z6 0.445 0.9413 99.69% 101 0.24 5893 0.129 0.114691 0.082 5.342612 PVB2014-004 (Lynne) z1 0.386 2.2642 99.81% 160 0.36 9498 0.112 0.114631 0.070 5.333782 z2 0.342 1.1049 99.70% 101 0.27 6047 0.099 0.114702 0.080 5.337770 z3 0.381 2.284 99.81% 158 0.36 9406 0.11 0.114538 0.065 5.324362 z4 0.474 2.5037 99.82% 172 0.38 9975 0.137 0.116244 0.069 5.384338 z5 0.353 3.5722 99.87% 228 0.39 13638 0.102 0.114741 0.063 5.335836 z6 0.334 0.8489 99.29% 42 0.50 2558 0.097 0.114650 0.106 5.327523 z7 0.487 2.4250 99.88% 284 0.23 15617 0.136 0.184909 0.059 12.61121 PVB2014-005 (Mountain) z1 0.243 4.4295 99.89% 266 0.41 16327 0.071 0.112339 0.066 5.107962 z2 0.248 3.7607 99.80% 144 0.64 8725 0.072 0.112666 0.067 5.137916 z3 0.223 1.6018 99.61% 75 0.52 4633 0.065 0.112805 0.079 5.150689 z4 0.219 2.1815 99.63% 77 0.69 4672 0.064 0.112715 0.076 5.143086 z6 0.408 1.7914 99.88% 248 0.18 14654 0.118 0.113012 0.078 5.168296 z7 0.418 2.0239 99.86% 218 0.24 12820 0.121 0.112706 0.068 5.140021 z8 0.362 2.0796 99.89% 266 0.20 15865 0.105 0.113053 0.080 5.172472 z9 0.348 0.3604 99.08% 32 0.28 1958 0.101 0.112598 0.159 5.132613 PVB2014-006 (Pelican River) z1 0.419 5.9757 99.84% 191 0.80 10915 0.122 0.114606 0.064 5.327151 z2 0.471 2.6986 99.68% 95 0.73 5386 0.137 0.114686 0.078 5.339604 z3 0.490 2.3009 99.55% 66 0.90 3714 0.142 0.114567 0.079 5.327679 z5 0.457 1.7836 99.72% 111 0.41 6466 0.132 0.114553 0.072 5.332348 z6 0.477 0.8156 99.72% 113 0.19 6539 0.138 0.114559 0.091 5.329281
60 Table 3-2 (continued).
Radiogenic Isotope Ratios Isotopic Dates
206 207 207 206 % err Pb % err corr. Pb ± Pb ± Pb ± 238 206 235 238 U coef. Pb U U (a) (f) (e) (f) (g) (f) (g) (f) (g) (f)
PVB2014-001 (Back Forty) z1 0.253 0.480030 0.216 0.922 2725.56 1.64 2638.9 2.38 2527.40 4.52 z2 0.129 0.325721 0.071 0.961 1833.38 1.17 1824.97 1.10 1817.61 1.12 z3 0.152 0.326298 0.080 0.915 1833.23 1.55 1826.40 1.29 1820.41 1.27 z5 0.143 0.313850 0.084 0.935 1833.55 1.29 1793.70 1.21 1759.62 1.30 z6 0.151 0.328332 0.091 0.926 1833.59 1.36 1831.83 1.28 1830.29 1.46 z7 0.164 0.326182 0.105 0.918 1831.44 1.43 1825.26 1.39 1819.85 1.67 z8 0.196 0.328844 0.134 0.893 1831.57 1.77 1832.21 1.67 1832.78 2.13 z9 0.187 0.313669 0.130 0.907 1833.30 1.59 1793.10 1.57 1758.74 2.00 PVB2014-002 (Bend) z2 0.349 0.337925 0.222 0.776 1872.08 4.07 1874.50 2.99 1876.68 3.62 z4 0.145 0.337211 0.083 0.94 1873.23 1.32 1873.24 1.24 1873.24 1.35 z6 0.143 0.33701 0.083 0.931 1872.32 1.31 1872.29 1.23 1872.27 1.36 z7 0.153 0.33722 0.092 0.914 1873.36 1.41 1873.32 1.30 1873.28 1.49 z9 0.185 0.336762 0.123 0.893 1873.07 1.68 1872.02 1.58 1871.07 2.00 PVB2014-003 (Horseshoe) z1 0.130 0.337559 0.071 0.961 1873.62 1.17 1874.3 1.11 1874.91 1.15 z3 0.135 0.337501 0.074 0.95 1875.00 1.23 1874.81 1.15 1874.63 1.21 z4 0.152 0.337389 0.088 0.897 1874.76 1.49 1874.41 1.30 1874.10 1.42 z6 0.162 0.337848 0.101 0.906 1875.04 1.48 1875.71 1.38 1876.31 1.65 PVB2014-004 (Lynne) z1 0.138 0.337468 0.077 0.947 1874.09 1.26 1874.29 1.18 1874.48 1.25 z2 0.154 0.33751 0.092 0.909 1875.21 1.44 1874.93 1.31 1874.68 1.50 z3 0.133 0.337144 0.074 0.964 1872.63 1.17 1872.78 1.14 1872.92 1.21 z4 0.139 0.335940 0.082 0.936 1899.24 1.23 1882.37 1.19 1867.11 1.33 z5 0.130 0.337274 0.072 0.967 1875.82 1.14 1874.62 1.12 1873.54 1.17 z6 0.184 0.337015 0.111 0.857 1874.39 1.91 1873.29 1.57 1872.29 1.81 z7 0.463 0.494648 0.451 0.992 2697.40 0.97 2651.06 4.36 2590.76 9.62 PVB2014-005 (Mountain) z1 0.130 0.329773 0.071 0.955 1837.60 1.20 1837.43 1.11 1837.28 1.13 z2 0.132 0.330745 0.072 0.955 1842.85 1.21 1842.40 1.12 1841.99 1.15 z3 0.148 0.331159 0.084 0.914 1845.09 1.43 1844.51 1.26 1844.00 1.34 z4 0.141 0.330933 0.076 0.930 1843.65 1.37 1843.25 1.2 1842.90 1.22 z6 0.145 0.331683 0.082 0.908 1848.40 1.42 1847.41 1.23 1846.53 1.32 z7 0.137 0.330764 0.080 0.941 1843.50 1.23 1842.75 1.17 1842.08 1.28 z8 0.143 0.331829 0.079 0.900 1849.07 1.45 1848.10 1.22 1847.24 1.27 z9 0.275 0.330604 0.195 0.825 1841.76 2.87 1841.52 2.34 1841.31 3.13 PVB2014-006 (Pelican River) z1 0.13 0.337121 0.071 0.962 1873.70 1.16 1873.23 1.11 1872.80 1.15 z2 0.144 0.337673 0.079 0.917 1874.96 1.40 1875.23 1.23 1875.46 1.29 z3 0.144 0.33727 0.077 0.920 1873.08 1.43 1873.31 1.23 1873.52 1.26 z5 0.141 0.337607 0.079 0.936 1872.86 1.30 1874.06 1.20 1875.15 1.29 z6 0.179 0.337394 0.119 0.892 1872.96 1.64 1873.57 1.53 1874.12 1.94
Notes: (a) z1, z2, etc. are labels for analyses composed of single zircon grains that were annealed and chemically abraded (Mattinson, 2005)
61 (b) Model Th/U ratio calculated from radiogenic 208Pb/206Pb ratio and 207Pb/235U date 206 (c) Pb* and Pbc are radiogenic and common Pb, respectively; mol% Pb* is with respect to radiogenic and blank Pb (d) Measured ratio corrected for spike and fractionation only; fractionation correction is 0.16 ± 0.03 (1σ) %/amu (atomic mass unit) for single-collector Daly analyses, based on analysis of EARTHTIME 202Pb-205Pb tracer solution (e) Corrected for fractionation, spike, common Pb, and initial disequilibrium in 230Th/238U; common Pb is assigned to procedural blank with composition of 206Pb/204Pb = 18.04 ± 0.61 %, 207Pb/204Pb = 15.54 ± 0.52%, and 208Pb/204Pb = 37.69 ±0.63% (1σ); 206Pb/238U and 207Pb/206Pb ratios corrected for initial disequilibrium in 230Th/238U using Th/U [magma] = 3.0 ± 0.3 (f) Errors are 2σ, propagated using algorithms of Schmitz and Schoene (2007) and Crowley et al. (2007) (g) Calculations based on the decay constants of Jaffey et al. (1971); 206Pb/238U and 207Pb/206Pb dates corrected for initial disequilibrium in 230Th/238U using Th/U [magma] = 3.0 ± 0.3
volcanic unit (Fig. 3-4, Table 3-2). One grain (z4) is younger than these two clusters, having a 207Pb/206Pb date of 1837.6 ± 1.2 Ma (Table 3-2).
Sample PVB2014-006, Pelican River VMS deposit (NAD 83 Zone 15, 790669 m. E., 5054054 m. N.): The sample was collected from an outcrop forming part of the host rock succession of the Pelican River deposit. The sampled rhyolite is aphyric and has been sericite altered. The rock was distinctly spotted due to the occurrence of secondary chlorite. Five zircon grains were analyzed. All analyses are concordant and overlap within error (Fig. 3-3f). The weighted mean of all 207Pb/206Pb dates is 1873.54 ± 0.61 Ma (Fig. 3-4), which is interpreted to be the crystallization age of the sample
Sample PVB2014-007, Flambeau VMS deposit (NAD 83 Zone 15, 646891 m. E., 5033946 m. N.): A sample of a coherent volcanic unit logged as a meta-dacite located in the hanging wall of the massive sulfide lens at Flambeau was collected. The sample originated from diamond drill core 22-136 at a down-hole interval of 24.7 to 31.6 m (81– 103.8 ft). No zircon grains were recovered from this sample.
62
Figure 3-3: Concordia plots of CA-TIMS U-Pb dates from zircon, using Isoplot 3.0 (Ludwig 2003). Error ellipses are at 2σ. Analyses in blue are not included in the weighted mean calculation. Zircons not shown: Back Forty (PVB2014-001) z1; Lynne (PVB2014-004) z4, z7.
63
Figure 3-4: Plots of 207Pb/206Pb dates from zircon analyzed by the CA-TIMS method, using Isoplot 3.0 (Ludwig, 2003). Error bars are at 2σ. Weighted mean dates are shown by the grey box behind the error bars. Analyses in white are not included in the weighted mean calculation. Zircons not shown: Back Forty (PVB2014-001) z1; Lynne (PVB2014-004) z4, z7.
64
Figure 3-5: Representative photomicrographs of zircon grains from the Mountain VMS prospect (sample PVB2014-005). A. Cathodoluminescence image. B. Backscattered electron image.
65 3.6. Discussion and Conclusions
The present study reports the first high-precision ID-TIMS U-Pb zircon dates of felsic volcanic rocks from the Penokean volcanic belt. The new geochronological data provide critical information on the absolute ages of the VMS deposits in the Pembine volcanic complex and represent a basis for the development of a chronostratigraphic framework for the belt. Based on the new U-Pb zircon dates and observed inherited ages, new constraints on the tectonic setting of VMS deposit formation in the Penokean volcanic belt can be derived, which has significant exploration implications.
Comparison to previous constraints on the age of VMS deposits in the belt
A comparison of previously available U-Pb ages with the results from the present study is shown in Figure 3-6. The precision of the new ID-TIMS U-Pb dates is significantly higher than previously published ages providing tighter constraints on the absolute age of formation of the Bend, Horseshoe, Lynne, and Pelican River VMS deposits. The ID-TIMS dates obtained on the felsic volcanic rocks from these deposits cluster tightly between 1874.99 ± 0.68 and 1872.94 ± 0.69 Ma.
The new ID-TIMS U-Pb dates are in agreement with previous age estimates for volcanic and synvolcanic intrusive rocks of the Pembine-Wausau complexes, establishing that the VMS deposits in the Pembine volcanic complex likely formed between 1890– 1860 Ma (Sims et al., 1989). An unpublished U/Pb zircon age for the host rocks of the Lynne deposit suggested that this deposit formed approximately 1870 Ma (R. Thorpe, pers. comm. to T. DeMatties, 1995; Schulz et al., 2008), which is similar to the U-Pb age of 1874.99 ± 0.68 Ma of the present study. Within error, the new ID-TIMS date of 1873.54 ± 0.61 Ma for Pelican River also overlaps with a U/Pb date of 1869 ± 6 Ma of a rhyolite porphyry from a location near Jennings, which is 14.8 km east of Pelican River (Sims et al., 1989).
66 The new U-Pb age constraints for the deposit formation in the Pembine volcanic complex are in agreement with Pb isotope model ages obtained on galena, suggesting that the Crandon, Flambeau, Pelican River, and Tomahawk deposits formed between 1900 and 1800 Ma (Afifi et al., 1984). Schulz et al. (2008) reported a U/Pb date of 1874 ± 4 Ma for the host rocks of the Back Forty deposit.
Figure 3-6: Comparison between previously reported U-Pb ages of volcanic and intrusive rocks of the Penokean volanic belt and new data of the present study obtained on felsic volcanic rocks. The data were filtered to exclude U-Pb zircon ages having large analytical uncertainties. Numbers in the figure are keyed to Table 3-1.
67 Possible resetting of U-Pb zircon ages
The rhyolitic host rocks of the Back Forty deposit yielded in this study a weighted mean 207Pb/206Pb age of 1832.98 ± 0.52 Ma (Fig. 3-4). This date is significantly younger than a previously obtained date for quartz porphyry (sample 108415) located in the footwall of the Back Forty deposit and syn-volcanic feldspar porphyry (sample 108408) that intrudes the deposit (Schulz et al., 2008). All data from both samples are discordant, and when pooled the zircon analyses from the two samples were interpreted by these authors to project along a common Pb-loss curve. Based on the assumption that Pb-loss was modern, a regression through the data anchored at zero resulted in an upper intercept date of 1874 ± 4 Ma (MSWD = 15).
This calculated age is significantly older than that measured for rhyolite host rocks in this study, but is more similar to age estimates for other deposits in the Pembine volcanic complex. The zircon grains investigated from Back Forty in this study were small and dark in transmitted light, possibly indicating that the grains were metamict. However, the individual 206Pb/238U dates are clustered and so are not indicative of partial resetting from an age similar to other deposit host rocks (Figs. 3-3 and 3-4). Therefore, it is likely that the data record a geologically meaningful age. One option is that the obtained date represents the crystallization age of the rhyolitic host. This could mean that the host rocks to the Back Forty VMS deposit are part of a distinctly younger volcanic complex. Previous research has shown that rocks within the Wausau volcanic complex potentially have similarly young ages (Sims et al., 1989; Fig. 3-6), although they are not known to host massive sulfide deposits. Alternatively, it could be possible that the host rocks of the Back Forty deposit are similar in age to the other deposits within the Pembine volcanic complex, but that a major thermal event caused the resetting of the isotopic system at around 1833 Ma.
As part of the present study, the raw data from the two Back Forty samples investigated by Schulz et al. (2008) were re-evaluated with regressions for data from each sample calculated separately. The data from the syn-volcanic feldspar porphyry sample
68 (USGS-408) yielded an upper intercept with concordia of 1873 ± 1.4 Ma (MSWD = 2.1; lower intercept = 0), which is consistent with the original estimate by Schulz et al. (2008). However, the data from the quartz porphyry in the footwall (USGS-415) yield different upper intercept dates, depending on whether or not lead loss is assumed to be modern. When the lower intercept is set to zero, the data produce an upper intercept with concordia at 1875 ± 16 Ma, but with a high MSWD value of 40. This high MSWD represents high scatter from the Pb-loss curve and therefore that Pb-loss in this zircon population may not be modern but the result of an older resetting event. If lower intercept is not anchored in a calculated regression, a statistically more robust upper intercept date is produced (1837 ± 9.4 Ma; MSWD = 1.06; lower intercept = -537 ± 170 Ma). This date closely matches the U-Pb date obtained in the present study for a rhyolite sampled in the same area. The reinterpretation of the data from Schulz et al. (2008) may lend further support to the occurrence of a major thermal event disturbing the zircon isotopic system at Back Forty.
The sample from the Mountain VMS prospect originates from the Waupee volcanic complex located east of the Wolf River batholith. Although no major VMS discovery has been made in the Waupee volcanic complex, the presence of this sulfide showing (Sims, 1989) of pyrrhotite with pyrite and sparse chalcopyrite and sphalerite provides evidence for syn-volcanic hydrothermal activity. The volcanic rock sample collected from the pyrrhotite showing was found to have a U-Pb zircon date of 1843.6 ± 0.6 Ma. However, the U-Pb dates of individual zircons grains are spread over about 9 Myr (~1849 to ~1838 Ma) although they are all concordant (Fig. 3-3). The zircon grains cluster at three age ranges. The two oldest grains have a weighted mean 207Pb/206Pb date of 1848.7 ± 1.0 Ma. The five intermediate aged grains are equivalent with a weighted mean 207Pb/206Pb age of 1843.6 ± 0.6 Ma. The youngest single grain has a 207Pb/206Pb date of 1837.6 ± 1.2 Ma. This finding indicates a number of possibilities for the crystallization age of the sample. However, the most likely possibility is that the cluster of five intermediate dates most closely represents the crystallization age. Figure 3-5 shows CL and BSE images from some of the zircon grains analyzed. The CL image does not show evidence of metamorphic rims on the zircon grains. However, the presence of cores, which may be
69 inherited, could explain the older two grains. The youngest age is possibly the result of a younger thermal event disturbing the isotopic system. Resetting may have occurred as part of the same event affecting the zircons contained in the felsic volcanic rock from the Back Forty deposit. More work is required to characterize the rocks surrounding the Mountain prospect.
Tectonic setting of VMS formation in the Penokean volcanic belt
Previous workers have suggested that the Paleoproterozoic volcanic rocks in the Pembine-Wausau terrane have been largely deposited on juvenile crust, suggesting that volcanism occurred in an intraoceanic arc setting, whereby oceanic arc crust would have occurred on both sides of the arc (Sims et al., 1989; Schulz and Cannon, 2007). Later suturing of the Pembine-Wausau terrane with the Superior craton is interpreted to have taken place along the Niagara fault zone, making this fault zone a terrane boundary (Schulz and Cannon, 2007).
In contrast, Van Wyck and Johnson (1997) argued that the volcanic rocks in the Pembine-Wausau terrane formed through continental margin arc volcanism. Based on Nd and common Pb isotope compositions, they suggested that the Superior craton extends south of the Niagara fault zone, implying that the volcanic rocks of the Penokean volcanic belt were deposited on Archean basement. As a consequence, Van Wyck and Johnson (1997) did not consider the Niagara fault zone to represent a major terrane- bounding structure. They suggested that the Eau Pleine shear zone represents a major suture separating the continental margin arc of the Pembine-Wausau terrane from the
Archean Marshfield crustal block because of a sudden change in Nd values of Penokean plutons on either side of the shear zone.
The present study for the first time reports the occurrence of inherited zircon within volcanic rocks from the Penokean volcanic belt. The zircon population from the Back Forty and Lynne host rock samples each included one Archean zircon grain (2725.56 ± 1.64 Ma and 2697 ± 0.97 Ma, respectively), and are interpreted to have been inherited.
70 The two dates fall within the age range seen in the Superior craton (Fig. 3-7), which suggests that they could be inherited from underlying Superior craton crust. These zircon dates demonstrate that the 1.87 Ga volcanism in the Penokean belt could not have entirely occurred on juvenile crust. These data therefore support models proposing the presence of Archean basement, at least at depth, south of the Niagara fault zone, or at least the presence of Archean basement blocks locally within the Pembine-Wausau terrane.
Figure 3-7: Comparison of the U-Pb ages of inherited zircon grains in the samples from the Back Forty and Lynne deposits with previously published geochronological data for Archean rocks occurring in the Marshfield terrane. Numbers in the figure are keyed to Table 3-1.
71 Comparisons to other volcanic belts of similar age
The Penokean orogen is part of a series of 1.91.8 Ga collisional belts formed during the assembly of the Columbia supercontinent (Zhao et al., 2002; Rogers and Santosh, 2009). Schulz and Cannon (2007) interpret the major contractional phase of the Penokean orogeny to be related to the convergence of the Archean Superior craton and Marshfield micro-continent.
The broadly time-equivalent suturing of the Archean Superior craton with the Archean Hearne-Rae craton formed the Trans-Hudson orogen, the largest and best- exposed Paleoproterozoic orogenic belt in North America (Corrigan et al., 2009). The Penokean volcanic belt as well as other districts of the Trans-Hudson orogen (e.g., Flin Flon: Syme and Bailes, 1993; Stern et al., 1995; DeWolfe et al., 2009; Hanson Lake: Maxeiner et al., 1999; Lynn Lake: Elliott-Meadows and Appleyard, 1991; Sherridon: Zwanzig, 1999; and Snow Lake: Skirrow and Franklin, 1994; Bailes and Galley, 1999) are some of the most important hosts to VMS deposits in North America.
Within the Trans-Hudson orogen, the volcanic rocks of the Flin Flon belt
(1.911.87 Ga) represent the most important hosts to VMS deposits. Studies of volcanic stratigraphy show two major phases of arc-related volcanism. The pre-accretion arc volcanism phase is marked by rapid subduction of oceanic lithosphere and extensive back-arc basin formation. Geochemically, the rocks of the Flin Flon belt show evidence for subduction and recycling of Archean sediments, which suggests the arc system was close to either the Superior craton or an Archean microcontinent. The post-accretion arc volcanism phase was marked by intra-arc extension at about 1886 Ma and mature calc- alkaline arc volcanism of andesitic to rhyolitic composition (Stern et al., 1995).
Broadly contemporaneous formation of Baltica resulted in the Svecofennian orogen (Nironen, 1997), which is at least 1200 km wide and stretches from Sweden through central Finland to beyond Estonia. Major VMS deposits within the Svecofennian orogen are located within the Skellefte (Allen et al., 1996a) and Bergslagen (Allen et al., 1996b)
72 districts. The Skellefte district consists of calc-alkaline basalt-andesite-dacite-rhyolite and tholeiitic basalt-andesite-dacite volcanic associations, komatiitic basalt, and syn- and post-volcanic granitoids. The depositional environment is deep-water subaqueous with evidence of strong extension and subsidence proceeded and/or accompanied by volcanism. The Skellefte belt shows evidence that it is an extensional intra-arc region with a continental or mature-arc crust basement. By contrast, volcanic rocks in Bergslagen consist of 90% rhyolite and show a transition in depositional environment from shallow-moderately deep subaqueous to deep subaqueous. The rocks appear to evolve from intense magmatism, thermal doming, and crustal extension to waning volcanism.
Collectively, the 1.91.8 Ga collisional belts of North-America and Scandinavia formed during the assembly of the Columbia supercontinent account for approximately 8% of the global VMS endowment (cf. data compilation by Franklin et al., 2005), highlighting that this 100 Myr of volcanism in the Paleoproterozoic is of outstanding economic importance.
Exploration implications
Knowledge of the geological make-up of the Penokean volcanic belt is currently limited as comparably few studies have focused on the tectonic, structural, and volcanological setting of its VMS deposits. The characteristics of individual deposits, especially the world-class Crandon deposit, are inadequately understood which is, at least in part, related to the fact that the exploration and delineation drill core is no longer available. The Penokean volcanic belt undoubtedly represents one of the most understudied, but highly endowed VMS belts in the world. The Back Forty deposit, discovered in 2002, which is outcropping at surface forming an easily recognizable gossan, highlights the significant under-explored potential of the belt.
The findings of the present study provide the basis for the development of a chronostratigraphic framework for the Penokean volcanic belt. It is demonstrated that
73 application of ID-TIMS dating of single zircon grains yields highly precise dates supporting more robust interpretation of crystallization ages. Importantly, the high analytical precision allows reliable resolution of the ages of felsic volcanic centers hosting VMS deposits, despite their antiquity (Paleoproterozoic).
The results obtained suggest that felsic volcanic centers hosting VMS deposits across the belt formed as a result of rapid, but short-lived extension between 1874.99 ± 0.68 and 1872.94 ± 0.69 Ma. Other felsic volcanic centers throughout the Penokean belt that are part of this favorable chronostratigraphic interval represent important targets for future exploration.
3.7. Acknowledgments
This study has benefited from discussions with M. Hitzman and T. Quigley on the geology of the Penokean volcanic belt. We also acknowledge help provided by E. Anderson during map compilation and interpretation of aeromagnetic data prior to field work. Logistical support in the field was provided by Aquila Resources. We also thank Aquila Resources for giving us access to the drill core from the Bend and Horseshoe deposits and thank G. Hudak for allowing us to sample drill core from the Lynne deposit. We are indebted to M. Schmitz for additional help during the laboratory research at Boise State University. This research was supported by a student research grant from the Society of Economic Geologists and a graduate student research grant from the Geological Society of America. Additional funding was provided by the Barrick Gold Endowed Scholarship Fund at the Colorado School of Mines.
74 3.8. References
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80 CHAPTER 4 CONCLUSIONS AND RECOMMENDATIONS FOR FUTURE WORK
The present study utilized a combination of geophysical, geochemical, and geochronological data to construct a new, detailed geology map of the Penokean volcanic belt and to understand its geologic and tectonic framework. The results of the study allow the following main conclusions:
1. Aeromagnetic and ground gravity data of the Penokean volcanic belt provide important information allowing extrapolation of the bedrock geology in areas with little to no outcrop. Through the interpretation of transformed total magnetic intensity data and compilation of existing geologic map data, it was possible to construct a new detailed geological map for the belt. Several blind mafic intrusive bodies were recognized. The geophysical data also allowed the identification of major faults throughout the belt and intrusions that may have acted as heat sources for syn-volcanic hydrothermal activity.
2. The Pembine-Wausau terrane of the Penokean volcanic belt can be separated into discrete geological domains. These domains have been defined using aeromagnetic and gravity data, but can also be distinguished based on whole-rock major and trace element geochemical data. The new geologic domains identified are referred to as the northern Pembine volcanic complex, the southern Pembine volcanic complex, and the Wausau volcanic complex.
3. High-precision ID-TIMS U-Pb zircon geochronology was conducted on seven rhyolite samples collected from the host rock successions of some of the major VMS deposits within the Penokean volcanic belt. The results revealed that four deposits, namely Horseshoe, Lynne, Pelican River, and Bend, were all formed at about 1874 Ma. This is evidence of a major period of volcanism occurring in an extensional setting at this time. The geochemical evidence suggests that the
81 deposits formed within a generally bimodal mafic host rock succession formed in a rifted arc or back-arc setting.
4. The Back Forty massive sulfide deposit is located east of the other deposits investigated in the present study. High-precision ID-TIMS U-Pb zircon geochronology of the host rhyolite yielded an age of about 1833 Ma, which is approximately 50 Myr younger than the host rock successions of the other deposits of the Penokean volcanic belt. There are two possible explanations for this apparent age. The first explanation is that this is a crystallization age and the host rocks to the Back Forty are part of a distinctly younger volcanic succession. Alternatively, a thermal event at 1833 Ma may have reset the U-Pb age.
5. The zircon population from host rock successions of the Lynne and Back Forty deposits contains zircons that have Archean ages and were presumably inherited from an Archean basement. The inherited zircon grains from the two deposits yielded ages of approximately 2700 Ma. These data support the idea the Superior craton extends beneath the Pembine-Wausau terrane south of the Niagara fault zone.
Research in the Penokean volcanic belt is still at an early stage and significant future contributions to the understanding of the tectonic setting of volcanism and the local volcanic environment of massive sulfide formation can still be made. The following are recommendations for future work:
1. It is recommended to conduct additional whole-rock major and trace element analyses. As sampling in the Penokean volcanic belt is limited by the availability of outcrops and access, only 35 geochemical samples were collected for the present study. The new dataset is of high quality and supplements existing geochemical data collected over the past decades. The existing chemical information is of variable quality, often limiting their use in geochemical discrimination diagrams that are based on subtle variations in trace element abundances. Expanding the
82 geochemical dataset using modern instrumentation and techniques is critical to future studies in the Penokean volcanic belt as there are significant gaps in the dataset, particularly in the Marshfield terrane. Finally, more geochemical information would help to confirm and strengthen arguments made in this study. For example, there are only a few analyses available for the mafic volcanic rocks of the Wausau volcanic complex. Data available at present suggest that the geochemical signature of these rocks is very different from mafic volcanic rocks elsewhere in the belt.
2. Additional high-precision U-Pb zircon geochronology of felsic volcanic rocks should be conducted in the Penokean volcanic belt. First, it would be important to obtain ID-TIMS U-Pb dates on host rocks to the other major massive sulfide deposits of the belt to further test the hypothesis that all massive sulfide deposits in the belt have formed during a single phase of volcanism in a comparable tectonic and volcanic setting. Particular emphasis should be placed in attempting to date the Flambeau deposit, which was not possible in the present study as no zircons were found in the sampled rhyolite. The Flambeau deposit is also located much farther west than the main group of deposits of the Penokean volcanic belt and could potentially have formed at a different time in a different volcanic setting. Second, more dates at the Back Forty deposit are needed to either confirm or challenge the date obtained in the present study. It is recommended to initially process several volcanic rock samples from the host rock succession of the deposit to find reasonably sized zircon grains for LA-ICP-MS dating. Although this method is less precise than ID-TIMS, LA-ICP-MS analyses would be more cost-effective to confirm the 50 Myr difference between the Back Forty deposit and the other deposits dated in this study.
3. The results of the present study have opened up the possibility that the volcanic rocks on the east end of the Penokean volcanic belt, particularly those hosting the Back Forty deposit, could have formed in a distinct volcanic setting at a different time. Therefore, future work should concentrate on the Back Forty area and the
83 Waupee volcanic complex. Future research in this area should include detailed geologic mapping, regional whole-rock major and trace element geochemistry, and additional high-precision U-Pb zircon geochronology. Regional scale studies covering the Back Forty deposit area are not yet available in the public domain and would represent a major contribution to the understanding of the geologic make-up of the Penokean volcanic belt.
84 APPENDIX A SUPPLEMENTAL ELECTRONIC FILES
The supplemental files give the complete whole rock major and trace element geochemistry results for the samples taken for this study and the legacy geochemistry data from the U.S. Geological Survey that were included for interpretations made in this study. Also included are geophysical maps of the Penokean volcanic belt with various transformations applied that were used to make geological interpretations.
Whole rock major and trace element Files containing locations, descriptions, and geochemistry files complete geochemistry results for the new and legacy data used in this study. All files are in Microsoft Excel 2010 format. All coordinates are in UTM NAD83 Zone 15. Geochemistry2014.xls Whole-rock major and trace element geochemistry for samples taken for this study. GeochemistryLegacy.xls Whole-rock major and trace element geochemistry for legacy data from the U.S. Geological Survey. Magnetic and gravity data Files containing magnetic and gravity maps with various transformation applied. The data that produced these images were used for making the geological interpretations in this study. All images have an overlay of the geologic domains interpreted for this thesis. MagRTP.pdf Reduced-to-Pole (RTP) aeromagnetic map.
MagLP80km.pdf RTP data that has been low-pass filtered to show wavelengths longer than 80 km.
MagUC15km.pdf RTP data that has been upward continued to 15 km highlighting longer wavelength features.
CBA.pdf Complete Bouguer anomaly map.
85 CBAHP80km.pdf Complete Bouguer anomaly that has been high- pass filtered to show wavelength shorter than 80 km which helped to remove the signature of the Wolf River batholith. CBAHP80kmLineaments.pdf Complete Bouguer anomaly that has been high- pass filtered to show wavelength shorter than 80 km with banded iron formation lineaments and Z-shaped lineaments of the Marshfield terrane.
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